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OCR for page 148
150
HYDROSTATIC
Density = 2.6 g/cm3
10 km
5 cm/yr
v > 1.6 X 10-7 cm/see
K > 3 X 10-7 cm/see
/ \
Figure 10.1 Simple model of hydrologic conditions in sediments
subducted beneath an accretionary wedge. Half arrows indicate
displacement along the dicollement zone. Fluid pressure in the
subducting sedimentary bed is hydrostatic at the deformation
front and lithostatic beneath the dicollement zone and accretion-
ary prism. Assuming a water depth of 5 km at the deformation
front and 1 km above the accretionary prism (given the modeled
geometry), a 10-km-thick accretionary prism, and bulb density of
2.6 g/cm3 within the prism (Bray and Kang, 1985), the hydraulic
gradient along the sedimentary bed is 0.306. If the plate conver-
gence rate is 5 cm/yr (i.e., 1.6 x 10-7 cm/s), sediments within the
subducting plate must have hydraulic conductivities (K) > 3 x
10-7 cm/s for fluids to migrate via Darcian flow up-dip along
sedimentary beds in a sea-level-based reference frame. Silt-
sized sediments have K in this range (Freeze and Cherry, 1979~.
Because fine-grained sediments are common in subduction zone
settings (Shepherd and Bryant, 1983), up-dip fluid escape along
sedimentary strata may often be untenable. In these cases if fluid
is to escape from the subducting plate, it may rise vertically, in
which case it must ultimately intersect the dicollement zone,
indicating the fundamental importance of the devolvement zone
in the hydrogeology of subduction zones.
reference frame (e.g., relative to the sea surface), fluid
flow in the sedimentary bed must exceed the subduction
rate to result in net migration toward the surface. For the
parameters described in Figure 10.1, the sedimentary bed
must be silt sized or coarser to permit Darcian flow. In
settings where subduction is more rapid than 5 cm/yr (e.g.,
the northeast Pacific during the late Mesozoic; Engebretson
et al., 1984), the permeability of subducted sediments
must be correspondingly higher to allow fluids to effec-
tively migrate toward the seafloor.
Because subducted sediments are mostly clay sized to
silt sized with low intrinsic permeability, fluids will either
be subducted into the mantle or must find an alternate path
of escape. Any resistance to fluid escape will cause fluid
pressures to rise further. Fluids may rise vertically until
they reach the decollement zone and associated faults,
which then may act as a fracture network for fluid flow.
Alternatively, low-angle, bedding-parallel natural hy-
drofractures may develop, as hypothesized along the Bar-
bados accretionary complex by Westbrook and Smith
(1983).
Moore et al. (1987) calculated from measured hydro
PETER VROLIJK AND GEORGIANNA MYERS
LITHOSTATIC logic parameters that fluid could only flow up-dip along
sandy sedimentary horizons of the subducting Atlantic
plate along the ODP Leg 110 transect in the Barbados
accretionary complex, an hypothesis that was supported
by geochemical and geothermal observations. These same
geochemical and thermal anomalies were also detected in
the decollement zone, suggesting that the decollement zone
and the sandstone beds have similar hydrologic roles but
with fluid flow in the decollement zone controlled by the
history of faulting.
The recognition of extensive networks of veins in an
cient rocks also provides evidence for high fluid pressures
(Clogs, 1984; VroliJk, 1986, 1987~. In ancient rocks syntec
tonic veins are interpreted as natural hydrofractures that
formed under conditions of high fluid pressure and low
effective stress (Secor, 1965~.
EVIDENCE OF FLUID PRESSURE HISTORY
FROM FLUID INCLUSION STUDIES
Examining ancient rocks is imperative in understanding
small-scale processes at levels deeper than about 1 km, or
the maximum depth of drilling in active margins. In con
trast, remote sensing techniques, such as seismic reflec
tion and electrical conductivity profiles, are useful for
deciphering relatively large scale variations in physical
character. Fluid inclusion analyses in syntectonic veins
have proven useful in examining the fluid pressure history
in a variety of tectonic environments.
One reason fluid inclusions are valuable is that fluids
reach thermodynamic equilibrium with each other far faster
than solid phases. If fluid pressures in a deforming rock
mass fluctuate, fluid inclusions trapped in continuously
growing crystals may preserve some record of the fluid
pressure history and thermochemical evolution. The ki
netics of the growth of quartz crystals, a common vein
filling mineral, suggest that variations on the order of days
may be preserved (Rimstidt and Barnes, 1980~. In con
trast, preservation of evidence of the fluid pressure history
within solid phases would require creating compositional
zoning patterns during mineral growth. Zoning must exist
on a scale large enough to escape chemical diffusion and
also be analytically identifiable, difficult criteria to fulfill
when conditions change rapidly. Another useful feature of
fluid inclusions is their common occurrence in veins that
appear at all scales. Because veins can be mapped at the
outcrop and microscopic scales, it is relatively easy to tie
veins to the development of structural fabrics in the rock.
Complex Fluid Pressure History
An example of the fluid pressure history recorded by
fluid inclusions in a subduction zone is presented by Vrolijk
OCR for page 149
FLUID PRESSURE HISTORY IN SUBDUCTION ZONES
(1987~. In this study syntectonic veins formed during
deformation in active fault zones were analyzed in rocks
from the Kodiak accretionary complex, Alaska. A transect
across half of each vein revealed that the density of meth-
ane in fluid inclusions varied dramatically during the growth
of the veins (figure 10.2~. Vrolijk (1987) suggested that
during the growth of each vein fluid pressures dropped
from the near-lithostatic values present during initial crack
opening to pressures as much as 45 percent lower (Figure
10.3~. The fluid pressure history was interpreted by using
the pressure-volume-temperature (P-V-T) characteristics
of methane in conjunction with independent determina-
tions of the quartz crystallization temperature, resolved by
analysis of coeval water-rich fluid inclusions. Moreover,
considering each fluid inclusion analysis as a single
paleofluid pressure measurement, justified on the basis
that fluid temperatures remained constant during crystal
growth, suggests that during the growth of veins fluid
pressures fluctuated widely, and fluid pressures close to
initial, fracture-forming, near-lithostatic values appeared
repeatedly. These successive high pressure pulses are
interpreted to reflect further widening of the fracture (Figure
10.4~.
