The role of water in the evolution of the continental crust Journal of the

Journal of the Geological Society The role of water in the evolution of the continental crust Bruce W.D. Yardley Journal of the Geological Society 2009; v. 166; p. 585-600 doi:10.1144/0016-76492008-101

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University of Leeds on 20 July 2009

© 2009 Geological Society of London

Journal of the Geological Society, London, Vol. 166, 2009, pp. 585–600. doi: 10.1144/0016-76492008-101.

Review The role of water in the evolution of the continental crust B RU C E W. D. YA R D L E Y School of Earth and Environment, University of Leeds, Leeds LS2 9JT, UK *Corresponding author (e-mail: [email protected]) Abstract: Aqueous fluids have a profound influence on the evolution of the crust, both as agents of chemical mass transfer and mineral reactions, and through modifying its rheology. This paper is particularly concerned with the composition and role of fluids through the crustal cycle of burial, metamorphism and uplift. Information now comes from both conventional geological studies and from geophysics, which has documented both the presence of fluid in active areas of the crust, and also its migration in response to faulting. This contrasts with the overall slow rate of some fluid processes: for example, devolatilization reactions are endothermic and limited by the rate of heat supply. There is no clear distinction between diagenetic and metamorphic fluids, and extensive equilibration with host rocks means that few chemical or isotopic characteristics survive to provide tracers for deep fluid origins. After the metamorphic peak, remaining pore fluid is consumed in hydration reactions resulting in strong, dry rock. Subsequently, infiltration may reintroduce fluids but these normally survive only briefly until consumed by hydration reactions. Metamorphic rocks are strongly overpressured during prograde metamorphism and therefore have a low permeability. Considerable doubt is therefore cast on some claims of widespread and pervasive large fluid fluxes accompanying metamorphism.

to determine about fluid behaviour in dynamic crustal settings today. The second part of the paper concerns the role of water in the metamorphic cycle of orogenesis, with the focus on crustal metamorphic processes. It is beyond the scope of the paper to do more than touch briefly on the cycling of water into the mantle during subduction and its return in magmatic fluids. It is well known that diagenetic and metamorphic processes are profoundly influenced by a fluid phase, even if present in only very small amounts, and this may operate from the nanoscale to the regional scale. Water plays a vital part in facilitating the reaction or recrystallization of minerals on the grain or grain boundary scale (David et al. 1995; Putnis et al. 2007a,b), and it is likely that most mineral reactions are extremely limited or impossible in a truly dry system over the normal range of crustal temperatures. In this paper, however, I will concentrate on the role of fluid on a larger scale, and in particular on the consequences of the presence or absence of a physically distinct aqueous fluid phase (as opposed to adsorbed water molecules in defects, for example), which can move in response to gradients in hydraulic head. Strictly, of course, hydraulic head is an abstract concept in deep crustal settings, which are strongly overpressured (i.e. fluid pressure exceeds that which would be generated by a column of overlying water extending to the Earth’s surface); however, it is of value in that it emphasizes that fluid flow does not always proceed downpressure (Ingebritsen et al. 2006). Its applicability to some crustal fluid problems is illustrated in Figure 1.

Aqueous fluids exert a fundamental control on crustal processes; indeed, the role of water in geology is just as pivotal as its role in biology. There are many well-documented examples of crystalline rocks spending extended periods of time at high pressures and extreme temperatures deep in the crust, while undergoing only very limited metamorphic change, except locally where water was introduced (Austrheim 1987; John & Schenk 2003). Likewise, where crystalline rocks at shallow levels undergo extensive transformation into assemblages stable at low T and P, there is infiltration of water (Taylor 1978; Moore et al. 1983; David et al. 1995). Fluids can dominate both physical processes through their influence on rheology, and chemical ones through their capacity to dissolve and precipitate minerals. In the uppermost crust we are used to the idea of fluid behaviour being dominated by intrinsic factors such as the porosity and permeability of host rocks, but with increasing depth and temperature the dynamic interactions between fluid and rock modify rock properties, so that an important objective in understanding crustal fluid processes is to identify which are the intrinsically independent variables, and which other variables respond to them. In this paper I shall explore some aspects of both the physical and chemical behaviour of crustal fluids with this distinction in mind. The main concern of this paper is the role of fluids in the formation and exhumation of crystalline rocks, but to understand this, it is important to also consider, at least briefly, the role of relatively shallow fluids present in sedimentary basins or circulating in the upper crust in response to igneous activity. These shallow fluids evolve into deep fluids in specific circumstances, and probably also form a reservoir into which deep fluids returning to the surface are dissipated. Furthermore, we know rather more about them than we do of fluids that are beyond the reach of drilling. In the first part I shall summarize what are essentially equilibrium aspects of fluids in different types of crustal settings, then I will try to put these in the context of what we are now able

The evidence for deep crustal fluid activity today Direct evidence from drilling and sampling of fluids comes from the top few kilometres of the continental crust and continental shelf, and generally from just a few hundred metres beneath the floor of the ocean basins. To supplement that, however, we can 585

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Fig. 1. The concept of hydraulic head applied to the crust as a whole. Vertical columns are notional pipes inserted into the crust, and have a water level indicated to represent the level to which water would rise. This stylized geological section shows to the left an area of crystalline basement reaching the surface, to the right a sedimentary basin with some deformation and metamorphism at depth. For hydrostatically pressured regions water level in the pipe corresponds to the water table or sea level irrespective of the depth to the base of the pipe (pipes A and C), whereas in overpressured regions the water level rises to a higher level, as the fluid overpressure gives rise to an excess hydraulic head. Typically, hydraulic head increases with depth in overpressured regions (pipes D and E). Very low hydraulic heads can arise in cooled crystalline rocks that are effectively isolated from sources of water and have the potential to consume water through hydration reactions (pipe B).

draw on a range of observations of rocks now exposed at the surface but displaying evidence of past fluid activity, and to an increasing degree on geophysical observations indicative of fluid activity at depth in the crust at the present day.

Drilled systems Sedimentary basins. Not surprisingly, we know most about the occurrence of deep groundwaters in sedimentary basins. Here, there is a progressive reduction in porosity with depth, reflecting both physical and chemical compaction. Although in shallow settings, fluid pressure is close to hydrostatic, higher fluid pressures (overpressures) become increasingly common at greater depths in sedimentary basins (Fig. 1), especially where there are relatively impermeable horizons to provide cap rocks. Bredehoeft & Hanshaw (1968) showed that in the Gulf of Mexico such overpressures developed as a result of the dynamic balance between the rate of sedimentation and the flow of fluid out of the sediment column. Because compaction is a function of effective stress in a fluid-saturated formation, rather than of vertical stress, overpressured rocks often have anomalously high porosities for their depth of burial. This is known as delayed compaction (e.g. Harrison & Summa 1991), and in turn provides the basis for the recognition of overpressured zones in seismic sections. In many continental margin basins, the aqueous fluids are essentially brines, and salinity commonly increases with depth (Hanor 1994). Salinity can be inherited from original pore waters or modified by meteoric water infiltration along seaward-dipping beds that crop out on land, by dissolution of evaporites, or by sinking of dense brines from the surface through the sediment column during times of low sea level and evaporation. In contrast, cation chemistry reflects diagenetic interactions with

