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OCR for page 64
4
Fluid Dynamics During
Progressive Regional Metamorphism
JOHN V. WALTHER
Northwestern University
INTRODUCTION
A major part of active tectonic processes is the evolu-
tion of continental margins with the progressive burial and
metamorphism of lavas and sedimentary strata to deep
crustal levels. The nature of fluid components in these
rock sequences buried to mid or lower crustal depths (10
to 40 kin) is poorly known but of utmost importance. The
presence of fluid changes the theological properties of the
rock and thus its response to stress. The fluid also has the
potential to carry heat and dissolved mineral components
and can therefore dictate the style of metamorphism, the
extent of chemical equilibrium, and much about the tex-
tural fabric observed in rocks. It has become abundantly
clear that the majority of sediments and rocks in active
tectonic processes in the upper few kilometers of the Earth's
surface have experienced large fluid fluxes. The nature of
the flow and its chemical and physical consequences are
addressed in other chapters in this volume. This chapter
addresses problems regarding the origin, flow dynamics,
and chemical consequences of fluid at mid-to-lower crus-
tal depths. Classically, to those who study the mineralogi-
cal changes in such rocks, the general process is termed
progressive regional metamorphism. All mountain fold
belts contain lavas and sediments that have been region-
ally metamorphosed.
64
FLUID-ROCK RATIOS AND FLUID FLUX
It is not surprising that large fluid fluxes can be re-
corded within the upper few kilometers of the Earth's
surface where the porosity and/or permeability of many
rocks are large and dramatic temperature and topographic
gradients can occur. Perhaps more surprising is the evi-
dence for large amounts of fluid flow occurring in deeply
buried rocks. Essentially this evidence consists of docu-
menting chemical changes in mineral assemblages during
metamorphism and calculating the minimum fluid volume
necessary to account for the observed changes, as shown
in Figure 4.1.
Perhaps the most straightforward calculation of this
type is documentation of the extent of reaction for a simple
decarbonation such as the production of wollastonite from
calcite and quartz:
CaCO3 + SiO2 CaSiO3 + CO
2
2.
At fixed pressure and temperature this reaction fixes the
fugacity of CO2 in the fluid. Thus, the concentration of
H2O or other components in the fluid needed to maintain
the fugacity of CO2 for the extent of production of wollas-
tonite can be calculated (e.g., Rumble et al., 1982).
The calculated minimum fluid necessary to produce the
observed chemical changes seen during metamorphism is
OCR for page 65
FLUID DYNAMICS DURING PROGRESSIVE REGIONAL METAMORPHISM
' t
/~/ ~-
:~? ;~1
-
Fluid flowing along
grain boundaries influences
phase equilibria
Instantaneous FRR<~0.01
1 1
Mineral
assemblage
after
meta
morphism
t
Mineral
assemblage
before
meta
morphism
termed the time-integrated fluid-rock ratio or more com-
monly the fluid-rock ratio (ERR). FRRs, calculated by a
number of approaches from mineral assemblages in basal-
tic or carbonate units, are as high as 2 to 20 (e.g., Ferry,
1976, 1980;Grahametal.,1983;Tracyetal.,1983~. That
is, 2 to 20 volumes of fluid have reacted with each volume
of rock during progressive metamorphism. Note that these
are minimum values since fluids in equilibrium with the
mineral assemblage are unrecorded by this technique.
Additionally, there is no guarantee that the fluid has reached
complete chemical equilibrium with the mineral assem-
blage, an assumption inherent in the calculation. How-
ever, given the high temperatures involved and the large
surface area to fluid volume in metamorphic rocks, the
equilibrium assumption is probably reasonable (Walther
and Wood, 1986~. Because the actual grain boundary
porosity of metamorphic rock is probably less than 0.1
percent, these high Fl(Rs indicate that fluid within the
flow porosity must have been replenished thousands of
times if the introduction of fluid does not expand the rock.
Fluid-rock ratios are often calculated based on changes
in the isotopic composition of oxygen in minerals. Unfor-
tunately, these changes are sensitive to fluid-rock interac-
tions at any point in the rock's history, that is, during its
burial to peak metamorphic conditions, at the peak of
metamorphism, or during its uplift to the Earth's surface.