One important observation drawn from this work is the
inextricable link between the formation and growth of
I'.'.:....\ ~
. . ;r .~
... \
. .
.j A- .
' · '. :.
.
. .
·. ·:- ·:
`..
~ 1; ' I ' ' ' ' '
o
030 032 034 036 038 0.40
b Density (g/c~n5)
Figure 10.2 (a) Drawing of quartz crystal from syntectonic vein
from a boudin formed during melange deformation. Dots indi-
cate distribution of fluid inclusions within quartz crystal. Cross-
hatched area at bottom indicates sandstone of vein wall. (b)
Methane densities of fluid inclusions plotted versus distance
from vein wall, indicating drop in density with crystal growth
(from Vrolijk, 1987~.
151
400
300
-
200
v,
In
~100
i:
0.40
V
. as/
30.38
0.36
034
0.32
, 0.30
0 200 400
Temperature PC)
Figure 10.3 Methane densities (g/cm3) in P-T space, plotted
from Angus et al. (1976~. Also plotted are high and low fluid
pressures from two samples of the Ghost Rocks Formation (stars)
and one sample of the Uyak Complex (circles), Kodiak accre
tionary complex, Alaska; note that high-pressure points for both
Ghost Rocks samples fall on same P-T point. Solid symbols are
interpreted as probable near-lithostatic fluid pressures, open
symbols as lowest fluid pressure in each vein. Square point
represents a later vein deposited in strike-slip fault of Ghost
Rocks Formation, indicating that fluid pressure drop in boudin
veins is not related to uplift (from Vrolijk, 1987~.
fractures and the fluid pressure history. This point is
illustrated in Figure 10.4, in which an inferred record of
fluid pressure fluctuations is plotted. The important parts
i.2 of this diagram include the following: (1) the initial fluid
pressure builds up to some value near lithostatic; although
, O ~ancient rocks contain no record of this stage, it is inferred
~from theory (e.g., Walder and Nur, 1984~. (2) The fluid
o.a ~pressure exceeds the least principal stress and the tensile
``, strength of the rock (e.g., Secor, 1965; Etheridge et al.,
~1984), creating a fracture. When the fracture forms, new
0.6 ~voids are formed, causing the fluid pressure to drop. Vro
- lijk (1987) suggested that this fluid pressure drop caused
04 ~silica to become oversaturated in the fluid, leading to
~quartz precipitation. Once the fracture exists, it is reutil
o~. ~ized as a fluid pathway and continues to accommodate
local extensional strain within the rock. (3) Fluid pressure
must repeatedly rise to widen the fracture, but with each
increment of growth the fluid pressure drops, creating a
cyclic fluid pressure history. (4) The fracture seals with a
final pressure decrease. Vrolijk (1987) hypothesized that
the gradual drop at the lowest fluid pressures represented
growth of an interconnected fracture network along which
fluids escaped to shallower levels of the subduction zone.
The presence of repeated high fluid pressures in these
rocks probably played an important role in determining
the style of deformation, following the ideas of Hubbert
OCR for page 150
52
400
300
loll
Oh
(A 200
Lid
CL
~ 100
B.Failure
C.Crock Lindens
Aim A. ~
'+ ~ Fluid Pressure
oL I
6~. OC'J
E.Residual
TIME --
Figure 10.4 Interpretive fluid pressure evolution in extensional
fracture. Fluid pressure values modeled after data presented in
Figures 10.2 and 10.3. A: Deformation first creates fluid reser-
voir by dilating rock mass and increasing porosity, then causes
fluid pressure to nse. B: Rock failure occurs along thrust faults in
melange matrix; where faults intersect boudins, extensional frac-
tures develop. Fluid pressure at this point equals least principal
stress plus tensile strength of rock. C: Fluid pressure drops to
some value near least principal stress (interpreted as lithostatic
pressure) as extensional fracture widens. D: Lowest fluid pres-
sure decreases during each increment of crack growth as inter-
connected fracture network grows toward increasingly shallower
levels. E: Fluid reservoir built up during initial dilatant defonna-
tion becomes exhausted, and local deformation wanes. Fluid
pressure equilibrates along fracture network as last voids are
sealed (from Vrolijk, 1987~.
and Rubey (1959~. The ubiquitous presence of faults,
fractures, and shear zones in all of the units on the Kodiak
Islands (e.g., Moore and Wheeler, 1978; Moore and All-
wardt, 1980; Byrne, 1984; Sample and Moore, 1987), in
contrast to the relative rarity of folds larger than a single
outcrop, suggests that high fluid pressures may have con-
tributed to strain being accommodated most easily along
fractures.
Consequences of High Fluid Pressure
Understanding the fluid pressure history of individual
fractures and how the fluid inclusion record within veins
reflects that history makes fluid inclusions useful for
paleobarometry and paleothermometry. Studying fluid
inclusions proves a useful analytical method because the
inclusions provide a direct record of the fluid phase, and
migrating fluids may often be the best medium for trans
PETER VROLIJK AND GEORGIANNA MYERS
porting heat and dissolved chemical constituents through
the rock.
