Please cite this article as: Zheng, Y-F., Chen, R-X., Regional metamorphism at extreme conditions: implications for orogeny at convergent plate margins, Journal of Asian Earth Sciences (2017), doi: http://dx.doi.org/10.1016/ j.jseaes.2017.03.009
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Regional metamorphism at extreme conditions: implications for orogeny at convergent plate margins
Yong-Fei Zheng*, Ren-Xu Chen,
CAS Kay Laboratory of Crust-Mantle Materials and Environments, School of Earth and Space Science, University of Science and Technology of China, Hefei 230026, China
metamorphism at low geothermal gradients of <10C/km, or to Buchan-type metamorphism at high geothermal gradients of >30C/km. Extreme pressures refer to those above the polymorphic transition of quartz to coesite, so that ultrahigh-pressure (UHP) eclogite-facies metamorphism occurs at mantle depths of >80 km. Extreme temperatures refer to those higher than 900C at crustal depths of 80 km, so that ultrahigh-temperature (UHT) granulite-facies metamorphism occurs at medium to high pressures. While crustal subduction at the low geothermal gradients results in blueschist-eclogite facies series without arc volcanism, heating of the thinned orogenic lithosphere brings about the high geothermal gradients for amphibolite-granulite facies series with abundant magmatism. Thus, UHP metamorphic rocks result from cold lithospheric subduction to the mantle depths, whereas UHT metamorphic rocks are produced by hot underplating of the asthenospheric mantle at the crustal depths. Active continental rifting is developed on the thinned lithosphere in response to asthenospheric upwelling, and this tectonism is suggested as a feasible mechanism for regional granulite-facies metamorphism, with the maximum temperature depending on the extent to which the mantle lithosphere is thinned prior to the active continental rifting. While lithospheric compression is associated with subduction metamorphism in accretionary and collisional orogens, the thinned orogenic lithosphere undergoes extension due to the asthenospheric upwelling to result in orogen-parallel rifting metamorphism and magmatism. Thus, the rifting metamorphism provides a complement to the subduction metamorphism and its operation marks the asthenospheric heating of the orogenic lithosphere. Because of the partial melting and melt extraction of the lower continental crust, contemporaneous granite-migmatite-granulite associations may serve as a petrological indicator of rifting orogeny that is built on precedingly accretionary and collisional orogens. The UHT metamorphic rocks have occurred since the Archean, suggesting that the hot underplating has operated very early in the Earth’s history. In contrast, the UHP metamorphic rocks primarily occur in the Phanerozoic, indicating that the thermal regime of many subduction zones has changed since the Neoproterozoic for the cold subduction.
1. Introduction Mountain systems are typically orogenic belts that are mainly composed of regionally metamorphic, sedimentary and magmatic rocks at convergent plate margins. Metamorphic rocks in orogens are often foliated with variable degrees of deformation and recrystallization. They are mainly composed of amphibolite-facies rocks (Miyashiro, 1961, 1973; Ernst, 1975; Brown, 2009), with very minor volumes of eclogite on the one hand (Carswell and Compagnoni, 2003; Chopin, 2003; Liou et al., 2009) and granulite on the other hand (Harley, 2008; Kelsey, 2008). Although both eclogite and granulite occur sporadically in orogenic belts, they record abnormal metamorphic pressure (P) and temperature (T) conditions on regional scales. In some extreme cases, either the pressure is above the coesite stability field (referring to ultrahigh-pressure, UHP), or the temperature is as high as 900-1100C (referring to ultrahigh-temperature, UHT). Such metamorphic rocks with extreme P-T conditions are commonly observed at ancient convergent plate margins (Brown, 2009, 2014; Liou et al., 2009, 2014; Hermann and Rubatto, 2014), marking operation of extreme tectonism in the regime of plate tectonics. During extreme tectonic episodes, crustal rocks are subjected to not only metamorphic dehydration but also partial melting (anatexis). This results in the generation of high-pressure (HP) to UHP eclogite-facies metamorphic rocks on the one hand, and the production of high-temperature (HT) to UHT granulite-facies metamorphic rocks on the other hand. In the latter case, felsic melts are extracted from granulites and then ascend into shallow crustal levels for emplacement (Brown, 2014; Kelsey and Hand, 2015; Harley, 2016). In order to understand how regional metamorphism under extreme conditions was operated at convergent plate margins, this paper presents a review on the petrogenesis of both UHP and UHT metamorphic rocks in orogens. Most of the ideas discussed in this paper have been presented before in a different context, and have had a long and complicated history. Nevertheless, the present emphasis is mainly placed on the petrogenesis of HT to UHT metamorphic rocks. In doing so, available observations and interpretations are integrated within the framework of plate tectonics. The results provide insights into tectonic settings of extreme metamorphism with respect to geodynamic evolution from subduction to rifting and vice versa. Mineral abbreviations are after Whitney and Evans (2010).
2. Metamorphic facies series 2.1 General category
The study of regional metamorphism can be tracked back to classic works on Scotland, Scandinavia and Pyrenees of Europe, sites of orogenic belts (Barrow, 1893, 1912; Goldschmidt, 1912, 1915; Eskola, 1915, 1920; Zwart, 1962, 1963). Later on it was extended to metamorphic rocks in North America (Turner, 1948, 1981) and circum-Pacific regions (Miyashiro, 1961, 1973; Ernst, 1975). Traditionally, regionally metamorphic rocks were categorized into two facies series (Eskola, 1920, 1929; Harker, 1932; Fyfe et al., 1958; Miyashiro, 1961; Zwart, 1969; Turner, 1981; Spear, 1993). One is the Barrovian facies series that developed at pressures higher than those for the Al2SiO5 triple point (Fig. 1), mainly involving the phase transformation between kyanite and sillimanite (Barrow, 1912; Harker, 1932; Barth et al., 1939; Spear, 1993; Bucher and Grapes, 2011). The other is the Buchan facies series that was produced at pressures lower than those for the aluminum silicate triple point (Fig. 1), with the phase transformation between andalusite and sillimanite (Barrow, 1912; Harker, 1932; Read, 1952; Spear, 1993; Bucher and Grapes, 2011). In the modern term of metamorphic petrology, the Buchan facies series shows metamorphic progression of zeolite prehnite-actinolite greenschist amphibolite granulite facies assemblages, and its common product is high-T/low-P metamorphic rocks (Holdway, 1971; De Yoreo et al., 1991). It develops at high geothermal gradients of >30C/km (Fig. 1), where the rates of heat conduction outstrip those of tectonic processes in the lithosphere to result in a much larger increase in temperature than in pressure (high dT/dP gradients). On the other hand, regional metamorphism also happens at low geothermal gradients of <10C/km, where the rates of tectonic processes outstrip those of heat conduction in the lithosphere to result in a larger increase in pressure than in temperature (low dT/dP gradients), corresponding to the Alpine facies series (Fig. 1). This series is composed of zeolite prehnite-pumpellyite blueschist eclogite facies assemblages, and its common product is low-T/high-P metamorphic rocks (Coleman et al., 1965; Ernst, 1971). Taken together, extreme metamorphism on a regional scale refers either to the Buchan-type metamorphism at the high geothermal gradients or to the Alpine-type metamorphism at the low geothermal gradients. In this context, either UHT or UHP metamorphic rocks are the common products of extreme metamorphism (Fig. 2). In general, extreme pressures refer to metamorphic pressures greater than 2.7 GPa, the minimum pressure required for the polymorphic transition of quartz to coesite at ~700°C.
This corresponds to lithospheric depths greater than thickened crust (>80 km), so that UHP metamorphism occurs at mantle depths. Pressures ranging from 1.0 to 2.7 GPa are defined as the HP metamorphism, and its products are mineral assemblages of blueschist to eclogite facies mineral assemblages (Fig. 3). Therefore, these HP to UHP metamorphic conditions are regionally produced at the low geothermal gradients of 10°C/km, with maximum values roughly corresponding to the metamorphic reaction from albite to jadeite and quartz (Fig. 3). Thus the resulted products belong to the Alpine facies series (Fig. 1). Extreme temperatures refer to metamorphic temperatures higher than those for the dry granite solidus, about 900C at medium to high pressures. Temperatures ranging from ~650C, the wet granite solidus, to ~900C are defined as the HT metamorphism, which takes place at low to medium pressures. The metamorphic products often show prograde phase transformation either from andalusite to sillimanite at pressures below the Al2SiO5 triple point, corresponding to the Buchan facies series that is produced at high geothermal gradients of >30C/km, or from kyanite to sillimanite at pressures above the Al2SiO5 triple point, corresponding to the Barrovian facies series that is produced at medium geothermal gradients of 10-30C/km (Fig. 1). Metamorphic assemblages for Barrovian facies series restrict to zeolite (prehnite-pumpellyite) greenschist epidote amphibolite amphibolite granulite facies. In high-grade metamorphic regions, the Barrovian facies series is characterized by the common occurrence of amphibolite-facies rocks, corresponding to regional metamorphism at temperatures in excess of 500C and pressures less than 1.2 GPa. Traditionally, amphibole in amphibolites was considered as a prograde metamorphic product from greenschist-facie rocks at elevated P-T conditions. However, more and more studies indicate that amphibolite-facies rocks are generally associated with granulite-facies rocks on the one hand and with eclogite-facies rocks on the other hand. In either case, the amphibolite-facies rocks can be produced by hydration of preexisting rocks at high temperatures, their underlying granulite-facies rocks would have undergone dehydration melting. For this reason, the Barrovian facies series is often associated with the Buchan facies series in high-grade metamorphic terranes. As a consequence, the two facies series often share their common products of HT to UHT granulite-facies metamorphism at medium to high pressures (Fig. 2). However, it has been puzzled for a long time how such P-T conditions are regionally produced at crustal depths of <60-80 km.
Extensive studies of metamorphic rocks in the past century have led to the consensus on the operation of regional metamorphism during different types of orogeny (e.g., Spear, 1993; Bucher and Grapes, 2011). Peak metamorphic P-T conditions have been obtained for regional metamorphic rocks from various orogens (Fig. 4). In doing so, the equilibrium is the basis of the metamorphic facies concept (Goldschmidt, 1911, 1915; Eskola, 1915, 1920). It is the sequence of metamorphic facies exposed along the erosion surface through a metamorphic belt, corresponding to the metamorphic facies series of Miyashiro (1961). Thompson (1955, 1957) provided the thermodynamic basis for the metamorphic facies concept, setting the ground for geothermobarometers in quantifying metamorphic P and T conditions. In doing so, thermodynamic equilibrium is assumed in metamorphic rocks. If equilibrium is achieved and preserved over significant portions of the mineral assemblage, as indicated by completely homogeneous mineral chemistry and straight boundaries between adjacent grains, all geothermobaromatric results are valid (e.g., Pattison et al., 2003; Zheng et al., 2011a). In this regard, the effects of any previous or subsequent processes are lost completely due to the very nature of equilibrium. As such, the dynamic evolution of a metamorphic rock is relegated to a secondary role.
2.2 High-pressure to ultrahigh-pressure facies series The discoveries in upper crustal metamorphic rocks of coesite (Chopin, 1984; Smith, 1984), the UHP polymorph of quartz at depths greater than 80 km, and of diamond (Sobolev and Shatsky, 1990; Xu et al., 1992), the UHP polymorph of graphite at depths greater than 120 km, have drastically changed the knowledge in the traditional plate tectonics model concerning the depth of continental subduction (Chopin, 2003; Liou et al., 2009; Hermann and Rubatto, 2014). All the currently described UHP metamorphic slices consist predominantly of supracrustal rocks of continental, and more rarely of oceanic, affinities. They have been the characteristic feature of collisional orogens in the Phanerozoic (Fig. 5) and their petrogenesis has a direct link to subduction orogeny (Chopin, 2003; Ernst, 2005; Liou et al., 2014; Zheng and Chen, 2016). Because these UHP metamorphic rocks and their lower-P counterparts are typically produced at the low geothermal gradients of <10C/km in the Alpine orogen, they are referred as to the products of Alpine-type metamorphism. Eclogites and blueschists are the typical products of regional metamorphism at the low geothermal gradients (Coleman et al., 1965; Ernst, 1971).
In the P–T petrogenetic grid (Fig. 3), the low dT/dP metamorphic space is occupied by two major facies of blueschist and eclogite. For the HP blueschist facies, temperature is generally lower than 500 °C, while pressures are bracketed between 0.6 and 2.3 GPa. The HP eclogite facies is defined in the quartz stability field and has a range of temperatures between 500 and 1000 °C, while pressures vary from 1.0 to 2.7 GPa. The UHP metamorphic series is composed of the other eclogite subfacies of coesite eclogite and diamond eclogite, respectively, in the coesite and diamond stability fields (Fig. 3). These UHP eclogite facies have a very large range of temperatures from 700 to 1200 °C and are essentially defined by pressures higher than the quartz/coesite transition, that is, >2.7 GPa for this range of temperatures. The most striking feature of HP to UHP metamorphic rocks is that their density increases dramatically from normal upper crustal values of 2.75 (granite) to 2.94 (gabbro) to values as high as 3.1 to 3.63, respectively. Thus, for any given depths greater than 100 km, mafic rocks of oceanic crust would be metamorphosed to form eclogites (Zheng et al., 2016), which is constantly heavier than peridotite in the mantle, whereas the felsic continental crust remains less dense than the mantle. Therefore, the difference in crustal density, depending on their bulk compositions, result in contrasting vertical mobilities of HP to UHP metamorphic slices during continental deep subduction. As the metamorphic product of mafic rocks at the low geothermal gradients of <10C/km, eclogite is typically composed of garnet + omphacite quartz/coesite kyanite rutile phengite lawsonite. Although omphacitic clinopyroxene is a diagnostic mineral of eclogite, it also occurs together with glaucophane, lawsonite, titanite and epidote in blueschist-facies metamorphic rocks. In either case, its growth is associated with the mineral reaction of albite = jadeite + quartz at HP to UHP conditions (Figs. 2 and 3), and P-T conditions for this reaction are just close to the geotherm of ~10C/km (Fig. 1). On the other hand, its disappearance marks transformation of eclogite to garnet pyroxenite at elevated geothermal gradients (Figs. 2 and 3). Water contents in Alpine-type metamorphic rocks have great bearing on the mass transfer from the subducting crust to the mantle wedge (Zheng et al., 2016). They are usually estimated from water bound in hydrous minerals and rock porosities (Hacker, 2008; van Keken et al., 2011). In the quartz stability field, whole-rock water amounts remain high in blueschist-facies rocks, of the same order of magnitude as in low-P greenschist-facies rocks, but decrease considerably in HP to UHP eclogite-facies rocks, because of the scarcity of hydrous minerals. In the coesite stability field, stable hydrous minerals are lawsonite, phengite and zoisite (e.g., Schmidt and Poli, 2003; Wei et al., 2010; Zheng et al., 2011b). Although these hydrous minerals
rare in UHP metamorphic rocks, it does not mean that they are absent under UHP metamorphic conditions. Furthermore, considerable amounts of water may occur in the forms of structural hydroxyl and molecular water in nominally anhydrous minerals (Zheng, 2009; Zheng and Hermann, 2014). In this regard, UHP rocks cannot be viewed as distinctively water-deficient. The extreme variability in water contents, coupled with variations in temperature and pressure, plays a significant role in dictating the melting potential and ability of deeply subducted crustal rocks at mantle depths (Zheng et al., 2016).
2.3 High-temperature to ultrahigh-temperature facies series Granulites are generally considered as outcropped lower crustal rocks (Harley, 1989; Bohlen, 1991). They are often exposed in cratonic regions and characterized by water-deficient HT to UHT parageneses that are primarily composed of nominally anhydrous minerals with little to no hydrous minerals (White and Powell, 2002; O'Brien and Rötzler, 2003). Although the temperatures of 650 to 1100C are defined for HT to UHT metamorphism (Harley, 1998; Pattison et al., 2003), crustal depths are confined for metamorphic pressures. In the P–T petrogenetic grid (Fig. 2), the high dT/dP metamorphic space is occupied by the two facies of amphibolite and granulite. While the HT granulite facies is regionally expanded to the UHT facies through the polymorphic transition from kyanite to sillimanite (Fig. 2a), the HP granulite facies is defined in the garnet stability field at pressures above ~1.0 GPa. The lower P–T limits of the granulite facies are usually placed around 0.5 GPa and 650-750C. Mafic HP granulite is separated from garnet pyroxenite by the polymorphic transition between calcite and aragonite, where P-T conditions for this phase change are just close to the geotherm of ~15C/km (Fig. 3). Thus, HT to UHT metamorphism may take place at geothermal gradients of >15C/km, despite its generation often at high geothermal gradients of >30C/km. Because HT to UHT metamorphism of crustal rocks may take place at variable depths from the lithosphere-asthenosphere boundary (LAB) to the upper crustal level, its low pressure products are Buchan facies series whereas its medium to high pressure products are Barrovian facies series (Fig. 1). Although amphibolite-facies rocks are common in both facies series, they are generally underlain by granulites. Furthermore, hydration prevails during amphibolite-facies metamorphism, whereas dehydration prevails during granulite-facies metamorphism. For this reason, the protolith of granulites is usually viewed as a source of aqueous solutions for their overlying amphibolites (sink). In fact, partial melting often takes place at temperatures above the wet solidus of crustal rocks, leading to the extraction of
hydrous melts from the underlying granulites on the one hand and metasomatism of the overlying amphibolite-facies rocks on the other hand. Therefore, HT to UHT metamorphic rocks occur in both Barrovian and Buchan facies series (Fig. 2), and they are the products of extreme metamorphism at the high dT/dP space. Petrologically, the UHT metamorphic facies is indicated by robust thermobarometry or the presence of characteristic mineral assemblages in appropriate bulk composition. However, it is not easy to constrain the peak P–T conditions of UHT metamorphic rocks and to establish the spatial and temporal scales of extreme metamorphism in a given orogen. In the past, quantification of the P–T values for such rocks primarily relied on conventional thermobarometry and univariant petrogenetic grids based on the results of experimental petrology in simplified lithochemical systems (e.g., Harley, 1989, 1998; Bohlen, 1991; Pattison et al., 2003). There are many limitations of these approaches, principally due to partial reequilibration of mineral compositions during cooling at HT conditions. These limitations can be overcome by applying the new calibration of single mineral thermometers, particularly Al-in-orthopyroxene (Hollis and Harley, 2003), Zr-in-rutile (Tomkins et al., 2007) and Ti-in-zircon (Ferry and Watson, 2007). Furthermore, pseudosections may be calculated using an internally consistent thermodynamic dataset and appropriate activity–composition (a–x) models for multicomponent systems, which approach the complexity of nature (e.g., Korhonen et al., 2014; Li and Wei, 2016). Coupling this advance with thermodynamic models for minerals such as sapphirine and osumilite that are typically found in the Mg–Al-rich residues of clastic metasedimentary rocks under UHT conditions, in addition to those for the common rock-forming minerals and melts, has enabled the use of pseudosection thermobarometry for these UHT rocks (e.g., White et al., 2001, 2007; Kelsey et al., 2004, 2005; Santosh et al., 2012; Korhonen et al., 2014). Because pseudosection thermobarometry uses both mineral assemblages and whole-rock compositions, it has intrinsic advantages over conventional thermobarometry (Powell and Holland, 2008). The use of mineral assemblages as the evidence for UHT metamorphism, as distinct from mineral chemistry for thermobarometers, circumvents some of the problems inherent in conventional geothermobarometry (e.g., Harley, 1989, 1998; Bohlen, 1991; Kelsey et al., 2003). As such, the UHT metamorphic facies are typically indicated by the presence of a diagnostic mineral assemblage in an appropriate bulk composition, such as assemblages with sapphirine + quartz, orthopyroxene + sillimanite quartz, osumilite and spinel + quartz (Harley, 2008; Kelsey, 2008; Brown, 2009). Such assemblages are commonly preserved in extremely Mg-Al-rich rocks of restitic origin. Occasionally widespread assemblages
like garnet + orthopyroxene, ternary feldspars, (F-Ti) pargasite or metamorphic inverted pigeonite are taken to be indicators of UHT metamorphism. Mineral assemblages characteristic of UHT metamorphism are generally found in pelitic rocks that have relatively high Al and Mg contents. Such Mg–Al-rich rocks are volumetrically rare in most metamorphic terrains. It was proposed that high magnesian bulk compositions may represent residual bulk compositions (e.g., Droop and Bucher-Nurminen, 1984; Raith et al., 1997) developed as a result of the removal of melt rich in Si and Fe. However, very little Fe and Mg partition into silicate melt (Kelsey et al., 2003), in which case the protolith of many diagnostic UHT mineral assemblages must have elevated Mg with respect to Fe (Kelsey et al., 2003). Nevertheless, the UHT metamorphism may also be diagnosed from psammitic and non-pelitic rocks (e.g., Brandt et al., 2003). Because the generation of UHT metamorphic rocks has bearing not only on orogenic processes but also on postorogenic reworking, these rocks have been intensively investigated by various approaches. The results indicate that UHT metamorphic events may take place in the geological history of Archean to Cenozoic (Fig. 6). Although the HT to UHT metamorphic rocks commonly occur at the convergent plate margins of Precambrian age, their petrogenesis has no direct link to subduction orogeny. Furthermore, there are still a number of problems in constructing the P-T paths of UHT metamorphic rocks. Most of applicants used either accessory mineral thermometry without any support from phase equilibria modeling or a pseudosection approach commonly in conjunction with Al-in-orthopyroxene thermometry but without confirmation from accessory mineral thermometry. On the other hand, the majority of geodynamic models for the thermal evolution of orogens do not generally predict such extreme temperatures (Clark et al., 2011). Because the HT to UHT metamorphism plays a fundamental role in the development and stabilization of continents, it is necessary to develop appropriate tectonic models to account for the extreme metamorphic conditions in orogens.
