1. metamorphic Ceol., 1995, 13, 41 9-430
Thermal perturbation during charnockitization
and granulite facies metamorphism in southern India
J . G A N G U L Y , ’ R. N. SlNGH’ A N D D. V. R A M A N A ’
I Department of Geosciences, The University of Arizona, Tucson, AZ 8572 I , USA
’National Geophysical Research Institute, Hyderabad 500 007, India
ABSTRACT
We have deduced the steady-state lithospheric geotherm at c. 1 Ga in the south Indian shield area using
the available data on the concentration of radioactive elements, and the f-T conditions of Proterozoic
mantle xenoliths in the south Indian kimberlites as constraints. The geotherm was adjusted back to 2.5 Ga
by keeping the surface temperature constant and calculating the temperature change at the top of
convecting upper mantle. The reduced or mantle heat flux, which was treated as an adjustable parameter,
was 20.9-21.3 mW/m2 at 1-2.5 Ga. Comparison of the calculated steady-state geotherm with the available
f -T data of the Archaean (c.2.5 Ga) charnockites and granulites from southern India suggests that the
granulite facies metamorphism in this region had resulted from a major thermal perturbation, which was
c. 400” C at 25 km.
Seismic tomographic and gravity data essentially preclude any significant magma underplating of the
granulitic crust in southern India. Previous workers have suggested that the formation of charnockites in
this region was associated with copious CO, influx from a deep-seated source, possibly the mantle. In this
work, we have evaluated both the transient and steady-state thermal effects of the heat convected by CO,
outgassing from upper mantle. It is shown that the thermobarometric array of charnockites and granulites
can be produced by the convective perturbation of the steady-state geotherm, and that a flux of CO, of
290 mol/m2 yr (corresponding to Darcy velocity of 20.30 cm/yr) for a period of 1 3 0 Ma was needed to
produce the required perturbation. This is c. 150 times the average CO, flux through the tectonically
active area of the Earth’s crust at the present time. There is, however, an uncertainty of a factor of 3 in
this value.
Seismic tomographic and gravity data independently suggest thickening of the crust beneath the
granulite terrane compared with the adjacent Dharwar craton. This suggests thermal perturbation due to
overthrusting as a major potential cause for the granulite facies metamorphism in south India.
Overthrusting of a 30-35-km-thick thrust block was needed to produce the required thermal effect. The
estimated thickness of the original crust from geobarometric and seismic tomographic data south of the
orthopyroxene isograd or ‘transition zone’ is compatible with the emplacement of a thrust block of this
magnitude. However, the latter fails to match the estimated pre-uplift crustal thickness at the transition
zone, if it is assumed that the crust has not thinned by non-erosional processes since the Archaean. Thus,
we propose a combination of overthrusting and CO, fluxing from a deep-seated source as the cause for
the formation of charnockites in this zone. The required focusing of CO, in this case is c. 40% of that
estimated in the model where CO, fluxing was considered to be the sole reason for thermal perturbation.
This combined thrusting-CO, fluxing model also helps explain the development of patchy charnockites
in the transition zone from amphibolite facies rocks.
Key words: Archaean; charnockites; Darcy velocity; geotherm; Proterozoic.
INTRODUCTION
Charnockites have been a subject of interest among
petrologists and geochemists since they were first described
by Holland (1900) as orthopyroxene-bearing rocks of
broadly granitic chemistry. Subsequent workers have
included closely associated orthopyroxenc-bearing rocks,
but of somewhat more basic chemistry, into the
charnockite series. In southern India, an orthopyroxene
isograd (Fig. 1) defines a continuous prograde southward
transition from amphibolite to granulite facies meta-
morphism (Pichamuthu, 1953, 1965; Hansen ef al., 1984a)
that culminated in late Archaean time at 2500-2700 Ma
(Ramiengar et al., 1978). Janardhan et al. (1982) and
Hansen et al. (1984a,b, 1987) have analysed the P-T and
fluid conditions attending the formation of charnockites.
They concluded that the charnockites formed under
relatively anhydrous conditions (fHz0
< 0.3f,,,.,) and that
the flow of CO, from either the mantle or lowermost crust
was responsible for the depletion of H,O. Although the
relatively anhydrous condition of the south Indian
419
420
I
J . CANCULY E T A L
dharwar Craton '
16'
14'
12'
Closepet
Granite
10'
Dharwar
I
76' E
Charnockite
80'
8"
84'
Fig. 1. Map of the south Indian shield (adapted from Janardhan el
ol., 1982) showing the distribution of major igneous and
metamorphic rock types along with the location of the kimberlite
pipes (P). Dhanvar craton and deep vertical shears postulated by
Drury & Holt (1980). K and N stand for Kabbaldurga and Nilgiri,
respectively. Unpatterned area represents Peninsular gneiss.
Numbers on the map are palaeodepths exposed, as deduced from
geobarometry.
charnockites, and of granulites in general, seems to be
incontrovertible, the nature of the geological process
leading to the H,O-depleted condition has been a topic of
lively debate (e.g. Lamb & Valley, 1984, 1985;
Bhattacharya & Sen, 1986; Vielzeuf & Vidal, 1990).
It is generally accepted that the P-T conditions of
granulite facies rocks, as inferred from phase equilibrium
data, do not represent conditions on the ambient
steady-state geotherm, but reflect transient conditions
resulting from thermal perturbation. In order to develop a
proper understanding of the relative importance of the
various geological processes in producing a thermal
perturbation, it is first necessary to quantify the extent of
the thermal perturbation in specific cases, and then
evaluate the quantitative requirements on the potential
geological processes to bring about the required thermal
effect. The purpose of this paper is to evaluate (a) the
probable magnitude of thermal perturbation during
granulite facies metamorphism in southern India, and (b)
the potential role of the different geological processes in
producing the required effect.