Vrolijk et al. (1988) and Myers (1987) used methane-
rich and water-rich fluid inclusions in syntectonic veins
from three units of the Kodiak accretionary complex,
Alaska, to investigate the fluid temperature history of vein-
forming fluids. The veins chosen for this study were
interpreted to have formed during deformation of sedi-
ments within the decollement zone between the subducting
oceanic and overriding North American plate. The princi-
pal conclusion of this study is that fluid temperatures
within the decollement zone (Figure 10.5) were substan-
tially higher than temperatures predicted by conductive
heat flow models (e.g., Oxburgh and Turcotte,1971; Ernst,
1974; Wang and Shi, 1984~.
Warm fluids in fault zones of the Kodiak accretionary
complex were suggested by Vrolijk et al. (1988) to arise
from the migration of fluids along faults faster than heat
dissipated from the fluid, although the presence of young
ocean crust during the formation of these units could not
be completely ruled out as a significant heat source. Other
potential heat sources, such as intruding magmas, enhanced
radioactive decay, and frictional heating, can be ruled out
by field observations. Plate reconstructions (Wells et al.,
1984; Engebretson et al., 1984) suggest that low thermal
gradients (e.g., 5° to 10°C/km) should have been produced
along the Kodiak margin because plate convergence was
fast. The hypothesis put forth by Vrolijk et al. (1988)
suggests that deformation within the decollement zone
allowed faults and fractures to open and permitted fluids
to migrate structurally up-dip. Fluid migration was suffi-
ciently rapid that heat advection outpaced conduction into
the surrounding rock (Figure 10.6~.
The decollement zone in this example appears to have
focused fluid flow along its surface. Because of the decol-
lement zone's apparent hydrologic importance, the term
tectonic aquifer is introduced in Figure 10.6. Aquifer is
used to highlight the observations that fluid flow is en-
hanced along the decollement zone, and the modifier tec-
tonic signifies the role that deformation along the decolle-
ment zone plays in increasing permeability, thereby allow-
ing the decollement zone to support enhanced fluid flow.
The implications of the hypothesis of warm fluid flow
along fault zones touch on the metamorphic history of
subduction zones. If fluid flow is short lived and episodic,
the isotherms in subduction zones may have complicated,
temporally variable shapes controlled by the growth of
faults and fractures (Figure 10.7~. On the other hand, fluid
flow may be pervasive and persistent, generating warmer
temperatures throughout the accretionary complex (Figure
10.7~. Fully resolving this problem will require informa-
tion from (1) further studies of the metamorphic history of
subducted and accreted materials; (2) studies and models
OCR for page 151
FLUID PRESSURE [IISTORY IN SUBDUCTION ZONES
1000
800
600
-
In
At:
400
200
i ~
BOUT \
5:
\\\\\~__- ~ ~
An/
/
or I ~_ ~400 500
O , . ,
0 100 200 300 400
TEMPERATURE (-C)
of the amount of water present in subducting plates, where
that water is released during subduction, and how the fluid
migrates within the subduction zone; and (3) coupled
hydrologic/thermal models that more realistically incorpo-
rate fluid migration mechanisms than has so far been at-
tempted.
In addition to thermal anomalies, chemical anomalies
also appear to be associated with fault zones. In the
Barbados Ridge complex, Moore et al. (1987) described
the importance of the decollement zone in focusing fluid
flow and in separating chemically defined hydrogeologic
regimes. In the ancient Kodiak accretionarv complex
Vrolijk ( 1986, 1987) distinguished veins formed in
melanges from veins formed in structurally coherent units
by comparing oxygen isotope ratios of vein-forming fluids.
Both examples suggest that fluids in fault zones migrated
from deeper structural levels in the subduction zone along
active faults, principally the decollement zone. Further
geochemical analyses of pore waters and vein-filling
minerals will serve to trace the origin and paths of fluids,
thereby more closely constraining the fluid migration his-
tory.
153
Figure 10.5 Rock P-fluid T measurements from veins formed
during melange deformation, Ghost Rocks Formation (Diamonds),
Kodiak Formation (Squares), and Uyak Complex (Circles), Kodiak
accretionary complex, Alaska. Plotted for comparison are a line
describing a thermal gradient of 20°C/lom (as compared to gradi-
ents 27°C/km from model calculations, e.g., Wang and Shi,
1984), assuming a sediment bulb density of 2500 kg/m3 and a
corresponding pressure-depth ratio of 1 kbar/4 km (Bray and
Karig, 1985) and the lower stability limits of typical blueschist-
facies minerals. Notice that if the Kodiak samples had been
subducted more deeply, they may not have developed the ex-
pected blueschist mineralogy, even though they formed in the
decollement zone of a subduction zone. Reactions ~ 1 ~
laumontite~wairakite + fluid, (2) laumontite~lawsonite +
quartz + fluid, and (3) wairakite~lawsonite + quartz are from
Liou (1971~. The glaucophane stability boundary (4) is plotted
from Maresch (1977~. The boundary between barroisitic and
actinolitic amphiboles (5) is drawn from Ernst (1979~. Nitsch
(1972) defined reaction (6), lawsonite + quartz~zoisite +
pyrophyllite + water, and the calcite' aragonite inversion (7)
follows Johannes and Puhan (1971~. Boundary pip marks the
transition from prehnite-pumpellyite facies (low temperature) to
prehnite-actinolite facies (pumpellyite + quartz~zoisite +
prehnite + chlorite + fluid), and pp2 limits the prehnite-pumpel-
lyite facies to the low temperature side and pumpellyite-actino-
lite facies to the high temperature side (prehnite + chlorite +
quartz' pumpellyite + tremolite + fluid); both reactions are
for the model metabasite system of Liou et al. (1985) (from
Vrolijk et al., 1988~.
