Download A model of intracratonic basin formation, entailing deep crustal

Geophys. J. R. asfr. SOC.(1980) 62,l-14
A model of intracratonic basin formation, entailing
deep crustal metamorphism
*
M . F . Middle ton Department o f G e o b 0 and Geophysics, Universtzy of.sydney,
NSW 2006, Australia
Received 1979 March 22;in original form 1978 November 14
Summary. Subsidence data from the Michigan and Illinois Basins in the USA,
md
&&
cooper Mi Eroaag2 B2ZirZ ir &8i??Zd'4#22?&2 2iv h&@82&)&2>
with a simple exponentially decreasing (or l/&, where t is time) subsidence
with time, Subsidence in these basins is observed to occur in two consecutive
stages which exhibit exponential-like decreasing subsidence with time. A
model which entails a period of heating followed by a period of cooling is
proposed to explain the observed subsidence. Subsidence due to deep crustal
metamorphism occurs during the latter part of the heating period, and is
followed immediately by subsidence due to thermal contraction during the
cooling period.
The metamorphism, which drives the initial stage of subsidence, is
assumed to occur at the greenschist facies to amphibolite facies metamorphic
boundary which occurs approximately along the 45OoC isotherm at depths in
the order of 10-20 km. This boundary is petrologically very complex, and
cannot be treated as a simple Clausius-Clapeyron phase transition curve.
Loss of volume at the metamorphic facies boundary, which is proportional to
subsidence, is approximated by determining the intersection of the changing
geotherm (during the heating period) with the greenschist-amphibole facies
boundary, defined by the curve: T = 52 t 350, where T is temperature in O C
and z is depth in km.This approach is valid as perturbation of the temperature field in the vicinity of the facies boundary during metamorphism is
shown to be negligible, and sufficient heat is shown to be available for metamorphism at crustal depths within 30 Myr of commencement of the heating
period.
The model indicates that subsidence of the form observed can be
attributed to deep crustal metamorphism followed by thermal contraction
of the lithosphere. The rapid reversal from heating to cooling assumed in this
model is justified by the observation of very slow subsidence suddenly
becoming very rapid, therefore a sudden change in the thermal regime is
needed to produce a subsidence model concordant with the data.
*Present address: Department of Geology, University of Melbourne, Park-,
Victoria 3052, Australia.
2
M.F. Middleton
Introduction
Hypotheses proposed for sedimentary basin formation include gravitational flow of the
crustal material (Bott 1971), thermal contraction (Sleep 1971; Turcotte & Ahern 1977),
mantle/crust phase changes (Joyner 1967; van der Lindt 1967), deep crustal metamorphism
(Falvey 1974), crustal extension (Artemjev & Artyushkov 1971), lithospheric flexure
(Walcott 1970, 1972) and combinations of these (Haxby, Turcotte & Bird 1976; Sleep &
Snell 1976; McKenzie 1978). These mechanisms may be considered as 'primary' mechanisms
of basin formation in the sense that such mechanisms may initiate basin subsidence. Sediment and water loading of a basin and their effect on subsidence may be considered a
natural consequence of a primary mechanism.
The thermal contraction hypothesis (Sleep 1971; Haxby et al. 1976; Sleep & Snell 1976;
Turcotte & Ahern 1977) is widely accepted as a principal mechanism of intracratonic basin
formation, but cannot alone account for available data from a number of such basins.
Published subsidence data from the Michigan Basin (Sleep 1971 ;Haxby et ul. 1976) and the
Illinois Basin (Sleep 1971) are reproduced in Fig. 1. Similar patterns of subsidence have been
observed in a deep borehole in the Michigan Basin, reported on by Sleep & Sloss (1978).
Subsidence data (in the form of sediment thickness above basement versus age) from the
Cooper and Eromanga Basins in central Australia, which have been described by Kapel
(1966), Gatehouse (1972), Wopfner (1972), Exon & Senior (1976) and Rumph (1978), are
shown in Fig. 2. Subsidence is observed from these data to occur in two discrete stages. (see
observed curve in Fig. 3).
Michigan
E
x
-1
I
I-
n
0
:\\
Sleep,1971
e9-
W
2-
3'450'
\
\=s,
Haxby et al.,1976 , R
'
'
'
\x
400
'
AGE,
' ' MY ' bp
'
'
3%
i
2t
1
400
300
A G E , MY bp
Figure 1. Published subsidence data from the Michigan and Illinois Basins. These data suggest that
subsidence in these basins does not follow a simple exponential decay with time, but rather occurs in two
stages each of duration of the order of 50 Myr.