In some cases isotopic exchange indicating substantial
fluid-rock interaction may be recording exchange that has
occurred near the Earth's surface either during burial or
uplift. Thus, they may not be recording FRRs during the
peak of metamorphism. Wickham and Taylor (1985;
Chapter 6, this volume) have suggested that high-grade
metamorphic rocks from the Pyrenees appear to require
65
Total
volume
of fluid
needed
FIGURE 4.1 A comparison of the amount
of fluid within a rock during metamor-
phism (left) and the fluid-rock ratio (FFR)
calculated from mineral equilibria (right)
(from Wood and Walther, 1986~.
the introduction of a large volume of seawater deep into
the metamorphic pile. Such a flux of fluid to great depths
in unlikely (Walther and Orville, 1982; and as discussed
below). The data are, however, also consistent with oxy-
gen isotope exchange nearer the Earth's surface.
Fluid-rock ratios are not a measure of the total time
integrated fluid flux (IFF) during metamorphism but must
be considered in the context of their ability to place limits
on the IFF. First, the volume of rock affected must be
known, but unfortunately it is often difficult to determine.
It is obvious, however, that the same total fluid volume
reacting toward equilibrium with a thin rock unit will
record a higher FRR than an identical unit that is thicker,
because the extent of reaction in the thinner unit is greater
per unit rock volume. These problems have been discussed
in more detail elsewhere (Wood and Graham, 1986; Wood
and Walther, 1986~. -
Fluid-rock ratios do not record the passage of fluid in
equilibrium with a rock. This has important ramifications.
Consider a typical divariant (sliding) reaction in a politic
rock undergoing metamorphism:
2HCl + CaCO3 + 2NaAlSi3O~
(calcite) (in play.)
CaAl2Si2O6 + 2NaCl + 4SiO2 + CO2 + H2O.
(in play.) (quartz)
Assuming that calcite and quartz are pure phases, at fixed
pressure, temperature, and chloride activity the composi-
tion of the plagioclase will dictate the equilibrium ratio of
CO2 and H2O in the fluid. Imagine a variety of composi-
tional layers, perhaps of sedimentary origin, with varying
plagioclase composition. A fluid passing through such
OCR for page 66
66
layering will react to adjust its CO2 to H2O ratio to ap-
proach equilibrium with each plagioclase crystal it makes
contact with. The.extent of reaction will depend then on
how far the plagioclase composition is from equilibrium
with the fluid. For the same IFF the recorded FRR will be
different depending on the variability of plagioclase com-
position. One may then be drawn to the wrong conclusion
that some beds experienced large IFFs while others did
not. Thus, it would erroneously appear that some beds
acted as metamorphic fluid aquifers while others were
aquitards. This could then lead to a model of the major
flow occurring along bedding when in fact it may not
(Ferry, 1987~. However, a number of stable isotopic stud-
ies have apparently demonstrated that bed-parallel fluid
flow occurs. What will be argued here is that it is not
differences in intrinsic permeabilities but differences in
the theological properties of the different layers that con-
trol fluid flow.
FLUID FLOW AT MID- TO LOWER-CRUSTAL
LEVELS
Fluid-rock ratios at depth are apparently similar to those
determined near the Earth's surface, where fluid convec-
tion is known to operate. It has, therefore, been proposed
that fluid convects to deep crustal levels (e.g., Etheridge et
al., 1983; Wickham and Taylor, 1985; Ferry, 1986~.
Arguments to the contrary have also been presented
(Walther and Orville, 1982; Walther and Wood, 1984;
Wood and Walther, 1986~. Let us assess some of the
evidence for the state and transport of fluid at mid or lower
crustal depths.