minerals, and Hanor (1980) has argued that formation waters can be seen as in chemical equilibrium with diagenetic minerals, notwithstanding their low temperatures. The chemical evidence for the evolution of pore waters has been summarized in a number of important papers (Rittenhouse 1967; Hanor 1980, 1994; Land 1995). Not surprisingly, extensive changes in fluid chemistry in relatively porous and permeable sediments are complemented by changes in rock chemistry. Land & Milliken (2000) have demonstrated significant changes in sediment composition accompanying illitization of Gulf Coast sandstones, and widespread diagenetic changes such as albitization of feldspars likewise take place in an open system at this stage. In contrast to most oilfields, drilling in the ocean basins has encountered less saline fluids than seawater, inferred to be due to the release of water from clay breakdown (e.g. Gieskes et al. 1990; Mottl et al. 2004). This contrasting result may reflect both the absence of dense, saline brines and the relatively rapid sedimentation rates in many such settings. Geothermal systems. A further source of knowledge about crustal fluids comes from drilling of geothermal fields associated with active volcanoes for geothermal energy resources (Nicholson 1993). Drilled geothermal systems are generally normally pressured, with convecting water circulating to depths of 2–4 km, corresponding to temperatures of c. 350 8C. Although it was formerly believed that geothermal waters were of magmatic origin, the first systematic stable isotope studies of geothermal waters (White 1957) demonstrated conclusively that the water itself was predominantly derived from local rainwater. This does not preclude the possibility that magmatic water is also present, albeit extensively diluted, and other tracers such as noble gas isotopes may have an essentially magmatic signature because they represent components that are much less abundant in meteoric waters. Geothermal convection systems develop in the permeable upper crust, making use of fracture permeability or the high intrinsic permeability of pyroclastic rocks or vesiculated lavas. The temperature structure in the upper part of high-temperature geothermal fields often closely tracks the boiling point–depth curve of water, increasing rapidly below the water table, but only slowly at greater depths (Fig. 2a). It has been pointed out by Fournier (1989) that at Yellowstone the base of the geothermal circulation system lies just a few kilometres above the roof of the magma chamber, and temperature must therefore increase rapidly in the intervening distance (Fig. 2b). Hence the geothermal system corresponds to a region of fractured rocks undergoing intense hydrothermal alteration at relatively low temperatures, with extensive metasomatic alteration caused by the large volumes of circulating water, but this lies above a volume of hotter rocks with a steep thermal gradient (i.e. a classic metamorphic aureole). If the aureole rocks were initially sediments, it is likely that fluid will be strongly overpressured in the aureole, as metamorphic dehydration and thermal expansion of water combine to make the contact metamorphic setting the one with the strongest tendency to develop overpressure through increased fluid volume, or geological forcing (Neuzil 1995). The steep thermal gradient makes it very unlikely that there is effective convection of fluid in the aureole. Circulating groundwater would be unable to penetrate such an aureole while it remained overpressured. Where the country rocks were initially more or less dry (e.g. earlier lavas), so that the contact aureole is less likely to become overpressured, then there is evidence that groundwater may be able to penetrate at high temperatures (e.g. Manning & Bird 1991); in this case, the steep temperature gradient may reflect both limited circulation and the low density of water at high temperatures close to the surface.

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Continental deep drilling. Deep drilling has been carried out at a few continental sites of contrasting tectonic setting. The Kontinentale Tiefbohrprogramm der Bundesrepublik Deutschland (KTB) demonstrated that the upper crystalline crust in a relatively stable setting is rather dry. Where fluid occurs in the deeper parts (.4 km), it evidently resides in poorly connected fractures rather than pervasively occupying grain boundaries (ELEKTB Group 1997; Mo¨ller et al. 1997), and is chemically distinct from shallow groundwater systems (Grawinkela & Stockhert 1997). This result is consistent with what is known of deep mine waters from shield areas, where deep groundwaters are also highly evolved fluids that have equilibrated with their host rocks despite the rather low temperatures (Frape & Fritz 1987). In contrast, detailed magnetotelluric investigations of the crust in the vicinity of the San Andreas fault suggest that, unsurprisingly, there is much more extensive infiltration of water in the vicinity of a major fault than in the crust at large (Bedrosian et al. 2004), with evidence for water penetration to several kilometres (Unsworth & Bedrosian 2004).

Geophysical evidence for crustal fluids today

Fig. 2. Geothermal gradients associated with magmatic geothermal activity. (a) Temperature–depth relationships for a high-temperature geothermal field. The boiling point–depth relationship is shown by curve 1 for a water table at the surface, and curve 2 where the water table is at some depth. Curve 3 is a typical geothermal gradient corresponding to the second case and initially lies close to the boiling curve so that rising water boils. At greater depths the temperature is more nearly constant, reflecting circulation of hot water. (b) Overall temperature gradient above a magmatic body. The upper part corresponds to a geothermal system as illustrated in (a), but there is a steep temperature rise in low-permeability material from the base of the geothermal field to the much hotter magma body, implying that conductive heat transfer predominates below the geothermal field. This region corresponds to a conventional metamorphic aureole, recording large temperature variations over short distances (based on Fournier 1989).

It is now becoming possible to test ideas about the fluid history of the crust deduced from geological observations of past events against direct geophysical observations. Investigations in a number of tectonically active areas have found direct evidence of deep crustal fluids through magnetotelluric studies (e.g. Wannamaker et al. 2002; Bedrosian et al. 2004) and through measurements of seismic velocities (Vp and Vp /Vs ratios) (e.g. ANCORP Working Group 2003). It seems clear that in tectonically active areas there may be areas of fluid saturation in the vicinity of reverse fault systems (Fuis et al. 2003), and in areas of anomalously high porosity in convergent zones (Hyndman 1988; Bangs et al. 1990). Not only has the presence of fluid been imaged in some specific settings, but transient behaviour suggesting fluid flow activity on a time scale of months to years has been reported; for example, from the vicinity of the Barbados Ridge de´collement at the base of the accretionary complex (Bangs et al. 1996). Direct measurements of fluid flow in the forearc offshore Costa Rica have identified correlations between fluid flow and seismic tremors, with three distinct episodes of tremors associated with anomalous flow within a 6 month monitoring period (Brown et al. 2005). Although the geophysical evidence for the presence of fluids in tectonically active parts of the crust is now well documented, there has been considerable controversy about other examples. A number of earlier studies of stable continental areas equated high conductivity, and in some cases seismic reflectors, with fluid in the deep crust (Hyndman & Shearer 1989). It will be argued below that this scenario is highly improbable because cooled crystalline rocks are not stable in the presence of water.

Geochemical tracers of modern deep fluid processes Notwithstanding the close correlation between water involved in upper crustal fluid processes and the local rainfall, there are several trace fluid components for which distinct deep signatures can be identified, and many studies have identified He, Ar and CO2 derived from mantle sources, primarily in areas of active magmatism but also in other areas where deep structures provide access to the surface for deep fluids (Kennedy & van Soest 2007; Wiersberg & Erzinger 2007). Furthermore, detailed correlation of geophysical and isotopic studies of the Cheb Basin suggests that seismic activity is able to dislodge gases from continental rocks, resulting in mixed mantle- and continent-derived gas

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fluxes at the surface following periods of tremors (Bra¨uer et al. 2003). The existence of fluids at the surface with distinctive mantle isotopic signatures for the gas components provides a clear demonstration that fluids released at depth are able to rise through the crust, and their association with magmatism and with major faults provides an indication of their major transport pathways. Nevertheless, considerable caution is needed in extrapolating these results to the behaviour of water. Deep-derived gases are able to travel independently of water because of their limited solubility, and their different wetting characteristics mean that they may not travel along the same pathways. However, a fundamental difficulty is that if deep-derived waters (as opposed to noble gases) do rise into the upper crust, they almost certainly represent such small fluid volumes relative to upper crustal groundwaters that they become diluted and impossible to trace. Although deep waters may be very difficult to detect at the surface, aqueous fluids released from subducting slabs almost certainly play a role in the generation of magmas from the overlying wedge, and some chemical signatures of arc magmas have been attributed to fluid transport of slab components into the mantle wedge (Elliott et al. 1997; Brenan et al. 1998; Kessel et al. 2005).

morphic or seawater origin on the basis of distinctive chemical and isotopic characteristics: for example, they define distinctive fields in terms of äD and ä18 O, and are readily distinguished in fluid-rich systems (Fig. 3a, Sheppard 1986). Palaeofluid compositions may be measured directly on fluid inclusions, or calculated from the compositions of hydrothermal minerals using appropriate fractionation factors. Where there has been any significant fluid–rock interaction, ä18 O values of the fluid reflect interaction with the host rock, not just the original source of the water molecules. Although äD is much more resistant to modification, because there is much less H in minerals than in fluid, extensive reaction can modify äD also, and exchange with hydrocarbons can lead to very low äD values. The formation waters plotted in Figure 3a are mostly from onshore basins in North America, and have light äD values probably indicative of infiltration of meteoric water, whereas Gulf Coast waters plot near seawater for äD. Almost all formation waters show some offset in ä18 O through mineral–fluid interactions, and fluids that have under-

Deep fluid processes from the geological record Studies of modern fluid processes are beginning to revolutionize our understanding of fluid processes in the Earth, particularly for their insights into rates, but much of our knowledge still comes from studies of past events on the geological record. Whereas some of the fluid events that are preserved there can be closely correlated to modern tectonic and volcanic processes, others have no obvious short-term analogues that might be detected from the surface today.

