Paleogeothermics: evaluation of geothermal conditions in the geological past

Lecture Notes in Earth Sciences Edited by Somdev Bhattacharji, Gerald M. Friedman, Horst J. Neugebauer and Adolf Seilacher

5 Paleogeotherm ics Evaluation of Geothermal Conditions in the Geological Past

Edited by GL~nterBuntebarth and Lajos Stegena

Springer-Verlag Berlin Heidelberg NewYork London Paris Tokyo

Editors Dr. GLinter Buntebarth Technische Universlt~.t Clausthal, Institut fur Geophysik Arnold-Sommerfeld-Str. 1, D-3392 ClausthaI-Zellerfeld, FRG Prof. Dr. La]os Stegena Institute of Environmental Physics, EStvSs-Unwersity Kun B61a T~r 2, H-1083 Budapest, Hungary

ISBN 3-540-16645-9 Springer-Verlag Berlin Heidelberg New York ISBN 0-38?-16645-9 Springer-Verlag N e w York Heidelberg Berlin

This work Is subject to copyright All rights are reserved, whether the whole or part of the material ~sconcerned, specifically those of translation, repnntlng, re-use of illustrat~one,broadcasting, reproduction by photocopying machine or similar means, and storage Ln data banks. Under § 54 of the German Copynght Law where copies are made for other than pnvate use, a fee Ls payable to "Verwertungsgesellschaft Wort", Munich. © Spnnger-Verlag Berhn Heidelberg 1986 Printed in Germany Pnntmg and bmdmg Beltz Offsetdruck, Hemsbach/Bergstr. 2132/3140-543210

PREFACE

During the l a s t decades, remarkable progress in heat flow studies has been made and a rough picture of the global surface heat flow density d i s t r i b u t i o n can now be drawn. Simultaneously, the question of over which time period the surface heat flow is constant arose. There is a big f i e l d of model c a l c u l a t i o n s , based on the changes in radioactive heat generation of the Earth, on plate motions, on s t r e t c h i n g hypotheses or on other ideas, which r e s u l t in geotherms in the geological past. Although these speculative paleogeotherms seem to be r e a l i s t i c e s p e c i a l l y in oceanic areas they do not belong to the scope of t h i s book. In continental areas however, i t is not possible to f i n d a simple time dependence of the surface heat flow density. However, petroleum research and tectogenetic studies are very interested in the geothermal h i s t o r y of sedimentary basins and other continental areas. To obtain s a t i s f a c t o r y r e s u l t s , a more or less d i r e c t determination of paleo heat flow dens i t y or geothermal gradient would be necessary to give more certain boundary cond i t i o n s f o r c a l c u l a t i n g o i l generation, and f o r c o n t r o l l i n g tectogenetic hypotheses. There are many methods a v a i l a b l e in the geosciences to determine temperatures in the geological past. Most of these models are able to estimate temperatures at which a mineral or a mineral assemblage was formed. These methods, however, are mostly unsuitable to reach the main goal of paleogeothermics in general, which is to determine the (regional) heat flow density v a r i a t i o n s during the geological past f o r bigger geological u n i t s , such as sedimentary basins. The methods applied most in sedimentary basins have been deduced from the degree of c o a l i f i c a t i o n of organic matter. Although much e f f o r t has been made to explain a n a l y t i c a l l y the organic metamorphism, the results found up to now have been i n s u f f i c i e n t . However, the widespread a p p l i c a t i o n of t h i s thermometer to estimate ancient thermal conditions is also r e f l e c t e d in the contents of t h i s very volume where the i n t e r p r e t a t i o n of the degree of c o a l i f i c a t i o n of organic matter plays an important r o l e .

As well as t h i s geothermometers, other methods are reviewed from a geophysical viewpoint which favours methods suitable to determine a paleothermal state of the upper crust. Further c o n t r i b u t i o n s of t h i s book deal with -

the h i s t o r y of the earth's surface temperature whose change provides an essential correction f a c t o r in heat flow density determinations,

-

isotope geothermometers and t h e i r a p p l i c a t i o n to various environments to evaluate thermal conditions in the past geological h i s t o r y ,

-

an a p p l i c a t i o n of the radiometric dating method to retrace the paleothermal condition of the Central Alps.

Most of the c o n t r i b u t i o n s were presented at the symposium "Paleogeothermics" which was held at the 18. General Assembly of the I n t e r n a t i o n a l Union of Geodesy and Geophysics, August 15-27, 1983 in Hamburg/FRG. I t has been the f i r s t

time that such a symposium has been organized by the I n t e r -

national Heat Flow Commission, and t h i s book presents an attempt to define paleogeothermics under the auspices of the I n t e r n a t i o n a l Heat Flow Commission.

G. Buntebarth I n s t i t u t e of Geophysics Technical U n i v e r s i t y Clausthal

L. Stegena I n s t i t u t e of Environmental Physics E~tv~s U n i v e r s i t y Budapest

CONTENTS

Preface

I

Contents

3

I.

Methods in Paleogeothermics BUNTEBARTH/STEGENA

2.

Temperature h i s t o r y of the earth's surface in r e l a t i o n to heat flow SHACKLETON

41

3.

Isotope geothermometers HOEFS

45

4.

Relations between c o a l i f i c a t i o n and paleogeothermics in Variscan and A l p i d i c foredeeps of western Europe TEICHMOLLER/TEICHMOLLER

53

The c o r r e l a t i o n of v i t r i n i t e in humic organic matter BARKER/PAWLEWICZ

79

5.

6.

7.

A comparison of two v i t r i n i t e paleotemperature gradients BUNTEBARTH/MIDDLETON

5

reflectance with maximum temperature

reflectance methods f o r estimating

Methods f o r paleotemperature estimating using v i t r i n i t e data: a c r i t i c a l evaluation VETU/D~VENY

95 reflectance 105

A reaction k i n e t i c approach to the temperature-time h i s t o r y of sedimentary basins SAJGO/LEFLER

119

Limits of a p p l i c a t i o n of the reaction k i n e t i c method in paleogeothermics LEFLER/SAJGO

153

Geothermal e f f e c t of magmatism and i t s c o n t r i b u t i o n to the maturation of organic matter in sedimentary basins HORVATH/DUVENY/LACZO

173

11.

Paleotemperatures in the Central Alps - an attempt of i n t e r p r e t a t i o n WERNER

185

12.

Geothermal studies in o i l f i e l d d i s t r i c t s of north China WANG/WANG/YAN/LU

195

8.

9.

10.

References

205

Subject Index

229

METHODS IN PALEOGEOTHERMICS

BUNTEBARTH,

G.* a n d L. S T E G E N A * *

* I n s t i t u t f~r Geophysik, T U C l a u s t h a l Arnold-Sommerfeld-Str. i, D-3392 C l a u s t h a l - Z e l l e r f e l d ,

F.R.

of G e r m a n y

** I n s t i t u t e of Geophysics, E ~ t v 6 s - U n i v e r s i t y K u n B ~ l a T ~ r 2, H-IO83 B u d a p e s t

Introduction An attempt is made to bring together geophysical, geological and geochemical methods bearing on ancient thermal conditions of the earth's crust. Methods are emphasized which are s u i t a b l e to estimate temperature gradients in the past, in order to evaluate the evolution of or merely the changes in the thermal regime w i t h i n the c r u s t . The a p p l i c a t i o n of the degree of c o a l i f i c a t i o n of organic matter has received p a r t i c u l a r a t t e n t i o n as a means of estimating the geothermal h i s t o r y of sedimentary basins because the degree of c o a l i f i c a t i o n is mainly influenced by the temperature of the environment and the time of exposure at t h i s temperature. Several empirical i n t e r pretation methods are reported which have been developed f o r s p e c i f i c basins and which are e s p e c i a l l y v a l i d f o r these areas. During crystal growth, l i q u i d s and other phases can be entrapped in the host c r y s t a l . These entrapped phases preserve the temperature and the pressure conditions which were present at the time of c r y s t a l growth. Chemical reactions are temperature s e n s i t i v e . Therefore, s o l u t i o n e q u i l i b r i a and isotope exchange reactions are applied to estimate paleothermal conditions, or to compare the calculated reaction temperature with the present thermal state in p a r t i c u l a r areas. A recent successfully tested method which deals with the transformation of minerals during diagenesis is reported. Clay minerals, z e o l i t e s and quartz polymorphs are transformed in sedimentary rocks of s i m i l a r composition at d i s t i n c t temperatures. Another method is reported which analyses the color a l t e r a t i o n of conodonts. This method is applicable f o r sedimentary rocks from the Late Cambrian to the T r i a s s i c period when the conodonts l i v e d .

Lecture Notes in Earth Sciences, Vol. 5 Paleogeothermics. Edited by G. Buntebarth and L. Stegena © Springer-Ver|ag Berlin Heidelberg 1986

Radiometric dating is the only method which y i e l d s a thermal h i s t o r y of c r y s t a l l i n e rocks. Because each r a d i o a c t i v e system has i t s own closure temperature, radiometric age determinations give the ages at which a rock cooled down to the respective c l o sure temperature. I.

Diagenesis of organic matter

Since organic l i f e grows on the earth, i t is included in the geological cycle. The remains of the organic matter are covered by sediments or deposited together with c l a s t i c d e t r i t u s . I f the circumstances are favourable, the organic matter is preserved and subsides within a sedimentary basin. During subsidence i t undergoes in creasing pressure as well as temperature, and both gradually a l t e r the o r i g i n a l m a t e r i a l . The a l t e r a t i o n of organic matter is known as diagenesis and process of c o a l i f i c a t i o n . There are two factors which govern predominantly the rank of c o a l i f i c a t i o n , which are the temperature in the depth where the organic matter existed during i t s h i s t o r y , a n d the time of i t s exposure. An i n t e r p r e t a t i o n of the degree of c o a l i f i c a t i o n based on the e f f e c t of temperature and time of exposure to that temperat u re, can be of l i m i t e d value only. More care must be taken on the o r i g i n of organic material and the f i r s t

steps in i t s structural and chemical decomposition

in d i f f e r e n t environments. The oldest coals which seem to be of plant o r i g i n are preserved in rocks of Algonkium age in North America. Several l o c a l i t i e s with coal embedded in a sedimentary sequence are known in the Lower Devonian. Since Middle and Upper Devonian, when plenty of plants grew on the continent and on the submerged shore, coal seams are more common. The most prominent bituminous coal deposits are of Carboniferous age in Europe and North America, and of Permian, Triassic and Jurassic age in South A f r i c a , Eastern A u s t r a l i a and India. Since Cretaceous, much more v a r i e t y in the f l o r a has been created which implies more heterogeneity in the plant remains from which the coaly matter o r i g i n a t e s . The coals are formed not only from d i f f e r e n t plant communities but also at d i f f e r e n t environmental conditions which are summarized by M. TEICHMOLLER & R. TEICHMDLLER (1981). I t is important, that the plants or t h e i r remains have to be deposited under conditions with r e s t r i c t e d oxygen supply. Usually, t h i s condition is present in swampy areas. I f a sedimentary basin with swampy areas subsides gradually, the organic matt e r can be deposited in layers of some thickness. A warm or temperate to cool climate with high humidity throughout the year is necessary to r e t a i n the condition favoured f o r organic deposition. There are a few peat-forming plant communities which grow in d i f f e r e n t swamp types, i . e . moss swamp, f o r e s t swamp, open reed swamps and p a r t l y submerged areas with water

plants. The most productive areas are f o r e s t swamps under t r o p i c a l c o n d i t i o n s . Economic~l coal

seams y i e l d from deposition in swamps, in general. As well as in coal

seams, organic matter is also present in a dispersed form in many minerogene sedimentary rocks. Plant remains in r i v e r deltas and on the shores of lakes and oceans, barks, other plant d e t r i t u s , and also coal which can be redeposited, can be covered by c l a s t i c sediments and buried. I f the environmental conditions are favourable f o r preservation, the organic substances undergo the diagenesis during the subsidence, and w i l l become coaly p a r t i c l e s l i k e the plant remains in swamps. However, there is a difference. The plant remains are exposed not only to the mechanical treatment during the transport by water, but also to the o x i d i z i n g atmosphere and to the bacter i a l a c t i v i t y at the surface which favours the preservation of e s p e c i a l l y r e s i s t e n t p a r t i c l e s . This means that the o r i g i n a l organic substance is not exactly the same as in seams. The composition of the organic matter bearing rocks is of some importance too. The organic matter is often oxidized in sandstones, e s p e c i a l l y in red-coloured ones, but is rather seldom in limestone. Usually clay and s i l t s t o n e s are the rock types from which the organic p a r t i c l e s can be observed and interpreted f o r paleogeothermal i n v e s t i g a t i o n s . There are a d d i t i o n a l factors i n f l u e n c i n g the composition of the organic substance which y i e l d s the coaly matter. Whereas organic deposits under t e r r e s t r i a l

and sub-

aquatic conditions are comparable, marine-influenced and calcium-rich swamps produce coals r i c h e r in ash, sulphur and nitrogen. These conditions imply that a d i f f e r e n t a c i d i t y of water may produce coals of same d i s t i n g u i s h a b l e properties, even with the same o r i g i n a l material. I t seems that the bacterial a c t i v i t y is a most important f a c t o r c o n t r o l l i n g the decomposition of plants and thereby at least the o r i g i n a l materials f o r the coals. Therefore, a l l environmental properties which favour or prevent bacterial l i f e also define the properties of the coal. V i t r i n i t e

is a most

common c o a l i f i c a t i o n product which is formed from organic deposits under some acid c o n d i t i o n . I f the environment is neutral to weakly a l k a l i n e , the bacterial a c t i v i t y is very high. Since the protein of the bacteria is also accumulated, the organic substances y i e l d hydrogen-rich bituminous products which form b i t u m i n i t e and weakly reflecting vitrinites Peat is the f i r s t

during subsidence (M. TEICHMOLLER & R. TEICHMOLLER, 1981).

stage in the diagenetic process of the organic matter. P e a t i f i c a -

t i o n can s t a r t a f t e r the b u r i a l of plant remains with the help of the bacteria, which are active down to some meters of depth. With continuing subsidence, the i n creasing overburden pressure causes the water to be squeezed out of the organic substances.