3. Tectonic settings for extreme metamorphism 3.1 Thermal regimes and metamorphic P–T paths It has long been recognized that the thermal structure of orogenic belts has a significant impact on the mineral paragenesis of metamorphic rocks (Oxburgh and Turcotte, 1970; England and Richardson, 1977). It was found that the internal thermal driving forces are associated with the movement of rock masses, with redistribution of heat inside orogens. This
explicit linkage between deep and shallow processes provides the first step in understanding how accretionary and collisional orogens may evolve and, as part of that evolution, how metamorphic rocks were exhumed. Oxburgh and Turcotte (1970) made the lithosphere-scale thermal models for thrusted regions, suggesting that perturbed geotherms would be produced to result in anomalously low-T crust at lithospheric depth, and produce temperature-depth (dT/dz) profiles in which T rises slower than the P increase with depth in cold subduction zones. By means of the subduction channel mechanism for exhumation of these cold rocks to shallow levels (Gerya et al., 2002; Zheng et al., 2013), the metamorphic record can be retained in these rocks if they did manage to reach the surface. In general, regional metamorphic rocks do not simply record peak mineral assemblages along the gradients mapped out by metamorphic field arrays, but instead record segments or points along P-T paths that reflect, in simple terms, the interplay between burial, heating and exhumation with time (Thompson et al., 1987; Zheng et al., 2011a). Lawsonite eclogites are a sound example that their peak mineral assemblages are not recorded (Wei and Clarke, 2011; Wei et al., 2013). As a consequence, P-T paths have become key pieces of evidence to constrain tectonic models for the behavior of orogens, especially when combined with geochronological data that can constrain the rate of P-T changes. The study of thermal modelling by England and Richardson (1977) made a conceptual advance in metamorphic P-T paths. The metamorphic field array, or piezothermal array, is the locus of maximum T points attained by rocks which travelled along P- T paths that are likely to be at a high angle (in the P-T space) to the preserved array itself. In particular, large UHP terranes generally experienced longer metamorphic periods under HP to UHP conditions than small UHP terranes (Zheng et al., 2009, 2013; Kylander-Clark et al., 2012). As a consequence, the peak T is attained after the peak pressure (Fig. 7). In this case, the P-T paths traverse to higher pressures first (i.e., burial dominates) and then heating occurs during the initial exhumation toward the surface (Zheng et al., 2011a). On the basis of field and petrographic observations, Zwart (1962) recognized that isograds might not be isochronous and that metamorphic rocks might record P-T histories that do not correspond to the final field array dT/dz gradient. Whereas isograds in progressive metamorphism were generally treated as broadly equivalent in age, the thermal modelling of England and Richardson (1977) suggested that higher-grade isograds and mineral zones are younger than the lower-grade ones. As a result, isograds could no longer be treated as fossil markers that define the P-T framework of metamorphic rocks in orogens. Instead they must
be considered along with evidence for the P-T paths of metamorphic rocks within the framework of tectonic evolution. In this respect, their study also laid the foundation for P-T paths as records of orogenic metamorphism that constrain the underlying tectonothermal processes (England and Thompson, 1984; Thompson and England, 1984). Therefore, understanding the record of regional metamorphism has become central to the development of tectonic models for plate behaviors, including subduction-zone evolution and collisional orogeny. The concepts introduced have changed not only the meaning, significance and use of isograds but also the viewpoint on the importance of metamorphism in plate tectonics. The tectonic history of a region, the evolutionary history of metamorphic rocks within that region, can be represented by the P-T-t path of those rocks. In a one-dimensional analysis, by which the crust is regarded as composed of columns of rocks with equal physical properties and between which there is no lateral heat transfer, this relationship is represented as a surface in the P-T-t space. The metamorphic histories of rocks are represented by lines on this surface. The projections of such rock histories onto the three two-dimensional planes of the P-T-t box represent the T-t paths described by the geochronologist, the P-T paths studied by the metamorphic petrologist and the P-t paths inferred by the tectonicist. The locus of peak P-T conditions is preserved by the mineral assemblages, which are represented by the metamorphic facies series (Miyashiro, 1961), metamorphic geotherm (England and Richardson, 1977), piezothermic array (Richardson and England 1979), P-T array (Thompson and England 1984), metamorphic field gradient (Spear et al., 1984) or set of metamorphic zones (Harte and Dempster, 1987). They all result from the intersection of P-T paths for individual rocks with the erosion surface (England and Richardson, 1977). Thus, each location along the metamorphic field gradient represents a unique point of pressure, temperature and time. Unless these points are strictly contemporaneous, otherwise the metamorphic field gradient is necessarily the locus of diachronous P-T conditions. In the traditional view of regional metamorphism, a single geothermal gradient is hypothesized for Barrovian-type metamorphic belts. As such, prograde metamorphism is indicated by a series of T-increasing processes, whereas retrograde metamorphism is characterized by a series of T-decreasing processes subsequent to the peak temperature with regardless of changes in pressure. P-T paths can be derived from individual rocks (e.g., Thompson and England, 1984) and minerals (e.g., Spear and Selverstone, 1983), and they can be related to the tectonic setting (e.g., England and Thompson 1984). Metamorphic reactions in the higher grade zones of the Barrovian type area were investigated by McLellan (1985), who also emphasized the importance of both subsolidus and supersolidus processes in the
generation of migmatitic leucosomes (McLellan, 1983, 1989). Further studies found that collisional orogens often contain UHP metamorphic rocks due to continental deep subduction. In such rocks, prograde metamorphism is characterized by a series of P-increasing processes, whereas retrograde metamorphism is indicated by a series of P-decreasing processes subsequent to the peak pressure with regardless of changes in temperature (Ernst, 1988). As a consequence, the derivation of P-T paths along the length and breadth of a collisional orogen enables us to unravel the three-dimensional reality of orogenic processes. A major development in metamorphic petrology during the past three decades is the realization that regional metamorphism is a series of dynamically evolutionary processes rather than a static process. The evidence for the dynamic evolution is recorded by mineralogical and geochemical disequilibria. As a consequence, one basic task in metamorphic petrology is to identify the mineral assemblages generated at peak metamorphic conditions and relate them to P-T-t history. Our ability to separate partially overprinted equilibrium mineral assemblages in partially disequilibrium rocks has enabled the use of thermobarometry to determine different sets of P-T data for individual rocks. To assess the peak P-T conditions, we have relied on thermodynamics and phase equilibria, and we have made the assumption that either the mineral assemblages, through use of a petrogenetic grid, or the mineral chemistries, utilizing thermobarometry, provide the record of metamorphic P and T conditions that are geologically significant. There are two fundamentally different types of P-T paths for regional metamorphic rocks, which are distinguished by relative timing of maximum T and maximum P. One is clockwise P-T paths, which achieved maximum P before maximum T. The other is anticlockwise P-T paths, which achieved maximum T before maximum P. The clockwise P-T paths are generated by two-stage processes: (1) crustal thickening due to subduction; and (2) crustal thinning during exhumation. Such evolutionary paths lead to metamorphic dehydration at temperatures below the wet solidus of crustal rocks during subduction, and the metamorphic peak normally postdates early deformation. Dehydration melting is insignificant in the subduction stage because of the low geothermal gradients, but it may become significant in the exhumation stage due to elevated geothermal gradients upon decompression. In this case, alkalic magmatism may be produced in association with amphibolite-facies metamorphism at the upper level (Zheng et al., 2009. 2015). The anticlockwise P-T paths are also generated by two-stage processes: (1) removal of the thickened orogenic root via a certain mechanism (e.g., slab breakoff, lithospheric delamination, and asthenospheric erosion); (2) underplating of the asthenospheric mantle for
the high heat flow into the thinned lithosphere. Such evolutionary paths lead to dehydration melting of the lowest crust, and the metamorphic peak also postdates early deformation. Since geothermal gradients are much higher in the second stage than in the first stage, the extent of dehydration melting is much larger in the late stage than in the early stage. As a consequence, HP granulite-facies metamorphism and adakitic magmatism occur in the first stage, whereas medium-P to low-P metamorphism and extensive magmatism occur in the second stage in association with orogenic collapse and crustal doming. Collisional orogeny includes two types of continent-continent and arc-continent collisions, and collisional metamorphism is substantially a result of crustal thickening at convergent plate margins. In general, the continent-continent collision results in deep subduction of one continent beneath the other to mantle depths for UHP metamorphism (e.g., Western Alps of Italy, Dabie-Sulu of China, and Western Gneiss Region of Norway). In contrast, the arc-continent collision may lead either to deep subduction of a continental block beneath an arc terrane for UHP metamorphism (e.g., Himalaya) or to shallow subduction of an arc terrane beneath a continental block for HP metamorphism (e.g., Appalachia). In either case, collisional metamorphism takes place in subduction zones, resulting in clockwise P-T paths, whose general characteristics are well understood as exemplified in the study of metamorphic rocks from the Appalachians (Armstrong et al., 1992), the Caledonides (Anderson et al., 1992) and the Himalayas (Searle et al., 1992). Furthermore, collisional orogens exhibit a general lack of coeval calc-alkaline magmatism (Hamilton, 1970; Molnar and Tapponnier, 1975; Molnar et al., 1987; Burchfiel et al., 1989; Rumble et al., 2003; Zheng et al., 2003). Clockwise P–T paths are characteristic of Alpine-type UHP eclogite-facies metamorphic terranes in continental subduction zones, where the maximum pressure is achieved at the same time or before the maximum temperature (Fig. 8). They pass through two kinds of polymorphic transition boundary with increasing pressure into the UHP regime during subduction. The first is the transformation from quartz to coesite and second is the transformation from graphite to diamond. After achieving the peak pressure, UHP slices may continue to heating without significant decompression (Fig. 7). In general, small UHP terranes such as Dora Maira, Lago di Cignana and Kaghan Valley underwent short-duration UHP metamorphism and were exhumed with synchronous decreases in both pressure and temperature, whereas large UHP terranes such as Western Gneiss Region, Dabie and Sulu underwent long-duration UHP metamorphism and were exhumed with near-isothermal decompression (Zheng et al., 2009, 2013; Kylander-Clark et al., 2012; Hermann and Rubotta,
2014). In either case, the exhumation P-T paths occur at higher temperatures than the subduction P-T paths. On the other hand, UHP terranes like Kokchetav, Erzgebirge and Eastern Greenland record both peak UHT and UHP metamorphic conditions (Fig. 8), suggesting slab breakoff for their exhumation (Zheng and Chen, 2016). The large UHP terranes often experience an increase in temperature during decompressional exhumation (Figs. 7 and 8). This results in a specific style of retrograde metamorphism for UHP assemblages, exhibiting granulite-facies dehydration on the one hand and the amphibolite-facies hydration on the other hand. The free water is derived not only from the breakdown of hydrous UHP minerals but also from the exsolution of structural hydroxyl and molecular water from nominally anhydrous UHP minerals (Zheng, 2009; Gong et al., 2013; Zheng and Hermann, 2014). The resulted amphibolite-facies rocks are the retrograde product of Alpine-type UHP metamorphic rocks during decompressional exhumation. In contrast, Barrovian-type amphibolite-facies metamorphism is coeval with HT to UHT metamorphism. In either case, the amphibolite-facies units overlie the granulite-facies units if there was no gravitational reverse (Zheng and Chen, 2016). Although the water in both cases is of internal origin, the former is derived from decompressional dehydration of UHP minerals during exhumation whereas the latter is produced by heating dehydration of HT to UHT metamorphic minerals during and subsequent to foundering of orogenic roots. Early studies suggested anticlockwise P–T paths for HT to UHT granulites (Thompson and England, 1984; Bohlen, 1987, 1991), in which the maximum temperature is achieved at the same time or before the maximum pressure. However, later studies indicate that the majority of these rock show clockwise P-T paths (Harley, 1989, 1998, 2008; Brown, 2001; Li and Wei, 2016). Nevertheless, there are still some rocks showing anticlockwise P-T paths (e.g., Zhao, 2009; Santosh et al., 2012; Korhonen et al., 2014; Tong et al., 2014; Xiang et al., 2014). There may be two stages for development of some HT to UHT metamorphic rocks: (1) the isothermal decompression of deeply subducted crust, and (2) a younger HT to UHT overprint due to introduction of an external heat source. Based on his combined metamorphic-structural studies in the Pyrenees of Europe, Zwart (1962) recognized plurificial metamorphism, the overprinting of at least two distinct metamorphic events in different tectonic regimes. In the case of the Pyrenees, the two metamorphic events identified were also of very different ages and related to entirely different types of orogeny (the Hercynian/Variscan in the Carboniferous and the Pyrenean in the Paleogene), rather than being two episodes during the thermal evolution in one orogeny. Therefore, it is possible that anticlockwise P–T paths are characteristic of plurificial
metamorphism in composite orogens. As such, it is critical to determine P-T-t paths for HT to UHT metamorphic rocks. While the P-T paths of regional metamorphic rocks can be deduced from petrographic, thermobarometric, and thermodynamic data, but their genetic interpretation is not unique in many cases. They may be either the result of a single cycle during collisional orogeny or the cumulative effect of several orogenic cycles. Although the rates of burial and uplift, with or without magmatic heating, may vary during a single orogenic cycle, single-cycle P-T paths are commonly represented as simple smooth curves. In contrast, multi-cycle P-T paths are a composite of information partially preserved from more than one path; an earlier clockwise P-T path may be superimposed by a later anticlockwise P-T path. This results in a more complex form for multi-cycle P-T paths in plurificial metamorphic terranes. In this regard, it is necessary to date plurificial metamorphic events in order to have a reasonable link to their tectonic settings. Furthermore, it remains to examine whether clockwise P-T paths are responsible for the Barrovian-type HT to UHT metamorphism whereas anticlockwise P-T paths are responsible for the Buchan-type HT to UHT metamorphism. Although much advances in petrological and geochemical thermobarometries have been made in the past three decades, difficulties are often encountered in retrieving the P-T conditions of either UHP or UHT metamorphic rocks. They are primarily caused by problems in the partial equilibration between metamorphic/peritectic minerals and the preservation of protolith minerals. This basically involves mineralogical processes such as recrystallization of protolith minerals during dehydration and anatexis, reactions between protolith minerals either during metamorphic dehydration below the wet solidus to produce metamorphic minerals and aqueous solutions or during partial melting on and above the wet solidus to produce peritectic minerals and anatectic melts, reactions of protolith minerals with metamorphic fluids (aqueous solutions and/or hydrous melts) to produce metasomatic minerals, and reactions of primary and/or secondary minerals with the fluids during their migration. Mineral dissolution and precipitation may occur in response to the access of metamorphic fluids or anatectic melts, and crystal-plastic processes activated by and facilitating deformation may also proceed contemporaneously with metamorphic and peritectic reactions. If the timescale of metamorphism is long enough for thermodynamic equilibrium to be attained or closely approached, the product mineral paragenesis is primarily dictated by the P-T conditions, the presence and nature of any fluid phase, and the compositions of protoliths. The textural features of metamorphic and peritectic products are determined in a complex way by the physicochemical conditions, the mechanisms of mineral
reactions, and the presence and nature of metamorphic fluids or anatectic melts.
