The lecture notes in earth sciences

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

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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 T<80 ° C, 20° C at 80 ° C120° C

15
The diagenesis of organic matter accelerates e x p o n e n t i a l l y with temperature. In the
whole process, the time which the layer under consideration passed away at maximum
temperatures, plays a decisive r o l e . HOOD et a l . (1975) worked out a model, in which
the period spent w i t h i n 15° C of the rock's maximum paleotemperature was taken into
consideration. For the maturation of organic matter, and i n d i r e c t l y , f o r the v i t r i nite reflectivity,

they created a scale of thermal maturity called the "level of

organic metamorphism" (LOM), which is c o n t r o l l e d only by the maximum temperature survived by the l a y e r , and by the " e f f e c t i v e heating time" spent by the layer w i t h i n

15° C of the rock's maximum temperature (Fig. 6). S t a p l i n ' s s i m i l a r scale (TAI,
thermal a l t e r a t i o n index) is based on microscopic structure v a r i a t i o n and the colouring of organic debris.
TMA x

500-

2.2
1,8

250-

1~

EA

20 -

aO0

:

1,1

200 -

0,8

300-

0,7

150-

25-

Ro
100--

:

!!

200 --

3050-

78
0.1

......

','o

i

i

i

i

i J i/

.

.

.

.

.

.

iI

I0
100
EFFEC11VE HEATING TIME, MILLIONS OF YEARS
( t i f f : T I M E WITHIN 15°C OF TMAx )

,

i

i

i

i

[ i

1000

Fig. 6. Relation of LOM and Ro to maximum temperature and e f f e c t i v e heating time
( a f t e r HOOD et a l . , 1975, modified)
PUSEY (1973) suggested that maximum paleotemperatures can be obtained accurately
from ESR (electron spin resonance) analysis of kerogen. The ESR is s e n s i t i v e to free
r a d i c a l s ; the number of free radicals increases as kerogen is subjected to increasing
temperatures, and kerogen free radicals are stable through geologic time. The ESR
geothermometer was c a l i b r a t e d by obtaining data from cores of T e r t i a r y basins believed to be a c t i v e l y subsiding and so s a t i s f y i n g the h i g h l y probable assumption
that samples from these basins are now at maximum temperature since b u r i a l . But ESR
signals are not only dependent on temperature but are also subject to v a r i a t i o n s in
kerogen type, diagenetic changes in kerogen, weathering and geologic time.


16
PRICE (1982) improves the idea that v i t r i n i t e

reflectivity

depends f i r s t

of a l l on

maximum temperature. A p l o t of Ro versus present temperature from a number of areas
that have not undergone major geologic m u t i l a t i o n , increases in a s t r i c t l y

linear

f a s h i o n ( r = 0 . 9 7 ) yet burial times f o r these d i f f e r e n t areas range from 0.3 to
240 Mao He suggests that some geochemical postulates are in e r r o r and that time has
little

e f f e c t on organic maturation. I t appears that v i t r i n i t e

reflectivity

can be

used as an absolute paleogeothermometer from 20 ° C to at least 400 ° C.
MIDDLETON & FALVEY (1983) propose, f o r s i m p l i c i t y , that maturation (C) and Ro are
related by the equation
In Ro = A + BC.
Empirical studies give A=-2.275 and B=0.177.
For maturation C, they accept Lopatin's o r i g i n a l idea with i n s i g n i f i c a n t modification
(AT= 10.2 instead of 10° C) and f o r s i m p l i c i t y use the logarithm of the previously
given integral
t
C=In~2T(t)/10"2dt
(as used by ROYDEN et a l . ,

1980 and DE BRAEMAEKER, 1983).

Equations combine to give an equation r e l a t i n g Ro to temperature as a function of
time:
t

(Ro)a = b~o exp[c T(t)] dt
where a=5.635, b = 2 . 7 . 1 0 -6 Ma- I and c =0.068 ° C- I .