The temperature during t h i s physical process may range between about

20 to 50° C. At the upper l i m i t of the temperature range, l i t t l e (van HEEK et a l . ,

methane is s p l i t o f f

1971), and the transformation from peat to brown coal is usually

reached in a depth range between 200 m and 400 m. At temperatures of about 70 to

I00 ° C C02 is released, and at temperatures of about 160 to 200 ° C, at which low v o l a t i l e bituminous coal gradually changes to semi-anthracite, large q u a n t i t i e s of methane develop. The rank of coal is determined in a general way by appearances and/or by i t s propert i e s , e.g. b r i g h t brown coal and gas coal. This q u a l i t a t i v e scale is not s u f f i c i e n t f o r a n a l y t i c a l i n v e s t i g a t i o n s . The composition of organic matter in sediments is 90 % kerogen and 10 % bitumen (hydrocarbon, r e s i n , asphaltene). The f r a c t i o n soluble in organic solvents, is called bitumen, whereas the other f r a c t i o n , insoluble in organic matter, is termed kerogen. There are methods to estimate the maturity by examining the soluble organic matter: percentage carbon in bitumen, carbon preference index (odd carbon number compounds to even carbon number), p a r a f f i n p r o f i l e , percentage wet gas. Other, more important methods, examine the kerogen as a maturation index. These methods are the kerogen a l t e r a t i o n index KAI, thermal a l t e r a t i o n index TAI, p y r o l y s i s , elementary CHO a n a l y s i s , and atomic H/C r a t i o . A l l these chemical rank parameters are not applicable in general f o r rocks with f i n e l y dispersed organic matter, because the chemical methods need some amount of organic p a r t i c l e s . The rank determination with microscope is successful. The method is not destructive f o r the sample, and is easy to apply. V i t r i n i t e

is the most common coal

maceral, and is the one taken in order to measure i t s o p t i c a l r e f l e c t i v i t y polished sample under o i l ,

at the

applying monochromatic l i g h t . This method is applicable to

both the coal from seams and the coaly p a r t i c l e s dispersed in sedimentary rocks. Vitrinite

reflectivity

is a r a t i o of the i n t e n s i t y of the r e f l e c t e d l i g h t and the

source l i g h t , expressed in percent, using v i t r i n i t e

(= woody kerogen) as the r e f l e c -

t o r . The value is often simply called Ro, % Ro, or % Rm the subscript "o" designates that the measurement was made in o i l , of Rmax, the maximum r e f l e c t i v i t y ,

and "m" means the mean r e f l e c t i v i t y ,

which should be applied at r e f l e c t i v i t y

instead values

above 4 % Rm. The r e f l e c t i v i t y nite/vitrinite

c o e f f i c i e n t gives a continuous scale f o r the c o a l i f i c a t i o n of humiwith values ranging from about 0.2 % up to more than 5 % (M. TEICH-

MOLLER, 1970). Huminite and v i t r i n i t e

are maceral groups of humous components,

where huminite is the precursor of v i t r i n i t e

in peat and brown coal. During the

progress in c o a l i f i c a t i o n huminite is converted i n t o v i t r i n i t e

between the c o a l i f i -

cation stages of dull and b r i g h t brown coal. I f some rocks are so poor in organic matter that concentrates must be prepared by chemical or physical methods, i t is much more d i f f i c u l t

to determine the correct

degree of c o a l i f i c a t i o n . The surroundings of the p a r t i c l e s are often helpful

to select the representative ones f o r measurement. The selection of the correct coal macerals, i . e . v i t r i n i t e ,

poses the greatest d i f f i c u l t y

in the determination of

the degree of c o a l i f i c a t i o n in rocks. For t h i s determination the so-called "kerobitumen" which can be found in bituminous shales is of some importance. The b i t u minous matter r e f l e c t s in the lower rank of c o a l i f i c a t i o n less than v i t r i n i t e ,

but

more in the rank of a n t h r a c i t e . The d i s t i n c t i o n between recycled and authochtonous organic matter is often d i f f i c u l t

in rocks, but nearly impossible in concentrates.

There are a l o t of problems a r i s i n g from the selection of macerals f o r measurements, which are described more d e t a i l e d e.g. in STACH et a l . (1982), ROBERT (1985), TISSOT & WELTE (1978). Besides the r e f l e c t i v i t y of v i t r i n i t e

in shales, sandstones and limestones with dis-

persed coaly p a r t i c l e s , the spectral fluorescence measurements on s p o r i n i t e has been introduced as an i n d i c a t o r of the degree of diagenesis. I f s p o r i n i t e is i r r a d i a t e d with u l t r a v i o l e t l i g h t (A=365 + 30 nm), a v i s i b l e fluorescence can be observed from yellow to dark red colour. However, the s p o r i n i t e fluorescence spectra are observed at low grades of diagenesis only, i . e . from the stage of peat to that of high v olat i l e bituminous coal (OTTENJANN et a l . ,

1974).

Both parameters, the r e f l e c t i v i t y of v i t r i n i t e

and the s p o r i n i t e fluorescence, are

used together to f i n d a more correct degree of diagenesis. The i n t e r p r e t a t i o n of the rank of c o a l i f i c a t i o n f o r paleogeothermics is based on the f a c t that the temperature is the most important f a c t o r that increases the degree of c o a l i f i c a t i o n , but the duration of heating must also be taken i n t o consideration. The influence of pressure, however, seems to be n e g l i g i b l e . Based on HUCK & KARWEIL (1955), LOPATIN (1971) gave a simple scheme f o r describing the degree of c o a l i f i c a t i o n . Supposing that the c o a l i f i c a t i o n process is to be treated as a f i r s t

order chemical reaction, the

Arrhenius' equation is v a l i d and the v e l o c i t y of the reaction (k) depends exponentially

on temperature: k = a exp(-E/RT)

(a: frequency-factor, E: a c t i v a t i o n energy, R: gas-constant, T: temperature in Kelv i n ) . Numerous chemical reactions double t h e i r reaction v e l o c i t y f o r each 10° C temperature growth, not f a r from room temperatures, because t h e i r a c t i v a t i o n energy l i e s around 54 kJ/mole. LOPATIN (1971) accepted t h i s value and suggested that the dependence of maturity on time is l i n e a r , and the dependence on temperature has an exponential character. Therefore, the v e l o c i t y of the " c o a l i f i c a t i o n " reaction can be w r it t en as k ~ 20"IT(t)

10

and the parameter which describes the rank of c o a l i f i c a t i o n t* C~

20"IT(t)dt

where T ( t ) is the temperature of the layer during the time i n t e r v a l dt, and t * is the time from the deposition of the layer t i l l

the present.

For practical reasons, LOPATIN introduced the sum instead of the i n t e g r a l , d i v i d i n g the whole temperature h i s t o r y of the layer i n t o 10° C temperature i n t e r v a l s . He then arbitrarily

chose the 100 to 110 ° C temperature i n t e r v a l (which is the mean domain

of o i l generation) as the base i n t e r v a l and assigned to i t an index value of n=O, to the 120- 130° C i n t e r v a l n = 2 , to the 90-100 ° C i n t e r v a l n = - 1 , and so on. The maturity parameter calculated in t h i s manner was called the Time Temperature Index (TTI), nmin TTI =

~(~tn)2 n nmax

where ~t n is the time i n t e r v a l (in Ma) the layer spent in the n-th 10° C temperature i n t e r v a l , and nmax and nmin are the n-values of the highest and lowest temperature i n t e r v a l s occurring in the thermal h i s t o r y of the layer. w

QC

n

20

Ma

I0

-9

-8 40 SO

- 7 - -6 .

0

0

Fig. I . LOPATIN's (1971) method f o r the calcul a t i o n of the Time Temperature Index f o r a layer l y i n g at a depth of 2300 m, aged 20 Ma. TTI is c h a r a c t e r i s t i c f o r the maturity of organic matter.

T , ~

-

70

110

I

120

2

Tim-Temperature

2 Depth, km Index ( 1-1"1)

/ITn : in M a

Fig. I demonstrates the method of c a l c u l a t i o n of TTI, f o r a hypothetical layer 20 Ma old and l y i n g at present at a depth of 2300 m. Let us suppose that the subsidence and burial h i s t o r y of the layer during geologic time was determined as shown by the curve of Fig. I . Let us then suppose that the present geothermal gradient is 50 mK/m, and the gradient was constant during the whole sedimentary h i s t o r y , as shown in Fig. I , by the horizontal s t r a i g h t geotherms. In t h i s case f o r the layer of Fig. I , TTI= ;5.2.

11

Based on 402 thermal maturity (Ro) data from 31 worldwide wells, WAPLES (1980) determined a correlation between TTI values calculated for each borehole from burial h i s t o r i e s , supposing the v a l i d i t y of present geothermal conditions during the geological past, and Ro values measured (Fig. 2).

1o

3

i~

30

~

30O

t o(x~o

I ~

30OO

3OOO0

1000000

~x~OeO

Inaa,

Fig. 2. Correlation between the Time Temperature Index of maturity and v i t r i nite reflectance Ro ( a f t e r WAPLES, 1980)

\

v,tnmte

\ \\ \

These antecedents make possible the paleo heat flow estimation for a borehole, by the following steps: -

Based on known ages of some sedimentary layers in the borehole, the sedimentary history for these layers is determined (Fig. 3, dotted l i n e s ) .

Mai

~Ma A 12 ~0 C8 DSE 4

2

0°

1 S'S

!

...... 13

'

JX\

tS

0epth.

Fig. 3. Sedimentary history of a borehole (HOD) in the Pannonian basin, calculated on the basis of the ages in the l e f t hand column, with and without correction of compaction (STEGENA et a i . , 1 9 8 1 )

12

Using porosity-depth functions and/or other considerations, the sedimentary h i s t o r i e s are corrected f o r the e f f e c t of compaction during the geological past (Fig. 3, s o l i d l i n e s ) (DU ROUCHET, 1980; STEGENAet a l . ,

1981; FALVEY & DEIGHTON,

1982). Based on present borehole temperatures, the geotherms f o r each 10° C round i n t e r val are constructed in the time-depth section (Fig. 4, l e f t ) with the present heat flow during the geological past. The constancy of heat flow during the past does not r e s u l t in p a r a l l e l and e q u i d i s t a n t s t r a i g h t l i n e s ; i t is possible to take i n t o consideration the probable changes with time and depth of thermal cond u c t i v i t y of the layers, with the aid of the l i t h o l o g y and burial h i s t o r y of the borehole.

A

°C

n

20

:: -

12

--%. IBOC 8

6E

Ma 4

2

0

°C 0

4o

5O 60 70

4

80 90

2

ICO

20 i ~"

0

120 130

!

140

4

40

_

50

2 T T I = 2 ' 6 ~ R o = 0"45

70

-

-

TTJ=I"2 ~ R o = 0 ' 4 0

80 3

1 ~ = 1 9 ~ R o = 0"69

-2

I 3 T;]=T4~R°

: 0"55

SO -1

:

170 180

i,

60

150 160

j

30 I

1

110

1,2

-9

i

v

30

.

100 0

4 TTI= 1 7 8 ~ R o = 1'35 !

TTI=26--R o = 0,74 11o t

190

120

200 210

S TflI1347~R 10 tl

220

2

o - 2'16

0epth,

130

I T I = 3 4 6 ~ R o = 1'58 3

km '8 T r t = 8 0 1 2 ~ R o = 3"17

1 9 0 / / ' / i " 6 T T I = 1 9 1 2 ~ R o = 2'32 210

Fig. 4. Calculated TTI values f o r the borehole HOD assuming that the heat flow density was constant through the sedimentary h i s t o r y ( l e f t ) , and that the borehole was heated up during the l a s t 5 Ma f o r the present heat flow value (STEGENA et a l . , 1981) A f t e r c a l c u l a t i n g the TTI values f o r each layer of the borehole the TTI-s are transformed to Ro values (Fig. 4), using the c o r r e l a t i o n of WAPLES (1980) (Fig. 2). These calculated R° values are compared with the Ro values measured in the borehole. The discrepancy between calculated and measured values is a t t r i b u t e d to the v a r i a t i o n s of heat flow during the geological past. Using plausible hypotheses, one makes a change in the past heat flow (Fig. 4, r i g h t ) and repeats the comparison of Ro values calculated from TTI-s and measured Ro values, t i l l between calculated and measured v i t r i n i t e

a good f i t

reflectances is achieved (Fig. 5).

13

Fig. 5 shows two boreholes of the Pannonian basin with heat flow h i s t o r i e s calculated independently. Both boreholes gave the same r e s u l t : the measured v i t r i n i t e reflectances are compatible with the assumption that the Pannonian basin has had a low heat flow ( ~ 5 0 mW/m2) before 5 Ma, and 5 Ma ago the heat flow began to in crease ( l i n e a r l y ? ) to i t s present value ( - 1 0 0 mW/m2). 0.2

0.3 0,4 0,50,6 O,O 1.0

2.0

3,0 4,0 6.0

O0,t

0,2

0.3 0.4 0,5 0,6 O.il 1.0

2,0

3.0 4,0 5.0

Ro,%

Ro %

1 2 W

U

3 4

Ma

em 5-

~16 14 16

4~ 2

0

I)epm, km

Fig. 5. The measured v i t r i n i t e reflectances in the borehole HOD and DER (both in the Pannonian basin) and the v i t r i n i t e reflectances calculated from the f o l l o w ing heat flow s t o r i e s : the heating-up of the boreholes began at co, 5, 2, I Ma ago ( a f t e r STEGENA et a l . , 1981). The above scheme serves better to understand the p r i n c i p l e s of the paleogeothermal c a l c u l a t i o n s , but does not present a f i n a l solution of the question. There are some fundamental problems in the o i l geochemistry which are not solved s a t i s f a c t o r i l y and which can influence the above sketched model. I t became usual to assume that increases in v i t r i n i t e

reflectance values were

v a l i d indicators of the extent to which organic matter maturated and o i l generat i o n had occurred (WAPLES, 1983). However, there is an uncertainty in some R o measurements, because the values have a wide spread, and sometimes i t is hard to d i s t i n g u i s h low r e f l e c t i n g r e s i n i t e and high r e f l e c t i n g fusunite from v i t r i n i t e s (HO, 1978). During the beginning of o i l generation, bitumen impregnations lower the v i t r i n i t e

reflectance. In a l l red-coloured rocks organic matter is oxidized;

in limestones v i t r i n i t e

is very r a r e l y preserved and i f i t occurs, the reflectance

value d i f f e r s from the value of v i t r i n i t e

in the same rank. RONSARD & OBERLIN

(1984) suggest t h a t , as with any other e l e c t r o n i c property of any s o l i d , r e f l e c tance depends on three parameters: chemical composition, atomic structure and microstructure. The same value f o r reflectance can thus be measured f o r materials d i f f e r e n t in t h e i r microstructure and chemical composition, which can be of d i f f e -

14 rent ranks or not. They suggest the use of transmission electron microscopy (TEM) by using successive heat treatment in an i n e r t atmosphere to 1000° C, which bett e r characterizes the maturation of organic materials. I t is generally supposed that pressure does not have a s i g n i f i c a n t e f f e c t on the maturation of organic matter and on the amount of hydrocarbon generated. I t is to be noted however, that the role of pressure in o i l generation has never been examined properly (WAPLES, 1983). The maturation of organic matter e x h i b i t s a very complex process, inv o lv ing a l o t of p a r a l l e l chemical reactions with various a c t i v a t i o n energies, and the whole process can hardly be described by a f i r s t - o r d e r k i n e t i c expression (SIEVER, 1983). This was also shown by pyrolysis experiments (CUMMINGS & ROBINSON, 1972). LASAGA (1981) has compiled a table of a c t i v a t i o n energies f o r geochemical reactions that shows a range from less than 4 kJ/mole to

more than 400 kJ/mole.

