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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
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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
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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
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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
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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~.
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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~. 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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.
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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. 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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~. 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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.
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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
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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: