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66 eming
2
3
~ 4
w
~
w
0 2
~
>
E 3
CIJ->::
4
5
6
1
m/m.y.
AGE Ma)
2
1
AGE Ma)
200
00
0 ~ _ . ~ ~ r ~ ~ ~ ~ ~
00
0
~
1
2
3
4
5
SSIVE
M RGINS
~ ~ ~ ~ ~ ~ = = ~ = = ~ = = ~
600
5
5
AGE
Ma)
400
.;;;;;;;::
Williston Basin
INTR CR TONIC B SINS
400
3
2
3
~
4
3
5
6 ~ ~ = ~ = : ~ = = ~
5
6
~
2
3
4
5
6
CE
FOREL ND
B SINS
9 ~ ~ ~ 1 0 0
500
4
2
3
~
5
6
AGE Ma)
Figure
9.1.
Representative tectonic subsidence histories for basins from different tectonic settings. The top graph shows the
slopes of a
range
of sedimentation rates after compaction and is provided for reference. After Angevine et al.,
1990.)
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100-50 m/m.y.) to a
passive
margin basin 20-10
m/m.y.). The lowest sedimentation rates -10 m/m.y.)
are found in intracratonic basins such as the Michigan,
Illinois, and Williston basins in North America Sleep,
1971; Schwab, 1976; Sleep et al., 1980). Strike-slip and
forearc basins are characterized by much higher rates
1000-100
m/m.y.). Foreland basins experience the most
varied sedimentation rates, but generally occupy the
middle ground. The highest sedimentation rates are
found in areas of rapidly prograding river deltas e.g.,
U.S. Gulf Coast basin), where sediment deposition can be
as much as 1000-5000
m/m.y.
Sharp and Domenico,
1976; Bethke, 1986; Bredehoeft et al.,
1988).
Geothermal gradients in sedimentary basins also vary
widely, from
as low
as 10 -15°C/km to as high as
50°-60°C/km. Part of this variation can be attributed to
differences in the background thermal state of the crust
on which the basin rests. However, the thermal proper
ties of sediments e.g.,
thermal
conductivity) and
physical processes acting within basins
e.g.,
sedimenta
tion and groundwater flow) are also important determi
nants.
Why do
some
basins
accumulate
sediment
much
faster than others? What controls temperature in sedi
mentary basins
and its variation between and within
basins? How does sedimentation itself affect the thermal
regime? The purpose of this chapter
is
to address these
questions by describing why and how sedimentary
basins form and the physical properties and processes
that control temperature within them.
FORMATION
OF
SEDIMENTARY BASINS
Sedimentation and Subsidence
A
sedimentary
basin is any downwarped area of the
continental or oceanic crust where sediments accumulate
and compact
with
burial into sedimentary rock. The
accumulation and removal of these rocks defines the life
cycle of a basin, from the initial event that creates the
basin through senescence, culminating in eventual uplift
and destruction.
A sedimentary basin forms when a topographic low is
created in the basement rock through either tectonic
subsidence or sedimentation subsidence, or both. Sedi-
mentation
subsidence
can be defined as the
downward
movement of
the
basement rock-sedimentary rock
contact in response to sediment loading e.g., a major
river delta), while tectonic subsidence
is
the subsidence
of
basement rock that occurs, or would occur, in the
absence of sedimentation e.g., the deep ocean basins).
In general, both tectonic subsidence and sedimenta
tion are necessary for the creation of a sedimentary basin.
Sediments accumulate only in topographic lows, thus a
basin must generally exist before the fill. Conversely,
sedimentation reinforces the tectonic subsidence that was
initiated
by
a basin-forming event. The load due to accu
mulated sediments is capable of increasing total basin
depth
by
a factor of
two
or three Turcotte,
1980).
The
9. Overburden
Rock
Temperature and
eat Flow
167
sediment tion
Pm
>
Pc
>Ps
sedimentary
rock
crust
mantle
Figure
9 2
Schematic illustration
of
isostatic subsidence
following crustal thinning and sedimentation Terms are as
follows:
c
= crustal density; Ps = sediment density; Pm =
mantle; h depth of compensation
relative importance of tectonic subsidence and sedimen
tation as driving forces for the creation of sedimentary
basins varies according to the circumstances involved.
Major river deltas e.g., Mississippi, Amazon, and Niger)
are primary examples in which sedimentation itself plays
a major role
in
forcing subsidence and increasing the
depth of a basin.