TISSOT(1969),

TISSOT & ESPITALIE (1975), TISSOT et a l . (1975), and JONTGEN & KLEIN (1975) have modelled the thermal a l t e r a t i o n of kerogen with a set of f i r s t - o r d e r rate equations, E d nki a i exp(- i T = -nki ~)

i = 1,2 . . . . 6

where nki is the mass function, a i is the frequency f a c t o r , Ei is the a c t i v a t i o n energy of the i - t h kerogen. I f i t is integrated over the thermal h i s t o r y of any horizon, the generated petroleum and the maturity of organic matter can be calculated. This process although giving a b e t t e r t h e o r e t i c a l approximation, is hardly applicable f o r paleogeothermal a p p l i c a t i o n s . LOPATIN (1971) tested his model on a very d i f f i c u l t

w e l l , MUnsterland I/FRG. Recali-

bration of Lopatin's method with l a r g e r and more r e l i a b l e data sets (WAPLES, 1980; KETTEL, 1981) has v e r i f i e d the general v a l i d i t y of the model i t s e l f ,

but has modi-

f i e d Lopatin's o r i g i n a l T T l - v i t r i n i t e r e f l e c t i v i t y c o r r e l a t i o n . LOPATIN & BOSTICK (1973) and LOPATIN (1976) l a t e r suggested some improvements to the o r i g i n a l scheme. LOPATIN (1976) used fewer and l a r g e r temperature i n t e r v a l s ; instead of ~T= 10° C, 1.37T 2 ~T = ~ . 3 7 T (T in Kelvin, E a c t i v a t i o n energy=42 kJ/mole). This formula gives 15° C f o r ~T at T50 ° C/km. STRASSER & WOLTERS (1963) determined a volumetric density of 2,8 g/cm3 f o r Devonian clayey sandstones in the MUnsterland I borehole at 5700 m. A f t e r a diagram from CISSARZ (1965: Fig. 2) t h i s value corresponds t o a r o c k t e m p e r a t u r e of 300 ° C. According to BREITSCHMID (1982) the top of the epi-zone (boundary between very low grade and low-grade metamorphism) corresponds to a v i t r i n i t e and to a rock temperature of borehole a v i t r i n i t e

reflectance of 5,5 % Rm

270 ° C in the Swiss Helveticum. In the MUnsterland I

reflectance of 5,5 % Rm is reached at a depth of 4700 m

(Fig. 12) which indicates a palaeogeothermal gradient of 57° C/km, facing a gradient of 33,6 ° C/km at the present time (HEDEMANN & R. TEICHMOLLER, 1966). A comparison of subsidence curves f o r the e a r l y T e r t i a r y of the Upper Rhine Graben and f o r the Ruhr Carboniferous (Fig. 17), with the temperatures measured in boreholes of the Upper Rhine Graben taken into consideration (Fig. 5), led to the

67 conclusion t h a t , i n the Ruhr Basin, the palaeogeothermal gradients reached 5 0 - 7 0 °C/km during the Westphalian (R. TEICHMOLLER, 1973). Geothermics of A l p i d i c foredeeps Northern Alps The molasse of the A l p i d i c foredeep north of the Alps is comparable to the molasse of the Subvariscan foredeep north of the Rhenohercynicum. Depths of burial and e f f e c t i v e c o a l i f i c a t i o n times are s i m i l a r in both cases (Fig. 17). However, the c o a l i f i c a t i o n pattern is very d i f f e r e n t . In the foredeep of the Alps, burial depths of 2,5 km (near Munich) to 5 km ( i n the Alps) were attained f o r the basis of the molasse (JACOB et a l . , uplift

1982). Folding and

began 8 to 12 m i l l i o n years before the present, in Late Miocene to Pliocene

times. As usual in foredeeps, f o l d i n g moved from the mountains (12 Ma) to the f o r e land (8 Ma). Fig. 14 shows a cross-section through the foreland and the northernmost part of the German Alps. In the north, the A l p i d i c foredeep is f i l l e d

with autochtonous, f l a t

l y i n g T e r t i a r y molasse sediments. Approaching the mountains in the south, the molasse is folded and imbricated. S t i l l

f u r t h e r south, in the Alps, t h i s imbricated, sub-

alpine molasse is overlain by imbricated strata of the Helveticum, Flysch and Austroalpine nappe p i l e s . N

non folded Molasse

folded Molasse

-

I

t

|

A........

~/",~...,_.~

"~.......

I........

LL ,

~.~

xi.

I

~L

S

Tegernsee Alps

Hausham syncline

---~.-.

Rupelian

G~an~bra, nk . ~

~Austloalpme

%.+ Locene

" ~ F l a m mkohle

...,..,,,o~. ~ . , . . ~

. ~oh.,pi.~

Mo,

....

~,,,,'~7,,c;e(ac~-~E77+;z,;.,T,~,Flammk.oh,!e ....

~m~.,,,.////////l//7~/~./p/77/z~f/,7/l/Tf//7~//.~. ~

600C

-6000m Ma 71k GIIk

mili,ofls ~f y~ars belor~ Lh, pleserll Flamalkohle

Gasllammkuhle

OL

_ _

lO[km I

I

i

i

after TEICHMOLLER1975, supptemenied Tectonics ol the Flysch afler SCHMIO1 THOM[ 1997 Tectonics ol the Molasse alter VEIT 1963 and M MULLER 1975 Ceahficahon of Ihe Flysch ariel MWOLF

Fig. 14. Cross-section through the subduction zone at the northern rim of the Alps with c o a l i f i c a t i o n p r o f i l e s ( a f t e r M. TEICHMOLLER & R. TEICHMOLLER, 1978). In the foreland as well as in the northern Alps, boreholes of the o i l industry permitted a study of the downward increase of v i t r i n i t e

reflectance. According to c o a l i f i c a -

t i o n studies of JACOB & KUCKELKORN (1977), the Miesbach I borehole did not encounter the boundary between sub-bituminous and bituminous coal (boundary between Braunkohle and Steinkohle according to the German coal c l a s s i f i c a t i o n ) u n t i l a depth of about

68

5000 m in the molasse (see Fig. 14). In t h i s borehole the present geothermal gradient is 23,5 ° C/km (JACOB & KUCKELKORN, 1977). JACOB et a l . (1982) evaluated c o a l i f i c a t i o n data from the Alpine molasse obtained by JACOB and other authors. They found that in the south, w i t h i n the folded and imbricated molasse, the c o a l i f i c a t i o n gradients are e s p e c i a l l y low, and even too low compared with the present geothermal gradients. Therefore, these authors assume a prekinematic c o a l i f i c a t i o n having taken place 10-20 km f u r t h e r south where the temperature gradients are lower at present and obviously have been lower in p r e - t h r u s t i n g times as w e l l . I t seems that in t h i s region the geothermal pattern of the northern Alps has not changed much since Pliocene/Upper Miocene times when the sediments of the present folded molasse reached t h e i r deepest level of subsidence and, thus, t h e i r highest c o a l i f i c a t i o n temperatures. In contrast to the northern Alps f u r t h e r west (e.g. near Oberstdorf, Allg~u and, p a r t i c u l a r l y , in Switzerland) where the main coalification

(as well as i l l i t e

diagenesis) occurred a f t e r imbrication due to former

t h i c k overburden with Penninic nappes (M. TEICHMOLLER & R. TEICHMOLLER, 1978; FREY et a l . ,

1 9 7 3 ) , a t t h e northern rim of the Bavarian Alps no post-kinematic c o a l i f i c a -

t i o n of the folded and imbricated molasse took place. Fig. 15 shows the r e s u l t s of JACOB et a l . (1982) with c o a l i f i c a t i o n gradients f o r the autochthonousaswell as the folded and imbricated molasse in i t s p r e - t e c t o n i c , non imbricated p o s i t i o n . The gradual decrease from the Anzing 3 borehole in the north ( l y i n g 23 km north of the sign f o r Anzing 3 in Fig. 15) with 0,09 % Rm/km, to Hausham in the south with 0,04%Rm/km, and f u r t h e r west, - from the Staffelsee I borehole in the north (0,06 % Rm/km) to the Egling I borehole in the south (0,03 % Rm/km), is evident. The present geothermal gradient in the Anzing 3 borehole in the north is 22,8 ° C/km, whereas i t is 22 °C/km in the V o r d e r r i s s l borehole (east of Garmisch-Partenkirchen) in the south (BACHMANN& MOLLER, 1981). The decrease of c o a l i f i c a t i o n gradients towards the south is obviously caused by a decrease of the geothermal gradients in the same d i r e c t i o n , - not only in the present time but also during Late Miocene/Pliocene times when maximal depths of subsidence were reached (M. & R. TEICHMOLLER, 1975; JACOB & KUCKELKORN, 1977; JACOB et a l . , 1982). This geothermal pattern is in accordance with the crustal thickness increasing from north to south, due to subduction. Fig. 16 shows, over a horizontal distance of 55 km, the dipping of the Mohorovi~i~ d i s c o n t i n u i t y from 33 km depth in the northern Alpine foreland (Anzing 3 borehole) down to 40 km depth beneath the northern Alps (Tegernsee 2 borehole). Thus, we must assume r e l a t i o n s h i p s between c o a l i f i c a t i o n gradients, geothermal gradients and the thickness of the crust in the foredeep of the German Alps.

69 ggk~ Anzing3 • /

11"E

/

~j
L-~'*~..% L

b

~- I

~ ' / ' ~.-----~-~;~,U

I%

+

bg/ I .7 V"

JG,

~

N

C

d

a

¢'---~ 2

"~

Y

~

L

:

~

ofalpinen~PMPleSsBAo=~C L ~ ' ~

~

~

-

i

~ ~

+ // c h b ~"" ~ b ,,~'~',.~

a

. , - - - ~~

~/

X -,'/i

t

~"

(-

,2. .'

.....

"~'+ + ~AR~ISCH ~ -

A-~

MIESBAC_H_~L0.08~~

f u ~

.

.

-- i . . . . . . . ~ "' i ~ g e SCH = Scheibum m A = Altenau ' •

_ A U S T R. .I. .A _

ROSENHEIM( ~)

MOLASSE

.~~vo...,s,(

. . . .

\ \x~:~ PARTENKIROHEN ~ % '(,/ \

~

+~r"----'~

S

0.09

~

k',

I

I

"~J"

JECH~..O..---O-06i

._/ ~i',

~-~-_

=~wOR=

_ _

....

,7.=

=

--

7

..=.

L " ' ' • horeholes ~ ...... p.rain mine i I i th thrusting urstingdistance di......

~.

---I~ +

.xisofOhatii. . . . . . gh

"1.0

~,,

FOREL.IFOLDED MOLASSE I-r u_ STAFFELSEE1 Mur~u N ] 2EGLING1 pre=klnem~ ~ MURNAU2 STAFFELSEE 1

KALKALPINE VOROERRI$S 1 proj

E6LIN61

pre.knem.pos,

ZONE Mlttenwold

Sc~omifz

S NN-

"

s*N,=oc

:.

" ~::]" I':!']~;#~,

:!600) with maximum b u r i a l temperature (Tmax in °C). These data are modelled by the l i n e a r regression equation: In(Rm) = 0.078 Tmax - 1.2 Tmax and Rm were compiled from over 35 systems, r i c h in humic organic matter to minimize the e f f e c t of chemical composition on Rmo The thermal maturation data span a range from early diagenesis to greenschist metamorphism overa Tmax i n t e r v a l from about 25-325 ° C and 0 . 2 - 4 . 0 % Rm. Burial h i s t o r y reconstruction indicates that the functional heating duration (elapsed time as temperature increases w i t h i n 15° C of Tmax) of these systems ranges from 10,000 yr to more than 10 m.y. The strong correl a t i o n of Tmax with Rm, i r r e s p e c t i v e of functional heating duration and in diverse geologic systems, indicates that increasing time at Tmax has l i t t l e

influence on

thermal maturation of sedimentary organic matter. Instead, uncertainty in correction of borehole temperature logs, Tmax determination, and Rm measurement explains much of the remaining v a r i a b i l i t y

not accounted f o r by the regression analysis. We did

not attempt to correct the measured borehole temperature to e q u i l i b r i u m r e s e r v o i r conditions because there is no consensus on which method to use, the necessary data is often unrecorded, and predictions made from our T -Rm c a l i b r a t i o n are compared max to uncorrected Tmax data. We found that Tmax is d i f f i c u l t to determine in sedimentary environments that have cooled because of the weak thermal imprint on the rocks in low temperature systems and poorly-known b u r i a l h i s t o r i e s . V a r i a b i l i t y in Rm measurement appears mainly due to operator or laboratory bias, increasing bireflectance with rank, and v a r i a t i o n in diagenetic h i s t o r y which causes reflectance suppression. These studies imply that Tmax controls Rm, making the r e l a t i o n s h i p useful as a maximum geothermometer, but that several physico-chemical and technical factors obscure the c o r r e l a t i o n . The problems in measuring Tmax and Rm, shown by the appreciable data s c a t t e r , make our c a l i b r a t i o n imprecise. However, a p p l i c a t i o n of t h i s geothermometer to systems where Tmax is well known shows that i t y i e l d s r e a l i s t i c paleotemperature

Lecture Notes in Earth Sciences, Vol. 5 Paleogeothermics. Edited by G. Buntebarth and L. Stegena © Springer-Verlag Berlin Heidelberg 1986

80 estimates. Other support f o r temperature control of Rm is documented from studies of metamorphic mineral assemblages and coal rank, and c r i t i c a l

t e s t i n g of temperature-

time-rank models in sedimentary systems. Introduction Maximum burial temperature and heating duration (or geologic time) are commonly reported as the most i n f l u e n t i a l controls in the thermal maturation of sedimentary organic matter (OM) (WAPLES, 1984). Recent studies suggest that the e f f e c t of heating duration is l i m i t e d , and r e l a t i v e l y soon (in a geologic time frame) a f t e r reaching maximum temperature, OM s t a b i l i z e s and ceases s i g n i f i c a n t reaction (BARKER, 1983; PRICE, 1983; and others). Under these c o n d i t i o n s , mean random v i t r i n i t e

reflectance

(Rm), a measure of thermal maturation, should be a function of the maximum temperature reached in the system. The question of whether heating duration is of continuizng importance in thermal maturation of OM can be addressed by p l o t t i n g a scatter diagram of the maximum burial temperature (Tmax) versus Rm f o r samples from sedimentary systems with a wide range of heating duration. A near-zero c o e f f i c i e n t of correl a t i o n (r 2) calculated f o r these data would indicate that these data vary randomly with respect to each other and that a t h i r d variable (presumably heating duration) has a s i g n i f i c a n t e f f e c t on the system. Conversely, a r2-value approaching one would indicate that Tmax and Rm are strongly dependent, and that heating duration cannot be a s i g n i f i c a n t f a c t o r in the thermal maturation of OM over extended geologic times. Our approach to c a l i b r a t i n g a maximum-recording geothermometer based on thermal maturation of OM is to compile published Tmax and Rm data from systems with a wide range of functional heating duration. Regression analysis of Tmax and Rm indicates whether t h i s system

can be adequately characterized by considering only these two

variables. This approach is necessary because of the uncertainty in sedimentary systems in determining a chemically meaningful heating duration (an estimate of reaction time), or a temperature of chemical reaction. Coupled with only a schematic understanding of the reaction mechanisms in OM (TISSOT & ESPITALIE, 1975) i t is d i f ficult

to define r i g o r o u s l y thermal maturation as a chemical system. The advantage

of t h i s approach is that the data is compiled from numerous independent measurements in diverse sedimentary systems. The l i m i t a t i o n s of our empirical method to modelling OM thermal maturation are t h a t : ( I ) maximum temperature reached in the system may be difficult

to determine; (2) correction of borehole temperature data to e q u i l i b r i u m

r e s e r v o i r conditions is poorly understood; and (3) Rm data may be subject to s i g n i f icant

e r r o r . Our model is a simple approximation of a complex chemical system. How-

ever, as shown by numerous studies (WAPLES, 1984), and by examples presented in t h i s paper, OM thermal maturation is successfully modeled by simple empirical functions.

81 Data and Results Maximum Temperature and V i t r i n i t e

Reflectance

Tmax and Rm data were compiled from over 35 systems (boreholes in sedimentary basins) undergoing conditions ranging from burial diagenesis to greenschist metamorphism. Only r i c h in humic OM were used because of the influence of chemical compositions on Rm. Maximum v i t r i n i t e

reflectance data was converted to mean random v i t r i n i t e

reflectance using the equation of TING (1978). The geothermal curve was determined by computing a geothermal gradient from bottom hole temperature (BHT) and the meanannual surface temperature. The reported sample depth was used to compute the system temperature at

that point by i n t e r p o l a t i o n from the geothermal curve. The T condition max in the system was indicated by geologic and s t r a t i g r a p h i c reconstruction (WAPLES, 1981) or comparative geothermometry (BARKER & ELDERS, 1981). The v a r i a b i l i t y

in

approach to determining Tmax and Rm makes t h i s a heterogeneous data set. We generally had to accept both the Tmax and Rm data without confirming t h e i r accuracy. Only systems where apparent maximum temperature existed or could be interpreted from other evidence are used to c a l i b r a t e t h i s OM thermal maturation model. Heating Duration Methods of computing heating duration f o r OM thermal maturation are attempts to e s t i mate reaction time during b u r i a l where temperature slowly changes, or s t a b i l i z e s , over an extended geologic time.