Fluid flow is obviously highly dependent on permeabil-
ity. Permeabilities of metamorphic rocks are not well
known. What is clear is that they increase dramatically in
laboratory experiments when fluid pressure equals rock
pressure (Brace, 1980~. Since most of the experimental
studies do not concern themselves with large fractures,
such permeability estimates are considered by many to be
minimal for the crust as a whole. Use of these permeabil-
ity values suggests that fluid flow during metamorphism
should occur under conditions of fluid pressure signifi-
cantly below rock pressure. That is to say that for the
anticipated fluid flux the permeabilities during metamor-
phism are great enough to allow fluid pressure to drop
below the rock pressure toward a more hydrostatic gradi-
ent. A hydrostatic gradient is required to promote convec-
tion. By hydrostatic pressure what is meant is the pressure
resulting from the density of an overlying column of fluid
as opposed to the much greater pressure resulting from the
density of an overlying column of rock (lithostatic pres-
sure). Fluid pressure less than lithostatic does not seem to
be the case when metamorphic rocks are closely examined
JOHN V. WALTHER
(Norris and Henley, 1976; Fyfe et al., 1978). Phase equi-
librium studies of devolatilization reactions observed in
metamorphic rocks generally seem to require fluid pres-
sure to be near lithostatic pressure. This observation is
confirmed by fluid inclusion studies where the trapped
fluid has the appropriate density for fluid pressure equal to
lithostatic pressure at the temperature of metamorphism.
These two divergent observations could be rectified if
the properties of the fluid phase were not those of the bulk
fluid. It has been suggested, for example, that the fluid
phase present during deep crustal metamorphism is not a
discrete fluid phase but rather an absorbed phase on min-
eral surfaces with presumably greater viscosity and lower
fugacity (Elliott, 1973; Rutter, 1976~.
Studies at room temperature indicate that multimolecu-
lar layers of H2O are absorbed on mineral surfaces. At
metamorphic temperatures and pressures where even struc-
turally bound water becomes unstable in hydrous miner-
als, it is reasonable to assume that absorption of H2O is no
more than a monolayer in thickness. If we examine the
amount of fluid trapped as fluid inclusions along healed
microcracks during metamorphism and redistribute it evenly
along the entire crack, the width of fluid in the crack is
generally about 200 ~ or 10 times the thickness of a
double monomolecular layer (Walther and Orville, 1982~.
Because this is the minimum thickness of the fluid film at
the time the crack healed, we may reasonably assume that
the properties of fluid flowing through such cracks are not
significantly modified by absorptive properties of mineral
surfaces, and hence the fluid phase can be considered to
have the properties of a bulk fluid.
Near the Earth's surface fluid pressure is controlled by
the extent of the overlying fluid column (hydrostatic pres-
sure) while the pressure exerted on mineral grains
(lithostatic pressure) is considerably greater owing to the
much greater density of the- minerals. This difference
between fluid pressure and rock pressure is maintained by
the effective crushing strength of the rock. At some depth
the closure of pores and the resultant decrease in permea-
bility causes the fluid pressure gradient to increase dra-
matically, so that fluid pressure equals lithostatic pressure.
Note that fluid pressure equal to rock pressure does not
imply that fluid is trapped in a static state but that it is
possible that the flow of fluid is balanced by permeability
changes near fluid pressure equal to rock pressure.
Figure 4.2 shows fluid pressure as a function of depth
determined from well bottom hole fluid pressure measure-
ments from a number of wells in the U.S. Gulf Coast.
Note that at a depth just below 3 km fluid pressure begins
to depart from hydrostatic, and at 5.5 km or at a lithostatic
pressure of 1.5 kbar the fluid pressure is very close to
lithostatic. The fluid pressure below 3 km is greater than
that exerted by an overlying column of fluid and is there
OCR for page 67
FLUID DYNAMICS DURING PROGRESSIVE REGIONAL METAMORPHISM
y
'_ 3
1 1 ~
BRAZORIA CO.. TEXAS
~ W
1 _\ '\
LITHOSTATIC
\ ~- GRADIENT ~ BOTTOM HOLE
2 \ \ FLUID PRESSURE
~- -x iFLUID PRESSURE
=._. ...' ._. . . .
4 _
5 _
HYDROSTAT C',\
GRADIENT \
i\
.w '\
\ ~ _
-do
6 _
0 0.2 0.4 0.6 0.8 1.0 1.2
PRESSURE, KBAR
FIGURE 4.2 Fluid pressure as a function of depth in a sedimen-
tary basin of the U.S. Gulf Coast. Note the crossover from
hydrostatic to lithostatic fluid pressure and its implications for
fluid convection (from Wood and Walther, 1986~.
fore considered "geopressured." There are various chemi-
cal and physical factors that influence the depth at which
geopressuring of the fluid occurs. The decrease of hydrau-
lic conductivity in clay layers is often considered the pri-
mary factor in determining the characteristics of the geo-
pressured zone (Bredehoeft and Hanshaw, 1968; Chapman,
1972).