On
the opposite extreme, the abyssal
plains of
the
oceanic basins are relatively
sediment
starved. They
owe
their existence to the cooling and
subsidence of oceanic lithosphere as it moves away from
the site of its creation at a mid-oceanic spreading ridge;
sedimentation is limited and plays an insignificant role in
determining total subsidence.
Isostasy and Flexure
What controls the subsidence of a sedimentary basin?
l f we assume that the lithosphere has no lateral strength,
the principle of isostasy applies. Isostasy is the funda
mental principle governing the
development and
evolution of topography
on the
earth s surface. A
succinct
mathematical
statement
of
isostasy is
that
density
p)
integrated
over
an
imaginary column
extending from the surface to the depth of compensation
remains constant:
hp z = constant
1)
where z is
depth and
h
the
depth of compensation,
commonly taken as near the base of the lithosphere, or
about
100 km
Figure
9.2).
More simply stated, equation
1 is a mass balance equation. The total mass of material
in
a column
between the surface and
the depth
of
compensation must be constant. If the mass or weight)
of the column increases, the column must sink, or isostat
ically subside. As
the column
sinks, relatively h igh
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McKenzie's model is often applied (and misapplied)
to estimate the timing of hydrocarbon generation. The
usual procedure is to ''backstrip sedimentary basin fill
for the purpose of separating tectonic subsidence from
the total subsidence. This is done by
applying
the
principle of isostasy and compensating
for
factors such
as
sediment compaction
and
changes in
sea level
(Steckler and Watts,
1978;
Sclater
and
Christie, 1980;
Sclater et al., 1980). The estimated tectonic subsidence
curve is then compared to McKenzie's theoretical predic
tions and a ''best value for the stretching factor found.
Once
is
known, heat flow can
e
estimated, tempera
ture calculated, and source rock maturity
predicted
(provided that the location of the source rock in the basin
fill
is known). This is a straightforward approach, but
there are many ancillary determinants that must also be
taken into consideration
if
meaningful estimates of the
thermal history are to be made. These
include the
depression of heat flow by
sedimentation (De
Bremaecker, 1983), the thermal conductivity of rocks
within the basin (Blackwell and Steele, 1989), the surface
temperature, and the possible influence of groundwater
flow. The relative importance of these intrabasin factors
grows with passing time as the influence of the initial
basin-forming event wanes.
Intracratonic asins
Intracratonic, or platform, basins form on continental
interiors e.g., the Michigan, illinois, and Williston basins
of North America; Figure 9.1). They are typically a few
hundred kilometers wide and contain a few kilometers
of flat-lying sedimentary rocks recording continuous
subsidence and sediment deposition over periods of time
greater than
100
m.y. (Sleep et al., 1980). Sleep 1971) was
the first to note that the subsidence of these basins was,
like oceanic basins, proportional to the square root of
time, with a time constant of about 50 m.y. This led to
speculation that the formation of these basins, like rift
basins, was
controlled by some
type of heating
or
thermal event followed by thermal contraction (Sleep,
1971; Sleep and Snell, 1976; Ahern and Mrkvicka, 1984;
Nunn
et al., 1984; Klein and Hsui, 1987).
For an intracratonic basin to be formed by thermal
contraction, isostasy requires that a considerable amount
of crustal erosion occur during the initial heating, uplift,
and thermal expansion phase. For example, i the basin
fill is 3 km deep, it would be necessary to first remove
about 1 km of the continental crust
through
erosion.
However, in many instances, there is little evidence that
this type of dramatic erosion ever occurred (Sleep et al.,
1980). Recognition of this problem has led to the
proposal of several alternative hypotheses. These include
(1) an increase in density of the crust due to one or more
phase transitions,
2)
rifting,
3)
mechanical subsidence
caused by an isostatically uncompensated excess mass of
igneous intrusions,
4)
tectonic reactivation along older
structures, or
5)
some combination of these or other
theories
see
review by Klein,
1991).
For example, Klein
(1991)
suggests
that intracratonic
basins in
North
9. Overburden Rock Temperature and Heat Flow 69
America initially underwent fault-controlled mechanical
subsidence in response to rifting. The initial phase of
basin formation was followed by thermal subsidence
and subsidence due to the isostatically uncompensated
mass of a cooled igneous intrusion.