HOODet a l . (1975) defined the e f f e c t i v e heating-

time f o r OM thermal maturation as the elapsed time when the system is w i t h i n 15° C of maximum temperature. MACKENZIE & McKENZIE (1983) have demonstrated that t h i s is a

reasonable measure of reaction time in sedimentary systems. We agree with t h e i r

conclusion except that thermal maturation of OM, given s u f f i c i e n t time, could s t a b i l i z e by reactions consuming a l l p o t e n t i a l l y cleavable bonds at maximum temperature. Experiments indicate that thermal maturation of OM proceeds by p a r a l l e l reactions that have a wide range of a c t i v a t i o n energy (Ea) (JONTGEN & KLEIN, 1975; TISSOT & ESPITALIE, 1975). For areasonable b u r i a l temperature, the wide Ea range over which OM thermal maturation occurs suggests that the reactions: ( I ) with a low Ea are complete and not generating products; (2) with a moderate Ea are generating s i g n i f icant

products; or (3) with a high Ea w i l l be slow and not complete reaction in

geologic time. The general c o r r e l a t i o n of Rm with T (NERUCHEV& PARPAROVA, max 1972; SUGGATE, 1982; and others) suggests that only a l i m i t e d suite of reactions control the thermal maturation of OM, and additional heating duration w i l l not make the slower (high Ea) reactions s i g n i f i c a n t in increasing Rm. Thus, a plausible heating duration estimate would be the elapsed time at maximum temperature necessary for the c o n t r o l l i n g reactions to approach completion. Thermal maturation of OM would s t a b i l i z e at t h i s point with respect to t h i s reaction temperature, and heating durat i o n is no longer a f a c t o r . We do not imply that thermodynamic e q u i l i b r i u m is

82

established in the system because OM thermal maturation reactions are i r r e v e r s i b l e (BLUMER, 1965; TISSOT & WELTE, 1984). The i r r e v e r s i b l e nature of thermal maturation reactions cause the rank of OM, and consequently Rm, to be set by the maximum temperature. Estimates of heating duration can emphasize elapsed time near Tmax because lower temperature reactions are not s i g n i f i c a n t in determining f i n a l OM rank. For example, OM thermal maturation, using the classic approximation that the reaction rate doubles f o r each 10° C increase, w i l l be approximately 1000 times f a s t e r at 150° C (approximate

cessation of l i q u i d hydrocarbon generation) than the same reaction at 50° C

(approximate i n i t i a t i o n

of hydrocarbon generation) (HUNT, 1979). The c o n t r i b u t i o n

to OM rank from low temperature reactions appears to be overwhelmed by reaction at higher temperature. These considerations indicate that the e f f e c t i v e heating-time of HOOD et a l . (1975) should be more l i m i t e d and not consider time elapsed during declining or stable temperature. A functional heating duration f o r OM thermal maturation is defined as the elapsed time while temperature increases w i t h i n 15° C of maximum temperature. Geologic time during temperature decline in the systems is not considered to increase Rm because OM thermal maturation is i r r e v e r s i b l e . Also, i f the temperature s t a b i l i z e s at near-maximum, t h i s time is not considered to increase e f f e c t i v e l y the thermal maturation of OM and is not included in the functional heating duration. The functional heating duration is s t i l l

a contrived estimate of reaction time.

There is no known method of determining a reaction time in sedimentary systems that is applicable to k i n e t i c equations. The control of OM rank by i r r e v e r s i b l e reactions at Tmax indicates that reaction time in sedimentary systems is the elapsed time at maximum temperature f o r the c o n t r o l l i n g reactions to approach completion. However, t h i s type of heating duration estimate is not e a s i l y measured in sedimentary systems where temperature has increased slowly during burial ( i n t h i s case, time at Tmax = O ) , or have poorly known burial h i s t o r i e s , making t h i s an impractical d e f i n i t i o n . The d e f i n i t i o n of the functional heating duration over a 15° C range near Tmax is a r e f l e c t i o n of imprecise geologic data. Regression analysis Proper designation of the dependent and independent variables for regression analysis of Tmax-Rm data and c a l i b r a t i o n of an empirical model are crucial (BARKER, 1984). In sedimentary basins, sample temperature is determined by the geothermal curve and depth w i t h i n the system. The independent variable (Tmax) is f i x e d by the i n v e s t i g a t o r by selecting a sample at some depth, and the dependent variable (Rm) is measured on OM concentrated from that sample. In t h i s case, estimates of Tmax from Rm must be

83

made from the regression of dependent variable (Y) and independent variable (X) (SNEDECOR & COCHRAN, 1967). Regardless of the s t a t i s t i c a l theory and i t s constraints, the selected regression curve should be a good model of the data and e f f e c t i v e l y predict i t s trend. Inspection of the data set suggests an exponential trend [Rm=a* exp(bTmax)] or i t s l i n e a r transform In (Rm)=a+b(Tma x) as used in t h i s paper. V i t r i n i t e reflectance in these sedimentary basins is highly dependent on maximum temperature (Fig. I ) . A least squares regression equation: In(Rm) = 0.0078Tmax - l . 2 computed from over 600 maximum temperature (in °C) and v i t r i n i t e

reflectance (in %)

data from these systems, indicates that about 70 percent (r 2=0.7) of the v a r i a b i l i t y in the v i t r i n i t e

reflectance data is explained by considering temperature alone.

a A

a A

a~

~X

~

xXxX

A A AXA,~ A A 9{,,z

×

LEGEND ~X

x

A

E 5O

i

1oo

i 150

diagenesis

a

Burial

X

Geothermal

i

200

i ~50

i

300

/ 350

T E M P E R A T U R E IN°C

Fig. I . General correlation of v i t r i n i t e reflectance with maximum temperature. Least squares regression analysis indicates that the variables are strongly related with correlation c o e f f i c i e n t (r 2 ) = 0 . 7 f o r a sample size (n)>600. The functional heating duration varies from 104 to 108 years in these systems. Sources of data used in t h i s figure are available upon request.

84

Heating Duration and Thermal Maturation The influence of heating duration is assessed using the LOPATIN (1971) model, which generates an estimate of OM thermal maturity by numerical i n t e g r a t i o n of heating duration in each 10° C i n t e r v a l over the system's b u r i a l h i s t o r y . The Lopatin model is modified in our analysis because the functional heating duration is computed using only the elapsed time in a small temperature i n t e r v a l near Tmax, making numerical i n t e g r a t i o n of time and temperature unnecessary. OM thermal maturity then equals elapsed-time m u l t i p l i e d by 2Tmax/10. I f temperature and heating duration both determine OM thermal maturation, then the range of Tmax across a l i n e of constant Rm would be due to heating duration. For OM thermal maturity to remain constand requires that the Tmax f a c t o r changes in an opposite manner to the heating duration f a c t o r . Tmax ranges over about 100°Cacross a given isoreflectance l i n e in our data (Fig. I) or over a f a c t o r of 1,000 in the Lopatin model. The functional heating duration would also range over a f a c t o r of 1,000. Using a functional heating duration of about I m.y. f o r b u r i a l diagenesis indicates a range up to I b.y. to compensate f o r the temperature f a c t o r . This is an unreasonably large e f f e c t f o r heating duration because geologic studies show that OM rank is only influenced by temperature a f t e r about 10 m.y. Geologic studies indicate that OM thermal maturation does s t a b i l i z e a f t e r about 106- 107 year (Table I) and in our model increased functional heating duration a f t e r s t a b i l i z a t i o n would produce n e g l i g i b l e increase in OM rank. Burial h i s t o r y recons t r u c t i o n of the systems we studied, and others with d i f f e r e n t OM types, shows that about 90 percent have been w i t h i n 15° C of maximum temperature f o r greater than 106 y r (Fig. 2). Thus, in most cases of burial diagenesis, heating duration at maxi-

~ii:

~

~

....

I

-

Fig. 2. Histogram of functional heating duration in selected sedimentary basins undergoing burial diagenesis. The functional heating duration in the nine qeothermal systems used in Fig. I (not plotted here) range from 104- 106 years. Data from BARKER (in press).

i~- ~

~i~ ~ - ~

~:,

I

~ 10'

10'

FUNCTIONAL

10' HEATING

DURATION

10 ~

10' (YR)

85

Table I . Published estimates o f the time required f o r the s t a b i l i z a t i o n o f OM t h e r mal maturation. I n s i g n i f i c a n t s t a b i l i z a t i o n time i n d i c a t e s t h a t OM thermal maturation was found t o be c o n t r o l l e d by temperature.

Reference

Stabilization Time (yr)

Seyer (1933) McNab et al. (1952) Vallentiyne (1964)

short 106 106

Tan (1965) Abelson (1967)

6.107 106

Abelson (1967

106

Notes

Petroleum generation above 200° C Petroleum generation. Complete amino acid decarboxylation at 100° C. Coal maturation. Methane generation from o i l shale pyrolysis would be complete in 106 yr at 115° C. Analysis of Los Angeles hydrocarbon generation data ( P h i l l i p i , 1965) sugests effective duration of the total heating) exposure is equivalent to roughly 2.10~ yr at 150° C. States that coal rank is a maximumrecording geothermometer. Coal and kerogen maturation. State that coal rank is a maximumrecording geothermometer. State that col rank is a maximumrecording geothermometer. "Geological time does not l i m i t the coalification process." Coal maturation. Coal maturation. Heating time not important in organic maturation. Coal maturation. Kerogen pyrolysis. Equilibrium reached in petroleum generation. The time factor for petroleum generation is of limited significance over geologic time and probably critical only for a short period. "A minimum temperature is needed to start any transformation of organic material with a particular activation energy. At such a temperature a certain length of time is necessary and sufficient to comple the reaction .." "The role of time is probably also exaggerated .. for the methods of Bostick and Lopatin." Influence of time is minor when compared to the effect of temperature in hydrocarbon generation from kerogen. Petroleum generation from kerogen. "Temperature remains c r i t i c a l : a source shale with Ro=0.8%can remain at that rank for many million years and never generate a drop of o i l . " Equilibrium in organic maturation no__tt established in this heating time. Organic metamorphism in hydrothermal bombs at 350° C. Coal maturation. Petroleum generation above 130° C. Kerogen maturation in liquid-dominated geothermal systems. Kerogen maturation.

l

Brooks (1970)

Insignificant (Waples,1984) Neruchev & Parparova (1972) I0o Lopatin & Bostick ( 1 9 7 3 ) Insignificant (Waples,1984) Bartenstein & TeichmUller (1974) Insignifcant (Waples,1984) Nagornyi & Nagornyi ( 1 9 7 4 ) Insignificant

Juntgen & Klein (1975) Hacquebart (1975) Ammosov et al. (1975,1977) Demaison (1975) Harwood (1977) Cornelius (1975) McTavish (1978)

107 107 to 108 106 108 Short about 107 Insignificant

Veto (1980)

finite

Veto (1980)

Time effect exaggerated

Barker, Colin (1979)

about 5.107

Sajgo (1979) Wright (1980)

106 Insignificant

Teichm~ller & TeichmUller (1981)

3.106

Price et al. (1981)

Short

Suggate (1982) Gretener & Curtis ( 1 9 8 2 ) Barker (1983) Price (1983)

106 Insignificant I0~ 106

86

mum temperature has been s u f f i c i e n t f o r the e o n t r o l l i n g reactions to approach comp l e t i o n and f o r the OM to s t a b i l i z e e f f e c t i v e l y with respect to temperature. Heating duration would have no f u r t h e r e f f e c t . Our i n t e r p r e t a t i o n is that the so-called "influence of time" is invoked to explain differences in temperature-Rm curves between comparable sedimentary systems - some now at T and others that have cooled. For instance, the Munsterland-1 well used max by LOPATIN (1971) to c a l i b r a t e his OM thermal maturation model has declined s i g n i f i c a n t l y from maximum temperature (Fig. 3) (PRICE, 1983). Temperature decreases causes Rm data to be s h i f t e d to the l e f t from t h e i r o r i g i n a l p o s i t i o n at Tmax (Fig. 3). This s h i f t is s i m i l a r to the e f f e c t increased heating duration would have on OM t h e r mal maturation. The remaining v a r i a t i o n in the Tmax and Rm data could be accounted f o r by considering functional heating duration but other s i g n i f i c a n t types of data v a r i a b i l i t y minimize the c o n t r i b u t i o n from t h i s source.

i [~

~ 8

[][] , 50

[]

[]

A

X

a

[]

x

SHIFT OF Rra DATA DUE TO COOLLNG

p .......

Sy .................

temperature

i 100

, 150 TEMPERATURE

i 200

~ 250

, 300

i 350

IN°C

Fig. 3. Scatter p l o t of Rm and temperature data including those not at maximum temperature. Rm and temperature data from systems in which temperature has declined are s h i f t e d to the l e f t from the systems now at Tmax shown in Figure 3. The Munsterland-~ well shown was used by LOPATIN (1971) to c a l i b r a t e his time-temperature index OM maturation model. Sources of data used in t h i s f i g u r e are available upon request.

87 Errors in V i t r i n i t e Vitrinite

Reflectance and Temperature Data

Reflectance

Operator bias in selecting v i t r i n i t e

f o r measurement is p o t e n t i a l l y the greatest

source of e r r o r during reflectance analysis. Bias can occur in reflectance measurements when m u l t i p l e v i t r i n i t e

populations are present in the sample r e q u i r i n g the

operator to select a single population f o r analysis. These mixed populations arise p r i m a r i l y from the admixture of recycled, p r e v i o u s l y - a l t e r e d v i t r i n i t e

with f i r s t

cycle OM during sedimentation. To minimize t h i s problem and compile a consistant data set, one microscopist should make the Rm determinations (although a single oper a t o r could s t i l l

c o n s i s t e n t l y s h i f t the Rm measurements). However, because our data

was in part derived from the published work of d i f f e r e n t microscopists, variable operator-bias is a s i g n i f i c a n t f a c t o r in causing data scatter. Rm measurements by t h i r t y microscopists on the s p l i t s of 19 d i f f e r e n t samples show that the range of Rm measurement can be up to +/- 0.4 % Rm in low rank OM (unpublished report, I n t e r national Commission on Coal Petrology, see BOSTICK, 1979). This wide Rm range arises from operato~ bias and differences between laboratories in processing, p o l i s h i n g , and photometer c a l i b r a t i o n . Even under an optimal scenario to minimize operator and laboratory bias - microscopists measuring the same uniformly prepared sample s u i t e , using the same microphotometric system, and s i m i l a r v i t r i n i t e Rm determinations s t i l l

selection procedures -

vary by up to -/+ 0.2 % at moderate to high rank (BARKER,

1983). Mixing of d r i l l

cuttings in the borehole can also introduce m u l t i p l e v i t r i n i t e

l a t i o n s in a sample. This is a major physical problem in v i t r i n i t e ysis

because d r i l l

popu-

reflectance anal-

c u t t i n g s are the only rocks available over a s i g n i f i c a n t depth

i n t e r v a l in most boreholes. Up-hole sloughing of rock from the sides of the borehole introduces a less mature v i t r i n i t e f o r the f i r s t - c y c l e

population, that i f abundant, could be mistaken

(lowest maturity) population indigenous to that sample. Up-hole

sloughing of rock can reduce the apparent Rm to the level of the l a s t casing depth, which in the worst case would be that of the near-surface rock. Rock-chip mixing in the d r i l l i n g

f l u i d and infrequent sample c o l l e c t i o n make d r i l l

c u t t i n g s a composite

sample (at best) of the d r i l l e d i n t e r v a l between sample points even when the sample depth is corrected f o r t r a n s i t time from d r i l l - b i t

to surface. This introduces uncer-

t a i n t y in where to determine the sample depth, thus causing e r r o r in T and Rm max data. I t also appears possible to severely a l t e r the rock chips by heat and pressure from the d r i l l

b i t (TAYLOR, 1983), though these are probably i d e n t i f i a b l e in a d r i l l -

c u t t i n g s sample. The less altered rock, however, may contain an a r t i f i c i a l l y

matured

OM, without v i s i b l y changing the petrographic character, making them undetectable.