THERMAL PERTURBATION D U R I N G
G R A N U L I T E FACIES M E T A M O R P H I S M
IN S O U T H E R N I N D I A
In this section, we will try to constrain the magnitude of
thermal perturbation during granulite facies metamorphism in southern India on the basis of the available data on
the P-T conditions of these rocks, and the inferred
steady-state geotherm at c. 2.5 Ga, which is the time of the
granulite facies metamorphism. Ganguly & Bhattacharya
(1987) determined the P-T conditions of equilibration of
mineral assemblages in peridotite xenoliths in the
Proterozoic kimberlite pipes in southern India dated at
1023 f 40 Ma (Paul et al., 1975; Basu & Tatsumoto, 1979).
Below we first deduce the steady-state geotherm at c. 1 Ga
using the P-T array of these xenoliths as one of the
constraints. This geotherm is then adjusted back to 2.5 Ga,
accounting for the temperature decay from 2.5 to 1 Ga.
Assuming that the radioactive heat production rate
decreases exponentially with depth, Z, and that the
thermal conductivity, K, density, p , and the isobaric
specific heat capacity, Cp, of the rock are constant, the
steady-state geotherm is given by the solution of the
equation
dZT + HSe-(Z'D)
K=0
(1)
dZ2
subject to the condition that at 2 >> D, K(dT/dZ) = q m r
where q mis the conductive heat flux from the mantle, H, is
the volumetric heat production rate at the surface, K is the
thermal conductivity of the mantle rocks, and D is a
characteristic 'scaling depth', which usually varies between
7 and 11 km (Pollack & Chapman, 1977). Integration of
Eq. (1) yields
where T(Z,) and T(Z,) are the temperatures at the depths
Z, and Z,, and K, is the constant thermal conductivity in
this depth interval. In any given region, the mantle heat
flux and the scaling depth are usually derived from the
'Birch-Lachenbruch' (B-L) relation, i.e. the observed
linear relation between the surface heat flux, q,, and the
volumetric heat production rate, H,, of the surface rocks.
The slope of the curve is interpreted to be equal to D ,and
the intercept at H, = 0, which yields a reduced heat flow,
q,, is interpreted to be the mantle flow, qm (see, for
example, Turcotte & Schubert, 1982).
Heat flow vs. heat production data for the Indian shield
are still quite scanty, with heat flow data from only three
localities (Kolar, Karadikuttam and Sargur), which have
been recently summarized by Gupta et al. (1991). The
mean surface heat production rate of these three areas is
1.75 X
W/m'. Gupta et al. fitted the q s vs. H, data to
derive D = 11.5km and q,=23mW/mz. Using the
concentration of radioactive elements in the Kolar gold
field, as reported by Gupta el al. (1991), the average decay
constant (A) of the radioactive elements is 2.38 X
IO-'OW/yr, which yields H, (1 Ga) = 2.22 X
W/m'
Schatz & Simmons (1972) presented in a graphical form
the thermal conductivity data along an average continental
shield geotherm. We have fitted these data to an analytical
form to obtain the following expression.
K=2.5W/m
at Z ~ 3 0 k m ,
and
K = 3.10 - 1.40 X 10-,(2) + 2.18 X 10-4(Z2)
- 5.34 X 10-'(Z3) at 30 5 Z 5 200 km.
T H E R M A L PE R T U R BAT10 N
The geotherm calculated numerically from Eq. (2) and
the above data, along with the constraint that T
( Z = O ) = Ts=298K, does not fit the array of P-T
conditions defined by the mantle xenoliths. For example,
at 49.5 kbar, the temperature on this geotherm is c. 100" C
higher than that defined by the mantle xenoliths. Thus, qm
was treated as an adjustable variable, so that the calculated
geotherm matches the P-T array of the xenoliths. The
results of the calculation, which correspond to qm=
20.92 mW/m2, are illustrated in Fig. 2. KZ was taken to be
the average value of K in the interval between 2, and 2,
with Z2- 2, = 5 km. In relating pressure to depth, pg was
varied as a function of depth as follows:
For 2 5 30 km, pg = 0.294 kbar/km;
for 30 < Z
5
100 km, pg = 0.285 + 2.86 X 10-4(Z)kbar/km,
and
100 < 2 5 200 km, pg = 0.304+ 1 X 10-4(Z) kbar/km.
These relations satisfy the values of pg at 100 and 200 km
depths given in Turcotte & Schubert (1982), and the mean
density of crustal rocks.
Turcotte & Schubert (1982, eqn. 7-208) derived a
relation for the cooling of a convecting upper mantle that
yields dT/& = -0.210(10-')T2 K/Ma. Inasmuch as the
lithosphere represents a thermal boundary layer above a
convecting mantle, the base of the mantle lithosphere
during Proterozoic time must lie below the maximum
depth (c. 185 km) recorded by the mineral assemblages in
the Proterozoic mantle xenoliths, since the P - T array of
these xenoliths reflects a highly superadiabatic gradient.
We assume that this depth constraint on the base of the
1500 1
\
A
0
Proterozoic Geotherm
W
( 1 Gal
-I
500
D U R I N C C H A R N OC KIT IZ A T l O N
421
mantle lithosphere was also valid at c . 2.5 Ga. Thus, using
the above relation, and assuming that the top of the
convecting mantle was at c.200km depth during the
Archaean, the temperature change at c. 200 km depth
should be c. 80" C from 2.5 to 1 Ga. From the present
average radioactive heat production rate and average
decay constant of radioactive materials in the near-surface
rocks of the Indian shield, as discussed above,
Hs(2.5 Ga) = 3.2 X
W/m3. Keeping & = 298 K, this
value of pH, requires q,=21.34mW/m2 to produce a
geotherm which is c. 80" C higher than the Proterozoic
geotherm at c. 200 km depth.