Similar Fluid Inclusion Studies
Fluid inclusion studies have proven useful in unravel-
ina tectonic and fluid histories in a number of regions and
in various tectonic settings. In the western Alps, Mullis
(1976, 1979) pioneered the use of methane plus water
fluid inclusions in tectonic studies. Large-quartz crystals
growing into cavities during Alpine tectonism trapped fluid
inclusions; these inclusions, like the Kodiak samples,
similarly record fluctuating fluid pressures. Mullis (1976)
interpreted these changes to have occurred on a longer
time scale than the Kodiak veins, and he correlated quartz
crystal growth stages and changing P-T conditions with
successive nappe stacking. Recently, however, Mullis
(1988) interpreted fluid pressure fluctuations recorded in
samples from the Apennines, Italy, as a record of the
expansion of crystal-filled cavities.
In another collisional tectonic setting, Orkan and Voight
(1985) used methane-rich and water-rich fluid inclusions
to determine P-T conditions of joint formation during the
Alleghany Orogeny in the Valley and Ridge Province of
Pennsylvania. This study systematically combined de-
tailed structural and kinematic data from joints and other
structural features with precise P-T measurements. Such
OCR for page 152
Deform ation
Front ~
Cooler Underthrust~
Material
_ ,
Figure 10.6 Heat redistribution by warm fluid migration in a
subduction zone. Fluid is liberated from subducted oceanic crust
and sediments and escapes laterally back toward the surface
along the dicollement zone. Because subduction acts as a con-
veyor belt to drag water to depth, and because the decollement
zone is an areally limited structural zone, fluid flow through the
decollement zone may be relatively rapid. Fluid escape thereby
serves to warm active fault zones at intermediate to shallow
levels. However, because no in situ physical or chemical process
is generating significant quantities of heat, the warming effect at
shallow levels must be balanced by a corresponding loss of heat
and minor cooling at deeper levels where fluids originate.
data will help clarify the sorts of problems regarding the
evolution of joints referred to by Engelder (Chapter 9, this
volume).
Fluid inclusions have also been profitably used in ex-
tensional tectonic environments by Parry and Bruhn (1986~.
By examining CO2 + H2O fluid inclusions in rocks ex-
posed in the footwall of the Wasatch normal fault in north-
ern Utah, Party and Bruhn (1986) suggested that fluid
pressures along the fault shifted from near-lithostatic to
near-hydrostatic as the footwall rocks were uplifted through
the brittle/ductile transition. In a related study, Parry and
Bruhn (1987) used the same CO2-H2O-NaC1 fluid inclu-
sion system to determine that rocks now at the surface
were once 11 km deep, thereby constraining the amount of
offset along the fault.
Future Prospects for Fluid Inclusion Research
Methane and water are increasingly being recognized
as the most important fluid constituents in subduction zones
(Ernst, 1972; Magantz and Taylor, 1976; Cloos, 1984;
Moore et al., 1987; Vrolijk, 1987~. In reconnaissance
examinations of syntectonic veins from accretionary
complexes in Japan, Washington, and Papua New Guinea,
methane-r~ch and water-r~ch fluid inclusions have been
recognized, suggesting that analysis of fluid inclusions
may prove useful in accretionary complexes around the
world.
PETER VROLIJK AND GEORGIANNA MYERS
Deformation
Front)
Ocean Crust ~
- - -300°C-~ ~ ~
-
Local Thermal
Effect
it,
a
\
Km
Deform ation
Front ~
~, _ , _
Ocea n Cr-ust - ~ ~
~~ 300°C-=~- - -
~ _~~ \
~ `~\
Regional Thermal
Effect
b
-
_ ~
'I\\
Figure 10.7 Extent of heating due to fluid escape in a subduction
zone. Half arrow indicates displacement along the decollement
zone. A: Local thermal effect. The gross temperature structure
of subduction zones can be described by conductive heat flow
models (the 300°C isotherm drawn here is taken from Ernst,
1970), but fluid escape along the decollement zone heats only
rock immediately adjacent the fault zone (i.e., heat conduction
out of the decollement zone is minimal). This thermal configu-
ration only occurs periodically during the life of a subduction
zone and results from episodic fluid flow. B: Extensive regional
thermal effect. Fluids migrating along the decollement zone
have a profound effect on temperatures in the subduction zone,
and isotherms are only mildly depressed as conduction of heat
from the decollement keeps pace with heat conducted into the
subducting plate. In this case fluid flow is more continuous than
in (A), and there is greater net heat flux upward by fluid flow.
The true thermal structure of subduction zones where fluid es-
cape is an important component probably lies somewhere be-
tween these two end-members. Temperatures may be somewhat
higher along the decollement zone, but heat conduction probably
smoothes out this anomaly (from Vrolijk et al., 1988~.
OCR for page 153
FLUID PRESSURE HISTORY IN SUBDUCTION ZONES
CONCLUSIONS
Subduction zones may represent the tectonic environ-
ment most strongly influenced by high fluid pressures.
Deformation, diagenesis, and metamorphism of subducted
water-nch sediments and rocks leads to increasing fluid
pressures because subduction effectively "bunes" maten-
als faster than fluid can escape through intergranular pore
spaces. Evidence from modern and ancient subduction
zone settings is beginning to uncover details of the fluid
pressure history and the chemical and thermal evolution of
fluids. However, this research remains in its infancy.
More information regarding the physical properties of
sediments is required to better understand fluid flow and
its effect on how sediments deform. Research into the
source, path, and migration history of fluids is required to
better understand the gross hydrogeology of subduction
zones and to determine how the fluid pressure history is
intertwined with hydrogeology.
Ancient rocks preserve- a record of phenomena occur-
nng at depths too great to be directly sampled from the
surface. Fluid inclusions in syntectonic veins offer a means
to examine processes occurring on a relatively short time
scale, much shorter than those that can be studied by most
other techniques. Moreover, several studies suggest that
the fluid inclusion method may be more accurate than
other techniques for determining P-T at incipient meta-
morphic conditions. Fluid inclusion research may be
broadly applied to a host of different tectonic problems
around the world.