The depth of onset of geopressuring has been observed
from as little as 45 m to depths greater than 8 km, although
in most sedimentary basins it is above 6 km. We might
imagine that in crystalline rocks in extensional environ-
ments hydrostatic gradients in the fluid may be maintained
to depths greater than 8 km. However, at a depth of 11 km
the difference between hydrostatic pressure and lithostatic
67
pressure is about 2 kbar. Although the crushing strengths
of rocks at the temperatures, pressures, and strain rates
,L ~1` appropriate for this depth are not well known, it seems
it. ~reasonable to conclude that hydraulic conductivity would
~ en be so greatly reduced by mechanical compaction that fluid
_ ~ ° below this depth can in general be considered to be close
O to lithostatic pressure. As mentioned above, this observa
lion is consistent with the evidence from fluid inclusion
studies that indicates that fluid pressure equals rock pres
. _ ._
sure during metamorphism.
° It would, therefore, seem that laboratory measurements
_ ~of the permeabilities of metamorphic rocks do not charac
~n ~terize the permeability of these rocks at the time they are
undergoing metamorphism, but overestimate them. It may
~be that the permeability is reduced by the inevitable re
<' crystallization that occurs when these rocks are subject to
high temperatures and pressures for time periods on the
order of millions of years during metamorphism. In any
event it seems that fluid pressure must be close to lithostatic
at mid- to lower-crustal depths. The imposition of a pres
sure gradient significantly greater than hydrostatic on the
fluid during progressive regional metamorphism means
that there is no reasonable way to transport the less-dense
fluid by flow downward (i.e., no convection of fluid can
occur). This means that the large integrated FRRs re
corded by some rocks must be recording fluid generated at
some greater depth.
FLOW MECHANISM WHERE FLUID PRESSURE
EQUALS LITHOSTATIC
Consider a rock initially devoid of fluid undergoing a
devolatilization reaction in response to increased tempera
ture during progressive metamorphism at mid-crustal lev
els. Fluid will be produced by each volatile-containing
mineral undergoing destruction by reaction. Depending on
the wetting characteristics of the fluid, it will either coat
the mineral surfaces or begin to collect in isolated pores at
mineral triple junctions (White and White, 1981; Watson
and Brenan, 1987), as shown on the left side of Figure 4.3.
Z
zap FIGURE 4.3 Fluid production at reacting
volatile containing minerals collecting in
isolated pores at mineral triple junctions
(left). On a larger scale with increased
fluid production, these must interconnect,
producing a fluid phase of some vertical
extent, z (right).
OCR for page 68
OCR for page 70
OCR for page 71
Representative terms from entire chapter:
fluid flow
68
In any event, continued devolatilization will eventually
build an interconnected three-dimensional fluid network
of some vertical extent, as shown schematically on the
right side of Figure 4.3. A static fluid at lithostatic pres-
sure would rise due to its lower density and therefore to its
buoyancy relative to the surrounding rocks if it was not
held by the tensile strength of the rock. Because rocks at
metamorphic conditions have low tensile strengths, the
fluid would become mechanically unstable and hydrofrac-
ture its way toward the Earth's surface. Similar arguments
have been used to explain the ascent of magmas from
depth even though the buoyancy forces are less. As the left
side of Figure 4.4 shows, the greater the amount of inter-
connected fluid space, the greater will be the difference
between the pressure on the fluid and surrounding rock
and thus the ability of the fluid to hydrofracture the rock.
The pressure difference between static fluid and rock is
given by
UP = gZ(prock - pfluid) ~ (4.1)
where g stands for the acceleration due to gravity, z is the
vertical distance of interconnectivity of the fluid, and PrOck
and Pfluic are the rock and fluid densities, respectively.
While the tensile strength of a rock acting to prevent
propagation of a fracture at a crack tip is poorly known,
we might imagine that under metamorphic conditions, with
the interplay of deviatoric stresses during active tecton-
ism, it is no more than 10 bars or so, particularly for
subcritical crack growth. This means that no intercon-
nected fluid network can exist statically that extends in the
vertical direction more than about 60 m. This severely
limits the possible vertical extent of fluid convection be-
fore the tensile strength is exceeded and fluid migrates
upward only. Over this short distance no convection is
possible for any reasonable temperature gradient and per-
meability.