The geotherm calculated with the above values of &, Hs
and qm, which is illustrated in Fig. 2 as a dashed line, is
taken to represent the steady-state ambient geotherm at
c. 2.5 Ga. Also shown in the figure are the P-T conditions
of charnockites, as determined by Janardhan et al. (1982),
Hams et al. (1984) and Hansen el al. (1984a,b). From
comparison of the P-T array of the granulite facies rocks
with the inferred steady-state geotherm during the
Archaean, it is evident that these rocks in south India must
have been associated with a major thermal perturbation,
which was c . W C at c.20km depth (corresponding to
c. 6 kbar pressure).
E V A L U A T I O N OF P O T E N T I A L CAUSES
OF T H E R M A L P E R T U R B A T I O N
The thermal perturbation during the granulite facies
metamorphism in southern India could have been due to
one or more of the following reasons: (a) convective heat
flux due to CO, streaming from a deep-seated source
(Harris et al., 1982; Chacko et al., 1987; Ganguly & Singh,
1988; Dunai & Touret, 1993), (b) tectonic processes (e.g.
England & Thompson, 1984; Connolly & Thompson, 1989;
Eckert, 1989) and (c) underplating of basic magma
(Lachenbruch & Sass, 1977; Bohlen & Mezger, 1989).
There is no evidence of 'magma fluxing', and as we show
below, crustal extension is also not a satisfactory model in
this region. We explore below the potential roles of the
above geological processes in terms of their quantitative
requirements, which would provide a framework for the
understanding of the reason for thermal perturbation
during granulite facies metamorphism in southern India.
co, flux
50
100
2po
150
km
, , , , , , ' , , , 1 1 1 1 1 ' , " ' , , , , ! , , , , , , , , ~ , , , , , , 1
0
20
40
60
80
P(kbar)
Fsg. 2. Steady-state lithospheric geotherms in the south Indian
shield at c. 1 Ga (solid line) and c. 2.5 Ga (dashed line). Crosses
and circles represent the P-T conditions of Proterozoic (c. 1 Ga)
mantle xenoliths and charnockites/granulites (c. 2.5 Ga),
respectively.
The notion of convective heat transfer by CO, as the cause
of temperature rise during metamorphism probably dates
back to the work of Schuiling & Kreulen (1979), who
suggested this effect to explain the thermal dome at Naxos,
Greece. In a comprehensive thermobarometric and fluid
inclusion study across an amphibolite-granulite transition
in West Uusimaa, SW Finland, Schreurs (1984) also
suggested that heat convected by CO, was responsible for
the nearly isobaric but sharp temperature rise associated
with the transition. Harris et al. (1982) have discussed the
422
J . CANCULY E T A L .
existence of crustal structures in southern India which
could facilitate fluid access to shallower levels of the crust.
From LANDSAT imagery and reconnaissance structural
analysis of the south Indian craton, Drury & Holt (1980)
also deduced the existence of a series of N-S shears prior
to the formation of charnockites. According to them, the
development of the pyroxene assemblages in high-grade
terranes is related to these shear belts, perhaps through the
influx of CO,. However, although the potential thermal
effect of CO, streaming from deep crustal or upper mantle
levels has been alluded to, there has been no attempt so
far to evaluate quantitatively the magnitude of CO, flux
that is required to produce the observed thermal effect.
Solution of energy balance equation
The magnitude of the fluid flux required to produce a
thermal perturbation can be calculated from the solution
of the energy balance equation. Assuming that the
radioactive heat production decreases exponentially with
depth, and the thermophysical properties of the rock and
fluid are constant, the energy balance equation can be
written as follows for a one-dimensional case
-dT
d2T
aT
pC, - = K , -7+ (pCPV4)(- + (1 - 4)Hse-"'D
at
az
az
(3)
where the subscript 'f' stands for fluid, K, is the effective
thermal conductivity of the rock with pore fluids, V is the
true velocity of the infiltraticfluid (positive upwards), 4
is the rock porosity, and p C p represents the weighted
average of volumetric heat capacity of rock and pore fluid.
The quantity V 4 is the volumetric fluid flux (i.e. fluid
volume per unit area per unit time, cm3/cmzs=cm/s) or
the so-called Darcy velocity. The fluid flow during
charnockite metamorphism was probably driven primarily
along fractures (Drury & Holt, 1980). However, as
discussed by Brady (1988) and Hoisch (1991), the thermal
consequences of flow along fractures are similar to those of
pervasive or percolating flow unless the fractures are
widely spaced.
The (C,,), and fluid flux ( p V 4 ) are assumed to be
constants. According to the continuity equation, the
second assumption implies that the fluid density is
independent of time. This simplifying assumption was also
made in several earlier works dealing with fluid flow (e.g.
Norton & Taylor, 1979, Brady, 1988). With regard to these
assumptions, we note from the data of Bottinga & Richet
(1981) that within the range of P-T conditions
encompassing those experienced by CO, rising from a
condition of 30kbar, 1200K to the crustal levels of
charnockite metamorphism, p and C, of CO, vary
approximately in the range of 0.93-1.53 g cm- and
0.37-0.36 cal g - ' deg-', respectively.
The transient and steady-state solutions obtained from
Eq. (3) are given in the Appendix. The initial condition for
the transient problem is defined by setting V = 0 (i.e. by
Eq. l ) , whereas the boundary conditions are taken to be
T=T,
atZ=0,
(4)
and
T=TL
at Z = L .
Instead of the second boundary condition, an alternative
boundary condition was used that at Z = L, the conductive
heat flux (K,dT/dZ) is constant and given by qm.