3.2 Regional metamorphism in relation to tectonics At the dawn of the age of plate tectonics, Miyashiro (1961) noticed significant differences in the phase assemblages of regional metamorphism in circum-Pacific regions, found contrasting geological occurrences, and inferred characteristic ranges of their formation conditions. He put forward the concept of paired metamorphic belts, involving an oceanward, narrow, low-T/high-P blueschist zone intimately intermixed with ophiolites, and a landward, broad, high-T/low-P realm associated with arc volcanic–plutonic rocks. It appears that Miyashiro (1961) was one of the first geologists to tie regional metamorphism with orogeny within the framework of plate tectonics. The paired metamorphic belts were then reinterpreted in the light of the subduction zone model as resulting from the contrasting thermal gradients (e.g., Oxburgh and Turcotte, 1970). Zwart (1967, 1969) noticed differences in metamorphic patterns and developed in intracontinental orogens, identifying a duality of metamorphic zones in Europe that were associated with differences in orogenic processes. Largely on the basis of available observations in Europe and North America, he proposed three types of orogens in relationship to tectonic setting and style: (1) Alpinotype orogens, which featured HP metamorphic zones and included high-P/low-T facies series. Granites and migmatites were rare or of only minor extent, whereas meta-ophiolitic and ultramafic rocks might be present and volumetrically important. Examples for it include not only the Alps itself but also its eastern extension, the Himalayan orogen; (2) Cordillerotype orogens, which were dominated by calc-alkaline igneous rocks and high-T/low-P metamorphic facies series. There is a general lack of ophiolites, and of migmatites and abyssal sediments. Examples for this type of orogens include not only the Cordillera itself but also the other circum-Pacific orogens containing continental arcs; (3) Hercynotype orogens, which were typified by shallow, low-P metamorphism with high temperature gradients leading to T-dependent metamorphic zones. They were characterized by an abundance of migmatites and granites and an absence of ultramafic and ophiolitic rocks, very broad high-T metamorphic zones and limited evidence for any exhumation near the metamorphic peak. Examples for this type of orogens include not only the Hercynian itself (typically the Pyrenees) but possibly also the Lachlan Fold Belt in Australia. After a great deal of studies in the past five decades, it turns out that the Alpinotype orogens are the product of collisional orogeny in the low geothermal gradients (Fig. 9b), and the
Cordillerotype orogens are the product of accretionary orogeny above oceanic subduction zones (Fig. 9a). Therefore, the two types of orogens are associated with subduction tectonism, and they are respectively consistent with collisional orogens (Dewey and Burke, 1973; Dewey and Kidd, 1974) and accretionary orogens (Coney, 1987, 1992). The accretionary orogens form above the boundary of oceanic subduction during continuing plate convergence in the absence of continental collision (Cawood et al., 2009). They vary according to whether the subduction boundary or trench is retreating, neutral or advancing. They may exhibit cyclic behavior in which a retreating trench changes to an advancing trench or in which an advancing trench changes to a retreating trench. Previously these orogens were called Pacific-type (e.g., Matsuda and Uyeda, 1971), Cordilleran-type (e.g., Coney et al., 1980) and Miyashiro-type (Maruyama, 1997). The collisional orogens form at convergent continental margins where ocean basins close and the subduction of oceanic slab steps back and flips (arc collisions), simply steps back and continues with the same polarity (block and terrane collisions) or ultimately ceases (continental collisions). Previously such orogens were called Turkic-type (Sengör and Natal’in, 1996) and Himalayan-type (Liou et al., 2004). Accretion of arcs and/or allochthonous terranes, some of which may be far-travelled, is a common feature of accretionary orogens. As a result, the collisional orogens are often superimposed on the accretionary orogens. Tectonically, Alpinotype orogens are elongate and relatively narrow with a high aspect ratio, where nappe structures are well developed at several structural levels and have been exposed because of the rapid (i.e., tens of millions of years) uplift and exhumation. Historically, rock assemblages typified by mafic blueschists and eclogites have been viewed as produced under relatively high-P/low-T conditions (Eskola, 1939; Fyfe et al., 1958). Eclogitic pillow lavas in the Zermatt–Saas-Fee area of the western Swiss Alps were described by Bearth (1959, 1967), demonstrating that submarine basalts, initially solidified at very low pressures, had been subjected later to HP metamorphic conditions. Miyashiro (1961) called attention to the restriction of high-P/low-T rock assemblages to long, linear belts sited oceanward from broad, more-or-less coeval, high-T/low-P assemblages beneath continental arc volcanics (paired metamorphic belts). With the advent of plate tectonics theory, it has been recognized early that HP terranes formed because of the subduction of oceanic lithosphere (Hamilton, 1969; Dewey and Bird, 1970; Ernst, 1970). The discovery of UHP index minerals such as coesite and diamond in eclogite-facies metamorphic rocks immediately led to the recognition that continental crust was subducted to mantle depths of >80 km for UHP metamorphism in collisional orogens (e.g., Chopin, 1984;
Smith, 1984; Sobolev and Shatsky, 1990; Xu e al., 1992). This was primarily based on the consensus in petrology that such index minerals are only generated by the lithostatic pressure at continental subduction zones (e.g., Ernst and Liou, 1995; Schreyer, 1995; Chopin, 2003). So far about 30 UHP terranes have been identified in collisional orogens (Liou et al., 2009, 2014). Consumption of an oceanic slab beneath a continental plate eventually introduced a salient of continental crust into the subduction zone and, attending collisional suturing, deepest parts of the down-going crust were subjected to UHP metamorphism at subarc depths (Ernst, 2005). The occurrence of deformed HP and UHP crustal rocks in collisional orogens testifies to the dynamothermal evolution of convergent plate boundaries (Maruyama et al., 1996; Zheng, 2012). Therefore, the Alpinotype orogens are petrologically characterized by the occurrence of Alpine-type HP to UHP metamorphic rocks along collided continental margins. A piece of orogenic eclogites can confidently serve as a guide to track subduction-zone processes at the low geothermal gradients (Chopin, 2003; Zheng et al., 2016). In the early stage of continental collision (subduction), temperatures on the top of subducting lithosphere are generally lower than the wet solidus of crustal rocks (Zheng et al., 2016). Nevertheless, metamorphic dehydration and deformation are remarkable, leading to development of a ductile layer between the subducting cover and the basement to facilitate exhumation of the deeply subducted crustal rocks along a continental subduction channel (Zheng et al., 2013). Metamorphic segregation and fluid extraction are syntectonic processes, and fluid migration pathways commonly relate to rock fabrics. In the late stage of continental collision (exhumation), temperatures on the top of subducting lithosphere become higher than the wet solidus of crustal rocks (Zheng et al., 2016). The breakdown of hydrous minerals below the wet solidus produces metamorphic fluids but that on and above the wet solidus yields anatectic melts. While decompressional dehydration of the UHP eclogite-facies rocks releases aqueous solutions for the overlying amphibolite-facies metamorphism, it leaves the underlying granulite-facies residues. The Hercynotype orogens are represented by zones of deformation, metamorphism and magmatism at high geothermal gradients, but they occur in the interior of continental blocks away from active plate margins. Because they are generally associated with extensional tectonics (Elston, 1984; Dewey, 1988), they are often referred to as extensional orogens. On the other hand, they are generally abundant in HT granulites and granites and thus referred to as hot orogens (e.g., Collins, 2002; Beaumont et al., 2011), some contain UHT granulite-facies metamorphic rocks and are referred to as ultrahot orogens (Chardon et al., 2009). In this regard, the Hercynotype orogens are also characterized by the occurrence of HT
to UHT metamorphic rocks. Although these rocks commonly occur at convergent plate margins, but it does not mean that they are the products of regional metamorphism during accretionary and collisional orogenies. Rare occurrences of mineral assemblages such as sapphirine-quartz, spinel-quartz and hypersthene-sillimanite-quartz were reported earlier in typical amphibolite-granulite terrains (Morse and Talley, 197l; Bondarenko, 1972; Caporuscio and Morse, 1978; Ellis et al., 1980). However, the recognition of UHT metamorphism in these rocks may be tracked back to Ellis (1980), who obtained 900-980C at 8-10 kb for osumilite-sapphirine-quartz granulites from Enderby Land in Antarctica. The occurrence of sapphirine–quartz assemblages in orthopyroxene–garnet granulites from the Anapolis–Itaucu Complex in the Brasilia Fold Belt indicates extreme P–T conditions at 1030–1050°C at ~10 kb, and high Al2O3 contents of 12.9–9.7 wt% in cores of orthopyroxene porphyroblasts even suggest the maximum T >1150°C (Moraes et al., 2002). So far about 50 localities that contain UHT mineral assemblages have been identified in the world (Kelsey and Hand, 2015). In some localities, the UHT assemblages are developed on regional scales, for instance in the Napier Complex of E. Antarctica, the Kerala Khondalite Belt and Madurai Block of the Eastern Ghats in India, the Khondalite belt in the North China Craton, and the Musgrave Block in central Australia. It appears that UHT metamorphic rocks are more common than UHP ones on Earth. With respect to the petrogenesis of HT to UHT granulites, however, a direct link to plate tectonics was not achieved yet, with the controversy not only in the heat source but also in regional stress regime between compression and extension. Traditionally, it is thought that the tectonometamorphic evolution of such orogens cannot be explained by conventional models of plate tectonics at convergent boundaries (e.g., Dewey and Bird, 1970; Tackley, 2000) and thus a specific geodynamic frame is required to transfer high heat flow from the mantle into the crust. Because they were generally generated in intraplate settings, intracontinental orogenies were often considered as the possible mechanism for their formation (e.g., Molnar and Tapponnier, 1975; Hand and Sandiford, 1999; Cunningham et al., 2009; Raimondo et al., 2010; Tugend et al., 2014; Wernert et al., 2016). Although oceanic and/or continental subduction is not active any more in intracontinental orogens, there are variable extents of tectonic inheritance from preexisting accretionary or collisional orogens (e.g., Dewey, 1988; Vauchez et al., 1997; Tommasi and Vauchez, 2001). Nevertheless, lithospheric thinning is temporally associated with the high-T/low-P metamorphism during the so-called intracontinental orogenies (Vielzeuf and Kornprobst, 1984; Weber, 1984; Wickham and Oxburgh, 1985; Sandiford and Powell, 1986). For this reason, rift tectonics is often suggested to play a role in
the formation of HT to UHT granulites (e.g., Thompson et al., 2001; Cunningham et al., 2009; Raimondo et al., 2014; Tugend et al., 2014; Tucker et al., 2015; Wernert et al., 2016). While extensional tectonic settings are evident for the high-T/low-P metamorphism, compressional regimes are common in high-T/low-P metamorphic rocks. This leads to ambiguous identification of continental rifts in collisional orogens. At present, it is still unresolved how HT to UHT metamorphic rocks are produced despite their occurrence also at convergent plate margins (Brown, 2014; Kelsey and Hand, 2015; Harley, 2016). While a piece of granulites does record metamorphic dehydration and partial melting at high to ultrahigh temperatures, such temperatures are considerably higher than lithospheric geotherms at crustal depths. Significant heating, accompanying uplift and erosion of the entire orogen, is common during post-subduction episodes of orogenic events. In this case, orogenic crust must be heated for partial melting. Crustal anatexis proceeds at different degrees, producing anatectic melts that may ascend and migrate into the overlying crust for felsic magmatism. Therefore, the HT to UHT metamorphism requires a larger increase in temperature than in pressure, so that the thermal perturbation is necessary in order to provide the high heat flow for crustal anataxis at shallow levels. The decay of radioactive isotopes in overthickened crust was suggested as a heat source for crustal anataxis (Holmes, 1931). As concluded by Clarke et al. (2011), however, the internal radiogenic heat production is insufficient for this purpose. As such, it is necessary to have external heat sources such as underplating of the asthenospheric mantle or its derived mafic magmas (Zheng and Chen, 2016). Initially, Barrow (1893) noticed that metamorphism of regional extent is generally associated with intrusive rocks of magmatic origin and thus regarded the latter as the heat source. For a long time, studies of crustal anataxis and granulite-facies metamorphism in orogens have ascribed the heat supply to contemporaneous mantle-derived mafic magmas (e.g., Wells, 1980; Sills, 1984; Lux et al., 1986; Wickham and Oxburgh, 1987; Huppert and Sparks, 1988; Voshage et al., 1990; De Yoreo et al., 1991). The occurrence of mafic intrusives adjacent to granulites was cited as geological evidence for this argument (e.g., Zhao, 2009; Guo et al., 2012). However, this hypothesis was challenged by Barboza et al. (1999) who provided the evidence that the emplacement of a given mafic igneous complex occurred after the regional granulite-facies metamorphism and thus was unlikely to supply the heat for the HT to UHT metamorphism. Furthermore, inspection of the mafic igneous rocks indicates that they are generally enriched in radiogenic Sr-Nd isotopes, differing from those of mid-ocean ridge basalts (MORB) derived from partial melting of the asthenospheric mantle (Zheng et al., 2015).