ACKNOWLEDGMENTS
The research reported here was supported by National
Science Foundation grants EAR 84-07720 and EAR 86-
08337 to J. C. Moore; by ARCO, Union, Sohio, and Mobil
oil companies; and by the U.S. Geological Survey. Many
ideas discussed here evolved from years of discussions
with J. C. Moore and J. C. Sample, although the authors
bear full responsibility for the paper's contents. W. G.
Ernst, H. Gibbons, and M. Reid kindly provided careful
and helpful reviews and comments.
REFERENCES
Angus, S., B. Armstrong, and K. M. de Reuck (19761. Interna-
tional Thermodynamic Tables of the Fluid State 5: Methane,
International Union of Pure and Applied Chemistry, Chemical
Data Series, No. 16, Pergamon Press, New York, 247 pp.
Bird, P. (1984~. Hydration-phase diagrams and friction of
montmorillonite under laboratory and geologic conditions, with
implications for shale compaction, slope stability, and strength
of fault gouge, Tectonophysics 107, 235-260.
155
Bray, C. J., and D. E. Karig (1985~. Porosity of sediments in
accretionary prisms and some implications for dewatering
processes, Journal of Geophysical Research 90, 768-778.
Brown, K. M., and G. K. Westbrook (1987~. The tectonic fabric
of the Barbados Ridge accretionary complex, Marine Petro-
leum Geology 4, 71-81.
Byrne, T. (1984~. Early deformation in melange terranes of the
Ghost Rocks Formation, Kodiak Islands, Alaska, in Me'langes:
Their Nature, Origin, and Significance, L. A. Raymond, ea.,
Special Paper 198, Geological Society of America, Boulder,
Colo., pp. 21-52.
Cloos, M. (1984~. Landward-dipping reflectors in accretionary
wedges: Active dewatering conduits? Geology 12, 519-522.
Colten-Bradley, V. A. (1987~. Role of pressure in smectite
dehydration: Effects on geopressure and smectite-to-illite
transformation, American Association of Petroleum Geolo-
gists Bulletin 71, 1414-1427.
Davis, D. M., and D. M. Hussong (1984~. Geothermal observa-
tions during DSDP Leg 78A, in Initial Reports, Deep-Sea
Drilling Project 78A, B. Biju-Duval and J. C. Moore et al.,
eds., U.S. Government Printing Office, Washington, D.C., pp.
593-598.
Davis, D., J. Suppe, and F. Z. Dahlen (1983~. The mechanics of
fold-and-thrust belts, Journal of Geophysical Research 88,
1153-1172.
Engebretson, D. C., A. Cox, and R. C. Gordon (1984~. Relative
motions between oceanic plates of the Pacific Basin, Journal
of Geophysical Research 89, 10,291-10,310.
Ernst, W. G. (1970~. Tectonic contact between the Franciscan
milange and the Great Valley sequence, crustal expression of
a Late Mesozoic Benioff Zone, Journal of Geophysical Re-
search 75, 886-902.
Ernst, W. G. (1972~. CO2-poor composition of the fluid attend-
ing Franciscan and Sanbagawa low-grade metamorphism,
Geochimica et Cosmochimica Acta 36, 497-504.
Ernst, W. G. (1974~. Metamorphism and ancient continental
margins, in The Geology of Continental Margins, C. A. Burk
and C. L. Drake, eds., Springer-Verlag, New York, pp. 907-
919.
Ernst, W. G. (1979~. Coexisting sodic and calcic amphiboles
from high-pressure metamorphic belts and the stability of
barroisitic amphibole, Mineralogical Magazine 43, 269-278.
Etheridge, M. A., V. J. Wall, and R. H. Vernon (1983~. The role
of the fluid phase during regional metamorphism and
deformation, Journal of Metamorphic Geology 1, 205-226.
Etheridge, M. A., V. J. Wall, and S. F. Cox (1984~. High fluid
pressures during regional metamorphism and deformation:
Implications for mass transport and deformation mechanisms,
Journal of Geophysical Research 89, 4344-4358.
Freeze, R. A., and J. A. Cherry (1979~. Groundwater, Prentice-
Hall Inc., Englewood Cliffs, N.J., 604 pp.
Fyfe, W. S., N. J. Price, and A. B. Thompson (1978~. Fluids in
the Earth's Crust, Developments in Geochemistry 1, Elsevier
Scientific Publishing Co., New York, 383 pp.
Gieskes, J., G. Blanc, P. Vrolijk, J. C. Moore, A. Mascle, E.
Taylor, P. Andreiff, F. Alvarez, R. Barnes, C. Beck, J.
Behrmann, K. Brown, M. Clark, J. Dolan, A. Fisher, M.
OCR for page 154
156
Hounslow, P. McLellan, K. Moran, Y. Ogawa, T. Sakai, J.
Schoonmaker, R. WiLkens, C. Williams, 1989. Hydrogeochem-
istry in the Barbados accretionary complex: Leg 110 ODP,
Palaeogeography, Palaeoclimatology, Palaeoecology 71, 83-
96.
Gill, J. (1981~. Orogenic Andesites and Plate Tectonics, Sprin-
ger-Verlag, New York, 400 pp.
Hall, P. L., D. M. Astill, and J. D. C. McConnell (1986~. Thermo-
dynamics and structural aspects of the dehydration of smec-
tites in sedimentary rocks, Clay Minerals 21, 633-648.
Hottman, C. E., J. H. Smith, and W. R. Purcell (1979~. Relation-
ships among Earth stresses, pore pressure, and drilling prob-
lems, offshore Gulf of Alaska, Journal of Petroleum Technol-
ogy 31, 1477-1484
Hubbert, M. K., and W. W. Rubey (1959~. Role of fluid pressure
in the mechanics of overthrust faulting, I: Mechanics of fluid-
filled porous solids and its application to overthrust faulting,
Geological Society of America Bulletin 70, 115-166.