However, Brady (1988) has demonstrated that this
boundary condition is not geologically reasonable since it
leads to very high temperature at Z = L even for very
moderate values of conductive heat and fluid fluxes.
Figure 3 shows the calculated temperature change as a
function of time and the Darcy velocity of CO, at depths
of 10 and 20 km, using & = 298 K, and L = 90 km so that
TL =850°C (Fig. 1). Model calculations by Tajika &
Matsui (1993) suggest that the average surface temperature
of the Earth changed very little in the last 3Ga. K, is
taken to be 2.5 W/m K, which is the value of thermal
conductivity of rocks at a depth of c.25km (Schatz &
Simmons, 1972). In the presence of pore fluid, the effective
thermal conductivity should be somewhat higher and the
Darcy velocity would increase in the same proportion since
it vanes directly with K, (see below). In order to develop a
feeling for the probable effect of pore fluid on K,, we note
that the thermal conductivity of marble with about 1%
porosity increases by 11% when it is soaked in water
(Clark, 1966). The porosity of mantle rocks under mantle
conditions is definitely much less than 1%. Thus, we
believe that the Darcy velocities cannot be significantly
higher than those illustrated in Fig. 3 due to the effect of
intergranular fluid on the thermal conductivity. It is
evident from Fig. 3 that if the inferred thermal
perturbations were solely due to the effect of heat
convected by CO, degassing from the mantle, then a
volumetric flux of >0.50cm/yr is required for a period
530Ma. If the V + were higher, then the time and the
time-integrated flux required to achieve the required
thermal perturbation would have been lower.
Figure 4 shows that the array of P-T conditions of
charnockites can be produced by the effect of convective
heat transfer by C02. The perturbed geotherms were
calculated from the steady-state solution of Eq. (3) (see
Appendix), assuming that the CO, had degassed from a
depth range of 90-105km, and setting + = O . The
perturbed geotherms correspond to b = 7 (V4 =
0.3 cm/yr), where b IS a dimensionless parameter defined
as
L(pCPV4)"ui,
b=
(5)
K, .
If the fluxing of CO, were the sole reason for the thermal
perturbation, then it is evident from Fig. 4 that it must be
derived from a depth 290 km (P 2 28 kbar).
Constraint on CO, flux from Darcy 's law
The question that arises at this point is whether a
volumetric flux of c. 0.50 cm/yr is permissible by the
T H E R M A L PERT U R BAT I0N D U R I N G C H A R N OC K IT1Z A T I0N
423
1000
(a)
2 = lOkm
800
0.8
-
600
0.6
0.5
__________
0
0.4
Y
b-
0.3
400
0.2
0.1
200
0
I
I
I
I
I
10
20
30
40
50
Time
(b)
(Ma)
1000
Fig. 3. Thermal perturbation of the 2.5 Ga
steady-state geotherm (Fig. 2) at (a) 10 km
commonly accepted values of the rock permeability.
According to Darcy’s law of fluid flow through porous
medium (Hubbert, 1940),
where K is the rock permeability, q is the fluid viscosity,
and q5 is the fluid potential which decreases upwards. As
discussed by Walther & Orville (1982), for fluid ofconstant
density, the driving force of fluid flow in the vertical
direction is provided by the pressure gradient in excess of
the hydrostatic gradient, which is given by ( p , - p J g ,
where pr is the rock density and g is the acceleration due to
gravity. Thus,
and, consequently,
v4 = 4
P S
-P
17
k
(8)
Etheridge ef 01. (1983) concluded that metamorphic
permeabilities are in the range of
to 10-’*m2, or
possibly even higher. This is in accord with the recent
compilation of rock permeabilities at various scales of
measurement (laboratory, borehole, regional) by Clauser
(1992). Brady (1988) suggested that rocks fractured by
volatile release were likely to have permeabilities greater
than 10-IRm2. Thus, for the purpose of present
calculations, 10-l8m2 is used as a reasonable to
conservative estimate of rock permeability. The viscosity
values calculated by Walther & Orville (1982) suggest that
at the P-Tconditions of interest in this work, the viscosity
of C 0 2 lies between 0.15 and 0.10 centipoise. Thus, using
p . = 2.93 g/cm3, which represents the density of hypersthene granulite (Daly ef a!., 1966), pco, = 1.2 g/cm3, which
is the mean density of CO, at the P-T conditions of
interest, gives v+ = 0.4-0.5 cm/yr, which agrees with the
minimum value of v+ required to produce the desired
thermal effect (Fig. 3). Thus, we conclude that although
metamorphic permeabilities are poorly known, the
424
J . CANCULY ET A L .
and upper mantle conditions. (The heat loss due to viscous
dissipation is not likely to be very significant for a gaseous
phase.) Consequently, non-adiabatic decompression of
CO, from 90 to 20 km would release an amount of energy
equivalent to pC,AT, which is 594-790J/cm3 or
22-29kJ/mol. The amount of heat that would be
convected by CO, in the absence of the JT effect is given
by pCp(V4)(dT/dZ)At, which is c.43kJ/mol for Af
c . 30 Ma and v 4 = 0.50 cm/yr. These calculations suggest
that the JT effect associated with the irreversible
decompression of CO, can reduce the magnitude of
required fluid flux by a factor of c. 1.5. Thus, accounting
for the JT effect, the required volumetric fluid flux or
Darcy velocity should be ?0.30cm/yr for a period of
5 3 0 Ma.
/
/
/
/
-
-I
/
Granulites
/
/
Archean Geotherm
(2.5 Go)
Required focusing of CO, through the southern
Indian craton
Fig. 4. Steady-stateperturbation (dashed lines) of the ambient
geotherm (2.5 Ga) due to degassing of CO, from depths of
90-125 km. using B = 7 (see text, Eq. S ) , which corresponds to
V& = 0.3 cmlyr.
volumetric CO, flux required to produce the inferred
thermal perturbation is consistent with their commonly
accepted values.