Although the subsidence history of intracratonic
basins is apparently consistent with some
type
of thermal
mechanism, the exact nature of the initial thermal event,
its subsequent evolution, and the role of other factors in
basin genesis and development are apparently not well
understood at the present time.
oreland asins
Foreland basins (Beaumont, 1981) are asymmetric,
wedge-shaped accumulations of sedimentary rock that
form adjacent to
fold thrust
belts. Migration of
the
fold-thrust sheet loads the lithosphere, causing isostatic
subsidence
underneath the
core of
the
orogen
and
flexural
downwarping in
the adjacent foreland. The
foredeep that forms next to the orogenic belt rapidly
fills
with sediment eroded from the adjacent mountains.
Sedimentation
amplifies flexural subsidence, and a
foreland basin is formed (Figure 9.1).
The foreland basin process continues until the forces
driving uplift and orogeny
cease. Erosion
then
dominates, reducing the weight of the mountain chain,
leading to uplift and further erosion. The life cycle of a
foreland basin is thus typically one of fairly rapid burial
and
subsidence followed
by
a much longer period of
uplift and erosion. Most source rocks buried by the
foreland basin
fill
probably
go
through a relatively short
heating
and maturation
phase, followed
by
a longer
cooling phase.
Thermal events play a minor role in the formation of
foreland basins. However, the thermal state of the lithos
phere influences its flexural strength, thereby exerting an
indirect control
on
the structural evolution of foreland
basins (Watts et al.,
1982).
Other Types of asins
Many other
types
of basins can
e
defined; these
types
are potentially as numerous as the heterogeneous crust
of the earth. Some of these include strike-slip, forearc,
and backarc (Figure
9.1).
Strike-slip or pull-apart basins
are formed by lateral movement along transform faults,
literally pulling the crust apart and creating a void that
fills with sediment e.g., the Los Angeles basin) (Turcotte
and Ahern, 1977; Turcotte and McAdoo,
1979).
Backarc
and
forearc basins
form
in back
of and
in
front
of
volcanic arcs, respectively, near subduction zones.
Backarc basins may form from active seafloor spreading
and rifting, in which case they exhibit high heat flow. In
other
cases,
backarc basins
are
apparently
passive
features that may merely represent trapped segments of
old oceanic crust.
Forearc basins
are the result of
sediments filling the topographic low created by
subduction.
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17 Deming
STRUCTUR L ND THERM L
EVOLUTION OF SEDIMENT RY
B SINS
Are the structural and thermal evolution of sedimen
tary basins linked? In some cases the answer is yes, but
on the scale of a petroleum system, this
fact
may have
little utility in attempts to understand the temperature
history of the basin
fill.
For example, the formation of
rift
basins is well understood through relatively simple theo
retical models that invoke an initial extensional event. In
these cases, it is possible to demonstrate with a fair
amount of confidence a link between the thermal
and
structural evolution of these basins. However, this may
be
of limited relevance in estimating temperature of the
basin fill at a time when
hydrocarbons
are being
thermally generated. The magnitude
of
the initial
thermal event associated with the creation of a rift basin
decays with passing time (Figure 9.3). Thus, by the time
sufficient overburden accumulates for hydrocarbon
maturation to begin, the influence of the initial basin
forming thermal event may be
relatively
small
in
comparison to other factors.
An
example of a
rift
basin in which the initial thermal
event had little influence on the maturation of hydrocar
bons
is
the Gulf Coast basin
of
the southeastern United
States. This basin formed by rifting in Late Triassic-Early
Jurassic time ( 180 Ma), but it was relatively sediment
starved up to about
40
Ma. Rapid accumulation of
sediments since that time has increased the burial depth
and temperature of the source rocks. Cretaceous and
early Tertiary age source rocks are estimated to presently
be in the oil generation window Nunn and Sassen,
1986). Because the thermal anomaly associated with the
rifting of the Gulf Coast basin has been decaying for the
last
180
m.y., the degree to which the lithosphere was
extended or rifted has a negligible influence on the
present-day thermal state (Figure 9.3). Factors such as
lateral
variations in
overburden thickness and
the
depression of heat flow by sedimentation have a greater
influence on source rock temperature. For example,
Nunn and Sassen 1986) estimate that present-day heat
flow in the Gulf Coast basin is depressed
-30
below its
equilibrium value by high rates of sedimentation.
On the scale of the petroleum system, the influence
of
initial basin-forming thermal events
is
thus of indirect or
limited importance in determining temperature of the
basin
fill
at
the
time
hydrocarbons
are
generated.