Ito, E., D. M. Harris, and A. T. Anderson (1983~. Alteration of
oceanic crust and geologic cycling of chlorine and water,
Geochimica et Cosmochimica Acta 47, 1613- 1624.
Johannes, W., and D. Puhan (1971~. The aragonite-calcite tran-
sition, reinvestigated, Contributions to Mineralogy and Pe-
trology 31, 28-38.
Koster van Groos, A. F., and S. Guggenheim (1984~. The effect
of pressure on the dehydration reaction of interlayered water
in Na-montmorillonite, American Mineralogist 69, 872-879.
Koster van Groos, A. F., and S. Guggenheim (1986~. Dehydra-
tion of K-exchanged mont-morillonite at elevated temperature
and pressures, Clays and Clay Minerals 34, 281-286.
Koster van Groos, A. F., and S. Guggenheim (19871. Dehydra-
tion of a Ca- and a Mg- exchanged montmorillonite (SWy-l)
at elevated pressures, American Mineralogist 72, 292-298.
Kulm, L. D., E. Suess, J. C. Moore, 13. Carson, B. T. Lewis, S. D.
Ritger, D. C. Kadko, T. M. Thornberg, R. W. Embley, W. D.
Rugh, G. J. Mossoth, M. G. Langseth, G. R. Cochran, and R.
L. Scamman (1986~. Oregon subduction zone: Venting, fauna,
and carbonates, Science 231, 561-566.
Langseth, M. G., G. K. Westbrook, and M. A. Hobart (l988~.
Geophysical survey of a mud volcano seaward of the Barbados
Ridge accretionary complex, Journal of Geophysical Research
93, 1049-1061
Liou, J. G. (1971~. P-T stabilities of laumontite, waikarite,
lawsonite, and related minerals in the system CaAl2Si2O~-
SiO2-H2O, Journal of Petrology 12, 379-411.
Liou, J. G., S. Maruyama, and M. Cho (19851. Phase equilibria
and mineral parageneses of metabasites in low-grade
metamorphism, Mineralogical Magazine 49, 321-333.
Magaritz, M., and H. P. Taylor, Jr. (19761. Oxygen, hydrogen,
and carbon isotope studies of the Franciscan formation, Coast
Ranges, California, Geochimica et Cosmochimica Acta 40,
215-234.
Maresch, W. V. (1977~. Experimental studies on glaucophane:
An analysis of present knowledge, Tectonophysics 43, 109-
125.
Moore, J. C. (1975~. Selective subduction, Geology 3, 530-532.
Moore, J. C., and A. Allwardt (1980~. Progressive deformation
PETER VROLIJK AND GEORGIANNA MYERS
of a Tertiary trench slope, Kodiak Islands, Alaska, Journal of
Geophysical Research 85, 5741-4756.
Moore, J. C., and B. 13iju-Duval (1984~. Tectonic synthesis
Deep Sea Drilling Project Leg 78A: Structural evolution of
offscraped and underthrust sediment, northern Barbados Ridge
complex, in Initial Reports, Dea-Sea Drilling Project 78A, B.
Biju-Duval, J. C. Moore et al., eds., U.S. Government Printing
Office, Washington, D.C., pp. 601-621.
Moore, J. C., and R. W. Wheeler (1978~. Structural fabric of a
melange, Kodiak Islands, Alaska, American Journal of Sci-
ence 278, 739-765.
Moore, J. C., A. Mascle, E. Taylor, P. Andreiff, F. Alvarez, R.
Barnes, C. Beck, J. Behrmann, G. Blanc, K. Brown, M. Clark,
J. Dolan, A. Fisher, J. Gieskes, NI. Hounslow, P. McLellan, K.
Moran, Y. Ogawa, T. Sakai, J. Schoonmaker, P. Vrolijk, R.
Wilkens, and C. Williams (19871. Expulsion of fluids from
depth along a subduction-zone decollement horizon, Nature
326, 785-788.
Mullis, J. (1976~. Das Wachstumsmilieu der Quarzkristalle im
Val d'Illiez (Walks, Schweiz), Schweiz. Mineral. Petrogr. Mitt.
56, 219-268.
Mullis, J. (1979~. The system methane-water as a geologic
thermometer and barometer from the external part of the Central
Alps, Societe Francaise de Mineralogie et de Cristallogra-
phie, Bulletin 102, 526-536.
Mullis, J. (1988~. Rapid subsidence and upthrusting in the North-
ern Appenines deduced by fluid inclusions studies in quartz
crystals from Poretta Terme, Schweiz, Mineral. Petrogr. Mitt.
68, 157-170.
Myers, G. (1987~. Fluid expulsion during the underplating of the
Kodiak Formation: A fluid inclusion study, M.S. thesis, Uni-
versity of California, Santa Cruz, 41 pp.
Nitsch, K.-H. (19721. Das P-T-XCO2-stabilitaetsfeld von lawsonit,
Contributions to Mineralogy and Petrology 34, 116-134.
Orkan, N., and B. Voight (1985~. Regional joint evolution in the
Valley and Ridge Province of Pennsylvania in relation to the
Alleghany Orogeny, in Guidebook for the 50th Annual Field
Conference of Pennsylvania Geologists, Bureau of Topographic
and Geological Survey, Harrisburg, Pa., pp. 144-164.
Oxburgh, E. R., and D. L. Turcotte (1971~. Origin of paired
metamorphic belts and crustal dilation in island arc regions,
Journal of Geophysical Research 76, 1315- 1327.
Parry, W. T., and R. L. Bruhn (1986~. Pore fluid and seismo-
genic characteristics of fault rock at depth on the Wasatch
Fault, Utah, Journal of Geophysical Research 91, 730-744.
Reck, B. H. (1987~. Implications of measured thermal gradients
for water movement through the northeast Japan a~ccretionary
prism, Journal of Geophysical Research 92, 3683-3690.