Intracratonic basin formation
3
COOPER
E
-0.2
-----O!
’a,
?
.*
I
260
I
I
240
220
AGE,MY bp
I
E
EROMANGA
0-
Y
m
m
-
*--•
E
m
\
1
0
._ 1c
-
I
al
E
2
location
&
_
C
2-
m
I
I
200
I
I
I
A G E , M Y bp
I
I
100
Figure 2. Subsidence data from the Cooper and Eromanga Basins in central Australia. These data, also,
suggest that subsidence occurs in two discrete stages for these basins.
Sleep (1 976) has explained two-fold subsidence as superposition of eustatic variation
(assumed by Sleep to be a sinusoidal variation with time) upon the essentially exponentialshaped thermal contraction subsidence curve. This explanation requires an unconformity
between the two stages of subsidence, which is not always observed. McKenzie (1978)
has also suggested an hypothesis which entails a two-fold subsidence: a rapid initial phase
due to rapid extention of the lithosphere, followed by a slower thermally contracting phase.
The subsidence model of McKenzie (1978), where a lithosphere of 100 km thickness and
crustal thickness of 30 km is extended by 1.I 1 times its original length (and other parameters the same as assumed by McKenzie 1978), is shown in Fig. 3 compared with an
idealized observed subsidence curve based on Figs 1 and 2. The shape of McKenzie’s curve is
inconsistent with the shape of the observed curve, and it would appear that McKenzie’s
hypothesis is not applicable to basins exhibiting subsidence of the form shown in Figs 1
and 2 .
An alternative explanation is that the initial stage of the two-fold subsidence is caused by
deep crustal metamorphism in basins of relatively young or unaltered crust. Falvey (1974)
has demonstrated that elevation of the geotherm during the rifting of continental margins
may produce subsidence of the margin due to deep crustal metamorphism. He has also
suggested that the formation of intracratonic basins may entail such a mechanism.
4
M. F. Middleton
TIME-
1
50
MY
100
Figure 3. An idealized ‘observed’ curve, adopted from Figs 1 and 2, compared with the subsidence model
of McKenzie (1978).
The deep crustal metamorphism subsidence mechanism implies a subsidence (with
sediment loading to sea level) of
where L is the thickness of lithology 2 metamorphosed from lithology 1, p1 is the density of
lithology 1, p2 is the density of lithology 2, p s is the density of the sediments and pm is the
density of the upper mantle rocks. Falvey (1974) has suggested that the metamorphism
Figure 4. The assumed metamorphic crustal model shown together with the assumed continental
geotherm G ( z ) . Curve A represents the assumed depth-temperature field of the low grade metamorphicgreenschist facies boundary, curve B represents the greenschist-amphibolite facies boundary and curve C
represents the amphibolite-granulite facies boundary.
Intracratonic basin formation
5
which drives basin subsidence predominantly occurs at the greenschis-amphibolite facies
phase boundary (see Fig. 4).
It will be demonstrated that this mechanism, coupled with the thermal contraction
mechanism, can produce a two-fold continuous subsidence. The present paper is concerned
with the development of a simple, preliminary subsidence model, based on deep crustal
metamorphism and thermal contraction of the lithosphere, which is concordant with
observational two-fold subsidence data from various intracratonic basins.
Metamorphic model assumed
The present subsidence model entails an approximation of the effects of deep crustal metamorphism, caused by elevation of crustal temperatures as a consequence of heating at the
base of the lithosphere. Fig. 4 shows the ‘idealized’ petrological crust to be assumed.
A complex continuum of metamorphic phase transitions is possible within the crust. In
order to obtain a tractable approximation to the volume of metamorphosed crust (i.e. L
in equation l), all relevant crustal metamorphic facies changes are considered to be concentrated into three ‘artificial’ boundaries:
(i) Transition from compacted sediments and low grade metamorphic rocks to greenschist
facies rocks (curve A in Fig. 4).
(ii) Transition from greenschist facies rocks to amphibolite facies rocks (curve B in
Fig. 4).
(iii) Transition from amphibolite facies rocks to granulite facies rocks (curve C in Fig. 4).