The potential to produce an interconnected fluid net-
work is great. The approximately 2 moles of fluid on
average given off per kilogram of politic composition rock
during medium- to high-grade metamorphism has the
potential to hydrofracture the rock many times. This amount
of fluid would occupy a volume of approximately 12 per-
cent under metamorphic conditions if it did not escape.
While we do not know the volume of fluid necessary to
produce an interconnected fluid pathway, it is probably no
more than a few tenths of a percent of the volume of the
rock. Thus, we might imagine that even on a local scale
hydrofracturing may have occurred a large number of times.
If we consider the passage of fluid released from minerals
lower in the metamorphic pile, it is not surprising that we
observe large numbers of healed fluid-filled microcracks
in metamorphic rocks.
JOHN V. WALTHER
The predominance of divariant (sliding equilibria)
devolatilization reactions during metamorphism suggests
that fluid is released from hydrous and carbonate minerals
more or less continuously in the metamorphic pile during
prograde metamorphism. It is possible that some fractures
stay open for extended periods of time fed by a continuous
supply of fluid. This situation is shown on the right side
of Figure 4.4, where we require a viscous pressure gradi-
ent, dPViS, to compensate for the difference in the pressure
gradient between the static fluid and the surrounding rock.
That is, the loss in hydraulic head by the upwardly flowing
fluid due to its viscosity compensates for the difference
between the hydrostatic and lithostatic gradients so that
fluid pressure in the flowing fluid equals lithostatic pres-
sure along the walls of the fracture, which in turn allows
the fracture to remain open:
dPvis = g ~ Ink - pfluid ~ · (4. 2)
Assuming fluid and rock densities of 0.9 and 2.8 g/cm3,
the viscous pressure gradient required as calculated from
Eq. (4.2) is about 2.0 X 103 dyne/cm3. We can model the
fluid flow in a microcrack as steady-state incompressible
viscous laminar flow through two parallel plates that ex-
tend in all directions to a very much greater extent than d,
the distance they are apart. Due to the viscous nature of
the fluid, the profile of the flow velocity across the crack
is parabolic, zero at the wall, and a maximum in the center.
The solution to this fluid mechanical problem is
2vmax d dPvis
,, ~ ~ A ~ 4 . 3
3 12v
where v and vmax are the average and maximum fluid
velocities, respectively, d is the fracture width, and v is the
viscosity of the fluid. Noting that the fluid flux, q, is equal
Pf>Priz+ /
Pf=Pr Ny I
Pf
FLUID DYNAMICS DURING PROGRESSIVE REGIONAL METAMORPHISM
to the cross-sectional area of the fracture opening times
the average velocity, we have
q= vdl =
HI d Pvis (4.4)
12v
where I is the length of the fracture opening.
Viscosities of supercritical H2O-CO2 mixtures are gen-
erally between 0.1 and 0.2 centipoise. For a given flux of
fluid we can calculate the length of cracks perpendicular
to flow versus their width per square centimeter.
Such calculations have been done (Walther and Orville,
1982; Walther and Wood, 1984) and indicate that the
widths of the fractures are in most cases less than 105 A.
Because of the cubic dependence of fluid flux on crack
width, it seems likely that changes in the flux of fluid are
accommodated by small changes in fracture width to
maintain the viscous pressure gradient at the value neces-
sary to keep the fracture open. Apparently, judging from
the extent of fluid inclusions along sealed fractures in
metamorphic minerals, if the fracture width falls below
about 200 A the crack will seal. Laboratory investigations
indicate that these cracks seal in a matter of days, at least
in quartz. It stands to reason that, if the fluid flux through
the metamorphic pile is not continuous, many generations
of these microcracks will form.
If such a mechanism of fluid flow operates, the useful-
ness of the concept of intrinsic permeability of a rock is
questionable. That is, the permeability is a dynamic func-
tion of the fluid flux through the metamorphic rocks. The
permeability of the rock is adjusted by the fluid phase to
accommodate the flux of fluid, so the fluid pressure is near
lithostatic during fluid flow. This is different from the
concept of reaction-enhanced permeability due to volume
loss of solids because of reaction- (Rumble and Spear,
1983~. What is argued here is that as a general approxima-
tion the permeability adjusts itself, so fluid pressure is
always close to rock pressure irrespective of the extent of
reaction.