Sources of aqueous crustal fluids There are three primary sources for the aqueous fluids that interact with crustal rocks at different stages in their development: water incorporated into sediments or other surface rocks and released during subsequent burial through expulsion from pores or metamorphic breakdown of hydrous minerals, surface waters (marine or meteoric) that penetrate the crust directly along fractures, and deep-derived water that is returned from the mantle or lower crust in solution in silicate melts, and released as they crystallize. To a first approximation, crustal fluids are often categorized as of metamorphic, surface or magmatic origin. The term connate water has been used to describe sedimentary formation waters that are believed to represent the preserved pore fluids from the time of deposition, but it seems likely that true connate waters are very rare except at very shallow levels, and so the term is not used here. Categorizations of fluid origins are often oversimplified: surface waters that enter the sedimentary column as pore waters probably merge continuously with water released from mineral lattices and are here linked with them as buried waters. Surface waters that penetrate pre-existing rocks are here termed infiltrated waters. The special case of infiltrated waters that flush sedimentary formations, replacing original pore waters, would be considered as buried waters during the later history, should the basin experience subsequent heating. Magmatic waters, released during the crystallization of melts, are also important for some crustal processes but are largely beyond the scope of this paper. Chemical and isotopic signatures of fluid origin. Fluids are commonly identified as being of meteoric, magmatic, meta-

Fig. 3. Isotopic and chemical tracers of the origins of fluids and their dissolved loads. (a) äD–ä18 O plot showing the meteoric water line and fields for magmatic and ‘metamorphic’ fluids, together with a range of sedimentary formation waters. After Sheppard (1986). (b) Plot of the halogen ratios Br/Cl and I/Cl showing the contrasting halogen signatures of seawater, magmatic fluids and brines derived by evaporation of seawater (Br/Cl . seawater) as opposed to those resulting from dissolution of halite beds (Br/Cl , seawater). Based on Bohlke & Irwin (1992), with additional data from Nissenbaum (1977), McDonough & Sun (1995), Banks et al. (2000) and Kendrick et al. (2006).

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gone extensive reaction with crustal rocks at elevated temperatures have ä18 O values that reflect their host and the temperature of interaction; these compositions were characterized as ‘metamorphic fluid’ by Sheppard (1986). This categorization does not distinguish between buried waters and infiltrated waters, as it is based simply on equilibrium with metamorphic rocks at modest crustal temperatures (Sheppard 1986). The different effectiveness with which fluid–rock interactions can modify stable O and H isotope characteristics of fluids demonstrates that the concept of ‘source’ for a crustal fluid is not always a simple one. As a fluid moves through the rock mass, it exchanges some components very readily, because they are abundant in the mineral phases, but others much less readily. These are known as the conservative components, and in addition to H include the halogens, in particular Cl and Br. Br and Cl are readily decoupled only by the growth or dissolution of halite, with brines formed from dissolution of halite having very low Br:Cl ratios compared with bittern brines left as halite precipitates (Fig. 3b). Bohlke & Irwin (1992) showed that the coupled Br/Cl and I/Cl signatures of fluid provide a useful discriminator between different fluids, with I values being very low where biological activity has been able to remove I from fluids, and rising as a result of fluid–rock interaction. Halogens are particularly important tracers because Cl is the main complexing ligand for most ore metals. In summary, different components in a fluid may have been picked up at different stages along the flow path, with äD, salinity and Br/Cl ratios most likely to reflect a fluid’s ultimate origins (Fig. 3b), whereas ä18 O and Sr isotope ratios are often indicative of the local environment. The same fluid may yield different information about its flow path by the use of different tracers: Pb-isotopes have been used to identify the sources of Pb in limestone-hosted deposits because the limestones have little potential to modify the Pb isotopic composition of the fluid as it passes through them (Doe & Stacey 1974). In contrast, Srisotopes would be unlikely to point further back along the flow path than the adjacent Sr-rich carbonate wall rocks.

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result of collapse of porosity. Except where impermeable seals result in the production of overpressures, porosities of sediments decrease rapidly with depth in sedimentary basins, so that for many clastic sediments, the bulk of the water present at depths greater than c. 5 km is likely to be structurally bound in phyllosilicates, although the salt content resides in the pore water almost exclusively. Figure 4 is a generalized scheme summarizing the effects of burial and heating on the amounts of pore water and bound water in three important rock types found in many metamorphic belts. It can be seen that there will be significant overall differences in the fluid history of basins characterized by different proportions of clay-rich rocks, volcanic products and relatively unreactive continental sandstones. The transition from diagenesis to regional metamorphism is a continuous one in the sense that original pore fluid is modified by addition of water released from minerals, and by exchange with minerals, but is probably accompanied by a fundamental change in the relationship between depth and fluid pressure. Where basin fluids remain connected to the surface through pores or cracks, fluid pressure approximates to a hydraulic head (Fig. 1), but when it is effectively sealed from the surface, in deep diagenesis, overpressures are developed. Such overpressures are, however, typically associated with anomalously high porosities for the depth, a situation that may lead to either rapid fluid loss if the seal is breached or reduction of porosity by slow chemical compaction in a chemically closed system. In appropriate settings, the onset of more or less ductile deformation, and in particular the development of cleavage, occurs next, at temperatures of ,250 8C. At this point there has been irreversible loss of porosity, certainly to below 1% but probably significantly lower than this in many instances (Manning & Ingebritsen 1999), and the main source of water during subsequent heating is the release of bound water from mineral lattices (Fig. 4). Such rocks are undergoing metamorphism, and there are

Fluids in the metamorphic cycle Progressive burial and metamorphism In this section, the intention is to consider the evolution of fluids in a sedimentary basin as it evolves into a metamorphic belt. In general, therefore, the starting material can be considered to be sediment pervasively wetted by pore waters occupying a continuous grain boundary film. Initially, the porosity is likely to be a major component of the rock volume and fluids evolve by interacting with minerals that are either stable in the presence of fluid or are transformed into phases that are able to coexist with fluid during the course of the fluid–rock interactions. Of course, relatively dry, high-temperature crystalline rocks can become caught up in such a basin, either as extrusive rocks or sills formed as part of the basin fill, or as tectonic slices of basement interleaved with sediments. Where these are massive, they may persist through extensive subsequent burial and metamorphism, but where they are permeable, such as pyroclastic deposits, they react with pore fluids very rapidly, as seen, for example, in the burial metamorphic provinces of southern New Zealand (Coombs 1954). Their hydrated products may then undergo progressive dehydration with further heating. Porosity, permeability and fluid pressure. During the initial stages of burial of sediments there is significant expulsion of water as a

Fig. 4. Water budget for metamorphism of some common lithologies. The curves denote the variation in the approximate water content of pelite, psammite and metabasite, with continuous lines indicating the content of mineralogically bound water, dashed lines the likely amount of free water in pores. It should be noted that porosity estimates for rocks undergoing active metamorphism are only general estimates and will be strongly dependent on the deformation history. It has been assumed that metabasites are extensively hydrated at low grades, although this is not normally the case for massive bodies.