88 Another s i g n i f i c a n t source of e r r o r in v i t r i n i t e

reflectance analysis is Rm suppres-

sion. Studies have shown that Rm can be suppressed up to several tenths of a percent by maceral association and differences in e a r l y diagenetic h i s t o r y (see review by PRICE & BARKER, 1985). Variation in the depositional and (or) diagenetic environment produces a hydrogen-rich OM that has lower Rm values than those expected from the thermal h i s t o r y of the sediment. This problem was minimized in t h i s study by considering only those systems rich in humic (type I I I ) Vitrinite

OM.

also becomes anisotropic or b i r e f l e c t a n t of moderate to high rank. The

physical e f f e c t is that reflectance becomes dependent on the microscope stage o r i e n t a t i o n . Bireflectance is i n s i g n i f i c a n t below about 1.2 % Rm, but the difference between the maximum and minimum reflectance increases to almost 25 % at 3.0 % Rm (STACH et a l . ,

1982). Bireflectance in high rank OM increases the range of the reflectance

histogram and often produces polymodal d i s t r i b u t i o n s , making the mean value less representative. Further, there is a convergence of the reflectance f o r v i t r i n i t e i n e r t i n i t e at about 2.5 % Rm which makes these two macerals d i f f i c u l t

and

to d i s t i n g u i s h .

Temperature Most Tmax data f o r systems now at peak temperature are from uncorrected BHTs t y p i c a l l y taken on a single logging tool run. A m i n o r i t y of the remaining temperature data are from corrected BHT measurements (the correction method sometimes unspecif i e d ) . A geothermal gradient is calculated from BHT by i n t e r p o l a t i n g between mean annual surface temperature and BHT. This l i n e a r approximation can be h i g h l y inaccurate because the temperature p r o f i l e can change with l i t h o l o g y (thermal conductivity),

subsurface f l u i d flow, etc.(DRURY et a l . ,

1984). Accurate determination of

present-day formation temperature also requires thermal e q u i l i b r a t i o n of the borehole before measurements. The extended borehole s h u t - i n time required to establish thermal e q u i l i b r i u m u s u a l l y means that borehole temperature is measured soon a f t e r drilling

is completed. D r i l l i n g disrupts the undisturbed temperature p r o f i l e because

cool f l u i d s are pumped down the borehole and are warmed up on ascent to the surface. The deep portions of the borehole are cooled and the shallow portions are heated during d r i l l i n g . drilling

The shallow portions of the borehole are exposed to more

f l u i d and cooling there is greater than at depth. BHT data are not usually

confirmed as e q u i l i b r i u m reservoir temperatures by repeated measurements over a sign i f i c a n t time i n t e r v a l making correction necessary. Our experience is t h a t the data necessary f o r c a l c u l a t i n g a correction of the log temperature to e q u i l i b r i u m format i o n temperature are not included in the borehole h i s t o r y reports. Further, attempts at correcting the BHT measurement to e q u i l i b r i u m conditions, although necessary, are often unsuccessful because d r i l l i n g

and measurement conditions vary widely and usu-

a l l y cannot be corrected by some uniform procedure (DRURY, 1984). Temperature cor-

89 rections in the order of 2 0 - 3 0 ° C are t y p i c a l f o r BHT data (HOOD et a l . ,

1975)

making t h i s e r r o r at least as s i g n i f i c a n t as that from determining Tmax. Tmax is a d i f f i c u l t

determination to make accurate, e s p e c i a l l y in systems that have

cooled. The burial h i s t o r y reconstruction method (see discussion by WAPLES, 1981) is widely used to estimate maximum burial depth (and temperature) during diagenesis. Temperature as a function of geologic time is calculated by using the e x i s t i n g or paleo-geothermal gradient applied to the depth-time curve constructed from the b u r i a l h i s t o r y . Lacking geologic evidence to d e t a i l thermal changes in the system, the c a l culated geothermal gradient is u s u a l l y assumed to remain constant through time w i t h out regard to heat-flow change, diagenesis, and l i t h o l o g y . Thus, the detailed geol o g i c analysis is reduced to time-temperature data by s i m p l i f i c a t i o n of the changes that can occur in the geothermal gradient. Unfortunately, the geologic record often w i l l not permit a more sophisticated approach and a rigorous d e f i n i t i o n of the timetemperature h i s t o r y . Tmax in systems t h a t have cooled may also be determined by using maximum geothermometers. The l i m i t a t i o n of t h i s method is that the T event must be recorded in the max rocks. Thermal events can be d i f f i c u l t to record in sedimentary systems because of the slowness of e q u i l i b r a t i o n reactions at low temperature. The short-term nature of some thermal events may also preclude them from leaving an imprint on the rock. Both of these conditions produce s i g n i f i c a n t changes in the OM rank (STACH et a l . ,

1982).

Discussion Three l i n e s of evidence support our regression analysis i n d i c a t i n g t h a t maximum temperature alone determines Rm and heating duration has l i t t l e

continuing influence on

OM thermal maturation: ( I ) tests of temperature-time-rank models in hypothetical or well-known sedimentary systems; (2) comparison of OM rank to mineral assemblages res u l t i n g from e q u i l i b r i u m reactions during metamorphism; and (3) a p p l i c a t i o n of t h i s empirical Tmax - Rm geothermometer to sedimentary systems where temperature is well known. Tests of Temperature-Time-Rank Models VETU (1980) tested published temperature-time models of OM thermal maturation in 45 sedimentary basins and found "the role of time is probably also exaggerated . . . f o r the methods of BOSTICK and LOPATIN." He considered that "a minimum temperature is needed to s t a r t any transformation or organic material with a p a r t i c u l a r a c t i v a t i o n energy. At such a temperature a c e r t a i n length of time is necessary and s u f f i c i e n t to complete the r e a c t i o n . . . " . WRIGHT (1980) in a s i m i l a r t e s t using hypothetical b u r i a l h i s t o r i e s concluded that "temperature remains c r i t i c a l :

a source shale with

go R =0.8 % can remain at that rank f o r many m i l l i o n s of years and never generate a 0

drop of o i l . "

Several other studies show that heating duration is not of continuing

importance in OM thermal maturation (Table I ) . Temperature-time models are poor predictors of heating duration in geothermal systems when compared to d i r e c t estimates of thermal event duration. BARKER (1979) found the c o r r e l a t i o n ( r 2 =0.8) between Rm and logged temperature in the central portion of the Cerro Prieto system, together with consistent temperature estimates from f l u i d i n c l u sion and oxygen isotope geothermometry, indicates that these rocks are now at maximum temperature. Application of t h i s data to KARWEIL's (1956) OM thermal maturation model predicts a heating duration of 5 m.y. f o r the Cerro Prieto system - a poor f i t

to a

heating duration of about 10,000 years indicated by f i s s i o n - t r a c k annealing studies (SANFORD, 1981), and the upper l i m i t of heating duration implied by reservoir rock age of about I m.y. Metamorphic Mineral Assemblages and Rank S t a b i l i t y of mineral assemblages r e s u l t i n g from hydrothermal metamorphism are temperature

dependent. The consistent occurrence of authigenic minerals at s i m i l a r temper-

atures in many rock types of d i f f e r e n t ages and d i f f e r e n t geothermal systems i n d i tes that phase changes r e s u l t from e q u i l i b r i u m reactions due to thermal metamorphism (ZEN & THOMPSON, 1974; BROWNE, 1978; WEAVER, 1979). The formation of e q u i l i b r i u m assemblages suggests that the reaction k i n e t i c s f o r these reactions are r e l a t i v e l y rapid and that heating duration should not be a s i g n i f i c a n t f a c t o r in hydrothermal metamorphism. BARKER et a l . ( i n press) found that certain Rm levels correspond with hydrothermal zones in several boreholes in the Cerro Prieto geothermal system. These mineral zones occur over a wide temperature range, but the range of the thermal stability

f i e l d and Rm is s i m i l a r between boreholes. This r e l a t i o n s h i p indicates that

temperature controls Rm in hydrothermal systems. KISCH (1969; updated by ZEN & THOMPSON, 1974) in a review of the available l i t e r a t u r e , found mineral zones associated with b u r i a l metamorphism correlated with coal rank in samples from various geographic l o c a l i t i e s and geologic times. T e r t i a r y sediments in many areas of the United States show clear c o r r e l a t i o n between pressuretemperature (P-T) dependent mineral metamorphism and kerogen metamorphism (VAN DE KAMP, 1976). S i m i l a r l y , STALDER (1979) has shown that clay mineral and z e o l i t e mineral assemblages correlate with OM rank in the Eocene-Oligocene Taveyannaz Sandstone (Europe) and i t s equivalents. SHIMOYAMA& IIJIMA (1976) showed that rank of Japanese T e r t i a r y coals correlate well with z e o l i t e zonation, and concluded, rank increase is e s s e n t i a l l y temperature dependent. LANDIS (1971) found that progressive g r a p h i t i z a t i o n is related to metamorphic grade as defined by mineral assemblages, and concluded that kerogen g r a p h i t i z a t i o n " . . .

is p r i m a r i l y dependent upon metamorphic temper-

91

ature; pressure and v a r i a t i o n in s t a r t i n g material presumably c o n s i t i t u t e secondary c o n t r o l s . " HOWER & DAVIS (1981) converted published coal rank and associated mineral assemblage data from various geologic times to consistent set of P-T, and v i t r i n i t e reflectance values. HOWER & DAVIS noted that considerable uncertainty was involved both in rank conversion, and in placing the rank data at a single p o i n t , whereas the metamorphic mineral assemblage d e l i m i t s on the P-T f i e l d . However, t h e i r Rm data can be plotted along simple, contourable patterns on the P-T f i e l d ,

i n d i c a t i n g that time

is a n e g l i g i b l e f a c t o r and can be ignored in modeling OM metamorphism. In summary, metamorphic mineral assemblages r e s u l t i n g from temperature and P-T dependent e q u i l i b r i u m reactions, in diverse metasedimentary systems with a wide range of heating durations, correlate with a s p e c i f i c coal rank.

This suggests that the con-

t r o l s on metamorphic mineral assemblages appear to control Rm. We i n t e r p r e t these data to indicate that Rm is p r i m a r i l y c o n t r o l l e d by temperature because s t a t i c pressure apparently has l i t t l e Applied Geothermometry:

influence on Rm (STACH et a l . ,

1982).

Case Studies

POLLASTRO & BARKER ( i n press) computed paleotemperatures from Rm, i l l i t e / s m e c t i t e r a t i o s in mixed-layer c l a y , and f l u i d i n c l u s i o n temperatures in the Wagon Wheel no. I borehole, Green River Basin, Wyoming (Fig. 4). Using our empirical c a l i b r a t i o n , they computed maximum temperature which is close to the temperature calculated from clay mineral data at about 190 ° C. Uncorrected borehole temperature at t h i s depth is now about 130° C. The slope of the present day borehole temperature measurements and the slope of temperatures interpreted from Rm are about 25 ° C/km. Fluid i n c l u s i o n homogenization temperatures in q u a r t z - f i l l e d fractures formed by l a t e T e r t i a r y deformat i o n are equal to present-day temperature. An Rm surface intercept of 0.33 % indicates about 1700 m of section has been removed (DOW, 1977). Decreasing the b u r i a l depth by 1700 m in a geothermal gradient of 25 ° C/km would decrease borehole temperatures by about 40 ° C, making i t consistent with the other geothermometric data. These data were interpreted to indicate that temperature has decreased by about 40 ° C, by u p l i f t and erosion rather than decline in geothermal gradient, before the fractures were TEMPERATURE IN~C SO

• ">kB, therefore B = I .

In the

studied isomerizations the rate constants are nearly equal (kA~k B) and ~ subsequentl y is close to 2. There are numerous ways to calculate the e q u i l i b r i u m constant. One of them is based on laboretory heating experiments. PUSTIL'NIKOVA et a l . (1980) applied pressurized heating experiments to study the isomerization of cholestane. The conditions have been referred to e a r l i e r . We must bear in mind the d i f f i c u l t i e s of t h i s method. The f i r s t

question is whether the mixture is r e a l l y at e q u i l i b r i u m .

To reach e q u i l i b r i u m is a long process, and in geological s i t u a t i o n s i t may require some m i l l i o n years. To speed up the process, PUSTIL'NIKOVA and her co-workers have used a temperature about 150-200 ° C higher than natural conditions. In r e p l i c a t e simulations, the e q u i l i b r i u m constant f o r 20S/20R, 5~(H), 14~(H), 17~(H) cholestanes has been found to be 0.82 and 0.98. Considering the dependence of the e q u i l i b r i u m constant on temperature we must be very careful i f we apply the values determined at high temperatures to natural conditions. From a k i n e t i c a l viewpoint, the presence of pressurized H2 and P t - c a t a l y s t is c r u c i a l , because i t can modify the mechanism of the reaction and therefore the rate parameters may change d r a s t i c a l l y . Van GRASS et a l . (1982) have computed the thermodynamic s t a b i l i t i e s

of 13 cholestane

isomers which were found in Petrov's isomerate (SEIFERT & MOLDOWAN, 1979) using a molecular model based on empirical energy functions. The e q u i l i b r i u m constant can be derived from the obtained composition of e q u i l i b r i u m mixtures of isomers. In t h i s case the inverse of the above-mentioned e q u i l i b r i u m constant (~) was found to be 0.817 at 298 ° K and 0.845 at 573 ° K. The r e l i a b i l i t y

of such model computations

depends on the r e a l i t y of the established models and on the consideration of a l l the possible i n t e r a c t i o n s . We have chosen a t h i r d way f o r the c a l c u l a t i o n of e q u i l i b r i u m constants, namely, s t a r t i n g from measurements on geological samples. The concentration r a t i o in geolo-

136 g i c a l samples o f the two isomers studied can be considered constant f o r a d u r a t i o n and over a c e r t a i n temperature. For example, in the case o f sterane i s o m e r i z a t i o n in a sample whose age is 7.82 Ma and has a temperature o f 410 ° K t h i s r a t i o has reached the e q u i l i b r i u m v a l u e , whereas in o l d e r samples the r a t i o s c a t t e r s i r r e g u l a r l y . The scale of f l u c t u a t i o n in the d e c i s i v e p a r t o f samples does not exceed the measurement e r r o r s (MACKENZIE, 1980; RULLKUTTER et a l . ,

1984) and t h i s means t h a t

the x, y and z values (Table I and equation 3) can be determined to an accuracy o f 0.04. The r e c i p r o c a l e q u i l i b r i u m constant was c a l c u l a t e d from the average of conc e n t r a t i o n values and, f o r example, i t was found 0.725 f o r the above mentioned sterane i s o m e r i z a t i o n , so in our case i t was found less than t h a t in van GRASS et a l . (1982) (0.817 at 298 ° K and 0.845 at 573 ° K). The i n t e g r a t i o n o f the r a t e - l a w equation sets some problems, since the r a t e constant behind the i n t e g r a l symbol has a temperature dependence, which f o l l o w s the c l a s s i c equation proposed by Arrhenius in 1889: (11)

k = A exp [ - ~R--~-]

where A is the pre-exponential factor~ ~H* is the a c t i v a t i o n e n t h a l p y , i . e . the height o f the p o t e n t i a l b a r r i e r between the reactants and products o r , in o t h e r words, the a d d i t i o n a l energy of the a c t i v a t e d complex with respect to the i n i t i a l s t a t e o f r e a c t a n t s ; R is the universal gas constant and T is the absolute temperature. According to more d e t a i l e d molecular t h e o r i e s the p r e - e x p o n e n t i a l f a c t o r has a temperature dependence which is however much less than t h a t o f the r a t e constant: (12)

A = X.

exp[ ~S T ] = A' T

whereX is the transmission c o e f f i c i e n t or the p r o p o r t i o n a l i t y constant, X ~ I but i t s value is u s u a l l y I - i t s d e v i a t i o n from one is p r o p o r t i o n a l to the reduced p r o b a b i l i t y of reformation of the reactants from the a c t i v a t e d complex - k is the Boltzmann constant; h is Planck's constant - the r a t i o

is termed the fundamental

frequency and i t s value at room temperature (T=300 ° K) is 6.1012 s - I , not f a r from the value of c o l l i s i o n entropy of r e a c t i o n and i t

frequency (

1013 s - I )

; ~S

thus i t is

is the a c t i v a t i o n

is an i n d i c a t o r o f the c o n f i g u r a t i o n of the a c t i v a t e d

complex. ~S * is u s u a l l y negative because A < -~h - . T This decrease in entropy is a consequence of the loss of t r a n s l a t i o n a l and r o t a t i o n a l freedom when reactants are combined to form the a c t i v a t e d complex.