Temperature of the sedimentary basin
fill is
more likely
to be sensitive to intrabasin factors such as thermal
conductivity,
groundwater
flow, sedimentation,
and
surface temperature (Table
9.1).
The following sections
discuss the importance of these four factors in more
detail.
M THEM TIC L DESCRIPTION OF
HE T TR NSPORT
Sedimentary basins are never in complete thermal
equilibrium, and groundwater flow may drastically
change the distribution of thermal energy within a basin.
Table 9 1 Factors Determining Temperature in
Sedimentary Basin Fill
Importance
Factor
Order)
Qualifications
Overburden thickness 1st
Always important
Heat flow 1st Always important
Thermal conductivity 1st
Always important
Surface temperature
2nd
Always important
Sedimentation
1st
>100 m/m.y.
2nd
100 m/m.y.
3rd
60 Ma)
Nevertheless, steady-state conductive heat transport is a
useful first order approximation that provides a starting
point from which one may later consider departures.
Fourier's law of heat conduction is
q=kg
2)
where
q is
heat flow, k is thermal conductivity, and
g is
the thermal gradient. Applying this to the analysis of
temperature within sedimentary basins, we obtain
T
=
T
+
q/k)
dz:
3)
where T
is subsurface
temperature, T
is the mean
annual surface temperature, and ru
is
thickness of the
overburden. Thus, heat flow, thermal conductivity, and
overburden thickness are of equal importance in deter
mining subsurface temperature. However, heat flow is
generally a more useful measure of the thermal state
of
sedimentary basins than temperature gradient alone
because the geothermal gradient,
g = q/k),
varies
according to thermal conductivity, which can change by
as much as factor of three or four among common rock
types.
A more generalized description of heat transport can
be obtained by considering departures from steady-state
conditions and including advection of heat by moving
fluids. The change of temperature with respect to time
)T j )t)
is then described by
pC )T
/iJt)
=d/dz[kz )T /iJz)]- VzPwCw )T /dz) +A
4)
where z is depth, p and C are the bulk density and heat
capacity, respectively, of a porous rock,
Pw
is fluid
density, Cw
is
fluid heat capacity, Vz
is
the Darcy velocity
of a fluid moving
through
a
porous
medium, k
2
is
thermal conductivity, and
A
is
radioactive heat genera-
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9. Overburden Rock Temperature and Heat Flow
173
100
a)
80
ATI
~ ~ H I N K P J
- T
1----
~
---- r _l_
NIN
I ETK DRP
60
SBE KOL I
----r---- IKP
40
20
LBN
0
A b)
A
SOH ETK JWD
SBE KOL
IN
NIN FCK NKPATI DRP
0
E
2
:::.
--
-
i
3
so
75
>
-
tO
Q
r
:1:
---
Q
5
.0
-
160
c
Q.6Q
50 km
l
7
2 0 0
B
•. ,
Figure
9.6.
Estimated
(a) shallow heat flow (in mW/m2)and (b) subsurface temperature (in 'C) from the North
Slope
basin,
Alaska . Dotted lines
are stratigraphic
units
of cross
section;
solid lines are
isotherms.
Numbers on trace of well bore are
corrected bottom-hole temperatures. Abbreviations for wells:
A
Tl, Atigaru Point
1;
DRP,
Drew
Point 1; EKU, East
Kurupa
1; ETK,
ast
Topagoruk 1;
FCK,
Fish Creek 1; IKP, lkpikpuk 1; INI, lnigok
1;
JWD, J. W. Dalton
1;
KOL,
Koluktak 1;
LBN,
Lisburne
1; NIN,
North
lnigok
1;
NKP, North
Kalikpik 1;
SBE, Seabee
1;
and SOH, South
Harrison
ay 1. After
Deming et al., 1992.)
For the North Slope basin, Deming et al. 1992) used
the
method
of variable bias a conceptually simple
algorithm designed to extract the maximum amount of
information
from
the data
while
simultaneously
averaging the noise. The method involves the sequential
estimation of temperature at different spatial locations
through a series of weighted least-squares regressions.
Based on an estimate of the magnitude of error in the
data, a decision
is
made that n
BHTs
are to be averaged
through an interpretive model. For each point at which
temperature
is
to be estimated, the algorithm searches
through
three-dimensional space until
it
locates the
closest
n
BHTs.
These are given substantial weighting in
the regression analysis; distant data are given much
lower weightings. Instead of all of the data from a basin
being
averaged simultaneously
only data in the
immediate vicinity of the estimation point are averaged.
If data density is locally high, local features of the
temperature field are thus resolved.
If
data density is
low, it is impossible to resolve detailed features of the
temperature field, and the data are averaged over a
wider area. Thus, the balance
between the need
to
resolve the temperature field and the need to reduce
noise
by
averaging is largely determined by the data
themselves. A complete description of this method is
given by Deming et al.
1990a).
Figure 9.6 shows
temperature
in the North Slope
basin estimated from the method of variable bias along a
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0
TERTIARY
..
.
.
I
I
..
.
GANNET
I
.
•I
..
.,
.
PREUSS
.:
2
- . .
I , •.. .••
I . - .. •
TWIN CREEK
E
NUGGET
--·.....1 ----1
.
.¥.
I
..
r-:-..;-•-- . J
.
.
NKAREH
·
. .
L--- •1
••
••
J:
3
I
...
·
I
THAYNES
I •
.
.
a..
L •
LLJ
..-7·
•
0
WOOD.
DJNW
I ¥
PHOSPHORIA
L __ . . ·
.
.
4
• •
: 1·
,
I
I
I
.
WE ER
I
.
I
.
I
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5
I•
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r
t
1
..
MADISON
.
.
·· ··
I
I
I
I 2 3 4
6
7
THERMAL CONDUCTIVITY (W
/m
K)
Figure
9.7. Thermal conductivity data,
Anschutz
Ranch
well
34-02, Utah-Wyoming
thrust belt. Dots
are matrix
conduc
tivities
measured
at 20°C in the laboratory; dashed lines
are estimated n
s tu
thermal conductivity. After Deming
and
Chapman,
1988.)
data measured on samples from the Anschutz Ranch 34-
02 well in the Utah-Wyoming thrust belt Deming and
Chapman,
1988).
The discrete points represent matrix
conductivities measured in the laboratory on drill chips;
the dashed lines show estimated in situ thermal conduc
tivities. The in situ estimates are lower than the labora
tory measurements due to the effects of porosity
and
temperature, and range from about 2 to 4 WIm K. The
wide scatter of
measurements
for
any
formation is
partially due to errors in measurement, but most of the
scatter can be attributed to changes in lithology
and
mineralogy.
The number of measurements needed to determine
the average thermal conductivity of a geologic unit to an
acceptable level of precision depends on its lithologic
heterogeneity. For some marine Paleozoic units that are
lithologically uniform over hundreds of kilometers, it
may be possible to make only 10-20 measurements for
an entire basin. However, large spatial variations in
thermal conductivity are more typical because most sedi
mentary rocks tend to have facies changes that occur
both vertically and laterally. It is therefore difficult in
most cases to collect enough data to estimate how the
thermal
conductivity of a geologic unit changes
throughout a basin.
To overcome this difficulty, concerted efforts have
been made to
estimate
thermal conductivity from
9.
Overburden Rock Temperature and Heat Flow 75
geophysical well logs. In many instances, strong correla
tions have been found between thermal conductivity and
one or more log parameters such as resistivity, seismic
velocity, and density Houbolt and Wells, 1980; Reiter et
al., 1980; Vacquier et al., 1988; Blackwell and Steele,
1989). In other cases, mineralogy has been estimated
from well logs, and the thermal conductivity of the bulk
rock
estimated
from
laboratory-derived
values for
different mine ralogies Brig a
ud
and Vasseur, 1989;
Brigaud et al., 1990; Demongodin et al.,
1991).
The limita
tion of all of these methods is the lack of an accurate
mineralogy log. Matrix thermal conductivity
is
deter
mined by mineralogic composition; correlations and
inferences found to be valid in specific instances cannot
be generalized. Thus, at the present time, there is no
simple algorithm for estimating thermal conductivity
from well logs that
is
demonstrably accurate. However,
well logs may prove useful in interpolating between
measurement sites when log parameters can be cali
brated by laboratory measurements.
CONTROLS ON TEMPERATURE IN
SEDIMENTARY BASINS
Heat Flow and Thermal Conductivity
Because the primary mode of heat transport in the
crust is conduction, both heat flow determined from
equation 2) and thermal conductivity measured
directly) are of first-order and equal importance in deter
mining temperature in sedimentary basin fill.
Heat
flow is inversely correlated to tectonic age
Vitorello
and
Pollack, 1980; Morgan, 1984)
and
is
depressed by sedimentation see later discussion). Heat
flow in young
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76 Deming
G
0
+ ~
..
.
=·
+
1000
2000
Elevation m)
3000
Figure
9.8.
Mean annual air temperatures +) from rnetero
logic data collected by author), and mean annual ground
temperatures •) estimated from extrapolation of borehole
temperature logs from the north-central Colorado
Plateau.
Made
by
Bodell and Chapman, 1982.)
Surface Temperature
Although it
is
often ignored, surface temperattli'e T
0
)
is
an important boundary
condition on geothermal
conditions. The
temperature
at the
earth's
surface is
determined by climate and has diurnal and annual cycles
that rapidly attenuate in the subsurface.
By
applying
equation 5, it can be seen that the annual variation of
temperature propagates no deeper than about
10m
into
the subsurface. Thus, the quantity of interest in geot
hermal studies (T ) is a long term
mean,
a fictional
quantity that
is
usually estimated by the linear extrapola
tion of a borehole temperature log to the surface. Obser
vations have shown that extrapolated borehole tempera
tures are closely related to mean annual air temperatures,
but that
ground
temperatures are
always
higher
by
about 2 -3°C (Figure 9.8). This discrepancy is commonly
attributed to the insulating effect of
snow cover
in
winter. However,
the
offset
between
mean annual
ground and air temperatures is also found at low latitude
sites (e.g., Howard
and
Sass, 1964). Mean
annual
air
temperature on
the
earth's surface is -16°C, varying
from -25°C at the equator, to --22°C at the poles (Gross,
1993). Air temperatures also decrease with elevation;
lapse rates typically range from -4 to -10°C/km.
Air and ground temperatures vary not only spatially
but also have short
and
long term temporal trends. From
about 1400 to 1900, air temperatures were about o.soc
colder than present day (the little ice age ). Before that,
from
about
1000 to 1400 A.D., there was a Medieval
warm period when air temperatures were about O SOC
higher than present day . Over the past 1 m.y., tempera
tures have fluctuated about 5-6°C as the glaciers
retreated and advanced in a series of ice ages. The last
such ice age ended about
10,000
yr ago as temperatures
rose
5--6°C,
coincident with the emergence of civilization
(Folland et al., 1990). Since Late Cretaceous time 70 Ma),
k, >k,
3:
k,< k,
5
Time
0
sediment
u::
c
ii
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9.
Overburden Rock Temperature and Heat Flow 77
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E
-
10
2
-
o
J
0
.> .>
lo..
.
t
q
.)
::: :l
-
t
c:
v
Q:
c
0
g:
-
10
3
c.
Q
..,;;;
'10-14 m2) aquifers and conspicuous signs of under
ground flow e.g., artesian wells) are not a prerequisite.
In areas of high relief and rugged topography, the
presence of groundwater flow is nearly ubiquitous,
making it difficult to obtain accurate estimates of back
ground thermal conditions in these locations.
Groundwater
moves
1) in response to potential
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78 eming
gradients Hubbert,
1940)
or
2)
as a result of free convec
tion. Strictly speaking, it is impossible to
define
a
hydraulic potential for a fluid whose density is variable.
However, the use of a hydraulic potential
or
pseudopo
tential is a useful concept
that
can
be used
to obtain
insight into geologic problems for which an exact answer
is unobtainable.
Common geologic mechanisms for creating potential
gradients
are
sediment
compaction and elevation
gradients. Groundwater
flow
driven
by sediment
compaction
and
pore collapse, however, is a relatively
inefficient mechanism for heat and mass transport unless
pore fluids are focused spatially or temporally Cathles
and Smith, 1983; Bethke, 1985; Deming et al., 1990b).
Flow velocities tend to be very low,
and
the total amount
of water is limited to the water contained in the original
sediments. Few direct observations of thermal anomalies
have been ascribed to compaction-driven flow. Bodner
and Sharp 1988) speculated tha t relatively high geot
hermal gradients associated with fault zones in the Gulf
Coast basin in south Texas may be due to the upward
flow of pore water along the faults. However, they were
unable to rule out alternative explanations for the high
thermal gradients, such as changes in thermal conduc
tivity.
In
contrast
to
compaction-driven
flow,
regional
groundwater flow over distances of 100-1000
km
due to
potential gradients arising from elevation differences has
been documented for several sedimentary basins. There
is virtually no doubt
that such flow systems exist e.g.,
Bredehoeft et al., 1982). Correspondingly, the thermal
regime must be disturbed, with the nature of the thermal
disturbance depending upon the
depth
and velocity of
groundwater
flow. In foreland basins,
the
following
pattern is typical. In the foothills of the mountain range
where water infiltrates at high elevations, the geothermal
gradient and surface heat flow are depressed as heat is
advected downward by moving groundwater. Near the
midpoint
of
the basin
axis of
the
basin fill), flow is
largely horizontal and the effect
on
basin temperature is
minimal. At the distal edge of the basin, flow is forced
upward by the basin geometry, leading to a high geo
thermal gradient
and
high surface heat flow. This pattern
has been
observed in
the
Western
Canadian basin
Majorowicz and Jessop, 1981; Hitchon, 1984;
Majorowicz,
1989);
the Kennedy, Denver, and Williston
basins of the Great Plains province of the central United
States Gosnold,
1985, 1990);
the Uinta basin in western
United States Chapman et al.,
1984);
the Great Artesian
basin in Australia Cull and Conley,
1983);
and the North
Slope basin in Alaska Deming et al., 1992). The thermal
anomalies
associated with these
flow systems
can
dramatically influence the
temperature-dependent
generation of oil and gas, and the flow systems them
selves may play a role in oil
and
gas migration Toth,
1988;
Carven,
1989;
Bethke et al.,
1991;
Meissner, 1991).
Free convection in sedimentary basins
may
conceiv
ably
arise from density gradients
due to
thermal
expansion or the presence of solutes. For free convection
to occur, the permeability of the porous medium must be
sufficiently high and a density inversion must exist, with
higher density fluid overlying less dense. In most sedi
mentary basins, however, free convection is probably
inhibited because the increase of salinity and density)
with depth overshadows the decrease in density due to
increased
temperature
and concomitant
thermal
expansion.
Relatively little is
known about the
occurrence
or
significance of free convection in sedimentary basins;
there are no direct observations of the process. However,
the presence of extensive quartz cementation in sand
stones found in the geopressured zone of the Gulf Coast
basin of the southeastern United States requires
much
higher fluid volumes than could possibly be supplied
by
pore
water
alone Land, 1991). One solution to this
discrepancy would
be
to invoke free convection within
compartmentalized cells in
the geopressured
zone.
Within each cell, the pressure gradient would be hydro
static and fluid would be free to circulate Cathles,
1990).
n
interesting new hypothesis
by Nunn 1992)
attributes
episodic subsidence in the Michigan basin to the cata
strophic release
of
heat by periodic
episodes
of free
convection in the upper 10
km
of the continental crust, as
originally envisaged by Deming 1992). At the present
time, however, the possible occurrence of free convection
in sedimentary basins and the continental crust remains
speculative. For example, little
is
known about the extent
to which fracture permeability exists in the crust and
what
stress states and tectonic processes could create
sufficient permeability to allow for free convection.
THERM L
HISTORY
OF A WELL FROM
NORTH
SLOPE
BASIN ALASKA
The importance of some of the factors that determine
temperature in sedimentary basins and thus source rock
maturation in the petroleum system) can be illustrated
by considering
an
example from the North Slope basin,
Alaska.
During the years
1977-1984,
the
U.S. Geological
Survey drilled
28
deep 1-6
km)
petroleum exploration
wells in the North Slope basin. See
Gryc,
1988;
Tailleur
and Weimer,
1987;
and Bird, Chapter
21,
this volume, for
comprehensive discussions of the exploration history
and regional geology.) A large amount of geologic and
geophysical data were collected as part of this drilling
program. Thermal
data
include thermal conductivity
measurements on core
and
drill chip samples Deming et
al., 1992), equilibrium ±0.1°C) temperature logs in the
upper sections
0--600
m, average) of 21 of
28
boreholes
Lachenbruch et al., 1987, 1988), and
23
reliable ±3°C)
estimates of formation temperatures at depths of about
1-4.5
km
obtained from extrapolation of series of BHTs
measured during geophysical logging
runs
Blanchard
and Tailleur,
1982;
Deming et
al., 1992).
Vitrinite
reflectance measurements were also made on core and
drill chip samples Magoon
and
Bird,
1988).