Given the thermal h i s t o r y of an organic sediment T ( t ) , t h i s equation can be used to
predict the v i t r i n i t e

reflectivity

of the sediment a f t e r a time t . Nor does WELTE&

YOKLER's (1981) equation add more to that formulated by LOPATIN (1971) and WAPLES
(1980):
Ro [%] = 1.301 l g ( T T I ) - 0 . 5 2 8 2 .
BUNTEBARTH (1978) t r i e d to calculate paleogeothermal gradients, as f a r as possible
without theory. I t is clear that a r e l a t i o n s h i p e x i s t s between the coal rank,
measured by the mean optical r e f l e c t i v i t y

of v i t r i n i t e

(Rm), and the integral of

depth and duration of b u r i a l . A c o r r e l a t i o n has been evaluated between the square of
vitrinite

reflectivity

and the b u r i a l h i s t o r y :


17

tI
Rm2 ~

z ( t ) dt
L)

(z depth, t time, t I means that the calculation can be r e s t r i c t e d to a part of the
whole burial h i s t o r y ) .
Furthermore, i t is clear that, in t h i s r e l a t i o n s h i p , the coal rank is proportional to
a function of the geothermal gradient,
tI
Rm2= 1.16.10 -3 exp(O.068 dT/dz)F z ( t ) dt.

Jo
Fig. 7 shows measured Rm values in some boreholes in the F.R.G. as a function of
tI
burial history (~ z ( t ) d t ) .
0
Geothermal gradients measured at present in the four boreholes in Fig. 7 allowed
the c a l i b r a t i o n of the empirical equation. The a p p l i c a b i l i t y of t h i s equation for
other areas is investigated in BUNTEBARTH & MIDDLETON ( t h i s volume).
2.0

"

I

Z[, Lower OIi~cene)
1.5

(~

Fig. 7. Relation between the mean v i t r i n i t e reflectance (Rm) and the integral of depth and time, in four boreholes of the Upper Rhinegraben
(BUNTEBARTH, 1979)

~,

~

Landau 2
( Upper Oligocene
/ to Pffocene) // Sandhiusen I

Anzlng 3

1.0

.5

O~

40

120

80
tl

zlt)dt

160

200
km Ma

Some case h i s t o r i e s :
Based on maximum measured v i t r i n i t e

r e f l e c t i v i t y data (HACQUEBARD, 1977) and burial

history of 28 wells lying in the Central Prairies Basin, Canada, MAJOROWICZ & JESSOP
(1981) estimated a lower average paleogeothermal gradient (27 mKm- I ) for the early
Oligocene time than the present day one (30.6 mKm-I in average) (Fig. 8). For the
calculation, they used the method of KARWEIL (1956) with BOSTICK's (1973) modifications and the method proposed by HOOD et a l . (1975).


18
AVERAGE OF
PALEOGEOTHERMAL
GRADIENT

24

MEAN

Fig. 8. Average of the paleogeothermal gradients
with the histogram of present geothermal gradients,
in the Central Prairies basin, Canada (MAJOROWICZ
& JESSOP, 1981)

2C-

f6z
w

f2-

i
t6

i

i

i

i

i

i

i

i
8

GEOTHERMAL GRADENT (mKm11

EGGEN (1984) worked with a l o t of v i t r i n i t e

reflectivity

data but present heat flow

estimations only. He stated that in the Viking Graben (Norwegian North Sea) the c a l culated paleo heat flow density (approx. 55 mWm-2) f i t s well with the present heat
flow estimation ( 5 0 - 6 0 mWm-2); on the flank of the Viking Graben, however, an
=2

average paleo heat flow close to 50 mWm

has to be assumed in order to obtain the
-2

observed maturity, while the present day estimation l i e s at 70 mWm

WANG JI-AN et a l . ( t h i s volume) found t h a t the c o a l i f i c a t i o n gradient increases from
0.25 to 0.65 % Ro/km from middle to e a r l y Eocene, in the western part of Liaohe o i l
field,

North China, and a c o a l i f i c a t i o n gradient of about 0.4 was determined in the

early T e r t i a r y sediments of the Central Hebei o i l f i e l d .

KARWEIL's (1956) and

LOPATIN's (1971) methods were used f o r paleotemperature reconstructions. In contrast,
RYBACH (1984) gives 0 . 0 9 - 0 . 0 5 % Ro/km c o a l i f i c a t i o n gradients f o r the Northern
Alpine Foreland (Molasse basin).
ROYDEN & KEEN (1980) predict R° values f o r the sediments of the Nova Scotia and

Labrador shelves, based on t h e o r e t i c a l l y derived thermal e v o l u t i o n , and on LOPATIN's
theory. A s i m i l a r work was carried out by ROYDEN et a l . (1980), f o r the Falkland
Plateau and f o r three places of the North A t l a n t i c .
BUNTEBARTH (1983,1985) estimated the paleotemperature gradient as well as the heat
flow density in a few sedimentary basins in the F.R.G.
In the Ruhr Basin, the foredeep of the Rhenish Variscan mountains, f o r which many
data are a v a i l a b l e (BUNTEBARTHet a l . ,

1982), the heat flow decreased during West-

phalian C from about 125 mW/m2 to about 105 mW/m2. Because of the low thermal cond u c t i v i t y of the coal, the temperature gradients reached mean values of 79 ° C/km
before, and 65 ° C/km a f t e r t h i s decrease in heat flow data obtained.


19
The thermal regime of the back-deep of the Rhenish Variscan mountains, that of the
Saar Basin, is nearly the same as that of the Ruhr Basin during the Westphalian
(BUNTEBARTH, 1983). A s i m i l a r high heat flow is indicated in other European Carboniferous basins, e.g. ROBERT (1985).
Within the Lower Saxony Basin, the Upper Carboniferous coal beds were heated by
i n t r u s i v e bodies during the Upper Cretaceous - about 200 Ma a f t e r sedimentation.
Model c a l c u l a t i o n s that take the cooling of the Massif of Bramsche into account,
indicate that temperature gradients between 60 and 80 ° C/km existed w i t h i n the coal
bearing s t r a t a . From model c a l c u l a t i o n s , p r i o r to the magmatic heating, the temperature gradient did not exceed 30 to 40 ° C/km during maximum burial (BUNTEBARTH,1985).
The paleogradient derived f o r the borehole Urach 3 (Swabian Alb) f o r t u i t o u s l y agrees
with the measured present day gradient. The paleogradient of 43 ° C/km corresponds to
Cretaceous to Lower T e r t i a r y times, because c o a l i f i c a t i o n ended p r i o r to the Upper
Tertiary uplift

(BUNTEBARTH & TEICHMOLLER, 1982).

The thermal regime of the middle Upper Rhine Graben changed during the T e r t i a r y .
Temperature gradients during the Lower T e r t i a r y were higher than those during the
Upper T e r t i a r y . The values ranged, r e s p e c t i v e l y , from 48 to 78 ° C/km, and from 34 to
50° C/km (BUNTEBARTH, 1978). A d i r e c t r e l a t i o n s h i p may e x i s t between volcanic a c t i v ity

in the Graben, and the high thermal gradients, both documented f o r the period

immediately a f t e r the opening of the Graben.
From these few data, i t can be t e n t a t i v e l y concluded that the Upper Cretaceous was
a time of widespread high thermal gradients. Furthermore, high gradients also existed during the Upper Carboniferous in northern Germany, and during the Lower Tert i a r y in southern Germany.
In contrast to the problems of v i t r i n i t e
to be r e a l i s t i c .

reflectivity

To avoid the d i f f i c u l t i e s
new idea:

reflectivity,

with v i t r i n i t e

enumerated, these r e s u l t s seem

McKENZIE (1981) proposed a

Some of the problems r e l a t i n g to the empirical r e l a t i o n s suggested by LOPATIN (1971),
WAPLES (1980) and others could be avoided i f chemical reactions i n v o l v i n g only one
molecular type which occur during the maturation of the organic material, were to be
i d e n t i f i e d . MACKENZIE & McKENZIE (1983) have investigated the rates of three react i o n s which occur before and during the early stages of o i l formation. Two of the

reactions are isomerization reactions, at C-20 in a sterane and at C-22 in a hopane
hydrocarbon; the t h i r d reaction converts C-ring monoaromatic to t r i a r o m a t i c steroid