Thermal effect of irreversible decompression of CO,
In the above calculation, we have neglected the
temperature change associated with the irreversible
decompression of CO,. Neglecting the effect due to viscous
dissipation, the temperature change associated with
irreversible and adiabatic decompression of a substance,
which is an isenthalpic process (dH=O), is given by the
well known Joule-Thompson (JT) relation. In its usual
form the J T relation ignores the effect of work done in
raising a substance against a gravitational force (e.g.
Denbigh, 1971). The latter effect was incorporated by
Ramberg (1972) to derive the following relation:
(9)
where g is the acceleration due to gravity, h is the height
and a is the coefficient of thermal expansion.
Spera (1981) has shown that the term Ta < 1 for CO, at
P 2 2 kbar and T = 900-1300" C, so that the temperature
of CO, gas increases if it undergoes irreversible adiabatic
decompression, in the absence of a significant change in
gravitational potential, under deep crustal and mantle P - T
conditions. For CO, ascending from the mantle, the
magnitude of the temperature increase would be
moderated by the second right-hand term in Eq. (6) since
dh/dP<O. Using the thennochemical data of C02 from
Bottinga & Richet (1981), there would be a net increase in
the temperature of CO, at a rate of c. 15-20" C per kbar of
decompression, if it ascends adiabatically in deep crustal
From the above analysis we conclude that (a) the
magnitude of fluid flux required to produce the observed
thermal effect seems permissible by the commonly
accepted values of rock permeability, and (b) allowing for
JT effect, the required flux ( p V 4 ) of CO, should have
been 20.40 g/cm2yr or 290 niol/m2yr for t 5 30 Ma. The
total exposed area of charnockite and granulite in southern
India and Sri Lanka is around 1.5 X 10skm2. If all these
rocks were produced by perturbation of the ambient
geotherm by the thermal effect of CO, degassing From
around 90km depth, then a total amount of CO, of
c. 18 X 10,' g (or c . 4 X 10'" mol) would have been required
for the limiting values of V + = 0.30 cm/yr and I = 30 Ma.
From a survey of the reported C 0 2 inputs into the
atmosphere from the Earth's interior, Bredehoeft &
Ingebritsen (1990)had taken a value of 3 X lo', mol/yr as
the present global flux of CO, from the Earth's interior.
There is an uncertainty of at least factor of 3 in this value.
However, they pointed out that this flux is focused through
only c. 1% of tectonicaJly active areas of the Earth. Thus,
on average, 0.6mol/m2yr of CO, is being outgassed
through 1% of the Earth's surface area, which are in the
tectonically active zones, as compared to c. 90 mol/m2 yr
calculated for the southern Indian shield. Thus, the
required CO, flux through the south Indian shield area is
c. 150 times higher than the present-day average CO, flux
through 1% of tectonically active areas. There is an
uncertainty of a factor of 3 in this value arising from that in
the estimate of present-day flux of CO,. Des Marais (1985)
suggested that the CO, flux during the Archaean was 3-6
times higher than the present-day flux. Thus, the required
focusing of CO, through the south Indian shield could
have been significantly lower than that calculated above.
We cannot, however, be more specific than this since we
have no idea of the total surface area of the Earth through
which the CO, outgassing took place during the Archaean.
Assuming that the carbon concentration of the magma
erupting at the ridges represents the carbon concentration
of the entire upper mantle, Des Marais (1985) estimated
the present carbon inventory in the upper mantle to be in
::
T H E R M A L PERTU R BAT ION DU RI NC C H A R N O C K l T lZ A T l O N
the range of 2.1 X 1OZ2-4.2X lou mol, and suggested that
between 0.9 and 14 times the present mantle inventory of
C 0 2 might have outgassed during the Archaean. This
implies that c.0.01-2% of the total C 0 2 that was
outgassed during the Archaean had to be focused through
southern India if the convective heat transfer by C 0 2 alone
were responsible for the inferred thermal perturbation.
We feel that the above calculations of the required
focusing of C02 through the Indian a a t o n to produce the
inferred thermal effect are too uncertain to constitute a
viable test for the validity of the CO, fluxing hypothesis
being the primary cause of the thermal perturbation, which
was required for the granulite facies metamorphism in
south India. We have set forth the quantitative
requirements which would provide the framework for
future work aimed at resolving this issue. The proper role
of CO, fluxing may, however, be better understood after
the other potential causes of the thermal perturbation have
been explored.
Stability and source of COz-rich vapour
in the Earth's mantle
The concentration of C02 in the vapour phase in
equilibrium with mantle peridotite lacking clinopyroxene is
governed by the reaction
MgCO,
425
Ambient Geotherm
1H)o
-
L 1000
!
3
"
c
;
900
800
I
I
10
20
X
I
I
30
40
co2
P. kbar
Fig. 5. Comparison of the equilibrium P-Tconditions of the
reaction Mg2Si0., + CO, = MgCO, + MgSiO, as a function of
vapour composition in the H,O-CO, binary system with the
2.5 Ga ambient geotherm in the mantle lithosphere in the south
Indian shield. Inset: displacement of the equilibrium pressure of
the above reaction as a function of Xco2 at 800" C. For = forsterite
(Mg,SiO,); Mgs = magnesite (MgC03);X = C 0 2 / ( C 0 2+ H,O).
+ MgSiO, = Mg2Si04+ CO,.
(a)
In the presence of clinopyroxene, the appropriate reaction
is
CaMg(CO& + 4 MgSiO, = 2MgzSi0,
+ CaMgSi,O, + 2C02,
(b)
which must lie at a lower pressure than (a).