Rimstidt, J. D., and H. L. Barnes (1980~. The kinetics of silica-
water reactions, Geochimica et Cosmochimica Acta 44, 1683-
1699.
Ritger, S., B. Carson, and E. Suess (1987~. Methane-derived
authigenic carbonates formed by subduction-induced pore-water
expulsion along the OregonlWashington margin, Geological
Society of America Bulletin 98, 147- 156.
Sample, J. C., and J. C. Moore (19871. Structural style and
kinematics of an underplated slate belt, Kodiak Islands, Alaska,
Geological Society of America Bulletin 99, 7-20.
OCR for page 155
FLUID PRESSURE HISTORY IN SUBDUCTION ZONES
Secor, D. T. (1965~. Role of fluid pressure in jointing, American
Journal of Science 263, 633-646.
Seely, D. R. (1977~. The significance of landward vergence and
oblique structural trends on trench inner slopes, in Island Arcs,
Deep Sea Trenches, and Back-Arc Basins, M. Talwani and S.
C. Pitmann, eds., Maurice Ewing Series 1, American Geo-
physical Union, Washington, D.C., pp. 187-198.
Shepherd, L. E., and W. R. Bryant (1983~. Geotechnical proper-
ties of lower trench inner-slope sediments, Tectonophysics 99,
279-312.
Shi, Y., and C.-Y. Wang (19881. Generation of high pore pres-
sures in accretionary prisms: Inferences from the Barbados
subduction complex, Journal of Geophysical Research 93,
8893-8910.
Shouldice, D. H. (1971~. Geology of the western Canadian
continental shelf, Canadian Petroleum Geology Bulletin 19,
405-436.
Suess, E., B. Carson, S. Ritger, J. C. Moore, M. Jones, L. D.
Kulm, and G. Cochrane (1985~. Biological communities at
vent sites along the subduction zones off Oregon, Bulletin of
the Biological Society of Washington 6 (special issue, The
Hydrothermal Vents of the Eastern Pacific: An Overview, M.
L. Jones, ed.), 475-484.
van Huene, R. (1972~. Structure of the continental margin and
tectonism at the Eastern Aleutian Trench, Geological Society
of America Bulletin 83, 3613-3626.
van Huene, R., M. Lan~seth, N. Nasu, and H. Okada (1980~.
Summary, Japan Trench Transect, in Initial Reports, Deep-Sea
Drilling Project 56 and 57 (part 1), M. Lee and L. Stout, eds.,
U.S. Government Printing Office, Washington, D.C., pp. 473-
488.
157
van Huene, R., S. Box, R. Detterman, M. Fisher, J. C. Moore,
and H. Pulpan (1985~. A-2 Kodiak to KuskoLwin, Alaska,
Continent/Ocean Transect, vol. 6, Geological Society of
America, Boulder, Colo., 1 sheet, scale 1:500,000.
Vrolijk, P. J. (1986~. Channelized fluid flow along melanges of
the Ghost Rocks Fm., Kodiak accretionary complex, Alaska
(abs.), EOS 67, 1205.
Vrolijk, P. J. (1987~. Paleohydrogeology and fluid evolution of
the Kodiak accretionary complex, Alaska, Ph.D. thesis, Uni-
versity of California, Santa Cruz, 232 pp.
Vrolijk, P., G. Myers, and J. C. Moore (1988~. Warm fluid
migration along tectonic melanges in the Kodiak accretionary
complex, Alaska, Journal of Geophysical Research 93, 10,313-
10,324.
Walder, J., and A. Nur (1984~. Porosity reduction and crustal
pore pressure development, Journal of Geophysical Research
89, 11539-11548.
Wang, C.-Y., and Y.-L. Shi (19841. On the thermal structure of
subduction complexes: A preliminary study, Journal of Geo-
physical Research 89, 7709-7718.
Wells, R. E., D. C. Engebretson, P. D. Snavely, Jr., and R. S. Coe
(1984~. Cenozoic plate motions and the volcano-tectonic
evolution of western Oregon and Washington, Tectonics 3,
275-294.
Westbrook, G. K., and M. J. Smith (1983~. Long decollements
and mud volcanoes: Evidence from the Barbados Ridge
Complex for the role of high pore-fluid pressure in the devel-
opment of an accretionary complex, Geology 11, 279-283.
Yamano, M., S. Uyeda, Y. Aoki, and T. H. Shipley (1982~.
Estimates of heat flow derived from gas hydrates, Geology 10,
339-343.
OCR for page 156
11
INTRODUCTION
Degassing of Carbon Dioxide as a
Possible Source of
High Pore Pressures in the Crust
JOHN D. BREDEHOE1?T and STEVEN E. INGEBRITSEN
U. S. Geological Survey, Menlo Park
Increased pore pressures, especially pore pressures
approaching lithostatic loads, change the state of effective
stress and greatly reduce the work necessary for tectonic
deformation (Hubbert and Rubey, 1959; Rubey and Hub-
bert, 1959~. A number of mechanisms have been proposed
that increase pore pressures (Hanshaw and Zen, 1965~.
One of the more interesting of the proposed mechanisms is
the movement of carbon dioxide (CO2) through the crust,
suggested by Irwin and Barnes (Irwin and Barnes, 1975;
Barnes et al., 1978, 1984~. The purpose of this chapter is
to investigate the possible role of CO2 as a source of high
pore pressure by examining the following question: Given
the current best estimate of the rate of CO2 degassing, how
low would the permeability have to be in order to generate
pore pressures approaching lithostatic values?
Barnes et al. (1978, 1984) compiled worldwide data on
the distribution of CO2 discharge from the crust. They
showed that most of the discharges are concentrated in
two areas of the world: (1) a narrow circum-Pacific belt
and (2) a broad mountainous area that extends across central
and southern Europe and Asia Minor. They went on to
note that the CO2 discharge areas coincide with areas that
are seismically active at the present time. Their clearest
statement of the role of CO2 in tectonic processes was
presented in Irwin and Barnes (1975~.