Also shown in Fig. 4 is the metamorphic scheme of Turner (1968), and a typical continental
geothem, G(z), which will be used in later computations.
The greater part of the metamorphism occurs at the greenschis-amphibolite facies
boundary, where albite-clinozoisite, muscovite-chlorite, calcite-actinolite, dolomitequartz assemblages are metamorphosed to hornblende, plagioclase, diopside, almandine,
grossularite, staurolite, kyanite and sillimanite. The facies transition is not a univariant
system or an olivine-spinel type transition (Gjveik 1972, 1973), and a simple thermodynamic treatment, involving the Clausius-Clapeyron relation, cannot be applied.
In the present treatment, the greenschist-amphibolite facies boundary alone will be
assumed to provide the impetus for metamorphic driven subsidence in intracratonic basins.
Although metamorphism will occur at the low grade and sediment-greenschist facies
boundary, it will only be slight for the changes in geotherm considered here and its effect is
neglected.
It can be seen immediately in Fig. 4 that the amphibolite-granulite boundary (curve C)
is not applicable to intracratonic basin formation, as the crustal geotherm does not
encounter it; this boundary may be relevant to higher temperature regimes such as young
rifted continental margins.
The greenschist-amphibole boundary between depths 10-30 km (see Fig. 4) is defined
by
T=mz+b
(2)
where z is depth, Tis temperature, m is a constant equal to 5.00”C km-’,and b is a constant
equal to 350°C. A significant density change (i.e. loss of volume) is experienced in transition
from greenschist facies rocks to amphibolite facies rocks; this loss of volume essentially
drives the subsidence. Fig. 5 shows histograms of mean density difference between greenschist and amphibolite rocks for varying silica concentration, and maximum density
difference for the same data from Joplin (1963).
M. F. Middleton
6
0.3k
;0.2 -
2
1
I
-
lam
I
2
0.l-
60
SiOp,
0-
1
II
50
60
I
Go2, X
X
Figure 5. Histograms showing the mean density difference between greenschist and amphibolite rocks for
varying silica concentration, and maximum density difference for the same data.
The thermal event
Metamorphism in the deep crust is assumed to be caused by heat conducted from the base
of the lithosphere, which has experienced a 'thermal event'. A thermal event in the present
paper is defined as an elevation (or diminution) of the temperature at the base of the lithosphere. Such a definition is appropriate for intracratonic basin formation hypotheses as the
near surface expression of very high temperatures, such as in mid-ocean ridge models, is not
common.
The effect of such a thermal event upon the crustal geotherm can be determined by
standard techniques; a one dimensional treatment is adequate in the present context because
the geometry of intracratonic basins essentially displays horizontal dimensions much larger
than the vertical dimension.
Carslaw & Jaeger (1959, p. 310) hctve shown that, if a temperature Tois maintained at
the base of a slab which has zero initial temperature, thickness H, and thermal diffusivity
k , then the temperature within the slab increases according to the relation:
x
m
T(z, t ) = To
[erfc
n=O
H(2n t 1) - z
(4kt)"'
- erfc
(3)
(4kt)'12
where t is time, z is depth below the top of the slab and erfc represents the complementary
error function. The lithosphere can be assumed to be a layer of thickness H with uniform
thermal diffusivity and stable temperature distribution at t = 0. In the present treatment the
temperature at the base of the lithosphere is maintained at To+ G(H) for the finite time
interval tl (and then maintained at G(H) for times greater than t l ) . The temperature
distribution in the lithosphere can be obtained from equation (3):
H(2n + 1 ) - z
1
[erfc
n=O.
(4kt)'12
m
T(z, t) = To
T(z, t )= To
x
m
n=O
+ erfc
[erfc
H(2n t 1) - z
(4kt)l12
H(2n + 1) t z
[4k(t - t l ) ]
]
- erfc
- erfc
t G(z),
H(2n t 1) + z
(4kt)' I 2
H(2n +- 1) t z
(4kt)l12
for t > tl.
]t G(z),
- erfc
fort < tl
H(2n t 1) t z
[4k(t - t l ) ]
(4)
Intracratonic basin formation
7
Displacement of the Earth’s surface, h(t), due to expansion and contraction of the
lithosphere under the influence of the thermal event (neglecting sediment and water
loading), can be calculated from the integral:
h ( t)=
j,”.