The pure end-member equilibrium (a) was determined
by Haselton ef al. (1977) at high P-T, which is illustrated
in Fig. 5. Assuming that the solid volume change, AV,, of
the pure end-member reaction is essentially constant, the
displacement of the equilibrium pressure of reaction (a) as
a function of the concentration of CO, in the vapour phase
can be determined according to the following relation
(Ganguly & Saxena, 1987, eqn. 4.33):
PAVZ 3 PAV;
T, X , ) - Inf~o,(P", T)] (10)
Here Po is the equilibrium pressure at the temperature for
the end-member reaction, P is the equilibrium pressure at
the same temperature when the vapour phase has a
composition X , , gotis the fugacity of pure CO, at P", T,
and fco2 is the fugacity of CO, at P, T, X,.
We assume that the vapour phase is essentially a binary
mixture of H 2 0 and CO,. Using the fugacity data of C 0 2
in the pure state and in the H 2 0 - C 0 2 mixture at high P-T
conditions (10-40 kbar, 800-1200" C) from Saxena & Fei
(1987), and V , at 1bar, 298K from Robie ef al. (1979) to
calculate AK, we have determined the displacement of
equilibrium (a) as a function of CO, concentration in the
vapour phase. These results, which are illustrated in Fig. 5 ,
- RT[lnfc~#',
show that a C0,-rich vapour phase would lead to the
formation of carbonate at the expense of forsterite at the
P-T conditions of the upper mantle in southern India,
which is incompatible with the fact that primary carbonates
are extremely rare in the mantle xenoliths.
The above result would seem to refute the hypothesis
for the origin of charnockite and related granulite facies
rocks in southern India by the dehydration and thermal
effects of CO, degassing from the mantle. However, it has
been shown recently by controlled decompression
experiments (Canil, 1990) that carbonates formed in the
mantle would 'rarely survive transport to crustal levels.
Carbonates coexisting with clino- and orthopyroxenes did
not survive even at the fastest decompression rate
investigated, which was 90km/h. In comparison, the host
magmas that transported the mantle xenoliths (kimberlites,
alkali basalts) are believed to have ascended at rates much
less than 70km/h (Mercier, 1979; Dawson, 1980 Spera
1980; Mitchell, 1986). Ganguly (1982) calculated that a
clinopyroxene crystal in a kimberlite xenolith from South
Africa cooled from 600" C to the emplacement temperature of 300°C in a period of c.22 days. The maximum
transport distance of the xenolith through this temperature
interval can be calculated by assuming that it moved
essentially along an adiabat (c.OSO/km). This results in a
distance of <600km, and consequently an ascending
velocity of (1 km/h, at least at T ~ 6 0 0 ° C . Primary
carbonates may be preserved at the ascending velocities of
mantle xenoliths when they form sufficiently coarse
aggregates (compared with those in experimental charges),
426
J . CANCULY E J A L .
as suggested recently by lonov ef al. (1993) to explain the
preservation of primary carbonates in peridotite xenoliths
from Spitsbergen.
Although the scarcity of primary carbonates in the
mantle xenoliths is not incompatible with the presence of a
C0,-rich fluid in equilibrium with a mantle peridotite, it is
important to address the source of such fluid in any model
of the formation of granulites through CO, streaming from
the mantle. Hansen et al. (1987) have discussed some of
the possible sources of CO, in the deep crustal and mantle
environments that might have flushed H,O out of the rocks
during granulite facies metamorphism. Newton (1988)
suggested outgassing of mantle hot spots for the origin of
heat-transporting mantle fluids. Green & Wallace (1988)
demonstrated experimentally that a carbonatite melt
occurs as a very small melt fraction in equilibrium with
pargasitic lherzolite at a range of upper mantle P-T
conditions. At P < 21 kbar and T = 1OOO-1100” C, this
melt gives off a fluid consisting of more than 80% CO,, if
the fluid were restricted to the binary C0,-H20 system.
Saxena (1989) showed that a vapour phase rich in CH,
could be in equilibrium with peridotites at the upper
mantle conditions. Green ef af. (1986) suggested a scenario
in which deep lithospheric fractures, such as which might
have guided the volatile release from the mantle during
charnockitization (Drury & Holt, 1980), would release
mantle volatiles rich in CH,, that would then interact with
oxidized crustal fluids to form CO,. Hopgood & Bowes
(1990) suggested crystallization of mantle-derived basic
magma as a source of CO, that might have been
responsible for the formation of patchy charnockites
associated with knuckle-shaped symmetrical folds in the
Proterozoic gneiss complex of Lake Baikal. However, as
discussed below, this model may not be viable for south
Indian granulites.
Finally we note that the presence of Proterozoic
kimberlites in the south Indian craton (Ganguly &
Bhattacharya, 1987) provides unequivocal evidence for the
presence of C-rich mantle fluid. Jackson ef al. (1988)
showed that CO, in fluid inclusions in incipient
charnockites from southern India is isotopically heavier
and more abundant than that in associated gneisses (see
also Santosh et al., 1991). On average, both graphite- and
quartz-bound inclusions recorded an increase of c. 1.5% of
6°C in charnockites relative to the adjacent gneisses.
Further, Dunai & Touret (1993) showed that the
high-density carbonic fluid inclusions in the mineral
separates from 2.5 Ga enderbite from the Nilgiri Hills (Fig.
1) have excess 3He, which must be mantlc derived. Thus,
the charnockites seem to have been infiltrated by
C-bearing fluid from a deep-seated source during their
formation. From Raman spectroscopic studies, Hess el al.
(1988) detected graphites in several samples of charnockites from southern India, in which graphite could not be
detected optically. This implies that the absence of
optically detectable graphite in samples of charnockites, in
which the mineral assemblage indicates fo, within the
graphite stability field, is not inconsistent with the presence
of C0,-rich vapour phase.