The subcontinental lithospheric mantle is cold and its partial melting also requires an external heat source (Zheng and Chen, 2016). Therefore, the other external heat source is necessary for the extreme metamorphism in preexisting orogens. The generation of magmas in the lower crust requires high heat flow, which is commonly manifested in continental rift settings where the thermal boundary layer is shallowed in the thinned lithosphere (McKenzie, 1978; Thompson, 1981). Lithospheric extension and its resulted orogenic collapse were considered as a possible mechanism for regional metamorphism and crustal anataxis at high geothermal gradients (Dewey, 1988; Menard and Molnar, 1988; Sandiford, 1989; Gibson and Irelandt, 1995). Backarc basins are often considered as a possible location for UHT metamorphism (e.g., Collins, 2002; Brown, 2006, 2007). This is primarily based on the observation that many of modern backarc basins are of thin lithosphere and high heat flow (Hyndman et al., 2005). It is known for a long time that backarc extension is generally associated with rollback of the subducting slab and its caused mantle upwelling (e.g., Elsasser, 1971; Sleep and Toksoz, 1971; Oxburgh and Turcotte, 1974; Molnar and Atwater, 1978; Uyeda and Kanamori, 1979; Garfunkel et al., 1986; Shemenda, 1993). For this reason, the slab rollback is suggested as a possible mechanism for UHT metamorphism (Pownall et al., 2014). Although backarc basins are common zones of high heat flow, the majority of UHT metamorphic belts were not located in backarc settings during their formation. Instead, they are common in accretionary and collisional orogens, where previous mountain-building systems were reactivated to result in bimodal magmatism and high-T/low-P metamorphism (Vielzeuf and Kornprobst, 1984; Weber, 1984; Wickham and Oxburgh, 1985; Sandiford and Powell, 1986). In accretionary orogens, UHT metamorphic belts are generally sandwiched by several hundreds of km thick passive margin sequences together with oceanic components in forearc settings, some of which were dragged down along a subduction channel prior to collisional orogeny and exhumation. On the other hand, collisional orogens containing UHT metamorphic rocks generally preserve a nearly flat Moho though they were thickened due to subduction of supracrustal rocks into the lower crust. The UHT metamorphism requires temperatures at the base of the thickened crust of >900 °C, but more commonly in excess of 1100 °C. Such high Moho temperatures indicate that these ultrahot orogens consist of an extremely thinned mantle lithosphere (Sandiford and Powell, 1991). This indicates that the lower crust of ultrahot orogens must be close to the asthenosphere at the time of orogeny. Furthermore, UHT metamorphic belts at convergent plate margins do not show any characteristic association with lithotectonic units formed in backarc settings. Therefore, the backarc model cannot satisfactorily explain
the formation of all UHT metamorphic rocks at convergent plate margins. Clockwise P-T-t paths commonly form during collisional orogeny at the termination of a Wilson-cycle. Although such paths are primarily related to subduction tectonism, they rarely develop in circum-Pacific orogens and cannot account for the features of isobarically cooled low- and medium-P granulite terranes in many Precambrian cratons. In contrast, broad zones of high heat flow occur in orogens that undergo lithospheric thinning and are associated with widespread felsic magmatism. These include the Basin and Range Province in the western United States and Mexico, the Taupo volcanic zone in New Zealand, and the vast circum-Pacific orogens, which are characterized by voluminous batholiths and repeated contraction-extension events. These orogens are underlain by granulites, but few are exposed, because crustal doming rarely occurs, though granulitic xenoliths in young basalts attest to their existence. Geological evidence and geodynamic modeling indicate that various orogens are not closed systems and their base may receive the high heat flow from the underlying asthenospheric mantle. Therefore, we have to look for the other mechanism to account for the formation of granulites, migmatites and granites in association with reworking of the orogenic crust. Collisional orogens generally develop at a time variably earlier than their extensional collapse, which may eventually lead, in some circumstances, to continental rupture for seafloor spreading. Wilson (1966) noted that the Central and North Atlantic oceans were nucleated upon the Appalachian/Caledonian orogenic belt, itself the site of earlier closed oceans. This concept has been extended to many sites of continental extension, some of which lead to continental rupture (e.g., Dewey, 1988; Vauchez et al., 1997; Tommasi and Vauchez, 2001). These include the Basin and Range Province upon the Laramide Belt in USA, parts of the East African Rift on the Mozambique Belt, the South Atlantic upon the Gariep Belt in Africa, and the Tasman Sea upon the Lachlan Belt in Australia. It is known that mid-ocean ridges are generated by seafloor spreading and high heat flow prevails along such sites, where mountain-building systems are primarily composed of MORB. Such basaltic magmatism is induced by decompressional melting of the asthenospheric mantle in response to its upwelling. The high-T/low-P metamorphism is common at mid-ocean ridges, where on-axis metamorphism is of much higher temperatures than off-axis metamorphism, and slow spreading leads to ridge metamorphism at higher temperatures than fast spreading (Manning et al., 1996, 2000; Nicolas et al., 2003; Bosch et al., 2004). Backarc basins are also generated by seafloor spreading. Some of them may develop beneath marginal arc terranes, leading to both growth and reworking of the juvenile crust. As such, the backarc spreading has superimposed on accretionary orogens, in
which HT to UHT metamorphism did take place. In summary, there are controversial interpretations of the tectonic regime for the petrogenesis of HT to UHT granulite-facies metamorphic rocks. The occurrence of these rocks at convergent plate margins has challenged our knowledge of regional metamorphism within the framework of plate tectonics. This may be ascribed to lack of the following two issues (e.g., Harley, 1998, 2008; Kelsey, 2008): (1) qualified approaches to determine the high temperatures for formation of these HT to UHT mineral assemblages; (2) reasonable understanding of the heat source for the extreme metamorphism at crustal pressures. Furthermore, lithospheric extension may take place at different stress regimes, resulting in different styles of basin sedimentation, regional metamorphism and rift magmatism. In particular, the absence of a reasonable geodynamic model capable of taking heat to crustal rocks at shallow depth for dehydration melting has prevented universal acceptance of the high heat flow from the asthenospheric mantle as fundamental to the HT to UHT metamorphism. While the trinity of continental drift, seafloor spreading and plate subduction does constitute the plate tectonics theory, the lithospheric extension in response to asthenospheric upwelling may provide an important complement to the tectonic development of orogenic collapse in intracontinental settings. In view of the above arguments, it appears that the high heat flow is prominent in locations of the thinned lithosphere, giving rise to a specific type of mountain-building systems not only in backarc basins and mid-ocean ridges but also in the accretionary and collisional orogens of intracontinental settings. In particular, orogenic belts with thickened crusts are the preferential sites of extension (e.g., Dewey, 1988; Vauchez et al., 1997; Tommasi and Vauchez, 2001). Thus, the Hercynotype orogens can be viewed as developing on the preexisting orogens, which underwent reworking at the high geothermal gradients. Furthermore, they are generally associated with rift tectonics. Therefore, we prefer to name such mountain-building systems as rifting orogens instead of the extensional orogens. As such, rifting orogeny instead of the extensional orogeny is suggested to develop in the thinned lithosphere in response to the high heat flow from the underlying asthenospheric mantle. In this context, the Hercynotype orogens are the product of rifting orogeny at the high geothermal gradients (Fig. 9c). As documented below, active continental rifting subsequent to thinning of the subduction-thickened lithosphere is the most possible mechanism for the HT to UHT metamorphism.