With the deviatoric stresses that operate during meta-
morphism and the differences in tensile strength of differ-
ent layers of rock, fracture production and, therefore, fluid
channels may develop along preferred layers that are at
some angle to the Earth's gravitational field. This would
give rise to layers that appear as "metamorphic aquifers"
and others that appear as aquitards, at least over limited
distances.
CHANNELIZATION OF FLUID FLOW
The interconnectivity, bifurcation, coalition, or distri-
bution of cracks in metamorphic rocks are largely un-
known. We do know that some fractures or fracture net
69
works have remained open long enough to experience the
passage of a considerable amount of fluid. In politic rocks
these fractures are often marked by quartz veins. At lower
temperatures calcite veins are dominant because calcite
has a higher solubility at low temperatures than does quartz.
The thickness of the vein represents the accumulated ef-
fects of mineral precipitation from the passing fluid. If
such major channelways of fluid escape are present, the
flow paths of these fluids must to some extent coalesce as
fluid is produced at each mineral in the rock that under-
goes devolatilization. Thus, there must be some flow of
fluid in intimate contact with minerals before fluid enters
a major channelway. Obviously, the extent of grain bound-
ary flow versus flow in major channelways controls to a
large extent the amount of fluid a particular rock unit may
experience during metamorphism and hence IFF. This in
turn dictates much about the textural fabric of the rock and
the approach to chemical equilibrium that may be ex-
pected.
Figure 4.5 shows a quartz segregation/vein in the Bund-
nerschiefer formation of Switzerland that is considered to
have formed during Lepontine metamorphism. If this
quartz segregation represents the cross section of quartz
deposited along a major conduit for fluid flow, we can
calculate the amount of fluid that must have passed to
cause the extent of quartz precipitation seen. Because
quartz is present in most of the mineral assemblages in the
Bundnerschiefer, it is anticipated that fluids responsible
for the deposition of quartz along the segregation/vein are
at quartz saturation.
Let us calculate the fluid necessary to precipitate a
quartz vein 50 cm in diameter, much like the one shown in
FIGURE 4.5 Quartz segregation thought to represent the cross
section of quartz precipitated during the lifetime of a major fluid
conduit during metamorphism. While the width of the fracture
was probably less than 10 ~m, the large amount of flow precipi-
tated substantial quantities of quartz.
70
Figure 4.5. For a distance of 1 cm along the vein this
amounts to 2000 cm3 of quartz or about 87 moles. At
500°C and 4 kbar with a geothermal gradient of 20° to
30°C~n, this requires 4 x 108 moles of H2O or the dehy-
dration of 7.6 x 10~° cm3 of average politic rock during
metamorphism if 2 moles of H2O are released per kilo-
gram of rock. This corresponds to a cube of petite 42 m on
a side.
Given the large volumes of politic and psarnmitic rock
in most metamorphic terrains, the fluid fluxes must be
high. For the whole of the metamorphic pile, fluid fluxes
of 1 x 1~° to 1 x 10~9 g cm-2 s~i have been calculated
(Walther and Orville, 1982), which means that an average
of 3 to 30 kg of fluid must pass through each square
centimeter of crust overlying the 400°C isotherm in each
million years of progressive metamorphism. The upward
flow of fluid may be even greater if significant fluid is
released from a subducting slab during metamorphism
(perhaps through a magma intermediary) or if mantle fluids
(Dawson, 1980) are significant.
Until it is determined to what extent fluid is channeled
along structural features such as faults, fold hinges, pri-
mary lithological contacts, and layers with low tensile
strength or, alternatively, flows at the scale of grain bounda-
r~es, the extent of heat and material transport and the very
nature of metamorphism will not be understood.
ACKNOWLEDGMENTS
This chapter was written while I was on sabbatical
leave at Universite Paul Sabatier, Toulouse, France. I
would like to thank Jacques Scholl for his warm hospital-
ity. Comments by J. M. Ferry and B. J. Wood led to
substantial improvement.
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Chapman, R. E. (1972~. Clays with abnormal interstitial fluid
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Dawson, J. B. (1980~. Kimberlites and Their Xenoliths, Sprin
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JOHN V. WALTHER
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Tracy, R. J., D. M. Rye, D. A. Hewitt, and C. M. Schiffries
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