137 In laboratory simulation experiments we can ensure the isothermal conditions - both in space and time - therefore we can eliminate the thermal dependence of the rate constant during the solution of the rate law equation. In such a case the integrat i o n becomes a simple m u l t i p l i c a t i o n . Unfortunately, the v a r i a t i o n of temperature with time and place is c h a r a c t e r i s t i c of geochemical processes. As the reaction rate is temperature-dependent through the exponential dependence of the rate constant on temperature, in cases of the same duration the d i f f e r e n t temperatures can produce considerably d i s s i m i l a r conversions. This feature of reactions can be e x p l o i t e d to elucidate the thermal h i s t o r y within a subsiding basin with help of the analyses of reactants and products of biomarker reactions. I f we know the rate parameters (A, AH and ~S*) of a reaction we can calculate the extent of conversion in the case of any thermal h i s t o r y - T= f(h,T) - or inversely, the thermal his t o r y can be reconstructed by the extent of conversion in a series of samples. For t h i s reason we attached great importance to determine the genuine rate parameters of the studied reactions. Knowing the temperature and age data of the samples (Table I ) , since the present temperature is maximum in the Pannonian basin and the age of the samples in the quickly subsiding basin can be determined accurately, and having the conversion data of the compounds of the three reactions (also Table I ) , we can give the rate parameters of the isomerizations of the sterane and the hopane, and the aromatizations of the monoaromatic steroid. These parameters may also be e x p l o i t a b l e to describe the thermal h i s t o r y of other basins. Some authors t r i e d to u t i l i z e the fact that the conversion depends on temperature e x p o n e n t i a l l y and on time l i n e a r l y during the solution of the integral equation of the rate law. For example, HOODet a l . (1975) suggested that the e f f e c t i v e heating time ( ~ e f f ) , the time which the sample spent w i t h i n 15° C of i t s maximum temperature, could be used to characterize the transformation of the v i t r i n i t e

instead of the

whole age of the sample. Others (NERUCHEV & PARPAROVA, 1972; SUGGATE, 1982) suggested that about one m i l l i o n years are enough f o r organic geochemical reactions (e.g. c o a l i f i c a t i o n ) to reach e q u i l i b r i u m . These models (e.g. HOODet a l . ,

1975) using

the present temperature and the v a r i a t i o n s of the geothermal gradient with time can be very elusive i f they are applied. The organic geochemical reactions go at any case in a given d i r e c t i o n from the viewpoint of f i n a l products. This may be some sort of e q u i l i b r i u m mixture of the reactants and the products, or, in the case of an i r r e v e r s i b l e reaction, the f i n a l products. On the other hand, the rate of each chemical reaction p r a c t i c a l l y increases with r i s i n g temperature. So, on the basis of the above i t is obvious that a conversion having proceeded at a higher temperature does not go back a f t e r a substantial temperature decrease. S t r i c t l y speaking, however, in the case of r e v e r s i b l e reactions such a reverse process can be imagined i f the lower temperature is favorable to the s t a r t i n g compounds in consequence of temperature dependence. But in the case of the isomerizations chosen, the temperature

138 dependence is so l i t t l e

that i t s influence can be neglected knowing the temperature

v a r i a t i o n of d i f f e r e n t basins. When the case is simpler, e.g. a constant rate of subsidence with a constant geothermal gradient or with a monotonic increase of heat flow density through a longer period, the influences of the temperature and the age on the rate parameters are c o n t r o l l e d by the extent of the a c t i v a t i o n energy of the reaction. Mathematically, the operation is correct only i f we take into consideration the temperature dependence of the rate constant and the time dependence of temperature during the course of i n t e g r a t i o n (from equations 9 and 12):

(13)

-

T

In(l-Ba)

-iT- exp [ T

] exp [ - ~ T - ] dr

We can generally describe the temperature of a subsiding and transforming sediment sample as a function o f ' t i m e and depth during the course of basin formation as (14)

T = f(~)

and the inverse function: (15)

T = g(T)

I n s e r t i n g these functions into the reaction integral we get: T

(16a)

- ~ In(1-Ba) =

exp[ --~--]

f(T) e x p [ ~ ] d ~

or T

(16b)

- ~- In(1-Ba) = ~-~ e x p [ T~S* ]

f

~H T g'(T) e x p [ - ~ T - ] dT

TO where g' is

g(T) T

Henceforth the above means of solving the reaction integral w i l l be called the method of absolute time (or in abbreviated form: MAT). This method is d e f i n i t e l y the best to determine the rate parameters (A, ~H* andeS ) from the measured extents of conversions and from knowing the temperature f u n c t i o n , T = f ( T ) ,

or backwards to

reconstruct the temperature h i s t o r y on the basis of the known rate parameters and the measured extents of conversions. However, the a p p l i c a t i o n of the method has some difficulties:

139

i)

The p r i m i t i v e f u n c t i o n cannot u s u a l l y be w r i t t e n in a closed from (or in an easy t r a c t a b l e form) in the case of the T= f ( T ) formula; the thermal h i s t o r y of the Pannonian basin was regarded as a simple case: (17)

T = To + a~

and the p r i m i t i v e f u n c t i o n may be given as:

(18)

fTg'(T)

-AH a-1 exp[ --RT- ] dT =

_ (AH ~ ,T,/*

+ (T)

1 ,AH ,2 ~T.t~ T ) In T+

CO

Z(-l)n+l

gt

~ 1

(T)

) \

T-n

n=l

U n f o r t u n a t e l y t h i s series converges r a t h e r slowly so i t s a p p l i c a t i o n is not convenient. ii)

There is no exact expressions f o r ~H

in an e x p l i c i t

form from the r e a c t i o n

i n t e g r a l , thus the knowledge of i t s approximate value is d e s i r a b l e . To solve the f i r s t

problem we can use numerical i n t e g r a t i o n methods, e . g . Simpson's

approximation or Romberg i n t e g r a t i o n which work r e l a t i v e l y

q u i c k l y and with an

accuracy according to choice. We can solve the second problem in two steps. F i r s t we determine AH in some way, and then we i n s e r t the approximate ~H

approximately

into the r e a c t i o n i n t e g r a l to

compute the value of AS . Having two parameters we can c a l c u l a t e the r a t e constants belonging to present temperature of the samples. From these values we c a l c u l a t e A H again and u t i l i z i n g

i t s new value we get a new set o f A S . We proceed with the

i t e r a t i o n up to o b t a i n i n g the l e a s t dispersion of the series of AS , The enthalpy of a c t i v a t i o n is o b t a i n a b l e by the d i f f e r e n t i a l

method in the simplest

way. The method was based on the f o l l o w i n g t r a i n of thought. As we have shown ( i n equations 12 and 13): T

(19)

-

l n ( 1 - g ~ ) = A'

T exp [ - - ~ - - ] dT 0

and c a r r y i n g out d i f f e r e n t i a t i o n

(20)

1 1-~B

on both sides of the equation:

d~ A'T exp [ -AH dr = --~T- ] = k

140 S u b s t i t u t i n g the f i n i t e

Asf o r ~ w ed~ ~

have determined AH* and A' (and thereby AS*)

approximately. The accuracy of determination f o r parameters was c o n t r o l l e d by the number of the a v a i l a b l e survey data in t h i s case apart from the accuracy of measurement. We can assess the r a t e parameters with the help of the e f f e c t i v e heating time (EHT) method. Applying t h i s method the r e a c t i o n i n t e g r a l can be given as (21)

- ½ ln(1-~)

= kAT

w h e r e a t is the e f f e c t i v e heating time and subsequently i t may be w r i t t e n (from equation 19) as (22)

- ~ In(1-g~) = A'T°exp[-~--~o ] ~

where T° is the present temperature. So we can r e w r i t e as (23)

In [ - 1-~-In ( I - ~ ) ] BTo

= I n A ' + InA~

AH -T"

I ~T

= x

This expression determines the equation of a s t r a i g h t l i n e in a coordinate system of xvs.

.

AH

can be obtained from the slope o f l i n e and the i n t e r c e p t y i e l d s A'

which then gives AS . The e r r o r of the EHT method can be assessed as f o l l o w s : .

(24)

Z-AT

i e _ AH ~ -~T-dT =

.

e

~H RTd~

0

+

j e _ AH* -RT-dT = I T-~T

on the basis of the mean value theorem of i n t e g r a l c a l c u l u s : (25)

I = (~-AT)e

~H ~H I~TT+ &T e

where T' is the temperature in the time i n t e r v a l in the i n t e r v a l (z-AT,T). The f i r s t

(0, T-AT), and T" is the temperature

term on the r i g h t - h a n d side is n e a r l y the same

as the expression on the r i g h t - h a n d side f o r the method o f e f f e c t i v e heating time (equation 22). The only d i f f e r e n c e is t h a t in t h i s case the present temperature has been replaced by T" which is d i f f e r e n t . not too long, then

I f the i n t e r v a l of e f f e c t i v e heating time is

141 (26)

T" = T° - f ( T ) A T

where using

0~0 OJl I.i~ 0

kl~

CO

~"

OJ

r~

%

-~

I.)

"E o

-~

%

~g

o

~

F-- e--~ O -~

~

~

~

o

,~

o~

~

o

~

N .$._ o

o o o

~

~

~

~

o

~

~

o~o 0

lo

o ¢'~

I I

0

0

¢0

I

I

(1J

4 -1

I/1

"~

•

t~

,~

t~

~,

>~

0

0

U'/

U~

~,

>~

O~

O~

O~

CI.

ZI.

02

0

0

0

CI.

~

U.--

'4--

t_)

t.)

0

(11

0

¢~

(1)

--~

0

0

0

0

(3-}

.$-

--~

(1~

--~

161 One type o f basin is the Pannonian-type basin in which r e c e n t l y the heating r a t e (y) is roughly 16° C Ma- I , trophe

thus a proved to be 8 ° C Ma-I before the date of the catas-

('Co).

The other type ( f o r s i m p l i c i t y l e t us term i t a basin with an extremely low r a t e of h e a t i n g ) bears r e c e n t l y a h e a t i n g r a t e of 1.6 ° C Ma-1. Thus, b e f o r e the t o, t h e r a t e of heating was only 0.8 ° C Ma-I in t h i s basin. With these two basins the ranges produced by nature are covered. C a l c u l a t i o n s were c a r r i e d out according to the r e l a t i o n s h i p (6). Results, i . e . the d i f f e r e n c e s between the t r u e conversion and the conversion c a l c u l a t e d a f t e r the recent rates of heating are p l o t t e d as a f u n c t i o n of the age of samples. First,

*

the r e a c t i o n of the sterane i s o m e r i z a t i o n type reactions (AH =90 kJmol

A=4.8.108-1013 Ma) are demonstrated in Fig. 2 f o r thermal h i s t o r i c a l of d i f f e r e n t

;

catastrophes

~o dates, in the Pannonian-type basin. Both in t h i s f i g u r e as well as

in the subsequent ones the l i m i t of r e l i a b i l i t y

-I

value of m=0.04 is shown which i n d i c a t e s the l i m i t

of the conversion measurements.

0.24

A H t = 90.0 kJ mot-1

022 • o=

To = 250 ° K

0.20- "~

i ~. /

: S o C M a -1

/

T : 16°C Ha'i

0.18-

o.16.

o

"~o: 1Ha

i~,

I

o lz o=

" '

L

.

005

o.o~.

~ o

~:,o".~'I

/ t!~,

/

~' /,

/

', \

,! 0.04 0.0 Z"

A = lO~3Ha-1

i

/

I ." II / ,, ! -/'

/" ,."

5

~

!

i

l

i

A :1011Md 1

0.14. o

O.10

#

i

/

l "

.;

precision of measurement

10

15 Time [ H a l

Fig. 2. Deviation of the c a l c u l a t e d conversions from the measured ones in the case of sterane i s o m e r i z a t i o n in a Pannonian-type basin i f the heat f l o w doubled one m i l l i o n years ago. Fig. 3 shows how the a r o m a t i z a t i o n type reactions of highest a c t i v a t i o n enthalpy (AH*= 150 kJmol - I ) react to the sudden doubling of the r a t e o f heating in the Pannon i a n - t y p e basin. Fig. 4 shows the behavior of a r e a c t i o n of r a t h e r low a c t i v a t i o n enthalpy (AH = 10 kJmo1-1) in a Pannonian-type b a s i n .

Such a r e a c t i o n ,

l y i n g on the boundary

162

A = 8 ,'Zgx ~014 M o I 0"141 m ~

o.~34g

& H i = 1S0kJ i'nol "1

/~

012~, G

To = ZS0°K

~

o . ~ >=

"'~7"h" .'CC,x'~,X,~_. ,. ~.," ', ", ' ~

__..~,-,b~ ,,, ,,';;,,',,'~',X',' 2 % reflectance values, as estimated from accepting the Lopatin-Waples method up to 250 ° C, 3 = Area characterized by 2 %>Ro > 1 . 3 % reflectances, 4=Area characterized by 1.3 %>Ro>0.65 % reflectances, 5= Position of the Ro =0.65 % r e f l e c t i v i t y horizon in sediments not influenced by the magmatic heat.

182 ago when the basin was 24 m i l l i o n years old and 2800 m deep. I t is supposed that no more subsidence and sedimentation occurred a f t e r the volcanic event. 1000° C o r i g i nal temperature was taken f o r the igneous m a t e r i a l . The results are shown in Fig. 6, which gives the present p o s i t i o n of l i n e s characterized by 0.65 %, 1.3 % and 2 % vitrinite

reflectances. I t can be seen that nothing happened f a r away from the v o l -

cano; notable maturity increase is confined to a volume below the thick volcanic strat a . The zone of influence is not very large and, in f a c t , our r e s u l t should be considered as a maximumestimate, because in r e a l i t y : i)

formation of a stratovolcano occurs in many cycles during severel hundredthousand years, and

ii)

lava flows are alternated by explosion of t u f f s , which have a much lesser temperature.

Anyhow, i t can be concluded that there is no real "telemagmatic" e f f e c t , the zone of influence of an igneous body is about the same volume as the body i t s e l f . The second model we have calculated is a large dyke which broadens downwards(Fig. 7). I t is a two-dimensional body with 1000° C o r i g i n a l temperature, which intruded the same sedimentary basin 17 m i l l i o n years ago. I t can be seen that the zone of in fluence above the top of the dyke is rather small. I t is because the magmatic heat

x

0

I ,

I

2 ,

I

3 ~

I

Z. ,

I

5 ,

I

{km) 6

,

L

7 ,

L

8

"0"65°I°

3

9

10

,

- ----O

65%

./ i

I

~

-t3 %

Fig. 7. Increase of maturity of organic matter next to the top of a large dyke as shown in the inset. Legend: I =R o = 2 % v i t r i n i t e r e f l e c t i v i t y i s o l i n e determined by the Bostick diagram, 2=Extension of the area characterized by Ro ~ 2 % reflectance values, as estimated from accepting the Lopatin-Waples method up to 250 ° C, 3=Area characterized by 2 % > R o > I . 3 % reflectances, 4=Area characterized by 1.3 %>Ro 0.65 % reflectances, 5 = P o s i t i o n of the Ro =0.65 % r e f l e c t i v i t y horizon in sediments not influenced by the magmatic heat.

183 affected here the cold and immaturated Mio-Pliocene part of the sedimentary succession. The influence is more s i g n i f i c a n t at the sides of the central peak of the dyke, where the Eocene-Oligocene sediments were already warmer and more matured at the time of i n t r u s i o n . But even here, the zone of influence is not very large. The elevated isoreflectance l i n e s drop down to the normal (undisturbed) level within a few k i l o meters from the dyke. This l i m i t e d influence of igneous bodies, however, can be important i f p o t e n t i a l source rocks would not otherwise reach the main phase of o i l generation. This is the case in North Hungary where p e l i t i c Paleogene (dominantly Oligocene) rocks are j u s t at the onset of the o i l - g e n e r a t i o n window. Therefore, we can expect that f l u i d hydrocarbons could have been generated at places where thick Paleogene sediments are associated with magmatic bodies of Middle Miocene age. Petroleum f i e l d s found so f a r appear to f u l f i l

t h i s p r e d i c t i o n , as is shown in Fig. 7.

Generally, we think that buried magmatic complexes in the Pannonian basin - p a r t i c u l a r l y the thick r h y o l i t i c f l o o d t u f f s and associated feeder dykes of Middle to Late Miocene age - can be important concerning hydrocarbon prospections. In addition to the f a c t that magmatic heat contributes to thermal maturation, i t should also be taken into consideration that hydrocarbons can migrate and be trapped in volcanoclastics. Conclusions The f o l l o w i n g main conclusions can be drawn from our study: i)

Conductive models combined with the Bostick diagram and the Lopatin method can be used to delineate the influence of magmatism on the maturation of organic matter.

ii)

Model calculations show that there is no real "telemagmatic" thermal e f f e c t . The zone of influence is confined to sedimentary rocks adjacent to the magmatic body which roughly encompasses the same volume as the body i t s e l f .

iii)

Maturity increase caused by a given igneous body strongly depends on the pree x i s t i n g thermal and maturity conditions in a sedimentary basin. The optimal s i t u a t i o n occurs in a not very young basin when f a i r l y matured sediments are driven into the o i l - g e n e r a t i o n window by the heating of igneous m a t e r i a l .

iv)

This condition probably prevailed at some places in the Pannonian basin during the intensive Middle to Late Miocene c a l c - a l k a l i n e volcanic a c t i v i t y .

Acknowledgements This work was i n i t i a t e d and supported by the Hungarian Geological I n s t i t u t e . We are p a r t i c u l a r l y grateful to Dr. A. Jambor f o r valuable advice and Dr. G. H~mor f o r permission to publish. We also thank Prof. L. Rybach (ETH, ZUrich) f o r discussion on thermal modelling.

PALEOTEMPERATURES IN THE CENTRAL ALPS, - A N ATTEMPT AT INTERPRETATION D. W E R N E R Institut

ffir G e o p h y s i k ,

ETH

Zfirich

Abstract Paleotemperature data contain important information about the u p l i f t h i s t o r y of a mountain range. A simple u p l i f t model f o r the Gotthard region (Central Alps) is presented. This model s a t i s f i e s the paleotemperature data but leads to u n r e a l i s t i c a l l y high temperatures at greater depths (lower crust, uppermost mantle). To improve the model a d d i t i o n a l crustal heat sources must be introduced which are assumed to be f r i c t i o n a l heat sources caused by crustal overthrusting. This thermal problem also appears in other subregions of the Central Alps. Further problems are related to the thermal t r a n s i t i o n zone between adjacent crustal blocks with d i f f e r e n t u p l i f t h i s t o ries. Introduction Paleotemperatures in conjunction with radiometric ages are an important tool to reconstruct the u p l i f t h i s t o r y of a young mountain range l i k e the Alps. The precondit i o n is that rock samples must be a v a i l a b l e which can t e l l us about i t s thermal history.

On the other hand, u p l i f t and denudation mirrors the temperature d i s t r i b u t i o n

in the earth's crust. This means that the "thermal memory" of such a rock sample contains information about the u p l i f t h i s t o r y of a whole crustal block. This thermal memory is based on a series of d i f f e r e n t processes', which are the concern of nuclear physics. For a l l these processes a so-called blocking temperature can be defined, which represents an upper l i m i t of a temperature i n t e r v a l in which nuclear accumulat i o n processes have taken place. The number of the nuclear events, then, is proport i o n a l to the time span between present, and the point of time at which the cooling rock sample has passed through the blocking temperature: f o r instance 500 ° C f o r muscovite and phengite Rb-Sr, 350° C f o r muscovite and phengite K-Ar, 300° C f o r b i o t i t e K-Ar, and 120° C f o r a p a t i t e f i s s i o n track ages (JAEGER et a l . ,

1967;

HUNZIKER, 1974; FREY et a l . , 1976; PURDY & JAEGER, 1976; WAGNERet a l . , 1977). From these benchmarks the thermal h i s t o r y of a rock sample during i t s path towards the

*) Contribution no. 483 of the I n s t i t u t fur Geophysik, ETH ZUrich

Lecture Notes in Earth Sciences, Vol. 5 Paleogeothermics. Edited by G. Buntebarth and L. Stegena © Springer-Verlag Berlin Heidelberg 1986

186

earth's surface can be reconstructed, and a simple formula may be used to estimate u p l i f t rates: u p l i f t rate = cooling rate/geothermal gradient. A more precise study, however, cannot be based on t h i s formula, because the geothermal gradient is dependent on depth and time and on the d i s t r i b u t i o n of radiogenic heat sources. In Fig. I , a simple sketch is shown which may i l l u s t r a t e the thermal changes caused by u p l i f t and erosion.

former s u r f a c e ~ eroded

layer~

it

~- T, A

ii A

later s u r f a c e ~

Ix.\\

Fig. I. Sketch of the temperature depth curves before and a f t e r a period of u p l i f t and erosion (T and T' resp.). A: radiogenic heat sources. Several methods have been used to describe the r e l a t i o n s h i p between u p l i f t and temperature f i e l d (CLARK & JAEGER, 1969; OXBURGH & TURCOTTE, 1974; ENGLAND,1978; WERNER, 1980,1981). In t h i s paper some d i f f i c u l t i e s

r e l a t e d to the i n t e r p r e t a t i o n of the

paleotemperature data in the Central Alps are discussed. The main problem is to f i n d an explanation f o r the considerably high paleotemperatures in the crust which does not agree e i t h e r with near surface geothermal observations, nor with the f a c t that in the uppermost mantle under the Central Alps r e l a t i v e l y low temperatures must be expected (WERNER & KISSLING, 1985). The problem may be demonstrated by considering the Gotthard region which is a subregion of the Central Alps. Taking a one-dimensional u p l i f t model the geothermal s i t u a t i o n in the Alpine lithosphere cannot be interpreted in a s a t i s f a c t o r y way. Additional heat sources within the crust must be introduced to f i t atures.

the paleotemper-

In t h i s paper a preliminary attempt is made to i n t e r p r e t e the high temper-

ature data by assuming that f r i c t i o n a l heating due to crustal overthrusting may play a geothermal r o l e . This means that we consider a model which combines onedimensional u p l i f t ,

and horizontal shear motions.

187 I t must be noted that at least two long-termprocesses are neglected here. The f i r s t one is related to the size of the region which is b u i l t of l a t e r a l l y l i m i t e d regions (crustal blocks) with d i f f e r e n t u p l i f t h i s t o r i e s and with l a t e r a l thermal exchange between adjoining crustal blocks and not of an i n f i n i t e one-dimensional area. The second process is the crustal thickening during the h i s t o r y of the Alps. Theoretical remarks The problem can be expressed by the d i f f e r e n t i a l equation describing the heat transport in a moving medium (e.g. LANDAU & LIFSCHITZ, 1978): (I)

~T I ~ = -7 grad T + cp _

div(KgradT)

+

A+½7

~v i

~v K 2

where T= temperature, t = time, V= v e l o c i t y of the medium, c = s p e c i f i c heat capacity, p = d e n s i t y , K=thermal c o n d u c t i v i t y , A = s t r e n g t h of radiogenic heat source, q = v i s c o s i t y , and x i = c a r t e s i a n coordinates. The f i r s t

term on the r i g h t hand side describes the convective portion of the heat

transport. The v e l o c i t y ~ is a given quantity representing a kinematic model of the tectonic h i s t o r y of the Alps, and in w r i t i n g ( I ) we have assumed that the medium is incompressible: (2)

div ~ = O.

The second term describes the conductive heat transport with a temperature dependent c o n d u c t i v i t y . The heat source A follows the material moving with v e l o c i t y ~ and s a t i s f i e s the c o n t i n u i t y equation (3)

~A/~t = -~ grad A.

The l a s t term in Eq. ( I ) describes the f r i c t i o n a l

heat sources assuming that the

medium behaves as a Newtonian f l u i d . To implement the boundary conditions we assume that f o r long-term processes the uplift

rate and the erosion rate are equal at a l l times. This means that the surface

remains a level plane with a constant temperature. In order to demonstrate the thermal problem in the Central Alps by a simple model, we assume that only two processes have taken place: a time dependent u p l i f t and horizontal motions. For the l a s t one a very simple kinematic model may be given by

188

(4)

Vx(Z) = vo

I I - ~ 2 arctg [ c ( z - z

o)]

1

where vo, c and z° are constants. This model describes the overthrusting of a crustal layer with constant thickness z o. Combining u p l i f t and overthrusting the geothermal model remains one-dimensional and Eq. ( I ) reduces to (5)

~T(z,t)~t _ _Vz(t) ~T(z,t)~z F c~I

[~

(K(z,t)

~

I ~Vx(Z't) 21 + A(z,t)+~ ~(z)(-- T ) ]

To evaluate Eq. (5), a simple f i n i t e difference method in space and time has been used. Paleotemperatures in the crust In the Gotthard subregion of the Central Alps considered here, three f i s s i o n track data (blocking temperature for apatite: 120° C; ages: 6.7, 7.2, 7.6 m.y.) and two Rb-Sr data (blocking temperature f o r b i o t i t e : 300 ° C; ages: 15.1~1.8 and 15.7+1.1 m.y.) are known (WAGNER et a l . , 1977). The errors of the blocking temperatures are in the order of +10%. The f i s s i o n track ages are topographically reduced values (BUCHLI & WERNER, 1985). From the degree of metamorphism in the Gotthard region f o l lows the original depth of the rock sample which amounts to about 15 km (WAGNER et a l . , 1977). The task now is to find a kinematic model such that the thermal history of the rock sample passes through the benchmarks of the paleotemperature data. For t h i s purpose the kinematic f i e l d (Vx,V z) which is the input of the calculation, must be modified to f i t

the data. Using Eq. (5), the entire temperature variations in space and time

have to be calculated. After each time step the burial depth z and the corresponding temperature Ts of the rock sample must be determined. Fig. 2 shows the u p l i f t history (a), the burial history of the rock sample (b), and i t s corresponding temperature history Ts (c) passing through the benchmarks of the paleotemperature data. The temperature curve I corresponds to a simple one-dimensional u p l i f t model shown in the upper part (a). The original temperature d i s t r i b u t i o n (26 Ma b.p.) to t h i s model is shown as curve 01 in Fig. 3. PI means the corresponding temperature d i s t r i b u t i o n at present. I t is obvious that the curves 01 and PI in Fig. 3 show u n r e a l i s t i c a l l y high temperatures in the lower crust and in the upper mantle whichdisagreeswith the cold body (lithospheric root) to be expected under the Central Alps. This body is characterized by increased seismic v e l o c i t i e s (PANZA & MUELLER, 1978; BAER, 1980) and can be interpreted as a negative mantle tem-

189

(up|lftll

m

I O!

II

25Ma i

0

l

i

i

t

i

i

2o 0

TIME~ 0

IO

5 i

O i

6

~

b

z 4oo

~

•

BIOTITE

Rb-Sr

~

•

FISSION

TRACKS

2~ o

~

3~'~.

,

¢

i

-

Fig. 2. a: Modelled u p l i f t h i s t o r y of the Gotthard region (Central Alps) which cons i s t s of 5 periods of d i f f e r e n t u p l i f t rates vz. b: The corresponding depth changes of a rock sample s t a r t i n g at a depth of 15 km 26 Ma b.p. c: The corresponding temperature history Ts of the rock sample. Curve I represents a simple one-dimensional u p l i f t model leading to u n r e a l i s t i c a l l y high temperatures at greater depths (curves 01 and PI in Fig. 3). Curve 3 corresponds to the same u p l i f t model but with additional heat sources in the crust, whereas curve 2 represents the model without heat sources.

o

i

20O ~

I

4OO I

I

600 I

800Oc

~

ITL

O

1

2

3

4

A rad 10~ W/m3

1o

20

~//

30

P3

'z

\

vZ

Fig. 3. Left: Original (0) temperature depth curves at the beginning of the u p l i f t h i s t o r y (26 Ma b.p.) and present (P) temperature depth curves. 01 and PI are unreal i s t i c , based on u p l i f t i n g only. Right: Assumed radiogenic heat sources (Ara d) at the beginning of the u p l i f t h i s t o r y , and f r i c t i o n a l heat sources ( A f r i c ) during the time span between 20 and 10 Ma b.p. (note the d i f f e r e n t scales). perature anomaly which corresponds to a thermally induced p o s i t i v e density anomaly. Such an anomaly agrees with g r a v i t a t i o n a l observations,on the one hand, and with a possible displacement h i s t o r y of the l i t h o s p h e r i c root, on the other. An adequate model (WERNER & KISSLING, 1985) leads to l a t e r a l temperature differences up to -450 ° C (at a depth of about 140 km) related to an undisturbed temperature d i s t r i bution outside the mantle anomaly.