4. Active continental rifting for Buchan-type HT to UHT metamorphism 4.1 Rift types and origins Rift tectonics is typical of divergent plate boundaries. It is evident not only in oceanic settings where mid-ocean ridges and backarc basins occur in linear zone, but also in continental settings where the association of crustal faults, basin sediments, volcanic sequences occur in linear zones. Continental rift zones are long, narrow tectonic depressions on the Earth’s surface where the lithosphere has been thinned significantly (Olsen, 1995; Ruppel, 1995). Thus, continental rifting can be described as a thinning process of the lithosphere. Nevertheless, it has two types of outcome. One is the breakup of continental lithosphere and thus the creation of new oceanic lithosphere, with formation of seafloor spreading centers at mid-ocean ridges and backarc basins. The other is the formation of sedimentary basins around failed rift zones, where the rifting process does not continue to completion for continental breakup and rifts become inactive before the breakup (Ziegler and Cloetingh, 2004; Buck, 2015). In either case, continental rifting is temporally and spatially associated with lithospheric extension, which may result in orogenic collapse for the exposure of metamorphic core complexes. Over the past five decades a large variety of rift features have been recognized and explained with different methods and different concepts. These features include for example aborted rift structures, anomalous topography or anomalously high velocity/high density bodies located in the lower crust. Explanations for anomalous features often link one mechanism with one observed rift feature, for example, the role of dynamic processes in controlling development of continental rifting (Ziegler and Cloetingh, 2004), the role of the Moho in extensional settings (Cloetingh etal., 2013), and the effect of volcanism in continental rifting and supercontinental breakup (Franke, 2013). Simultaneously with thinning of the lithosphere along active rift zones, the asthenospheric material may migrate along the base of the thinned lithosphere and rise towards the deformed crust where it is attempting to break through. This migration only ceases when the material that is still at the base of the lithosphere (at depths between 200 km and 10 km) reaches thermal equilibrium. As such, one mantle anomaly can laterally flow over significant distances along the thinned orogenic lithosphere, after being risen to the base of the lithosphere. When the thermal equilibrium is reached, the mantle material glued to the base of the lithosphere at shallower depths corresponds geometrically and location-wise to high-velocity/high-density bodies. In general, lithospheric thinning may result from one of the following two distinct rift
mechanisms (Sengör and Burke, 1978; Turcotte and Emerman, 1983; Cloetingh and Wortel, 1986; Houseman and England, 1986). One is the horizontal extension of continental lithosphere, in which far-field stresses generated within, or at the boundaries of, the lithosphere creates the extensional field. This leads to passive rifts in which both the crust and the lithospheric mantle are simultaneously stretched right from the start of the rifting process. The other is the action of a heat source on the base of the lithosphere, resulting in thinning of the lithospheric mantle by thermal erosion in the early stage. Deformation, metamorphism and magmatism also take place in this stage, but the continental crust has been thinned considerably. This generates active rifts in which uplift, metamorphism and magmatism take place much more extensively in a late stage. The classification of rifts into active and passive ones is of dynamic nature as it is based on the forces that initiate rifting (Burke and Dewey, 1973). This classification provides a genetic link to the major plate tectonics with respect to their origin. An active rift is produced by an external force (e.g., upwelling of the asthenospheric mantle) below the lithosphere (Fig. 10a), where lithospheric thinning is achieved by local buoyancy forces. This is primarily caused by impinging of the heat mantle on the base of the lithosphere subsequent to the lithospheric thinning by foundering its mantle part. It also generates the tensional stresses able to cause stretching of the upper lithosphere but compressing of the lower lithosphere. In contrast, a passive rift results from lithospheric extension by regional stresses at lithospheric boundaries (Fig. 10b), where lithospheric thinning is produced by horizontal, in-plane far-field forces, possibly due to large-scale plate interactions, which initiate and drive lithospheric extension. The major differences in thinning mechanism between passive and active rifts are the direction and position of forces acting on the lithosphere, leading to a series of differences in the products of lithospheric thinning (Olsen, 1995; Ruppel, 1995; Ziegler and Cloetingh, 2004). In general, active rifts show the structure of axial asymmetry whereas passive rifts exhibit the structure
of
axial
symmetry.
Active
rifting
is
characterized
by
small-scale
deformation-associated metamorphism and magmatism in the early stage, and crustal doming and grabening as well as large-scale metamorphism and magmatism in the late stage. Conversely, passive rifting exhibits graben formation and lacustrine sedimentation in the early stage, followed by volcanism in a later stage. In addition, active and passive rifts often show the different sequence between volcanic and sedimentary rocks in the resulted basins (Sengör and Burke, 1978; Cloetingh and Wortel, 1986). Volcanism precedes sedimentation in active rift zones but succeeds sedimentation in passive rift zones. In either case, the mantle upwelling is a
response to thinning of the lithosphere either by vertical foundering of the previously thickened lithosphere or by lateral extension of the normal or cratonic lithosphere (Buck, 2015). The volcanic nature of rifted margins is usually considered as a function of mantle temperature and the structure of the lithosphere prior to rupture (Reston and Morgan, 2004). The passive margins of the North Atlantic Ocean show a gradual transition from the magma-poor property in Newfoundland–Iberia to the magma-rich property in East Greenland–Hatton Bank, which was ascribed to an increase in the mantle temperature from ∼1200C in the south to ∼1400C in the north (White et al., 1987). However, the rate of lithospheric thinning has a significant influence on the rate of conductive heat transfer from the underlying asthenosphere to the thinned lithosphere. Rapid detachment of the thickened lithospheric root by either breakoff or delamination would induce decompressional melting of the asthenosphere, giving rise to magma-rich margins with volcanic rocks of juvenile composition. In contrast, sluggish detachment through convective erosion of the asthenosphere would bring it to contact directly the base of the thinned lithosphere, leading to magma-poor margins with volcanic rocks of ancient composition. Therefore, the style of lithospheric thinning has the capacity to cause the great differences in the generation and composition of mafic melts along the rifting continental margins. In addition, the maximum temperature reached during Buchan-type metamorphism is primarily dictated by the extent of lithospheric thinning and thus the structure of rift zones. The evolution of rift tectonics is controlled by a number of factors and parameters such as plate divergent rates, coupling and decoupling between crustal and lithospheric configurations, and evolution of asthenospheric thermal anomalies and associated mantle dynamics (e.g., Huismans and Beaumont, 2003; Burov, 2007; Liao and Gerya, 2014). Because of the crustal and lithospheric thinning during continental rifting, the structure of rift zones is dictated by the extent of rheological coupling between different lithospheric layers. This is particularly so when rifting affects a pre-existing orogen. In orogenic settings, subduction and suture zones formed during the amalgamation of continents, providing the possibility of lithospheric scale reactivations by rifting (e.g., Dunbar and Sawyer, 1988; Sokoutis et al., 2007). Such conditions are common in extensional backarc positions, where asymmetric rifting reactivates suture zones and nappe contacts where an overthickened, hot and weak lithosphere tends to be detached (Le Pourhiet et al., 2004; Tirel et al., 2008; Huet et al., 2011). A given scale convective behavior can be created by the interaction between the asthenosphere and lithosphere beneath orogenic regions (Huismans et al., 2001; Burov, 2007; Cloetingh et al., 2013). Sinking of the detached lithospheric mantle is coupled with upwelling of the
asthenospheric mantle, and the thinned lithosphere is subject to upward heat and stress from the asthenosphere. Therefore, the thermal structure of rift zones is dictated by a combination between convective and conductive heat transport effects at the asthenosphere-lithosphere interface.
4.2 Rift metamorphism in orogens Although intracontinental settings are inactive away from modern subduction zones, it does not preclude the possibility that they were convergent plate margins in the geological history (Zheng et al., 2015). In other words, they are portion of ancient orogens. The HT to UHT metamorphic rocks are common in orogens of all ages (Fig. 6), but particularly those of Precambrian ones (Brown, 2007; Kelsey, 2008). However, in most (if not all) of these orogens, overall increases of pressure in metamorphic rocks are in much smaller magnitudes than those in temperature. Thus, heating is associated with thinning rather than thickening of the orogenic lithosphere. Some of these orogens show the evidence that prograde low-P facies metamorphism is later than the peak episode of continental thinning at variable timescales from <1 Myr to >10 Myr (Kelsey and Hand, 2015; Harley, 2016). In such areas, the thermal anomaly during the lithospheric thinning would have largely continued, whereas the syncompressional deformation and high-P facies metamorphism takes place at an early time. Usually, the heat source is hypothetically provided by mantle-derived mafic magmas emplaced by underplating at the crust–mantle boundary (e.g., Huppert and Sparks, 1988; Voshage et al., 1990). However, it is rarely observed that the hot to ultrahot orogens contain contemporaneous basalts resulting from decompressional melting of the asthenospheric mantle. Nevertheless, extremely weak lithospheres, and especially those consisting of a two-layer crust with Moho temperatures largely above 800 °C, develop distributed stretching strain in response to the upwelling of the asthenospheric mantle beneath the thinned lithosphere. Thermal erosion of the lithospheric mantle beneath the ultrahot orogens leads to a dramatic decrease of its viscosity that becomes negligible compared to that of the overlying crust and not significantly greater than that of the underlying asthenosphere. In this regard, the hot to ultrahot orogens are steady-state systems that achieve three-dimensional mass redistribution in different stages: (1) transpressive shortening during syn-convergent orogenies and lithospheric stacking; (2) upward, gravity-driven flow of crustal slices in subduction channels; (3) detachment of the thickened mantle lithosphere to result in the over-thinned mantle lithosphere beneath the orogens, (4) heating of the thinned lithosphere for HT to UHT metamorphism and bimodal magmatism. The fourth stage is associated with lithospheric extension at high geothermal
gradients. In particular, anticlockwise P–T paths for the products of extreme metamorphism require the external heat to be provided for crustal anatexis. Lithospheric extension in association with continental rifting was suggested to cause anomalously high base-crustal heat flow, crustal and upper mantle melting, an attenuated series of isograds and high-T/low-P facies metamorphism that is characterized by the occurrence of andalusite-sillimanite assemblages (Vielzeuf and Kornprobst, 1984; Weber, 1984; Wickham and Oxburgh, 1985). Very high metamorphic gradients of 80-100C/km were inferred for granulites in the Pyrenees (Wickham and Oxburgh, 1985). This mechanism was further documented by Sandiford and Powell (1986) in terms of evidence for granulite-facies metamorphism during continental rifting and the P-T paths associated with such metamorphism. Although this mechanism was consistent with measurements of high heat flow in modern continental rifts (Edwards et al., 1978; Lachenbruch, 1979; Morgan, 1982, 1983), it did not gain popularity in the past three decades. This is largely because the metamorphic effects of continental rifting on orogenic roots are not thoroughly understood yet. In many highly extended rifts on Earth, tectonic removal of the upper crust exhumes lower crust rocks, giving rise to metamorphic core complexes. The development of lithospheric thinning is actively compensated by mantle rocks of anomalously low density, as indicated by low seismic velocities (Abers et al., 2002). This is illustrated by the Woodlark Rift in Papua New Guinea, which is the only place on Earth today where continental rifting is transformed into seafloor spreading, low-angle normal faults are seismically active, and UHP eclogite-facies rocks of Pliocene age occur on the surface. A metamorphic core complex there has also exhumed blueschists above sea level at near plate tectonic dip slip rates since the late Pliocene in the continental rift zone between the Australia and Woodlark plates (Little et al., 2007). Nevertheless, there may be different mechanisms for exhumation of high-grade metamorphic grades in this area: the UHP rocks are exhumed through subduction reversal along the paleosubduction channel, whereas the metamorphic core complex is exposed through thinning of the overlying crust as a low-angle unroofing process. In the latter case, the rift tectonics allows the upper continental crust to accommodate tens of kilometers of lithospheric extension in response to thinning of the upper crust. It is difficult for the andalusite-sillimanite facies metamorphism to result solely from the lithospheric extension during passive rifting, because this process does not transfer sufficient heat from the asthenospheric mantle into the shallow crust for the high-T/low-P metamorphism. In contrast, lithospheric thinning after foundering of the orogenic root leads to
shallowing of the thermal boundary layer, allowing for high heat flow from the asthenosphere into the thinned lithosphere (Houseman et al., 1981; Platt and England, 1994). Prior to the lithospheric thinning, there is mainly lateral pressures from the laterally convective asthenosphere beneath the thickened lithosphere, resulting in deformation, metamorphism and anataxis at higher pressures. As soon as the orogenic root is removed, the asthenospheric mantle upwells to the base of the thinned lithosphere. This gives rise to uplifting pressures that exceed the lithostatic pressure, leading to extension of the thinned lithosphere in its upper part but compression in its lower part. While deformation is weak in this stage, high heat flow is conductively transferred into the lower crust of the thinned lithosphere for Buchan-type metamorphism. It also leads to crustal doming for the exposure of metamorphic core complexes in accretionary and collisional orogens. Local occurrences of high-T/low-P facies metamorphic rocks in metamorphic core complexes suggest that lithospheric thinning is heterogeneous beneath hot to ultrahot orogens, as results from the detachment model (Wernicke, 1985; Lister et al., 1986). The occurrence of HT to UHT granulite-facies rocks at ancient convergent plate margins indicates that high geothermal gradients were attained despite intracontinental settings. Very high metamorphic temperatures at shallow depths, for example, 800°C at 0.5 GPa, have been documented in many orogens. Lithospheric thinning is prominent in such orogens, where continental rifting was active and free water was produced from the underlying metamorphosed rocks. As a consequence, the high heat flow was acquired from the asthenospheric mantle along the thinned lithosphere beneath the hot to ultrahot orogens. Thus, active continental rifting is a potential mechanism for the HT to UHT metamorphism in the thinned orogenic lithosphere. Rift tectonics in continental regions is usually indicated by the occurrence of extensional basins and basaltic rocks with oceanic island basalts (OIB)-like geochemistry. It is rarely linked to regional metamorphism and granitic magmatism. On the other hand, HT to UHT metamorphic rocks are common in the accretionary and collisional orogens of intracontinental setting. Because the high heat flow acted on the base of the thinned lithosphere, the lower crust of orogenic lithosphere underwent dehydration melting at low pressures. In fact, rift tectonism takes place not only at convergent plate margins such as accretionary and collisional orogens but also at divergent plate margins such as mid-ocean ridge and backarc basin. Therefore, the rifting orogeny is a feasible mechanism for the development of HT to UHT metamorphism. As such, the active continental rift model has a great potential to explain the extreme metamorphism in hot to ultrahot orogens.