Magma underplating
Anovitz & Chase (1990) explored this mechanism for the
specific case of granulite facies metamorphism in the
Grenville province, and concluded that it provides ‘a poor
explanation for the observed P-T-t
data’. From
teleseismic tomographic data, Rai ef al. (1993) concluded
that the south Indian granulite province exhibits ‘crustal
thickening coupled with lower velocity in the crust and
upper mantle’, which is in contrast to the expected general
enhancement of the average crustal velocity and density
beneath a metamorphic terrane if it were underplated by
basic magma. Thus, magma underplating seems to be an
unsatisfactory explanation also for the thermal perturbation during granulite facies metamorphism in southern
India.
Thrusting and crustal thickening
Crustal thickening due to thrusting leads to an increase of
the maximum temperature in the P-T-f paths of rocks as
these are uplifted from various depths in response to
erosion (England & Richardson, 1977; England &
Thompson, 1984). These, in turn, lead to a metamorphic
‘field gradient’ (i.e. array of maximum P-T conditions
recorded by the mineral assemblages) which is at a higher
temperature than the ambient geotherm. For a onedimensional case, the change in temperature at a given
depth is given by
where k is the thermal diffusivity, H ( 2 ) is t h e radioactive
heat production rate per unit volume of the rock at the
depth 2, and V is the uplift velocity, which is taken to be
constant and positive.
We have calculated the P -T path of rocks uplifted from
depth in response to erosion immediately following a
thrusting event using the interactive, one-dimensional
numerical program 1DT by Haugerud (1986). The initial
thermal profile is a ‘saw-tooth’ profile resulting from the
Archaean geotherm (Fig. 2) as a result of an instantaneous
thrusting along a horizontal plane. The boundary
conditions are T = 298 K at 2 = 0, and either a constant
basal temperature or a constant basal heat flux. As before,
H was assumed to decay exponentially with depth with a
scaling depth of 11.5 km. For use in the above program,
the H vs. Z plot was linearized in several segments so that
it closely matched the exponential variation.
The calculated field gradients associated with thrust
blocks of 25 and 35 km thickness, initial crustal thickness
of 35 km, constant basal mantle heat flux, and uplift rate of
0.1 mm/year are illustrated in Fig. 6 (solid lines). In order
to evaluate the effects of boundary conditions and uplift
rate, the field gradients resulting from a 35 km thrust with
a constant basal temperature and 0.1 mm/yr uplift rate
(short dashed line), and a 25 km thrust with constant basal
heat flux and 0.05rnm/yr uplift rate (long dashed line)
were also calculated. An uplift rate of 0.1 mm/yr yields
THERMAL PERTURBATION D U R I N G CHARNOCKITIZATION
Fii. 6. Perturbation of the ambient 2.5 Ga geothem due to
overthrusting of blocks between 25 and 35 km thickness along
horizontal plane. Solid lines: constant basal heat flux
(21.3 mW/m*), and constant uplift rate of 0.1 mm/yr. Short
dashed line: constant basal temperature and constant uplift rate of
0.1 rnm/yr. Long dashed line: constant basal heat flux and
constant uplift rate of 0.05 mm/yr. For other symbols: refer to
Fig. 2.
a cooling rate of c. 2" C/Ma, which is almost independent
of the depth, whereas that of 0.5 mm/yr yields a cooling
rate of c. 12" C/Ma at 20 km depth, which slightly increases
at shallower depth. The dependence of cooling rate on
depth increases rapidly with increase of the uplift rate.
From a survey of the available data on cooling rate of
granulites (e.g. Cosca et al., 1987; Anovitz el al., 1988;
Chakraborty & Ganguly, 1992), we feel 0.1-1 mm/yr to be
the probable range of uplift rates of the granulite facies
rocks of southern India. This range of uplift rates produces
cooling rates of c.2-2O0C/Ma at 20km depth. Comparison of the calculated field gradients of granulites for
different values of uplift rate, thickness of the thrusted
blocks, and 'basal' boundary conditions with the P-T
conditions of the granulite facies rocks from southern India
(Fig. 6) suggests that a 30-35-km thrust is capable of
producing the inferred thermal perturbation for these
rocks.
From analysis of travel time residuals and 3-D seismic
images, Rai et a f . (1993) concluded that the south Indian
granulite province is characterized by a thick, low-velocity
crust together with a lower-velocity upper mantle down to
c . 200 km depth relative to the crust and mantle beneath
the Dharwar granite-gneiss craton (see Fig. 1). The
inferred present-day crustal thickness beneath the granulite
terrane is c.35-37 km, which is, on average, 2-4km
thicker than that of the Dharwar craton. As discussed by
Rai et af. this is in accord with the crustal thickness map of
south India calculated by Subba Rao (1988) from 3-D fit
(prism shape) to Bouguer anomaly data of south India,
which shows 5-6 km thickening of the granulite terrane
427
compared to the Dharwar craton. From the above
dissimilarity of the crust/upper mantle character between
the granulite and Dharwar granite-gneiss terranes, Rai el
al. (1993) concluded that the south Indian granulite
terrane represents an 'allocthonous' zone.