190 The curves 01 and PI in Fig. 3 are not only u n r e a l i s t i c f o r greater depths but also with respect to near surface geothermal observations. Considering, f o r instance, a local area within the Gotthard region where extremely high radiogenic heat sources have been observed (Rotondo g r a n i t e , KISSLING et a l . , 1978) the surface heat flow amounts to 86 mW/m2 (WERNER, 1985). S i m i l a r results have been found by 80DMER(1982). This means that there is no evidence of remarkable high heat flow values in the Gotthard region, or in the whole of the Central Alps. A better model can be found by introducing a d d i t i o n a l heat sources in the crust which are simulated here as f r i c t i o n a l heating. In order to f i t

the benchmarks these sour-

ces should have existed between 20 and 10 Ma b.p. That means, the model of Eq. (4) is r e s t r i c t e d to a time span of 10 m i l l i o n years. The assumed parameters are: v o= 4.5 mm/y, zo= 10 km, c = 0 . 2 5 . The t o t a l horizontal displacement, then, amounts to 45 km. The v i s c o s i t y at 10 km depth is assumed to be 1024 Poise. This simple overthrusting model, however, leads to contradictions from a dynamic point of view: i t cannot be explained as a r e s u l t of g r a v i t y s l i d i n g processes. Assuming that horizont a l motions are only caused by g r a v i t y g l i d i n g , the v i s c o s i t y must be remarkably reduced and cannot be in the order of 1024 Poise. In this case, the thermal e f f e c t of f r i c t i o n a l heating can be neglected, and cannot be helpful in i n t e r p r e t i n g the crustal paleotemperatures in question. Taking our simple overthrusting model, the c u r v e 3 i n Fig. 2c shows the r e s u l t i n g temperatures Ts which s a t i s f i e s the benchmarks again, but corresponds to more r e a l i s t i c temperatures in the lower lithosphere (curves 02 and 03 in Fig. 3). The same case but without a d d i t i o n a l crustal heat sources is shown f o r comparison in Fig. 2c (curve 2) and in Fig. 3 (curve P2). In the framework of one-dimensional modelling the o r i g i n a l temperature d i s t r i b u t i o n s 01 and 02 are assumed to be steady state d i s t r i b u t i o n s which may not be quite r e a l i s t i c f o r the Alpine region. The r i g h t part of Fig. 3 shows the assumed d i s t r i b u t i o n s of the heat sources. The near surface radiogenic heat of the Rotondo g r a nit e is extremely high (KISSLING et a l . , 1978). At the beginning of the u p l i f t h i s t o r y these rocks were situated at a depth of 15 km. The high values (3.77x 10-6 W/m3) are not considered as representat i v e f o r the formerly uppermost crustal layers. A value of 1.7x 10-6 W/m3 has been assumed f o r the eroded layers. I t is obvious that t h i s model of radiogenic heat production must be a speculative one. Another point is u n s at is f a c t or y , namely that the one-dimensional u p l i f t model cannot describe the crustal thickening which is connected with the thickening of radiogenic heat sources. Fig. 3 also shows the f r i c t i o n a l heat sources from the horizontal motions based on Eq. (4). As mentioned, these heat sources are l i m i t e d to a time span between 20 and 10 Ma b.p.

191 Such an i n t e r p r e t a t i o n of the paleotemperature data may be questionable but i t demonstrates the geothermal problem in the Central Alps. The model shows that paleotemperatures are not only i n d i c a t o r s of the u p l i f t h i s t o r y , but also indicators of addit i o n a l geothermal processes. Other subregions in the Central Alps The same thermal problem as in the Gotthard subregion can also be found in other subregions w i t h i n the Central Alps (Fig. 4, p r o f i l e s I to 5). Most of these regions require a two-dimensional treatment because of t h e i r l i m i t e d size. The degree of metamorphism and the paleotemperature data c l e a r l y show that d i f f e r e n t crustal blocks (subregions) with d i f f e r e n t u p l i f t et a l . ,

h i s t o r i e s must be distinguished (WAGNER

1977; WERNER, 1980). A r e l a t i v e v e r t i c a l movement between adjacent crustal

blocks leads to a l a t e r a l heat exchange. Therefore the kinematic and geothermal model must be handled at least as a two-dimensional one. An attempt has been madeto construct such models along p r o f i l e s crossing the Central Alps (Fig. 4). The calcul a t i o n s f o r a l l the p r o f i l e s led to the same r e s u l t as discussed above: additional crustal heat sources must be introduced in order to obtain r e a l i s t i c temperatures at greater depths. The r e s u l t i n g u p l i f t

h i s t o r i e s are shown in Fig. 5 (BUCHLI & WERNER,

1985).

. . . . . . . . INSUBRIC LINE ~ ............. CENTOVALLI LINE ) 4o5 : SAMPLE NUMBER / :::::;::

/

J

BERGELL INTRUSION "459o394_-11>"

• '03 Li~"#~ Go'~J:'"....s--

/ ~MPLON/ 164i

--

#' ~

'~-="~-.:,. 'tJ

f __I

_~ J /~1 ,. ~

/ J~r%~,

J

~ )

/ ~.

I \

J

"~138 | 1189 'LIlOS J ~ # ~ l TICINO leC4INO~.~e137/ / ~ - " - " L~B.'ER"CIELL 75 -/,: ,' : : : } . . . . . .

..... _-,_-_'~..~ . . . . . . . . . . ~"~+*'~" . . . . .

T"'-,. ,:;(............ : ; : ; ' " I P Y - - - 3 ),82

/

f ' ~

\

o.,"

#e405 I• •u ~506 m2 572 MONTE ROSA . l i l : ~ ~'~ / i '~ • " ~ " / (

4P"

l

dr"

!

4

375 ~

~ ~ O SOUT[IE;~N A'LPS ~I~ II ......... i

10 20 30 II u, L J ""

Fig. 4. Location map of the Central and Southern Alps with d i f f e r e n t p r o f i l e s f o r d i f f e r e n t u p l i f t h i s t o r i e s (BUCHLI & WERNER, 1985)

192 V(mm/y) il

MONTE ROSA H~lSkm

!l s,_ 30

20

-

~

10

O

V (mnvy)

,, oo_

H~20 km

il

i 4 z 1 5 km

" T--O

Nn2S

1•

_

N=25k

3o

20

1o

0

o

30

:~0

1-0

b

ao

2o TIME (Ma)

-,o

0

o,°

3o

2o TIME (Ma)

'io

b

Fig, 5. Modelled u p l i f t h i s t o r i e s f o r subregions from west to east ( l e f t ) , and from north to south ( r i g h t ) a f t e r BUCHLI & WERNER (1985). H means the total u p l i f t . For location see Fig. 4.

"M~F I

I~,IN

NUMBER:

I I~1~ II I

TIME (Ma)

30

~~,

25

SI MPLON

MONTE

"',,,D

/ 300o QQ"~m n . . . . , C . . . . . ,

15

5

wmmmm~mmm~nm~T~U,l[.~ mm n ~ m D -

\

t~÷9

1~oo

\

uo,t.~.t

_~,

SOUrCe

~oc ;~. . . . . ~ . , . . , . . . . . . . .

~J~

•

I*

I

o

3o

2~

1'o

;

o.m

DISTANCE FROM THE CENTOVALLI'LINE

lb

2b

Fig. 6. Cooling history for two crustal blocks, Monte Rosa and Simplon (see p r o f i l e I in Fig. 4) with d i f f e r e n t u p l i f t h i s t o r i e s . The point of time at which the rock samples should pass the blocking temperatures of 300 ° C and 120° C are indicated by the broken l i n e s . Times at which rock samples (÷) passed the blocking temperatures 300 ° C and 120° C resp. are also indicated. To obtain a better data f i t , f r i c t i o n a l heating between the crustal blocks has been introduced, indicated by the solid lines (BUCHLI & WERNER, 1985).

193 A p a r t i c u l a r r e s u l t of these c a l c u l a t i o n s is shown in Fig. 6 f o r the t r a n s i t i o n zone between the adjacent crustal blocks Monte Rosa and Simplon, which are separated by a deep-reaching tectonic lineament (Centovalli l i n e ) . I t should be expected that paleotemperature data from locations near the Centovalli l i n e r e f l e c t the thermal t r a n s i t i o n zone. In order to study the thermal s i t u a t i o n within t h i s zone, four parameters must be taken in consideration: I. time, 2. depth of rock samples, 3. temperature of rock samples, 4. distance of rock samples from the Centovalli l i n e . In Fig. 6 only three parameters can be seen: time, temperature, represented by modelled lin e s f o r 300 ° C and 120° C, and distance between the id e a liz ed boundary of two blocks (Centovalli l i n e ) and the locations where the rock samples come from. Furthermore, i t is assumed that the two u p l i f t h i s t o r i e s are v a l i d f o r each block as a whole, excepting a narrow f r i c t i o n a l zone (width of the Centovalli f a u l t in the order of 100 m). The i n t e r e s t i n g stage now, is to compare the broken lines in Fig. 6, with respect to the 300° C paleotemperature data. The broken lines represent a c a l c u l a t i o n r e s u l t which is based only on thermally conductive contact between the two crustal blocks. Taking, f o r instance, the point of time 29 Ma b.p. at which the u p l i f t of the Monte Rosa block was greater than the u p l i f t of the Simplon block (see Fig. 5), we must except higher temperatures in the Monte Rosa than in the Simplon. The lower temperatures in the Simplon should influence the near Simplon parts of the Monte Rosa. This means that a near Simplon rock sample of the Monte Rosa block should pass the 300° C temperature point e a r l i e r than a " t y p i c a l " rock sample from the Monte Rosa f a r from the Simplon. The data, however, show a reversed tendency. Rock samples from l o cations near the Centovalli l i n e reach the blocking temperature l a t e r than expected from the conductive thermal model, and i t seems that the f r i c t i o n a l zone between the blocks acts as a heat source (BUCHLI & WERNER, 1985). A s i m i l a r behaviour can also be found in t r a n s i t i o n zones of other crustal blocks within the Central Alps, f o r instance between the Ticino block and the Southern Alps. F r i c t i o n a l heating due to d i f ferent v e r t i c a l motion may be a possible explanation of this phenomena. In a l l cases, the paleotemperature data in the Central Alps contain not only i n f o r mation f o r the d i f f e r e n t u p l i f t h i s t o r i e s , but o f f e r new problems concerning the r e l a t i o n s h i p between tectonics and geothermics.

GEOTHERMAL

S T U D I E S IN O I L FIELD D I S T R I C T S OF N O R T H C H I N A

W A N G JI-AN,

W A N G JI-YANG,

YAN S H U - Z H E N

and L U X I U - W E N

I n s t i t u t e of Geology, A c a d e m i a Sinica P . O . B o x 634, Beijing, C h i n a

Abstract In North China, T e r t i a r y sediments give the main o i l - g e n e t i c series. The mean value of t e r r e s t r i a l

heat flow density has been considered to be 6 0 - 6 5 mW/m2, and the

geothermal gradient in T e r t i a r y sediments usually ranges from 30 to 40 ° C/km in the region studied. Supposing that the onset of o i l generation l i e s at about 90 ° C, the upper l i m i t of the depth of o i l - g e n e r a t i o n is at about 2000 to 2500 m depth. Recent paleogeothermal studies using v i t r i n i t e

reflectance, clay and authigenic minerals,

as well as other methods showed that in Eocene the geothermal gradient has been higher than at present. Some results were obtained and discussed. Introduction North China is r i c h in o i l resources and recently became one of the main resourcebases f o r energy supply in China. In some large-scale Mesozoic-Tertiary sedimentary basins such as Lower Liaohe, Central Hebei, Northern Shandong, o i l - g a s deposits of commercial i n t e r e s t have been found in many places. The North China Plain ( i n c l u d i n g Lower Liaohe) g e o l o g i c a l l y is a large-scale Meso-Cenozoic basin of f a u l t i n g - d e p r e s sion o r i g i n which developed on the Pre-Cambrian basement (ZHANG WEN-YOU et a l . ,

1982).

The s t r u c t u r a l framework of the region has been formed during the "Yintze" tectonic cycle (225 to 195 m.y.). Pre-Cambrian rocks are widely exposed in the mountainous area surrounding the P l a i n , and Paleozoic sedimentary strata of platform type as well as Mesozoic c l a s t i c and volcanic rocks of t e r r e s t r i a l o r i g i n are i n t e r m i t t e n t l y distributed

in

the periphery of the Plain. Within the Plain, on a series of rises and

depressions of Pre-Cenozoic s t r a t a , Cenozoic sediments are l y i n g . The b l o c k - f a u l t i n g movements i n i t i a t e d since Mesozoic have also been continued in the Cenozoic. In spite of some differences in Cenozoic sedimentation h i s t o r y , in the early Eocene (Sahejie, e s p e c i a l l y early Sahejie), the North China Plain as a whole subsided r a p i d l y . At Oligocene (Dongying), the amplitude of subsidence decreased gradually, but in some places such as Lower Liaohe and p a r t l y Shandong, the thickness of sediments has s t i l l been recorded as t h i c k as 1000 m or even more. By the end of early T e r t i a r y , the Plain had come to the end of intensive subsidence, and hence, a thickness of several hundred meters f o r the Miocene (Guantao) and Pliocene (Minghuazhen) sediments are

Lecture Notes in Earth Sciences, Vol. 5 Paleogeothermics. Edited by G. Buntebarth and L. Stegena © Springer-Verlag Berlin Heidelberg 1986

196 usually observed in most areas with the exception of the Bohai Bay area, where the Pliocene sediments were found as thick as 1000 m or more. Since Quaternary, the North China Plain has gradually stepped into a stage of peneplaination development; the thickness of Quaternary sediments is less than 400 m. The v e r t i c a l crustal movement, dominated in Cenozoic era and accompanied by the b l o c k - f a u l t i n g processes, has been regarded as the main form of tectonic a c t i v i t y in this area. Meanwhile, f o l d i n g process developed weakly, Early T e r t i a r y b a s a l t i c v o l canism, characterized by fissure eruption and associated with sediments, was extensive in the northern part of the North China Plain. Since l a t e T e r t i a r y , the magnitude and i n t e n s i t y of b a s a l t i c volcanism decreased s i g n i f i c a n t l y and the Quaternary volcanism has only been found in some local places. In the North China Plain there exists two types of o i l - g a s deposits of commercial interest: I.

The so-called "buried h i l l "

type of deposits with mainly Pre-Cenozoic reservoirs

of e a r l y Paleozoic to Sinian ( l a t e Proterozoic) carbonate rocks. 2.

Oil-gas deposits with Cenozoic reservoirs of various coarse c l a s t i c rocks.

The t o t a l thickness of Cenozoic sediments in some strongly subsided depressions amounts to 7000-8000 m or even more, and in most areas a thickness of 3000 to 4000 m s t i l l

remains. Apparently, the Cenozoic sediments, e s p e c i a l l y the e a r ly

T e r t i a r y ones, have been considered to be the main o i l - g e n e t i c series because of the more favourable geological and geothermal settings: the abundant source of organic m a t e r i a l ; the appropriate temperature conditions f o r petroleum maturation; and the r e l a t i v e l y stable tectonic environment. Main geothermal features Recent studies indicate that the o i l f i e l d d i s t r i c t of North China is characterized by a r e l a t i v e l y high geothermal s e t t i n g . The mean heat flow density value of the North China Plain has been thought to be 63 mW/m2 with the i n d i v i d u a l values ranging from 61 to 74 mW/m2 (Geothermal Res. Div., 1979a,1979b~ WANGet a l . , WANG, 1982) (Fig. I ) .

1981; DENG&

In many o i l f i e l d s of North China, a higher geothermal gra-

dient has been observed in comparison with other o i l f i e l d s in China (Fig. 2). Based on numerous temperature logs, a temperature map f o r the depth of 2000 m (Fig. 3), and a map of geothermal gradient fo r the whole Cenozoic sediments (Fig. 4) were recently compiled. The temperature at a depth of 2000 m is usually about 75 to 85 ° C. A geothermal gradient of 30 to 40 ° C/km covers most (70 %) of the area with higher values in the northern Shandong to the South-East of Bohai, and with lower values at the piedmonts.

197

Fig. 1. Histogram of heat flow densities in North China

H e a l flow ~lgn~ly I m W / m

30

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500

z)

70 I

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Fig. 2. Geothermes f o r some o i l f i e l d s in China. Solid l i n e s : North China; dashed: Sichuan Basin; dotted: Shanxi Basin

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