4.3 Active continental rift model Traditionally, orogeny has been associated with compressional tectonism, resulting in crustal thickening and foreshortening, low-angle faulting, folding, dynamic metamorphism, and calc-alkalic magmatism (Dewey, 1988). Plate subduction is generally considered as the cause of compressional regimes for accretionary and collisional orogenies. Substantially, the compressional regimes only prevail in the stage of coupling between subducting slab and mantle wedge, during which subduction zones have the lowest temperatures. The exhumation of HP to UHP metamorphic rocks indicates that the compressional regimes have been relaxed for subduction channel processes at elevated geothermal gradients (Zheng et al., 2013). Furthermore, operation of arc volcanism suggests heating of the mantle wedge base by the asthenospheric mantle in response to the initial rollback of the subducting slab (Zheng et al., 2016). At this stage, the tectonic regime in oceanic subduction zones have changed from compression to extension regime. The generation of backarc basins marks the climax of lithospheric thinning above the subduction zones (Dewey, 1988). If we view these barckarc basins as a kind of active rifts due to upwelling of the asthenospheric mantle, some of them have developed beneath continental arc terranes if rollback of the subducting slab is associated with trench retreat. This leads to reworking of the juvenile continental crust for magmatism and metamorphism. As such, the active rifts have superimposed on accretionary orogens. A corollary of this tectonic model is the rifting orogeny, during which tension prevails for mountain building in intracontinental orogens. In contrast to subduction orogeny, the thickened crust is thinned and extended along preexisting accretionary or collisional orogens. Low-angle faults are normal (younger-on-older) detachment surfaces rather than thrusts, resulting in crustal doming and the exposure of metamorphic core complexes. Folding is less prevalent than in subduction orogeny; such folds result from gravity sliding, detachment faulting, or magma intrusions. Regional metamorphism takes place at high temperatures but lower pressures, involving gneisses lineated parallel to the direction of maximum extension, not foliated normal to maximum compression. Calc-alkalic magmas are produced on an unprecedented scale; granitic batholiths are so shallow that their cupolas blistered and burst, sometimes forming hundreds of calderas during the great ignimbrite flareup. At first glance, the concept of rifting orogeny may seem paradoxical. Thinning of the lithosphere is expected to cause subsidence rather than uplift. In fact, regional subsidence does occur to result in rift basins. On the other hand, it also builds mountains at divergent plate boundaries such as mid-ocean ridges and backarc basins by basaltic volcanoes.
Nevertheless, basalts in continental rift zones generally show OIB-like trace element compositions rather than MORB-like ones. Although bimodal igneous rocks are often considered as the products of rift magmatism, felsic rocks are much more common than mafic rocks in continental regions. Other extensional mountains have been interpreted as resulting from the rise of metamorphic core complexes. In this case, the HT to UHT metamorphic rocks may be found along collapsed orogens. Therefore, an active continental rift model has the capacity to explain the HT to UHT metamorphism at a regional scale. This model is further constrained below in three ways. First, the HT to UHT metamorphism is contemporaneous with thinning of the thickened orogenic lithosphere. The thermal pulse from the underlying asthenospheric mantle is substantial to the lithospheric extension, different from the case in passive rifts. While the lithospheric thickening during the compression lowers the average T/P ratio, the foundering of the thickened orogenic lithosphere results in the thinning for rifting. Given the widespread occurrence of HT to UHT metamorphic rocks in orogens with a wide range of ages and tectonic settings, and the fact that commonly thick, mature sedimentary sequences, which may reflect substantial pre-rift subsidence due to subduction, are affected by this extreme metamorphism, it is very likely that the metamorphism was contemporaneous with active continental rifting. Second, the amount of heat added to the lithospheric base is large and continuous. Although the steady-state conductive geotherm is generally convex towards the temperature-axis, the high heat flow into the base of the orogenic lithosphere subsequent to convective thinning of the asthenospheric mantle may result in the multiplication of temperature on any given level. The non-steady-state geotherm is instantaneously attenuated such that it transects the kyanite-sillimanite stability field at the lower crust level and the andalusite-sillimanite stability field at the upper crust level. Because the continental rifting is not only instantaneous but also progressive, there is no thermal relaxation in the further heating stage of the thinned lithosphere. The strain rate is low during the rifting, and the thermal pulse is fast due to a high conductivity, effective heat transfer mechanisms, or energy gain through endothermic reactions. In addition, the metamorphic geotherm generally reflects peak T/P ratios. For the kyanite-sillimanite facies metamorphism, the temperature at the base of the lower crust is high during the active rifting; this is caused by the conductive heat transfer from the underlying asthenospheric mantle (Houseman and England, 1986; White et al., 1987; McKenzie and Bickle, 1988). Third, the HT to UHT metamorphism is generally associated with the relatively short
timescale for thermal pulse. The conductive pulse of a thermal perturbation in a system with a length-scale L and a thermal diffusivity k follows an exponential decay law, with a time constant of L2/2k (Carslaw and Jaeger, 1959). Thermal pulse of a homogeneously thinned lithosphere with L = 150 km and k = 10-6 m2/s, would have a time constant of 72 Myr. With L = 100 km and k = 10-6 m2/s, the time constant would only be 32 Myr. If the heat transfer rate in the lithosphere is increased due to shallowing of the thermal boundary layer, the thermal pulse would be even faster. Partial melting of the crust and the lithospheric mantle, and subsequent melt ascent, despite temporary increase in crustal temperatures, would reduce the timescale for thermal pulse even further. Nevertheless, the heat anomaly would be reduced as a result of the difference between the latent heat of melting and the latent heat of fusion, endothermic reactions and strain. The extreme metamorphism should therefore coincide with active rifting or postdate it by less than few tens of Myr. Prior to the rifting orogeny, subduction-thickened lithospheric roots become heated at first. This process is associated with deformation, metamorphism and anataxis at higher pressures. Temperatures are high at the bottom of orogenic lithospheres, being generally the order of 1000-1200°C. Subsequent to foundering of the orogenic roots, crustal anatexis develops more extensively in the areas of active continental rifting in response to upwelling of the asthenospheric mantle, where high heat flow from the asthenospheric mantle acts on the thinned lithosphere. As a consequence, the HT to UHT metamorphic rocks are generated in association with extensive felsic magmatism but minor mafic magmatism. Substantially, the partial melting of crustal rocks may happen if temperatures exceed the wet solidus of these rocks. The high temperature may be reached by orogenic collapse in response to lithospheric thinning, which can be induced by either passive rifting or active rifting. In either case, upwelling of the asthenospheric mantle follows foundering of the orogenic roots. The latter tectonism may proceed via one of the following three mechanisms: (1) delamination of the lower continental crust (Bird, 1979); (2) convective removal of the subduction-thickened lithosphere (Houseman et al., 1981); (3) breakoff of subducting lithosphere at the ocean-continent transition (Davies and von Blanckenburg, 1995). The delamination model implies that the asthenosphere comes in direct contact with the crust and it is therefore expected that abundant crustal melts occur in connection with the surface uplift (Bird, 1979). Mantle xenolith studies in several regions around the world show that the uppermost mantle lithosphere is significantly younger than the overlying crust, indicating that this process does happen under some conditions. Jull and Kelemen (2001) have even suggested that the lowermost (eclogitized) crust may delaminate together with the mantle
lithosphere. It is also likely that only the lower part of the cratonic mantle lithosphere is delaminated into the asthenospheric mantle. In either case, the lithospheric delamination is a fast process and thus results in decompressional melting of the asthenospheric mantle. The convective removal model assumes that the uppermost part of the mantle lithosphere is so viscous that its sinking rate is geologically irrelevant, despite its high density. The viscosity of the lowest part of the mantle lithosphere, on the other hand, approximates that of the asthenosphere. This part, where heat is being transported mainly by conduction (and therefore part of the thermally defined lithosphere), has a negative buoyancy and a viscosity comparable to that of the asthenosphere. This part of the mantle lithosphere may be removed from the rest of the mantle lithosphere by convective erosion in the surrounding asthenosphere and it may ultimately sink (Houseman et al., 1981; Sandiford, 1989; Platt and England, 1994; Houseman and Molnar, 1997; Molnar et al., 1998). As such, the asthenopsheric erosion may laterally remove the orogenic lithosphere at different depths. Nevertheless, this is a slow process and thus does not result in decompressional melting of the asthenospheric mantle. The slab breakoff model hypothesizes detachment of the subducting oceanic lithosphere from its followed continental lithosphere at mantle depths during continental collision. As soon as the oceanic slab is detached at the ocean-continent transition and sinks into the asthenospheric mantle, subduction of the continental lithosphere terminates and becomes exhuming wholly by buoyancy (Davies and von Blanckenburg, 1995). At the same time, rapid upwelling of the asthenosphere occurs along the locus of slab detachment, inducing not only rapid heating of the lithospheric base but also partial melting of the asthenospheric top. This results in high geothermal gradients in the lowest lithosphere on the one hand and the generation of basaltic melts with MORB-like composition on the other hand (von Blanckenburg and Davies, 1995). While the orogenic lithosphere is thickened during oceanic subduction and subsequent continental collision, its thinning is abruptly achieved by the slab breakoff. With regardless of which mechanism operates, a partially molten layer is produced to decouple the weak crust from the underlying mantle lithosphere, enabling the ascent of deeply buried crust (e.g., metamorphic core complex). In general, lithospheric mantle-derived mafic magmatism is common, but asthenospheric mantle-derived magmatism is absent if continental rifting did not run into rupture. In particular, the slab breakoff at mantle depths has the capacity to result in both UHP and UHT metamorphism at the same time. This is illustrated by UHP terranes like Kokchetav, Erzgebirge and Eastern Greenland, which record both peak UHT and UHP metamorphic conditions (Fig. 8). On the other hand, the