158
There is a continuing debate about the nature of vola-
tiles in the mantle. Gold (see, e.g., Gold and Soter, 1980)
argues that the carbon in the mantle is present largely as
methane. On the other hand, most of the geological
community reports the observations of carbon emanating
from the mantle as CO2 (see, e.g., Leavitt, 1982; Gerlach,
1988~. The current consensus of the geological commu-
nity seems to be that CO2 is the more likely form of
volatile carbon in the mantle. For the purpose of our
simple experiment we have assumed a source of free CO
either in the mantle or within the crust.
Carbon dioxide is thought to come from three different
sources: (1) organic material, (2) metamorphism of marine
carbonate rock, and (3) degassing of the mantle (Barnes et
al., 19841. Each source is thought to have a different ratio
of carbon isotopes. The evidence for the isotopic compo-
sition of mantle-derived CO2 comes from analyses of fluid
inclusions from volcanic rocks erupted along oceanic
spreading ridges. Moore et al. (1977) reported 6~3C values
ranging from - .7 to -5.8 Boo for CO2 inclusions from
basalts in the Pacific. Pineau e! al. (1976) found similar
values (SAC of -7.6 + 0.5 ~) for fluid inclusions in
tholeiitic rocks from the mid-Atlantic ridge. Carbon diox-
ide with 6~3C in the range of - .7 to -8.0 9~ is thought to
indicate mantle-derived CO2, although the carbon isotope
ratios alone do not provide an unambiguous indication of
--2
OCR for page 157
SOURCE OF HIGH PORE PRESSURES IN THE CRUST
a mantle denvation. 6~3C in the range of - .7 to -8.0 Boo
can also be derived by mixing an organic CO2 source,
typically -20 Boo, with a marine carbonate CO2 source,
typically 0 tow. However, in the case of the oceanic basalts
a mantle source seems to be indicated as carbon isotope
data is combined with other geologic information in order
to interpret the source of the CO2.
Magmatic degassing may be a major source of free CO2
within the crust. Harris (1981) showed that the solubility
of CO2 in tholeiitic basalts was strongly pressure depend-
ent; thus, CO2 that is in solution in magma at great depth
will exsolve as the magma migrates upward in the crust
and the pressure is reduced. Gerlach (1986, 1988) sug-
gested that most of the CO2 is degassed from plutons at
depths of 5 to 10 or 12 km. Basaltic magmas associated
with hot spots, such as Kilauea, have a higher CO2 content
than mid-ocean ridge basalts (Gerlach, 1988), so that they
will tend to degas at higher pressures and greater depths.
GLOBAL FLUX OF CO2
To estimate the permeability necessary to maintain near-
lithostatic pore pressure, one must first estimate the flux of
CO2. A number of investigators have made estimates of
the CO2 flux from deep in the Earth's interior to the sur-
face. Leavitt (1982) summarized earlier estimates; we
have extended Leavitt's summary to include more recent
data. The various estimates are summarized in Table 11.1.
They range from 1.7 x 10~° to 2 x 10~3 moles of CO2 per
year; most estimates are in the range of 10~2 moles per
year. The scatter in the estimates is not as wide as one
159
might expect given the nature of the data.
If one assumes that the ultimate source of all carbon in
the Earth is the mantle, one can place the current estimates
of CO2 flux in perspective by calculating a constant flux
rate that would generate the known carbon reservoir.
Current estimates of the global carbon reservoir are sum-
mar~zed by Sundquist (1985~. One of these estimates
(perhaps the best) is presented in Table 11.2. Assuming a
constant rate of outgassing over 4.5 billion years yields a
rate of outgassing of 2 x 10~2 moles of CO2 per year.
These are grossly simple assumptions; however, the calcu-
lation suggests that the current estimated rate of CO2
outgassing would approximately account for the global
carbon reservoir. Other workers (e.g., Marty and Jambon,
1987; Gerlach, 1988) have also commented that the cur-
rent rate is sufficient to generate the total carbon reservoir.
There are obvious complications, for example, differ-
ing rates of degassing in the geologic past and recycling of
carbonate rocks in subduction zones. Des Marais (1985)
noted that a rate of 10 x 10~2 moles of CO2 per year from
the mantle would generate the global carbon inventory in
approximately 700 million years and went on to point out
the subduction tends to recycle carbon from the crust into
the mantle. He estimates that perhaps half of the current
flux is recycled carbon.
MODELING CO2 FLUX IN THE CRUST
We have investigated the effects of a CO2 flux at mid-
crustal depths utilizing a numerical simulation model. The
model simulates the simultaneous transport of mass and
TABLE 11.1 Reported CO2 Inputs into the Atmosphere from the Earth's Interior
C(~2 Released
Reference CO2 Source (10'2 moles/yr)
Borchert (1951) Igneous and metamorphic 6.7
Igneous only 1.3
Rubey (1951) Total "excess CO2" since Earth's origin, 0.5
including that from hot springs
Plass (1956) CO2 "released from the interior of the Earth" 2
Li (1972) Total CO2 at Earth's surface 1.0
Libby and Libby (1972) Volcanic CO2 0.017
Buddemeier and Puccetti Hawaiian estimate 2.0
(1974)
Anderson (1975) Outgassing of oceanic crust 0.23
Baes et al. (1976) Volcanoes, fumaroles, and hot springs 1.7 to 8.3
Leavitt (1982) Volcanic eruptions 0.15
Javoy et al. (1982) Mid-ocean ridges 20
Des Marais (1985) Mid-ocean ridges 1 to 8
Marty and Jambon (1987) C/3He, mid-ocean ridges 2.0
Gerlach (1988) Mid-ocean ridges 0.3
Representative terms from entire chapter:
fluid inclusions