[T(z, t ) - G(z>ldz+ h,,,
(5)
where 01 is the coefficient of thermal expansion, and hois the initial displacement. The model
at this stage can only account for subsidence in a thermally contracting basin after a
substantial thickness of uplifted crust has been eroded away. This is essentially the same
hypothesis proposed by Sleep (1971).
Anomalous heat flow A q at depth z in the lithosphere during the heating phase ( t <t l )
can be obtained by the equation:
q z , t )= K
a
T(z, t),
az
where K is thermal conductivity. Combining equations (3) and (6), and performing the
differentiation :
-
+ exp [I- [H(2n + 1) + zI2/4kt]l}.
(7)
Movement of the metamorphic facies boundary
-4s the geotherm in the crust increases due to heating from the thermal event, the
metamorphic facies boundary (described by equation 2) moves closer to the Earth’s surface.
This movement can be determined from the assumptions concerning the metamorphic
nature of the crust and the temperature increase due to the thermal event described above.
The initial depth of the facies boundary z’(t=O) is the intersection of the initial
geotherm G(z) (see Fig. 4 ) with the facies boundary curve (equation 2). The depth of the
facies boundary z’(t) during the heating phase is the intersection of the geotherm (defined
by equation 4 ) with the curve described by equation (2). Retrograde metamorphism is
assumed not to occur, and for times greater than t l the depth of the phase boundary is taken
as z’(tl).
The thickness of the greenschist facies rocks metamorphosed to amphibolite facies rocks,
L(t), can be written:
L(t)=z’(O)- z ( t ) .
(8)
Fig. 6 shows the behaviour of L ( t ) with time for thermal events entailing Toof 100,200 and
3OO0C in a 100 km thick lithosphere with thermal diffusivity of 32 km2 Myr-l; G ( z ) and the
facies boundary are defined in Fig. 4 and equation (2), respectively. Fig. 7 shows the rate of
change of this thickness with time, A L / A t , determined numerically using finite differences.
Given that the densities of greenschist and amphibolite are pg and pa, respectively, then
the Earth’s surface experiences a downward displacement, Z(t), according to the equation:
Z(t)=
Pa-pg
~
L (t)
(9)
pg
where L ( t ) is the thickness of the layer of amphibolite facies rocks metamorphosed from
greenschist facies rocks after time t.
8
M. F. Middleton
I
TIME,M Y
Figure 6. Thickness of the metamorphosed layer, L ( t ) ,versus time for To= 100, 200 and 300'C.
,
TIME,MY
Figure 7. The rate of change with time of the thickness of the metamorphosed layer, a L ( t ) / A t ,versus
time for To= 100,200 and 300'C.
Subsidence model
Subsidence will occur in two stages. Initially, subsidence will be due to deep crustal metamorphism, and will continue as long as crustal temperatures are increasing. Then, subsidence
due to thermal contraction of the lithosphere will occur for times greater than t l .
Equation (9) describes the downward movement of the Earth's surface, Z(t), due to
metamorphism during the heating stage of the thermal event. Offsetting this, however, is the
tendency for uplift expression in equation (5). Total displacement of the Earth's surface
Intracratonic basin f o m t i o n
9
(i.e. the initial surface when t = 0), taking these two factors into account and neglecting
erosion, is
for h(t) > Z ( t )
(i) S ( t ) = Z(t) - h (t),
(ii) S ( t ) =
Pm
~
Pm - P s
[l(t)- h (t )] ,
for h ( t ) < Z(t)
where negative S ( t ) is uplift, and pm is the density of the mantle, p s is the density of
sediment, the ratio Pm/(pm - pS) is the isostatic loading factor for deposition occurring to
sea level, and h(t) is the displacement described in equation (5).
Subsidence due to thermal contraction can be obtained from equation (5) and is given by
where pm,ps and h( t ) are as defined previously, It is assumed that the elevation is at sea level
for the commencement of this stage of subsidence, which is to be expected if it is preceded
by subsidence due to deep crustal metamorphism.
Therefore, movement of a datum (the Earth’s surface at t = 0) for times less than t l
is described by equation (lo), and the movement for times greater than tl is described by the
sum of equations (10) and (1 1). Instantaneous deposition to sea level is assumed to occur
after a short initial period of uplift during which erosion is assumed negligible.