Systematic geobarometric data of the south Indian
charnockites and granulites between 1238" (i.e. near the
orthopyroxene isograd or the transition zone) to 12"N
latitude by Hansen et al. (1984) show a monotonic increase
of pressure southward (see Fig. l), corresponding to
palaeodepths of the exposed rocks from 1 7 i 2 to
25 & 3 km. This implies an original crust increasing in
thickness southward from c. 53 f 3 km near the transition
to 61 f 4 km at 12W, if we assume that the crust has not
been subjected to any thinning process other than erosion
since the Archaean. The proposed thrusting of 30-35 km
leads to an overall crustal thickness of 65-70 km, if the
original crust was 35 km, as assumed above in calculating
the thermal effect of thrusting. Thus, it seems very likely
that the thermal perturbation south of the 12"N latitude
was primarily due to crustal thickening due to thrusting, if
the estimated crustal thickness were due to overthrusting
of a single crustal block. However, the estimated crustal
thickness near the transition zone falls c. 10 km short of
what is needed to produce the inferred thermal effect,
unless the crust was thinned by that amount by
non-erosional process (e.g. by subhorizontal lower crustal
flow) since the Archaean. It should, however, be noted in
this context that the thermal effect of thrusting has been
somewhat overestimated since we assumed instantaneous
thrusting.
SUMMARY AND CONCLUSIONS
We have demonstrated above that the granulite facies
metamorphism in southern India required a major thermal
perturbation. The relatively thick crust and allochthonous
nature of the south Indian granulite terrane, as discussed
above, strongly suggest that overthrusting had played a
significant role in this process and was primarily
responsible for the thermal perturbation south of the
transition zone. However, overthrusting alone seems to be
an inadequate mechanism for producing the required
thermal effect at the transition zone. The relatively lower
velocity of upper mantle rocks beneath the granulite
terrane argues against magma underplating as a significant
cause of the thermal perturbation. Thus, it is suggested
that the charnockitization and, in general, the granulite
facies metamorphism at the transition zone were due to a
combined effect of thrusting and fluxing of CO, from
deeper levels. Although this combined process was not
modelled in detail, the calculations presented in Figs 3 and
6 suggest that a Darcy velocity o r volumetric flux of
c. 0.20 cm/yr, combined with the effect of crustal
thickening to c. 55 km (20 km thrust sheet), would have
been adequate to bring about the granulite facies
metamorphism in the transition zone. Allowing for the
JT effect, the required CO, flux in the combined process is c.O.l3cm/yr, which is much less than what is
428
I. C A N C U L Y E r
AL.
required t o produce granulite facies metamorphism
completely by convective heat transport through CO, flux,
and implies a focusing of c . 6 0 times (with an
uncertainty of a factor of 3) the average focusing of CO, at
the present time through the tectonically active regions of
the Earth’s crust. This is probably not an unreasonable
degree of focusing, given the highly sheared nature of the
south Indian granulite terrane, especially near the
transition zone (Fig. 1). This model also helps explain the
development of patchy granulites in the transitional zone
such as Kabbaldurga and Ponmundi (Pichamuthu, 1965;
Janardhan el al., 1982; Hansen er al., 1987) through
heating of the rocks adjacent t o the shear planes and zones
of weakness. Since ‘tectonic heating’ must be pervasive, it
cannot, by itself, lead t o the development of patchy
granulites from amphibolites. However, influx of CO,
through zones of weakness in rocks, which were already
heated to upper amphibolite facies conditions, can provide
additional heat in local domains to lead to the
development of patchy charnockites.
ACKNOWLEDGEMENTS
T h e senior author’s travel t o India in connection with this
research was supported by a grant from the International
programme of the U.S. National Science Foundation.
Thanks a r e d u e Drs J. Brady, S. Sen, J. Connolly, T.
Hoisch and S. Chakraborty for reading earlier drafts of the
manuscript and helpful suggestions, which have substantially influenced the final version. Constructive formal
reviews by Dr S. Peacock and an anonymous reviewer are
gratefully acknowledged. We are thankful to W. Erickson
for help with the use of 1DT program of Haugerud.
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non-dimensional form using (A2) and
kt
I* = L2
as
where b and c are as defined in Eq. (A4).
Using the following transformations (Ozisik, 1980)
T*(z*, I*) = Z* + q z * , t*)e-bZ'Re-b2'*'4,
Received 20 Ocrober 1992; revision accepted 19 October 1994.
-=a8
ar*
APPENDIX
(A91
Eq. (AS) reduces to
a2e
+
+
a ~ * ~
bjebZ'R
b2r*/4
(AW
Also, the boundary conditions reduce to
Solutions of energy balance equation
8=0
8=0
I. Steady -slate solution
The steady-state form of Eq. (3) is
K, d2T/dZ2+ (pCPVd),dT/dZ
+ (1 - 4)Hse-"D
= 0. (Al)
atZ*=O
atZ*=1
(All)
The initial condition is obtained by solving the steady-state form
of Eq. (3) (with advection term as zero) with the boundary
conditions as defined by Eq. (4). We get the initial condition as
8(Z*, 0) = [cd2(1 - e-z'u) - cd2(l - e-"d)Z*] ebZWR. (A12)
The above equation is reduced to non-dimensional form using
Equation (AlO), with boundary conditions ( A l l ) and initial
condition (A12), can be solved by using the method of separation
of variables (Ozisik, 1980). The solution is
z * = Z/L
T* = ( T - Ts)/( TL - Ts)
as
I
T*(z*,
d2T*
dZ*2
+
+
e-
=
t * ) = Z*
+ 2 C sin(nrrZ*)e-bZ'R
*
n =O
dZ*
where
H A 1 - 4)L2
KJTL - Ts)
Then the boundary conditions given by Eq. (4) reduce to
c=
T*=O
T*=l
atZ*=0
atZ*=l
where
The solution of the Eq. (A3) and (A4) is
T*(Z*) =
(1 -e-bZ-)
l-e-h
An = nrr
+-1 cd2
-bd
where d = D/L. This equation has been used to derive Fig 4.
II. Transient solution
The energy balance equation (Eq. 3) is transformed
in
Equation (A13) has been used to derive Fig. 3 using the
parameters given in the text.