20
hydrocarbons. All three reactions were assumed to be f i r s t

order and monomolecular;

the isomerization reactions are r e v e r s i b l e , with a rate of conversion of the R to
the S form of 1.174 and 1.564 resp., while the aromatization reaction was assumed to
be i r r e v e r s i b l e .
This method excels by i t s clear t h e o r e t i c a l (thermodynamic) p r i n c i p l e s , the Arrhenius
equation is c e r t a i n l y v a l i d f o r these reactions. The problem, however, is that f r e quency f a c t o r and a c t i v a t i o n energy cannot be determined in the laboratory or only
very inaccurately, because of the slowness of the reactions. Because of t h i s ,
McKENZIE's (1978) stretching theory f o r the evolution of sedimentary basins was used
to c a l i b r a t e the reactions. This theory involves a thermal h i s t o r y , which can be
derived s u f f i c i e n t l y accurately from the burial h i s t o r y . Based on chemical analyses
of North Sea and Pannonian Basin cores, and using more or less determined or hypot h e t i c a l stretching models, the kinematics of the three reactions were determined
in Table I.
Table I.

Rate parameters of three reactions (MACKENZIE & McKENZIE, 1983)
Frequency f a c t o r

(s-l)

A c t i v a t i o n energy
(kJ mol -I )

Isomerization of steranes

6.10 -3

91

Isomerization of hopanes

16.10 -3

91

Aromatization of steroid H C - s

18.1014

200

Fig. 9 shows the results f o r the Pannonian basin, which are in a certain agreement
with the t h e o r e t i c a l curves derived with the assumption that the stretching r a t e , B ,
is 2 (SCLATER et a l . ,

1980). A s i m i l a r study was carried out by HOFFMANNet a l .

(1984) f o r the Malakam Delta, Kalimantan, Indonesia, and SAJGO et a l . (1983) f o r the
Pannonian basin. A r t i c l e s of SAJGO & LEFLER ( t h i s volume) give d e t a i l e d information
about the a p p l i c a b i l i t y of some marker reactions to paleogeothermal determinations.


21

I
~37'

0.6:

12s~

113'o

0
.lo1"
nS'o
~oo'~

Arom.

1
Arorn.
0'61

2a

2O 23

24

OZo

o

1

~

2a.

14e

Arom.
0

Atom.

123

3 ,,1

0

20

40

60
t

~0

10~

Ma

Fig. 9. The extent of sterane and of hopane isomerization as a function of s t er oid
hydrocarbon aromatization f o r samples of the borehole HOD in the Pannonian basin.
The curves are calculated t h e o r e t i c a l l y , based on the thermal h i s t o r y of the basin
from McKENZIE's (1978) stretching theory. The basin is assumed to have been formed
by sudden extension ( B = 2 ) , 15 Ma ago. The marks on the curves are present temperatures at 5° C i n t e r v a l s (above) and present depths at 200 m i n t e r v a l s (below)
(MACKENZIE & McKENZIE, 1983). The l e f t lowest diagram shows the approximate thermal
h i s t o r y which belongs to various ~ values.


22
2.

Fluid i n c l u s i o n thermometry

In nearly a l l minerals, whether ores, rock forming minerals or others, small amounts
of f l u i d s are entrapped in the host crystal which preserve the physical and chemical
conditions of the surrounding medium during the time of the crystal growth. I t is
generally assumed that no subsequent change in the entrapped material takes place
(LEMMLEIN, 1956; ROEDDER, 1967).
At the time of entrappment, the f l u i d i n c l u s i o n is a homogeneous phase c o n s t i t u t i n g
mainly of water, s a l t (in general sodium chloride) and some amount of carbon dioxide,
and also s i l i c a t e melt. Since the thermal expansion of the f l u i d / m e l t is greater
than that of the mineral, a vapour bubble is formed w i t h i n the c a v i t y when the temperature decreases.
The f l u i d inclusions are formed by crystal growth when the advancing faces, edges and
corners of the growing crystal are disturbed (primary i n c l u s i o n s ) , by f r a c t u r i n g and
healing of c r y s t a l s during mechanical disturbances or by overgrowth of a crystal
(secondary inclusions) (PAGEL & POTY, 1983).

From rock or mineral samples, t h i n sections (-80pm) are prepared which are mounted
on a glass plate and polished on both sides. The t h i n section is heated up under a
microscope by using a heating stage (e.g. OHMOTO & RYE, 1970). At a certain temperature,

the bubble disappears in the i n c l u s i o n . The heating is then reversed to

cooling u n t i l the entrapped f l u i d / m e l t becomes an inhomogeneous phase again. The
temperature at t h i s condition is measured and called the homogenization temperature.
This p a r t i c u l a r temperature is related to the temperature of formation. However, the
pressure of formation has to be involved. Generally, an increase in pressure requires a greater temperature to complete homogenization (ROEDDER,1967; SIGURDSON,
1974; POTTER, 1977; BOWERS & HELGESON, 1983a,b). The pressure correction is d i f f e rent f o r s i l i c a t e melt inclusions with shrinkage vapour bubbles ( o n l y - 2 0 ° C/kb)
and f o r more compressible f l u i d s with water and carbon dioxide (ROEDDER, 1982).
Since the f l u i d inclusions contain a substantial concentration of sodium c h l o r i d e ,
the pressure correction can be applied, i f t h i s concentration is known. The thermodynamic properties of aqueous solutions are affected remarkably by the concentrat i o n of sodium chloride. The pressure correction reported by POTTER (1977) is based
on :olumetric properties of the NaCI-H20 system. The correction by BOWERS & HELGESON
(1983a,b) works at pressures above 50 MPa and high temperatures from 350 to 600 ° C;
additional to the graphs of the ternary system, FORTRAN programs are given to generate the pressure correction (BOWERS & HELGESON, 1985) which is based on a modified
REDLICH & KWONG (1949) equation of state matching the pressure-volume-temperature
data reported by GEHRIG (1980).


23

Prior to the pressure c o r r e c t i o n , the sodium chloride content must be analyzed. The
s a l i n i t y of the f l u i d in inclusions can be estimated by the depression of the freezing point during cooling. The higher the s a l i n i t y ,

the higher is the depression of

the freezing p o i n t . ROEDDER (1962) reports the freezing point data f o r pure solutions
of sodium c h l o r i d e . Applying t h i s data to f l u i d i n c l u s i o n s , the s a l i n i t y of the brine

can be given in sodium c h l o r i d e equivalent only, since the f l u i d consists of other
c o n s t i t u e n t s , and t h i s mixed s a l t s o l u t i o n causes some uncertainty in the determinat i o n of the sodium chloride content (ROEDDER, 1976). The freezing point of the i n c l u sion is determined during cooling of the sample using cooled nitrogen gas. When the
i n c l u s i o n is frozen, the gas flow is reduced so that the c r y s t a l s begin to melt. The
melting temperature of the l a s t ice crystal determines the freezing point of the i n clusion. The pressure is taken e i t h e r as the actual pressure corresponding to the
b u r i a l depth of the sample, or as the paleo-pressure which is estimated from the
sedimentary h i s t o r y .
CURRIE & NWACHUKWU (1974) and MAGARA(1978) used t h i s p r i n c i p l e f o r determination of
paleo-geothermal gradients in Canadian Cardium sandstone as f o l l o w s : Thin sections
were made from f r a c t u r e - f i l l i n g

materials (mainly quartz) of sandstone cores from 5

boreholes of a single r e s e r v o i r . Those quartz f i l l i n g s

that contained f l u i d i n c l u -

sions were heated and microscopically observed. The ranges of homogenization temperatures and calculated paleogeothermal and measured present geothermal gradients are
shown in Table 2.
Table 2. Ranges of homogenization temperatures and calculated geothermal gradients
of Cardium sandstones (by CURRIE & NWACHUKWU, 1974 and MAGARA, 1978) .
Well

Homogenization
temperature (°C)

Maximum burial depth
geothermal gradient
(mK/m)

A

45- 108

38

33

B

46- 100

35

25

C

50 - 85

33

33

E

51-

84

36

33

F

51-

88

33

31

Near present
Present
gradient (mK/m) gradient
(mK/m)

32

I t was supposed that the highest homogenization temperature of a core refers to
the quartz f i l l i n g

formed at maximum temperature and the lowest homogenization

temperature o r i g i n a t e s from quartz f i l l i n g s
peratures.

formed at lower ("near present") tem-