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a number of reasons for supposing that they are also overpressured but now have intrinsically low porosity and permeability. First, in contrast to sedimentary rocks, there is normally a complete lack of porosity-filling cement textures. Second, apart from burial metamorphic terranes (which are arguably sedimentary basins with a reactive volcaniclastic fill in place of quartzofeldspathic detritus), regionally metamorphosed rocks have undergone the development of cleavage or schistosity, with associated porosity loss. Third, veins occupying extensional fractures can occur in even the highest pressure metamorphic rocks, and not uncommonly contain metamorphic index minerals such as kyanite (Read 1932, and many subsequent descriptions) indicative of similar metamorphic pressures to the wall rocks. Even eclogite minerals may occur in extensional veins (e.g. Widmer & Thompson 2001). Finally, thermobarometric calculations for carbonate-free rocks below the onset of melting generally give similar temperatures for volatile-absent and dehydration equilibria only if fluid pressure is taken as close to lithostatic (see discussion by Spear 1993). In the regime of diagenesis, porosity and permeability are typically considered intrinsic properties of a rock, but with increased temperature in metamorphism it is likely that they respond to the physical environment to a much greater extent, as the fluid-poor fluid–rock system approaches a mechanical equilibrium in which fluid pressure is equal to lithostatic pressure. This requires that fluid escapes at the rate it is produced. Because devolatilization reactions are strongly endothermic, it follows that their overall rates are determined independently by the heat flux, and as a steady state is approached, permeability evolves to match the reaction rate (Yardley 1986). Rocks that contain little reactive material (such as psammites) and those that are heated slowly will have lower permeabilities than those rich in reactive minerals (such as pelites) or heated rapidly. Numerical estimates of metamorphic permeability based on such arguments arrive at very low values (c. 10
18 to 10
19 m2 ) for even the most rapidly heated rocks (Yardley 1986; Hanson 1992). Only where reactions proceed rapidly following overstepping (Waters & Lovegrove 2002), can higher permeabilities develop. Although such conditions may be short-lived, they almost certainly account for the development of some important rock types, including skarns (Yardley et al. 1991). A further distinction between metamorphic belts and sedimentary basins is that bedding in metamorphic rocks is commonly inclined as a result of folding. Contrasting gradients in fluid pressure and rock pressure along an inclined bed lead to mechanical instabilities, especially if the rocks have low tensile strength. Consequently, fluid will escape along upward-migrating fractures (Yardley 1986) or in a more general sense as porosity waves (Connolly & Podlachikov 2004); only if beds have a low intrinsic permeability can strong fluid overpressures be retained. In detail, of course, permeability contrasts between beds almost certainly occur in metamorphism, so that fluid flow is probably focused along specific beds. The nature and composition of metamorphic fluids. Sediment pore waters are the starting materials for prograde metamorphic fluids. In metamorphism, rocks have very low porosities, but what small amount of pore water they do contain has evolved continuously from formation water (Yardley & Graham 2002) and there is no reason to suppose that at any intermediate stage the pore fluids are totally removed and then replaced. Metamorphic waters commonly contain high proportions of dissolved salts and/or gas species. As metamorphism progresses, three types of changes may take

place in the fluid. First, the concentration of dissolved gases, notably carbon dioxide, can evolve as a result of mineral–fluid interactions involving organic matter or carbonate minerals and through changing gas solubility in water. Many low-grade pelitic rocks contain small quantities of carbonate minerals and/or graphite, which typically break down releasing carbon dioxide under greenschist-facies conditions. It is possible that this phase of carbonate breakdown is linked to increasing miscibility between water and carbon dioxide fluids with increased temperature (Fig. 5). This is because the activities of both H2 O and CO2 are

Fig. 5. Phase diagrams for H2 O–CO2 –NaCl fluids after Liebscher (2007). (a) 500 8C and 50 MPa (0.5 kbar); (b) 800 8C and 200 MPa; (c) 500 8C and 500 MPa. (a) is labelled to show fields in which vapour (V), liquid (L) and solid (S) phases occur; in (b) and (c) the continuous line separates the field of a single water-rich supercritical fluid from the two-phase field in which CO2 -rich ‘vapour’ coexists with brine.

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high when they occur in separate immiscible phases, but are reduced as they become mutually soluble, thereby promoting devolatilization reactions. For example, the elimination of calcite from low-grade pelites can be described through a reaction with kaolinite component of clay to produce epidote: 8CaCO3 þ 3Al4 Si4 O10 (OH)8 ¼ 4Ca2 Al3 Si3 O12 (OH)

(1)

þ 10H2 O þ 8CO2 : This reaction will be driven further to the right as increased mutual solubility of H2 O + CO2 reduces the activity of both. Second, the cation composition of the fluid evolves as a result of mineral interactions, whereas the dominant anion, chloride, is extremely conservative and normally remains in the fluid phase throughout. Mineral reactions typical of diagenesis, in which metastable grains are replaced (e.g. albitization of plagioclase) give way at higher temperatures to progressive cation exchange reflecting continuous temperature changes; for example, the proportion of K to Na increases with T in fluids in equilibrium with two feldspars. Third, the concentration of elements such as Si and Al increases progressively with P and T. Although these lithologically important components do not show such a strong direct correlation with chloride as most cations, their solubility is still influenced by fluid salinity although the effect is much less than those of P and T (Newton & Manning 2000; Manning 2006; Shmulovich et al. 2006). Although these components represent only a small part of the dissolved load of most crustal metamorphic fluids (except at very high pressures), they become significantly mobile, as demonstrated by their importance in vein minerals throughout metamorphism, and are redeposited in response to a return to lower grade conditions without requiring wall-rock interactions. In contrast, most of the metal chlorides that are important in metamorphic fluids are at concentrations that would remain in solution even under surface conditions were it not for the presence of minerals with which they react. It is commonly assumed that metamorphic fluids must become progressively more dilute during metamorphism, through the release of bound water from mineral lattices into a very small pore volume, following the trend reported from ocean drilling, noted above. The evidence of fluid inclusions, and to some extent from Cl contents of minerals, is that this dilution does not in fact take place. The literature compilation of Yardley & Graham (2002) demonstrated that metamorphic fluids may be very saline, and are seldom less saline than seawater. There is a distinct difference between fluids from sequences that were originally deposited on continental margins and those formed in ocean basins and accretionary prisms, probably reflecting whether or not evaporites had formed during the depositional history. It seems therefore that, once present, salt is as hard to rinse out of the continental crust as it is to remove it from any other porous object, although the mechanism by which salt is retained during dewatering is not understood. To summarize: within the crust, aqueous fluids may interact with rocks in a variety of ways although being physically distinct from, and notably significantly less dense than, their host rocks. Their cation composition is dictated primarily by ion exchange in a chloride medium, and so the dissolved load is very different from the rock or melt chemistry in both the absolute and relative amounts of a very wide range of elements (Yardley 2005). At crustal pressures at least, the concept of a distribution coefficient between fluid and minerals is essentially meaningless for most elements, because fluid concentrations are dictated by chlorinity,

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and this is an independent variable that can vary by more than an order of magnitude (Yardley & Graham 2002). Multi-phase fluids. Oil, water and gas are of course immiscible in sedimentary basins, but the degree of mutual solubility between them increases with temperature. As metamorphic conditions are approached, organic species are less important and it is the relationships between water and carbon dioxide fluids that are of most general significance for deep crustal fluids. If pure, these become completely miscible at a temperature of c. 265 8C for pressures between c. 1.5 and 1.9 kbar (Takenouchi & Kennedy 1964). The critical curve rises to slightly higher temperatures with increased pressure, but converges with the critical point of water at lower pressures. However, modest salt contents, which are commonplace in crustal fluids, can extend the P–T range of immiscibility between these and other H2 O– non-polar gas pairs to embrace most crustal conditions (Fig. 5) (Bowers & Helgeson 1983; Duan et al. 1995; Liebscher 2007), so that two fluid phases may coexist throughout the crust, wherever salt and carbon dioxide (or another non-polar gas species) are both present in significant amounts relative to water. These effects are particularly marked in carbonate-bearing sequences, where hydration–decarbonation reactions serve to increase fluid salinity and hence extend immiscibility (Trommsdorff & Skippen 1986; Heinrich et al. 2004; Heinrich 2007). Where two fluids are present, the more abundant phase (normally that produced by continuing devolatilization reactions such as reaction (2)) can flow out of the rock leaving the immiscible phase behind (Yardley & Bottrell 1988). This is because there is a threshold proportion for each phase below which it no longer forms a continuous film through the pore space, and so the rock is no longer permeable to it. This can be true of either the wetting or the non-wetting phase (Fig. 6), although in detail the threshold levels will vary with wetting characteristics. The combination of immiscibility and the existence of a minimum threshold below which either phase will cease to flow means that coexisting fluid phases can persist throughout a period of extensive reaction, and where this happens the chemical potentials of both fluid species remain high. Conditions of two-fluid phase behaviour are likely to be more important in metamorphism than has been demonstrated to date, simply because many metamorphic petrologists make the assumption that the fluid

Fig. 6. Schematic representation of the occurrence of two fluid phases in the pore structure of a sediment or metamorphic rock. In (a) the wetting phase (dark grey) is more abundant and forms a continuous film so that it is able to move through the rock in response to gradients in hydraulic head, although flow is constricted by the presence of bubbles of the nonwetting phase forming partial obstructions in the pores. In (b) it is the non-wetting phase that is more abundant and is able to move; here drops of the wetting phase on grain surfaces likewise constrict the flow paths and restrict migration of the non-wetting phase.

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phase consists only of H2 O and CO2, in the absence of fluid inclusion evidence of salinity. For example, Heinrich et al. (2004) and Heinrich (2007) have shown that fluid immiscibility has played an important role in determining the reaction history in most localities where contact metamorphism of carbonatebearing sequences has taken place. Similarly, many metasedimentary sequences have pore fluids whose salinity is sufficiently high for any CO2 liberated by oxidation of graphite or breakdown of minor carbonates to remain immiscible to at least upper greenschist-facies temperatures. Fluids at very high pressures. At the conditions of the release of fluids from a subducting slab and the generation of arc magmas, silicate rock is much more soluble in water than under crustal conditions, and its solubility is no longer dominated by chloride salts. At normal crustal pressures, aluminosilicate material contributes only a very small part of the dissolved load of all but the most dilute fluids, but at pressures between 10 and 20 kbar, significant concentrations of dissolved silica and aluminosilicates (Newton & Manning 2000; Zotov & Keppler 2000; Manning 2006; Shmulovich et al. 2006) become present. At high temperatures in the crust, aqueous fluid promotes melting and dissolves into the melt, so that with further heating, under granulite-facies conditions, hydrous minerals may break down directly to produce water-bearing melt through the process of dehydration melting. Eventually, at pressures corresponding to perhaps 60–200 km, according to composition, silicate systems reach a ‘second critical end point’, at which aqueous fluids and silicate melts become miscible (e.g. Kessel et al. 2005). These considerations are of fundamental significance for understanding the release of fluids and/or melts in deep subduction, and their contributions to arc magmas, but are beyond the scope of the present paper because silicate melt–fluid miscibility is not widely attained in the crust except probably in some evolved magmatic fluids enriched in fluxing elements such as F or B.

Chemical change to rocks during prograde burial and metamorphism In the early stages of burial, pore fluid chemistry evolves through diagenetic mineral fluid interaction, but whereas at early stages of diagenesis the pore water, if sufficiently saline, provides a significant reservoir for many cations, by late diagenesis the mass of all but the most conservative solutes present in the pore fluid is insubordinate to the quantities of the same elements in mineral matter, as a result of the reduction of porosity (Fig. 4). Hence, early in diagenesis, some aspects of the bulk composition of the solid matter in a sediment can be changed simply through reaction with pore waters that are subsequently expelled, whereas after more extensive porosity reduction, many pore volumes of fluid must flow through a layer to achieve metasomatic changes to rock composition. Fluid flow is also more likely to occur early in diagenesis, because it is inhibited once overpressure develops. It follows that diagenetic systems are more likely to be chemically open than metamorphic ones, and features such as dolomitization of calcite and albitization of feldspars bear testimony to extensive diagenetic metasomatism when sediments remain permeable and porous (Milliken et al. 1994; Land 1995; Land & Milliken 2000). Ever since the development of the facies concept by Eskola, it has been accepted by most metamorphic petrologists that metamorphism itself normally takes place without further change in the rocks’ chemical compositions, apart from loss of volatiles (C, H2 O, CO2 ). Exceptions (i.e. metasomatic rocks) do occur, for

example in skarns, and result in the development of rocks whose bulk compositions correspond to the mineral or assemblage of minerals that make up the metasomatic rock, rather than to an original sedimentary or igneous composition. In general, metasomatic rocks have been viewed as restricted to specific settings associated with igneous activity, deformation, or the juxtaposition of reactive lithologies, rather than typical of metamorphism. They are, of course, ubiquitous in hydrothermal ore bodies, where they are distinct from normal metamorphic products. In the modern literature, a number of reserchers have proposed models involving continuous large-scale fluid flow through rocks undergoing normal regional metamorphism, although other studies have specifically argued against such processes. In this review, I take a sceptical view of claims of large, continuous and widespread fluxes of fluid through metamorphism for two main reasons. First, the existence of large fluid overpressures in rocks undergoing metamorphism, documented above, is incompatible with large fluid fluxes, as they will lead to irreversible fluid loss at a rate very much greater than that of fluid production. Second, rocks that undergo extensive metasomatism develop simple mineralogies through the effect of the ‘Korzhinskii phase rule’. Korzhinskii (1958) pointed out that in metasomatic settings where some chemical components can be considered ‘perfectly mobile’ (i.e. their chemical potential in the pore fluid is a reflection of the conditions at the fluid source because so much water is infiltrating the rock mass of interest), then they should not be included in the count of components when applying the phase rule to calculate how many phases can occur in the system at equilibrium. Thus strongly metasomatized rocks, such as vein walls, have fewer phases than their precursors, and those phases have rather constant compositions even if they are solid solutions. Such effects are ubiquitous in ore deposits but, although known in metamorphism, they are rare, localized, and arise from specific events within the metamorphic history. Etheridge et al. (1983) suggested that there might be large fluid fluxes during metamorphism arising from the development of convection cells. This hypothesis is now considered improbable, because of the widespread evidence for near lithostaticpressured fluids in metamorphism (above). The fluid pressure gradient within a convection cell must be hydrostatic, and it follows that if the mean pressure is lithostatic, the top of the cell will have fluid pressure greater than lithostatic, because for a hydrostatic gradient fluid pressure reduces upwards less strongly than lithostatic confining pressure. Likewise, at the base of the cell, fluid pressure will be less than lithostatic pressure. Hence such a cell is very limited in its possible vertical extent, because if it attains a height of more than a few tens of metres it must migrate upwards as a porosity wave. Even short vertical connections within or between overpressured layers lead to irreversible fluid loss through hydrofracturing or the migration of porosity waves. John Ferry has championed the idea of large fluid fluxes in metamorphism for many years, and has developed an approach (e.g. Ferry 1983, 1987, 1994) based on calculating the amount of fluid infiltration required to allow metamorphic reactions to proceed despite producing fluids whose composition would otherwise be expected to stop further reaction. For example, the reaction calcite þ quartz ¼ wollastonite þ CO2

(2)

can begin at relatively low temperatures if the pore fluid is water, but if the CO2 is able to accumulate so that its partial pressure rises, further reaction will be inhibited. If porosity is small, very

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little reaction can occur before the reaction stops, hence the wollastonite produced is likely to be undetectable. On the other hand, a continuous flux of water that washes the CO2 away can permit reaction to proceed to a much greater extent. Ferry’s approach is to calculate the water fluxes required to permit the observed extent of mixed volatile reactions, but it does depend on a number of assumptions and these are now in doubt for many situations. It is a fundamental assumption that the fluid is a single-phase H2 O–CO2 mixture. In addition, possible diffusional exchange between contrasting adjacent lithologies, such as that which gives rise to bimetasomatism, is excluded from consideration. All reactions are assumed to take place very close to equilibrium with fluid pressure remaining close to lithostatic pressure irrespective of the magnitude of the fluid flux. All of these assumptions can be questioned. The evidence for fluid immiscibility and two-phase flow in metamorphism has been discussed above. If the two fluid components can move independently of one another because they are concentrated in different phases, the basis of the calculation is invalid. Ignoring cross-layer diffusion or flow is likewise an extreme limiting position, and calculated fluxes are reduced if it takes place (Evans & Bickle 1999). Indeed, the recent work of PennistonDorland & Ferry (2006) recognized this and documented evidence for diffusion across layers accompanying water infiltration. Additionally, reaction does not take place at equilibrium, by definition, and the effect of considering reaction rates under nonequilibrium conditions is likewise to reduce the calculated fluxes significantly (Lasaga et al. 2000). Thus we can be sure that the fluxes calculated by Ferry and co-workers are maximum figures and are probably significant overestimates. Indeed, it is clear that the approach has been applied to, and produced flux estimates for, situations where metamorphism has not been driven by infiltration at all (Heinrich et al. 2004). On the other hand, many of the reaction phenomena described in the publications of Ferry’s group do require an element of infiltration at specific stages in the metamorphic history to account for their development, although not necessarily as a regional-scale feature affecting all lithologies. More recently, an alternative line of argument for significant fluid fluxes through rocks undergoing metamorphism has been developed by J. Ague and colleagues, based on direct comparisons of rock compositions where the same formation occurs at different grades. Ague and colleagues have suggested that, far from being essentially isochemical, as demonstrated in classical studies, much metamorphism involves chemical change. Ague (1994) showed that the Wepawaug Schist of Connecticut exhibits systematic chemical variation with grade. Subsequently, Breeding & Ague (2002) deduced significant chemical changes in the Haast Schists of New Zealand, based again on comparing the composition of metasediments from different grades that had similar immobile element signatures. Interestingly, this result directly contradicted the larger-scale approach taken by Cox (1993), who compared the bulk composition of high-grade Haast Schist with lower-grade schist and precursor Torlesse Group supracrustal units, and concluded that, when quartz veins were integrated back into the outcrop, there had been very little chemical change on a scale of more than tens of metres, apart from loss of volatiles. The difficulty in interpreting such outcrops can be appreciated from Figure 7, a photograph of highly deformed and veined chlorite-zone pelitic rocks of the Haast Schists, on the shore of Lake Hawea. The effects documented by Ague and colleagues are subtle and difficult to correlate with specific metamorphic events. It is possible that they represent the net effect of background fluid loss from deeper to shallower

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Fig. 7. Veined and folded chlorite zone schist, Lake Hawea, New Zealand. An early quartz vein, intensely folded, is flanked by dark, chlorite-rich material too depleted in silica to represent an original rock composition, but it is unlikely that the vein and its dark selvedges would together exactly recreate the composition of the original sediment. Hence longer-range transport has taken place as well as very local segregation.

levels. Unfortunately, natural variability within sedimentary sequences is often large, and high-grade rocks do not always have stratigraphic equivalents exposed at lower grades, leading to conflicting interpretations as for the Haast Schists. Some ‘immobile elements’ such as Zr are introduced into the original sediment in distinct particles with different physical characteristics from the dominant minerals accounting for most other chemical characteristics, hence there may be significant original variations in element ratios. Clearly, this is an important and challenging line of research that merits more attention. The role of veins is critical to the various interpretations, and these are discussed further below.

Patterns of fluid flow during metamorphism In parallel with the debate about the extent to which metamorphic rocks may be metasomatized by migrating fluids, a number of approaches have been taken to investigating fluid fluxes on a relatively local scale (Oliver 1996). Particular approaches have included investigating the distribution of meta-

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somatic rocks, studies of cross-layer flow through the induced reaction progress or isotopic shifts, and studies of quartz veining. Metasomatic rocks. In a few instances where there is extensive metasomatic activity, fluid flow paths can be deduced from lithological patterns. This is most notably possible where skarns develop in marbles. Decarbonation reactions trigger the development of secondary porosity, and in an overpressured environment this leads to further fluid infiltration, which will typically lead to further reaction and renewed porosity generation. An example is the development of coarse diopside skarn beds from dolomitic marble beds interbedded with quartzite in the Connemara Marble Formation of Ireland described by Yardley et al. (1991) (Fig. 8a). In a BSEM image (Fig. 8b) single diopside crystals are seen to have cores of very pure diopside (XMg ¼ 0.97) with corroded outlines. These are overgrown by oscillatory zoned diopside that is significantly richer in Fe (XMg ¼ 0.88) and develop a euhedral form once more. The reaction history of this rock can be interpreted as follows (Fig. 8c): initial reaction of dolomite with isolated quartz sand grains led to the formation of diopside and CO2, and created a secondary metamorphic porosity. The new porosity allowed infiltration of overpressured water from nearby schist beds, resulting in the breakdown of diopside to tremolite. However, because of the convex-upward shape of the equilibrium curve for this reaction, continued infiltration of water can lead to very water-rich fluids and reverse the reaction, stabilizing diopside again. This accounts for the new growth of oscillatory zoned diopside over the corroded core. The higher Fe content of the later diopside probably reflects the relatively high levels of Fe in chloride fluids in the crust. Yardley et al. (1991) mapped the distribution of the diopside-rich unit and demonstrated that the high fluid fluxes that it marked had been predominantly layer parallel. Similar skarn bed developments are common in metamorphosed dolomitic marble sequences, and probably arise through rather short-lived processes at specific points in the metamorphic history, where there is positive feedback between reaction and fluid flow. It is unlikely that they represent longterm fluid flow patterns. Tracers of cross-layer flow. Bickle & McKenzie (1987) first developed the application of advection–dispersion theory to metamorphic problems and showed how cross-layer flow results in the migration of chemical fronts at different rates according to fluid and rock compositions. The telescoping of such fronts is an indication of layer-parallel flow: the development of metasomatic sides. Skelton et al. (1995) used carbonation of metabasite assemblages in greenschist-facies rocks of SW Scotland as a sensor of fluid flow and were able to deduce the regional flow patterns from local fluxes. Although there is no doubt that crosslayer flow must take place, most studies emphasize layer-parallel flow as predominant, as would be expected if permeability varies with lithology during prograde metamorphism. Quartz veins in metamorphic rocks and their significance for regional fluid flow. Quartz veins are ubiquitous in many metamorphic rocks, and some have argued that these form as a result of the decrease in quartz solubility with fall in temperature, and so arise as fluids migrate down-grade over long distances (Walther & Orville 1982). Alternatively, however, they may simply result from local segregation. Many metamorphic quartz veins (e.g. Fig. 9) show little depletion of the vein components in their immediate wall rocks, indicating that they form in a system that is open on a scale that is much larger than the diffusional length scale. Some have

Fig. 8. Diopside skarn beds formed through infiltration of water into transient porosity generated by reaction of dolomite with quartz in specific beds (Creggs quarry, Connemara, Ireland). Note coin for scale at the base of the lower diopside bed. (a) Outcrop with diopside beds (formerly dolomite with quartz detritus) cut by quartz–feldspar veins and interbedded with thin quartzite layers. (b) BSEM image of diopside showing a dark core of nearly pure diopside with corroded outlines, overgrown by euhedral, oscillatory zoned overgrowths of more Fe-rich diopside. Field of view is 1.2 mm. (c) Schematic T–XCO2 diagram to illustrate the reaction sequence: initial heating in the absence of water allows dolomite + quartz to persist until they break down to form diopside, generating significant porosity. This is believed to have given rise to infiltration of water, first corroding the diopside as it reacts to tremolite + calcite + quartz, and then causing these phases to themselves react to form a new growth of diopside. Fe and some silica are introduced in the aqueous fluid.

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Fig. 9. Examples of regional metamorphic quartz veins with neither quartz depletion in the immediate vein margins nor evidence of disequilibrium with the wall rocks, both from the Maam Valley, Connemara, Ireland. (a) Vein quartz cuts a quartz-rich psammite in the lower part of the image, but as it passes into the overlying pelite the vein is composed solely of andalusite (delimited by the upper two arrows). The lowest arrow indicates an andalusite crystal in the quartzrich part of the vein. It should be noted that there is also no evidence of alumina depletion next to the andalusite vein, although the spatial distribution of andalusite clearly requires a local origin. (b) Late metamorphic quartz vein in quartzrich psammite.

argued that unless the composition of veins and their immediate selvedges adds up to that of the protolith, an origin by infiltration of external fluids is required (Penniston-Dorland & Ferry 2008). Alternatively, others have suggested that in the process of deformation and fracturing that accompanies vein formation, fluid may be move around on a scale of tens to hundreds of metres within a capped body of overpressured rock, in response to transient pressure fluctuations induced by cracking. This strain-induced movement would provide an effective mechanism for segregation, so that veins may be derived in part from their immediate margins and in part from a much larger volume of the apparently unaltered host (Oliver et al. 1990; Yardley & Bottrell 1992). Key lines of evidence come from oxygen isotopes and wall rock metasomatism. Several studies have shown that metamorphic quartz veins have oxygen isotope compositions that track their wall rock’s composition, even if the immediate margins are not an adequate source of quartz, which supports a segregation origin with transport over a distance of metres to tens of metres (Yardley & Bottrell 1992; Cartwright et al. 1994; Li et al. 2001; Richards et al. 2002). Likewise, lack of metasomatic reaction in the wall rocks of veins such as those in Figure 9 suggests that the vein fluid was close to equilibrium with the local lithologies and P–T conditions; otherwise reactions such as the alteration of feldspar to muscovite (or vice versa), seen in the walls of ore veins, should be evident. It is relevant to consider why veins arise in the first place. They are commonly assumed to form in cracks developed in response to dehydration reactions, but in the field metamorphic veins are almost invariably associated with deformed rocks. Sites of contact metamorphism experience the highest geological forcing of permeability within the crust, because of the combined effects of fluid generation by reaction and thermal expansion of existing fluids (Neuzil 1995), but veins are rare or absent in most metamorphic aureoles. On the other hand, if veins are segregation features that form in fractures that arise when an overpressured rock is subject to deviatoric stress, then veining will correlate with deformation episodes rather than porphyroblast growth, and this is indeed often the case. Oliver et al. (1990) have argued that variations in mean stress within a metamorphic belt, reflected in different structural styles, have led directly to different patterns of metamorphic fluid flow. They showed that metasediments in the low-P, high-T Mary Kathleen belt of Queensland have generally experienced little fluid influx and

retain primary contrasts in oxygen isotopic composition on the scale of centimetres. However, the terrain is also cut by veins that are concentrated in metabasites and calc-silicates, and Oliver et al. (1993) showed that these have a uniform, externally buffered stable isotope composition, and probably derive from fluid released by magmas crystallizing at depth. Although I believe that it is clear that most metamorphic veins are formed predominantly by segregation, there are also some examples of metasomatic alteration in vein selvedges that are indeed consistent with an infiltration origin, the veins acting as channelways for fluid derived by dewatering at higher grades (Yardley 1986; Van Haren et al. 1996; Fig. 10). It is likely that there is a gradation between segregation veins and those formed by infiltration, as any overpressured system will have a tendency to leak, even if the dominant fluid movement is local straininduced recirculation. Metamorphic veins may indeed be linked to pulses of through-going fluid of the type imaged by the seismic studies referred to above.

Retrograde metamorphism From the point of view of the distribution of fluid and the fluid pressure regime in the crust, the key change is not the transition from diagenesis to metamorphism, but the transition from progressive heating to progressive cooling during the metamorphic cycle. This is because once metamorphic rocks begin to cool, retrograde reactions that consume water replace prograde reactions that release it. It is clear from thermodynamic calculations that these reactions are able to reduce fluid pressures to very low values simply by absorbing water into hydrous minerals (Yardley 1981; Frost & Bucher 1996; Yardley & Valley 1997; Stober & Bucher 2004). It is therefore at this point (i.e. the onset of cooling during an orogenic cycle) that there is a fundamental change in the crustal fluid regime, from pervasive high-pressure fluid in grain boundary networks to more or less dry rocks that absorb any fluid that infiltrates them through retrograde reactions and can contain fluid only in isolated fractures insulated from high-T minerals by a lining of retrograde growth. The kinetic data of Schramke et al. (1986) for the retrogression of andalusite + K-feldspar assemblages to muscovite + quartz in the presence of excess water suggests that the rate of retrogression of highgrade assemblages in the mid- to lower crust is so rapid that the lifetime of free water once it has penetrated a fracture in cooled

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Fig. 10. Rare example of metamorphic quartz veins accompanied by marked metasomatic alteration of the wall rocks, Knockaunbaun, Connemara, Ireland (Yardley 1986). In this outcrop only small areas of the pale grey host muscovite-rich schist (Ms schist) are present. Within a few tens of centimetres of the vein quartz this schist is transformed into a yellow–brown metasomatic rock dominated by fine staurolite (St). There is also extensive development of black tourmaline (Tm) around the vein quartz.

crystalline rock is likely to be of the order of years (Fig. 11). Water can persist in cooled basement rocks only if they have been completely hydrated by earlier infiltration events, or if the water-bearing fractures are lined with inert mineral precipitates, such as quartz, that isolate the water from the surrounding rocks. Despite ubiquitous evidence for retrograde fluid activity in crystalline rocks at the surface, the source of retrograde fluids is often impossible to deduce because their extensive reaction with the host rock resets the potential geochemical indicators (Frape & Fritz 1987). Near-surface systems can often be clearly related to deep penetration of groundwaters, but for retrogression at

Fig. 11. Estimated times for the persistence of water in fractured andalusite–K-feldspar schist at 200 MPa (2 kbar), calculated from the data of Schramke et al. (1986). The calculation is for the time it would take for a 0.1 mm thick film of water to be consumed, and the calculation uses the experimental reaction rates, assuming the surface area of andalusite is the limiting factor.

greater depths evidence is generally lacking. In exceptional cases, as where hydrocarbons are introduced (Munz et al. 1995, 2002), an origin for infiltrating retrograde fluids from overlying or adjacent sedimentary basins can be demonstrated. The near surface is of course the most voluminous and effective reservoir of potential infiltrating fluids, but an origin from within an active metamorphic pile may also be possible in some situations.

Fluids and deformation The metamorphic cycle summarized above in simplified form (for example, magmatic interactions have not been considered) has important implications for crustal deformation. Much deformation takes place via mechanisms that depend on the involvement of a fluid phase to permit solution transfer, enhance diffusion or lower the effective normal stress across surfaces. Water-weakening effects have been well known for many years, but when this knowledge is combined with an understanding of the behaviour of fluids through the rock cycle, some systematic trends appear. Water can have a very large effect on rock strength, especially at elevated temperatures, but it is the metamorphic cycle that controls the availability of water. As we have seen, the stage of progressive burial and diagenesis of sediments, continuing into metamorphism, is accompanied by the presence of a pervasive fluid phase, typically overpressured relative to the hydrostatic head. Under these conditions, once moderate temperatures are attained, rocks are easily deformed because deformation mechanisms requiring mass transfer are possible through the medium of water, and we find examples of ductile deformation affecting slates and even sedimentary rocks that have never been at temperatures greater than 250 8C, as well as higher grade metamorphic rocks. It is notable that in metamorphic belts of all ages, pervasive ductile deformation occurs either during the heating cycle or at approximately the peak of metamorphism; in other words, pervasive ductile deformation takes place in rocks that contain a pervasive fluid. It may be accompanied by brittle deformation because of the embrittling effects of high fluid pressures noted above. In contrast, in the absence of free water, silicate rocks are

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very strong. It follows that the loss of a water phase that occurs in the middle and lower crust as metamorphosing rocks begin to cool must result in a marked increase in rock strength, and hence an end to pervasive deformation. This pattern seems to be broadly consistent irrespective of the maximum temperature that crystalline rocks have attained: orogenic belts typically exhibit multiple episodes of ductile deformation that accompanied progressive metamorphism, but after peak temperatures were attained, deformation is limited to discrete faults and shear zones. Granulite-facies rocks can spend long periods of time at even amphibolite-facies temperatures without pervasive deformation, and rocks that have not exceeded greenschist-facies temperatures will similarly stop deforming as soon as they cool, despite the great difference in the maximum temperature attained. Temperature is often thought of as the main factor controlling crustal strength, but it is unlikely that it outweighs the presence or absence of water. The prograde evolution of high-grade metamorphic rocks that experienced pervasive ductile deformation throughout their heating history can be investigated today only because in general they stopped deforming pervasively as soon as they began to cool (Yardley 1981). If this change in deformation style is the result of a change from a wet to a dry environment at the onset of cooling, driven by mineralogical changes rather than by a change in the stress regime, it would explain why so many orogenic belts show the same deformation history. Of course, metamorphic and plutonic rocks can sometimes undergo extensive and ductile retrograde deformation, but this is partitioned into shear zones, which at all but the highest temperatures show evidence of retrograde hydration reactions (McCaig 1997). The coupling of strain and water infiltration is illustrated in Figure 12, which shows schematic strength–depth curves for wet and for dry crust. Clearly, there is a significant depth range in the mid- to lower crust over which there is a large strength difference between wet and dry silicate rocks. As long as the deep crust is dry it can sustain very large stresses, possible approaching those for brittle failure as recorded in the upper crust. However, if water is able to penetrate along brittle cracks, it will result in a significant weakening of the wet material, probably leading to strain partitioned into the wet region. Because the strain will be accompanied by retrograde hydration reactions, this weakening is transient and will decay as the introduced water is taken up into mineral lattices. Recent models for crustal rheology often assume that rocks are wet and weak throughout the mid- to lower crust (Chen & Molnar 1983) and that the main parameter affecting rock strength is temperature. Although petrological studies, such as those cited above, have cast considerable doubt on the assumption of a wet lower crust that is implicit in this model, only recently has the geodynamics community recognized the failings of the model, through the recognition that earthquakes do indeed occur in the lower crust in a number of parts of the world (Jackson et al. 2008). Thus the petrological record of exposed rocks can now be reconciled, once allowance is made for scale, with the conclusions of geophysical studies. We should consider the direction of temperature change as a key variable controlling rheology, together with temperature itself. In these terms the history of an orogenic belt can be broken down into periods in which rocks are heating, are water-saturated and weak, and are able to undergo pervasive ductile deformation, and periods in which they are cooling, are dry and strong, and behave in a brittle manner with localized ductile deformation only at times and sites of water infiltration. The mineralogy of the crust, and its response to temperature change, is therefore able to dictate to

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Fig. 12. Schematic representation of the variation in the strength of the crust with depth and temperature, according to whether water is present or absent. Much of the lower half of the continental crust probably corresponds to the intermediate region in which rocks are strong when dry but become very weak if infiltrated by water. From Yardley & Baumgartner (2007).

a significant degree the wetness, and hence strength of the crust itself for much of the time. The cooling history of orogeny is particularly interesting because it offers the possibility of rapid transitions between contrasting rheologies, at least at a local scale. When dry crust, which is strong and therefore may be subject to large stresses, is infiltrated by water, which is able to reduce the strength locally along the infiltration pathways, there will be an immediate response, with strain partitioned into the hydrated regions. The duration over which such rheological changes persist will depend on the rate at which infiltrated water reacts with the host rocks and is converted to hydrous minerals, but judging by the hydration rates discussed above (Fig. 11) may be for periods of months to years after an infiltration episode. This implies that crustal earthquakes, if they permit infiltration of fluid into dry crystalline crust, will lead to transient weakening and creep, which may then itself feed back into the earthquake cycle. On the crustal scale, it may be that there are sufficient such shear zones for their transient weakening to account for a scarcity of deep crustal earthquakes even away from sites of continuing heating, thereby reconciling the geophysical evidence for weak basement with the geological evidence for very limited deformation of much of the lower crust through geological time.

Synthesis: an overview of fluids in the crust That water-rich fluids are important in the behaviour of the crust has been well understood for many years. What I have attempted to show here is not only that such fluids play a role in both chemical and physical processes, but also that these roles are

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intimately related. It is a fortunate consequence of this interrelationship that both the fugacity of water and many features of the composition of crustal fluids are constrained by a combination of mineral–fluid interactions and the physical limits imposed by rock strength. Hence the number of independent variables in crustal fluid–rock systems is rather limited, except in very dynamic systems where flow rates are fast relative to reaction rates. In most crustal settings, the main independent variables are temperature, lithostatic pressure (or depth) and salinity. It has been demonstrated here that there is a fundamental change in the occurrence of free aqueous fluid when rocks pass from being buried and heated to being cooled. During heating, fluid is pervasively released through compaction and reaction, and so a pervasive fluid film is present in all but the most impermeable and unreactive lithologies. In contrast, rocks that are below their original temperature of formation will react with any available water to produce hydrous minerals, resulting in crust that is essentially dry except for brief periods following an infiltration event. Where water does occur, it will not do so pervasively but will be confined to the pathways through which it has been introduced. These different patterns of fluid distribution have profound implications for rheology: heating rocks are weak, cooling rocks are strong, irrespective of the absolute temperature. Despite many speculations in the literature about large metamorphic fluid fluxes in the deep crust, it seems extremely unlikely that these are a common feature of deep crustal processes, even in areas undergoing heating. During burial, strong fluid overpressures preclude recirculation of fluid, and so large fluxes are possible only where there is strong focusing into specific layers. Such processes are of considerable interest and importance, but they represent brief, intense periods of activity with positive feedbacks between reaction and permeability, not background processes over geological time. Less extreme focusing may account for some chemical modification of original sedimentary lithologies during high-grade metamorphism under conditions where the dissolved load of the metamorphic fluid is high. The widespread preservation of high-grade mineral assemblages at the surface today likewise attests to the limited degree to which water permeates cooled and stabilized crust. The ideas presented here have been developed through discussions and disagreements with many colleagues. I would particularly like to thank L. Baumgartner, M. Bickle, C. Graham, W. Heinrich, B. Hobbs, G. Lloyd, B. Jamtveit, N. Oliver, R. Sibson, A. Thompson, J. Valley and V. Wall for their inputs, and S. Cox for showing me parts of the Haast Schists I would not otherwise have reached. The paper was written in part while I was a guest at the Geoforschungzentrum Potsdam, and N. Claiden drew many of the figures. G. Dipple and M. Bickle provided some penetrating comments at the review stage.

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Received 8 August 2008; revised typescript accepted 10 March 2009. Scientific editing by Rob Strachan.