One of the wells from which temperature, thermal
conductivity,
and
vitrinite reflectance data are available
is the
Ikpikpuk
well latitude 70.46°N,
longitude
154.33°W) Figure 9.11). The history of burial and
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180 emin
g
Table 9.2. De pos i
tional History, lkp
ikpuk Well, North
Slope o f Alaska
Thickness
Age of Dep
osition or Event
Sedimen
tation Rate
Ge
ologic Unit(s) or Ev
ent
(m) (Ma)
(m/m
.y.)
Basement at
surface
350 and old
er
En
dicott Group
235
350-3
35
16
lisburne Group 94 6
335-258
12
Sa
dlerochit Group
320 258-23
5
14
Shublik Fo
rmation/
Sa g River Sand
stone
168 235-2
08
6
Kingak Sha
le
721 208-133
10
Nondeposit ion/er
osion (?)
0
133
-122
Pebble
shale unit
76
122-116
13
Tor ok Formatio
n
1304
116--106
130
Nanushu
k Group
86
9
106--95
79
Colvil le Grou p/
Sag
avanirktok Forma
tion
-1000
9
5-55 ?)
2
5
Erosion
-100
0
55-1 ?)
Gub
ik Formation
10
1-Q
10
Vitri
nite Ref le
ctance(
Ro)
interval
heat perm
eability
0.2
1
- 2
c
-
Q)
3
4
5
Nanushuk
Torok
ebble
s h l e ~
\
Sadleroch
it
-
- ,
\
\
\
\
f low mW/m
2
10-
1
4
m2
4 0
31
10
44
.79
153
4.5
392
2.9
-
293
Lisburn e
40
60
80 100
\
\
\
\
\ BHT
\
\
\
61
.16
ndicott
bas
ement
T
emperatu
re Equiva
lent (°C)
F
igure 9.12. Avera
ge vitrinite reflect
ance determined from
a piecewise
linear regression
on
measuremen
ts (solid line), and
predicted for s
teady-state condu
ctive hea t flows o
f
40
,
60,
80, and 100 mW
/m2 (dashe
d lines), lkpikpuk
well, North Slo pe
basin, Alaska. Also shown
are
interval heat flow s ca lculated from vitrinite reflectance
and
thermal conductM ty data alon
g
with average perm
eabilites me asur
ed parallel to bed
ding.
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sedimentary
basins that could be responsible for fluid
circulation at the
Ikpikpuk
well is free convection. Free
convection is normally inhibited in
sedimentary
basins
because increasing salinity with depth
more than
offsets
the decrease of fluid density due to thermal expansion.
However,
original brines
in
the
North
Slope
basin
have
apparently
been flushed
out by
the regional
ground
water flow system (see analyses
of
formation waters by
Kharaka
and
Carothers, 1988;
Woodward,
1987).
t
may
therefore
be
conceivable
that
the
Ikpikpuk well
is located
in the descending section of one
or more
convection cells.
f
so, other wells
in
the
North
Slope
basin may be
located
in
upwelling sections
of
the same
or other
convection
cells. Vitrinite reflectance
data
from
these wells should
show the opposite pattern to that found in the Ikpikpuk
well-high paleothermal
gradients near
the
top of the
well
and
low paleothermal gradients near the bottom.
Although it is possible to demonstrate that a partic
ular hypothesis groundwater flow) is consistent
with
data from the Ikpikpuk well, the proposed hypothesis
3)
cannot
be shown
to be
the
only
one that
satisfies the
observations. At our current level of understanding, the
estimation of any thermal history is a complex problem
that usually cannot
be
brought to a unique conclusion.
Data errors
and the
difficulties inherent in inferring pale
otemperatures from geothermometers such
as
vitrinite
reflectance
make it
difficult to reconstruct
the
tempera
ture
history
of
potential
source
rocks accurately.
Similarly,
the potential of different mechanisms
(e.g.,
sedimentation, groundwater flow,
and
conductive heat
refraction) to
lead
to identical
thermal
anomalies
makes
it
difficult
to uniquely designate
specific
physical
processes as important factors
in
specific instances.
Acknowledgments I would
like to thank
my
colleagues and
friends at the
Branch of
Petroleum
Geology
in the U.S. Geolog-
ical
Suroey
at
Menlo
Park California. Ken Bird provided the
estimated
burial history for the Ikpikpuk well
and
Les Magoon
made
substantial contributions that improved
the
manuscript.
Jud
Ahern George
Klein Dan McKenzie Jeffrey Nunn
Gerard Demaison Peter van
de
Kamp and John
T.
Smith
reviewed the manuscript
and
made suggestions for its improve-
ment.
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