Fig. 8 shows this movement with time for a metamorphic density contrast, (pa - p g ) / p g ,
of 0.1 1 and To of 100,200 and 300°C with the parameters H = 100 km,k = 32 km2 Myr-l,
pm = 3.3 g CIII-~, p s = 2.4 g ~ m - 01~ =, 3 x
OC-’
and t l = 100 Myr; L ( t ) , which gives Z(t)
via equation (91, are plotted in Fig. 6 . Similarly, Figs 9 and 10 shows subsidence with time
for (pa - p J p , of 0.09 and 0.07, respectively.
I
I
1-
E
Y
-
u
Y
Lo
2-
P,
-P
4
= 0.11
3-
I
0
time,
I
I
!
I
100
I
MY
I
I
I
I
I
2 00
Figure 8. Movement of a datum (Earth’s surface for t = 0) versus time for To of 100,200 and 300°C with
( p a - pg)/p, of 0.11. Erosion is neglected during the period of uplift which occurs for the initial
20-30 Myr.
10
M.F. Middleton
t i m e , MY
I
,
I
Figure 9. Movement of a datum (Earth’s surface for t = 0) versus time for Toof 100,200 and 300°C with
(0, - P g ) / p g of 0.09.
Figure 10. Movement of a datum (Earth’s surface for t = 0) versus time for To of 100, 200 and 300°C
with ( p a - p g ) / p g of 0.07.
The subsidence models of Figs 8-10 show continuous two stage subsidence similar to
that observed in the earlier mentioned intracratonic sedimentary basins. The larger the
thermal event (To) the greater the magnitude of subsidence in the second phase, which is
essentially independent of the metamorphic density contrast (pa - p g ) / p g .The magnitude
of the metamorphic driven subsidence is strongly dependent on the metamorphic density
contrast. Clearly, from Fig. 10, metamorphic subsidence is very small for crusts with the low
density contrast of 0.07 across the greenschist-amphibolite facies boundary. Of course, if
Intracratonic basin formation
11
no contrast exists, that is, if the crust is essentially all amphibolite facies or higher grade
metamorphics, no subsidence will occur.
The present method of determining subsidence is subject to the condition that the
thermal regime in the vicinity of the facies boundary is not significantly perturbed by the
energy losses due to the metamorphic reactions. It is necessary and sufficient to demonstrate
that
(i) the energy supplied by the thermal event is sufficient to drive the metamorphic
reactions, and
(ii) the energy loss as a consequence of the reactions does not alter the thermal gradient
in the vicinity of the facies boundary.
The endothermic effects of metamorphism
The endothermic effects of the greenschist facies to amphibolite facies metamorphism
can be simply demonstrated to have negligible effect on the subsidence model.
Turner (1968, p. 381) has suggested that the heat energy necessary to convert greenschist
facies rocks to amphibolite facies rocks by progressive dehydration is between 30 to 40
cal g-’. For the present demonstration, 40 cal g-’ will be assumed appropriate for the energy
absorbed during metamorphism.
E N E R G Y L O S S V E R S U S E N E R G Y SUPPLIED
An estimate of heat flow loss across the metamorphic boundary can be made based on the
relation (Carslaw & Jaeger 1959, p. 284):
where K 1 = aTl/az and K 2 = aT21az are the heat fluxes across the phase boundary from the
lower medium and upper medium, respectively, p is the mean density of the media involved
in the transformation, dL/dt is the rate of movement of the phase boundary, and Q is the
amount of heat absorbed (cal g-’). K 1 and K 2 are approximately equal. The heat loss per
unit time is equal to the right hand side of equation (12).
The heat supplied can be determined from equation (7). The difference
where Z’ is the depth of metamorphism at time t and L ( t ) is the thickness of the metamorphosed layer at time t , indicates by its sign if sufficient heat is available for total
metamorphism as a function of time. Negative values of equation (13) indicate no metamorphism or partial metamorphism, and positive values indicate total metamorphism. Fig.
11 shows a plot of equation (13) versus time for To of 100, 200 and 3OO0C with L ( t )
[-z’(O) -z’(t)] and A L ( t ) / A t being given in Figs 6 and 7, respectively. The initial depth of
the facies boundary z’(0) is 24 km (see Fig. 4).
Sufficient heat (positive value in Fig. 11) is available to produce total metamorphism for
times greater than 30 Myr after commencement of the thermal event (t = 0). As subsidence
due to deep crustal metamorphism does not occur until after 30 Myr from the commencement of the thermal event (see Figs 8-10), the energy condition that sufficient heat be
available is accommodated.
M. F. Middleton
12
L
/
0
- 0.0 u05
30
Figure 11. A plot showing the difference in heat energy available, Aq, and heat energy absorbed, Q B i ,
at the moving greenschist-amphibolite metamorphic facies boundary versus time for To of 100, 200
and 300°C. No metamorphism or partial metamorphism occurs for times less than 30 Myr when insufficient heat is available for all metamorphic reactions to occur; for times greater than 30Myr,
sufficient heat is available for all reactions necessary for totrl metamorphism of greenschist facies rocks to
amphibolite facies rocks.
P E R T U R B A T I O N O F T H E GEOTHERM
Carslaw & Jaeger (1959, p. 293) have derived an expression for temperature perturbation,
AT(z, t ) ,due to an instantaneous plane heat ‘sink’ (or source) moving at a rate i ( t )= dL/dt,
and absorbing heat at a rate QDL (t):
AT@, t ) =
Q
t
L(t)dt‘
2~(nk)”’/~ (t - t’)’j2
- exp [-
(z +Z’)’/4k(t
{exp [-(z -z’)2/4k(t - t‘)]
- t’)]]
,
(14)
where t is the time after which the heat sink began to move, z’ is the depth to the sink at
time t , z is the depth of observation and c is the specific heat of the metamorphosed
medium. The rate L ( t ) is defined for the present study by the curves in Fig. 7, and the
integration in equation (14) can be performed numerically using the i ( t ) functions in that
figure.
Fig. 12 shows the temperature perturbation AT(z’, t ) at the facies boundary versus time
for a thermal event with Toof 2OO0C,i(t) defined in Fig. 7, z’=z’(O)-L(t) (L(t)from Fig.
6 ) , Q of 40 cal g-’, c of 0.25 cal g-’ ‘C-’ and k of 32 km2Myr-’. Clearly, the perturbation
to the geotherm, which is less than 0.04”C, is insignificant.
Therefore, the energy exchange at the facies boundary during metamorphism has
negligible effect on the above method for determining L ( t ) ,and consequently the subsidence
W).
Intracratonic basin f o m t i o n
13
To = Z O O 0
.04/
Figure 12. The temperature perturbation, AT(z’, t ) , at the moving greenschist-amphibole metamorphic
facies boundary versus time for Toof 200°C. The disturbance to the temperature field is negligible.
Conclusions
The present model of intracratonic basin formation (Figs 8-10) adequately describes the
two phase subsidence observed in some American and Australian intracratonic basins (Figs
1 and 2). Also, according to the model, heating in the absence of unmetamorphosed deep
crustal rocks will cause uplift alone.’ If insufficient metamorphism occurs to override uplift
due to pre-existing high metamorphic grades,; then no subsidence will occur; under these
conditions, sufficiently high Tomay cause plateau uplift.
Flexure of the lithosphere has been neglected in this model; Haxby et al. (1976), and
Sleep & Snell(l976) have treated the problem of unification of flexural and thermal models.
Nevertheless, point-wise Airy isostasy is an adequate assumption in many young basins.
A finite difference approach involving equations (2), (4), (13) and (14) may be used to
find the thickness of the metamorphosed layer L(t). However, the method described herein,
which neglects the endothermic effects of metamorphism (equations 13 and 14) in
determining L ( t ) , is considerably simpler, more rapid, and suitable for use on a programmable calculator. Moreover, the analysis has shown that the endothermic effects are
essentially negligible, and equations (1 3) and (14) are not necessary to estimate L (t).
Although this hypothesis of intracratonic basin formation is based upon the highly
idealized concept of an ‘artificial’ greenschist-amphibolite facies boundary, model curves
are sufficiently concordant with observational subsidence data from a number of basins to
indicate that deep crustal metamorphism may be a viable mechanism in the formation of
many basins. More precise definition of specific changes in particular basins, as well as
eustatic cycles, is necessary before a confident match can be made between deep crustal
metamorphic models and observed subsidence data in any particular locality.
Acknowledgments
Drs D. A. Falvey, G. H. Packham and D. W. King of the University of Sydney and Dr M. G.
Langseth of the Lamont-Doherty Geological Observatory of Columbia University offered
useful suggestions in the preparation of this manuscript.
References
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of rifting,J. geophys. Res., 76,1197-1211.
M.F. Middleton
14
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