9. overburden rock, temperature, and heat flow_d. deming

8/17/2019 9. Overburden Rock, Temperature, And Heat Flow_D. Deming

1/22


2/22

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.)


3/22

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


4/22


5/22

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.


6/22

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-


7/22


8/22


9/22

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


10/22


11/22

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

.

I

5

I•

I

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


12/22

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

15

/


13/22

9.

Overburden Rock Temperature and Heat Flow 77

-

.

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


14/22

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


15/22


16/22

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.


17/22


18/22


19/22

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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9. overburden rock, temperature, and heat flow_d. deming

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