delamination may take place at a long time after the termination of either accretionary or collisional orogeny. Likewise, orogenic roots can be convectively removed by the asthenospheric erosion at a time considerably later than previous compressional orogenies. In this case, lithospheric thinning allows for shallowing of the thermal boundary layer due to the conductive heat transfer from the underlying asthenospheric mantle into the thinned lithopshere. As a consequence, early HP assemblages are overprinted by HT to UHT metamorphism at reduced pressures. Because the lithospheric thinning is associated with orogenic collapse, it leads to denudation of deep crust and thus crustal doming with diapiric uplift of metamorphic core complexes. In addition to the occurrence of UHT metamorphic rocks in accretionary orogens where UHP metamorphic rocks have not been identified (Brown, 2014; Kelsey and Hand, 2015), UHP metamorphic rocks are often overprinted by UHT metamorphism in collisional orogens (Kotková et al., 2011; Liu et al., 2015). Common examples include the Caledonides in Greenland, the Erzgebirge in Germany, the Kokchetav in Kazakhstan, the North Dabie in China, and the Rhodope in Greece. This suggests the tectonic switch in continental subduction zones, where the UHP metamorphic regime of compressional tectonism is transformed to the UHT metamorphic regime of extensional tectonism (Zheng and Chen, 2016). The HT to UHT metamorphism postdates the HP to UHP metamorphism, which is a common feature in the internal zones of the Variscides. This has been documented in the Bohemian Massif, the Schwarzwald, the Western Alps, the Spanish Central System (Díaz-Alvarado et al., 2012 and references therein), and the Ossa Morena Zone (Pereira et al., 2009). Except for the slab breakoff for the coeval UHP and UHT metamorphism (Fig. 8), the subduction of supracrustal rocks to lower crustal and even upper mantle depths for HP to UHP metamorphism and the HT to UHT superimposition of the HP to UHP metamorphic rocks may be caused by two independent, causally unrelated events. Available studies indicate that the timescale of UHP metamorphism in the coesite stability field may vary from <5 Myr to >15 Myr, depending on the size of UHP terranes (Zheng et al., 2009, 2013; Kylander-Clark et al., 2012). On the other hand, the timescale of UHT metamorphic events may vary from <10 Myr to >30 Myr (Kelsey and Hand, 2015; Harley, 2016), but the cause for the variation is unclear. Furthermore, time intervals between the UHP and UHT metamorphic events are variable in different orogens, depending on which mechanism operates for the asthenospheric upwelling. Almost no interval occurs in the case of slab breakoff, which results in coeval UHP and UHT metamorphism at mantle depths (Fig. 8). Long intervals are necessary if the previously thickened lithosphere is thinned for heating
by underplating of the asthenospheric mantle. Although the lithospheric thinning may be caused either by gravitational delamination or convective removal, the UHT event significantly lags behind the UHP event. In either case, a combined study of petrology and geochronology is necessary in order to decipher the plurificial metamorphism whereby the HP to UHP metamorphic rocks in collisional orogens were superimposed by the HT to UHT metamorphism at a later time. In view of the above arguments, we develop an active continental rift model for the petrogenesis of HT to UHT metamorphic rocks. Before continental rifting, lithospheric roots developed during accretionary to collisional orogenies. These roots show the negative buoyancy relative to their adjacent lithospheres (Fig. 11a). The convective removal of the subduction-thickened lithosphere may proceed at least in two stages (Fig. 11b): (1) orogenic roots are gradually foundered by convective erosion of the asthenospheric mantle. Lateral flow of the asthenospheric mantle induces the forced heating of orogenic roots, resulting in contemporaneous deformation, metamorphism and anatexis to small extent. At the end of the foundering of the orogenic root, the mantle part of the lithosphere is often thinner than at the start of the orogenic process; (2) the orogenic lithosphere undergoes slow heating due to upwelling of the asthenospheric mantle subsequent to the foundering. At the same time, uplifting pressures from the upwelling asthenosphere act on the base of the thinned lithosphere, where the upward pressures exceed the lithostatic pressure to cause crustal doming. Active continental rifts develop in this stage along preexisting orogenic belts, resulting in contemporaneous metamorphism and anatexis to large extent; deformation is weak to even absent in the metamorphic and magmatic rocks of this stage. After completion of the continental rifting, isostatic gravity drives lithospheric balance between the thinned lithospheric mantle and the upwelling asthenospheric mantle to result in orogenic collapse (Fig. 11c). Lithospheric rifts are sited, preferentially, along orogens that were thickened during accretionary or collisional orogeny. Because such orogens contain structural heterogeneities in continental lithosphere, they are susceptible to thinning by removing orogenic roots. This eventually results in orogenic collapse, which occurs especially where convective thinning of the thickened lithosphere leads to shallowing of the thermal boundary layer beneath the orogens. Therefore, the rifting orogeny offers a feasible mechanism not only for why oceans cyclically close and reopen in roughly the same places, but also for why large crustal thicknesses in accretionary and collisional orogens are returned to normal ones without very
much erosional denudation but with the widespread preservation of HT to UHT metamorphic assemblages and bimodal magmatic rocks.
5. Crustal anatexis during extreme metamorphism Petrological evidence for partial melting of crustal rocks is prominent in the evolution of orogens, but its origin is variable depending on anatectic regimes. Anatectic processes are recorded by metamorphic mineral parageneses that indicate decompression with increasing temperature for HP to UHP metamorphic rocks but near-isobaric heating for HT to UHT metamorphic rocks. Although regional melting during rifting orogeny may be dictated by a number of factors, the anatectic regime of near-isobaric heating is substantial to rifting metamorphism. Therefore, there are different types of anatectic reactions in the lower crust of mafic composition (Fig. 12). Partial melting of crustal rocks is closely associated with extreme metamorphism at elevated temperatures. It is much more common in the HT to UHT metamorphic rocks than in the HP to UHP metamorphic rocks. Dehydration and melting may be two sides of one process if the breakdown of hydrous minerals results in their own melting. Otherwise they are two coupled processes, liberating water for amphibolite-facies overprinting and hydration melting of the overlying crust, respectively, below and above the wet solidus of crustal rocks (Sawyer, 2010; Zheng and Chen, 2016). The thermodynamic parameters, such as lithostatic pressure, temperature, fluid activity, bulk compositions of crustal rocks that are partially melted, and so on, constitute critical factors that vary strongly according to tectonic settings (Zheng et al., 2011b; Weinberg and Hasalová, 2015). Their values and their variations promote or hinder the crustal anataxis during extreme metamorphism. During regional melting, furthermore, the transport of felsic melts out of the source region takes with water, which has consequences for the evolution of the lower crust and hydration reactions in the middle to upper crust once the transported water is exsolved on crystallization of the melts. Therefore, it is substantial to understanding the physicochemical mechanism of crustal anatexis if we want to decipher the geodynamic regime for regional metamorphism at extreme tectonic conditions. It is common for granulites to coexist with migmatitic and magmatic rocks in orogens (e.g., Schmitz and Bowring 2003; Hollis et al., 2004; Kemp et al. 2007; Pownall et al. 2014). However, the fact that, deep in the continental crust, either mafic or felsic rocks become partially melted, is not obvious in view of the geothermal gradients of continental lithosphere
(Zheng et al., 2016). On the other hand, field observations indicate that crustal rocks underwent partial melting, yielding migmatites and granulites (Vielzeuf et al., 1990; Sawyer et al., 2011). These rocks constitute a large part of the lower to middle crust of eroded orogens. The petrogenetic relationship between granulite, migmatite and granite has been the subject of hot debate in decades (Vielzeuf et al., 1990; Solar and Brown, 2001; Dostal et al., 2006; Korhonen et al., 2015; Wei, 2016; Wei and Zhu, 2016). It is common in nature that partial melts still reside in parental rocks without significant evolution (anatectic melts), or they have separated from parental rocks with considerable evolution (magmatic melts). The former results in migmatites whereas the latter left residues as granulites (Waters, 1988; Vielzeuf et al., 1990). The resulted migmatites and granulites are generally composed of residual, peritectic and anatectic minerals in different proportions, and the resulted granites are often composed of magmatic, peritectic and residual minerals in different proportions. Nevertheless, the granulites often contain small amounts of anatectic melts, and thus they are not completely composed of the residual minerals. Likewise, the granites often contain small amounts of residual and peritectic minerals, so that they are not completely composed of the magmatic minerals. Granulites are generally exposed in uplifted deep-crustal terranes and commonly contain segregations of feldspars and quartz, the components of crustal rocks that melt at the lowest temperature in the presence of H2O. The classical experiments of Tuttle and Bowen (1958) showed that melting temperatures of common quartzofeldspathic rocks are lowered by high H2O pressure to a temperature range that prevails in HT granulite-facies metamorphism. This led to the concept of granite formation as the culmination of regional metamorphism (Brown and Fyfe, 1970). Subsequent studies indicate that the temperature scale of crustal metamorphism
extends
from the
water-saturated solidi of ∼650–700C
to
the
water-undersaturated solidi of 750-900C. For nearly a century, studies of granite petrogenesis and crustal evolution have been guided by the concept of partial melting in the continental crust (Brown, 2013). However, the H2O is restricted in amount (or in its activity) to limit the extent of partial melting. The distribution of temperature and H 2O contents in felsic systems is only compatible with derivation of the magmas by fluid-absent partial melting reactions at high-temperature, granulite-facies conditions (Clemens and Watkins, 2001). On the other hand, fluid-present partial melting reactions are of widespread occurrence and critical importance in the processes of granulite-facies metamorphism and crustal differentiation (Weinberg and Hasalová, 2015). The processes of crustal melting and melt extraction were suggested by Fyfe (1973) for the granulite-granite connexion. Further studies from field observations (Phillips, 1980; Tracy
and Robinson, 1983; Waters and Whales, 1984) and thermodynamic modellings (Thompson, 1982; Powell, 1983; Grant, 1985) indicate that the partial melting of deep crustal rocks through the breakdown of hydrous minerals (muscovite, biotite, amphibole) and the subsequent extraction of granitic melts results in the geochemical differentiation of continental crust, thereby the removal of H2O and other hyperfusible components left a refractory granulite-facies lower crust. Petrological studies indicate that the temperature range of crustal melting, 750–900C, is attained in high-grade metamorphism. Studies of experimental petrology confirmed that partial melting of metatonalites yields granitic melts of metaluminous composition (Rutter and Wyllie, 1988) whereas partial melting of Al-rich metasediments produces granitic melts of peraluminous composition (Vielzeuf and Holloway, 1988), and that the residual minerals are mafic-aluminous silicates such as pyroxene, garnet, cordierite, and sillimanite characteristic of the granulites. This led to the dehydration melting model, which can explain certain trace-element signatures of the exposed high-grade terranes (e.g., the Lewisian complex in Scotland), including depletion, relative to upper crustal rocks of the same general major element compositions, of the LILE like Cs, Rb, Ba, and, to a lesser extent, K, and the radioactive elements U and Th (Heier, 1973; Pride and Muecke, 1980). Such processes are important in long-term survival of the continents in that it creates a refractory underpinning to continental masses (Collerson and Fryer, 1978). Substantially, a wide range of crustal rocks may undergo dehydration-driven incongruent melting reactions at temperatures of 850–1100 °C (Zheng et al., 2011b; Weinberg and Hasalová, 2015; Gao et al., 2016). The crustal rocks that had experienced granulite-facies metamorphism are dominated by nominally anhydrous minerals, leaving refractory restites with H2O deficiency in the lower continental crust (Aranovich et al., 2014). It is generally assumed that the H2O available comes entirely from micas and amphiboles initially present in metamorphic rocks of the lower continental crust. If the peritectic reaction of amphibole + plagioclase + quartz = orthopyroxene + melt is used to define the beginning of crustal anataxis for granulite-facies metamorphism, the lower P–T limits of the amphibolite-granulite facies boundary are placed around 0.8-0.9 GPa at ~800-850C. Felsic melts produced are markedly peraluminous compositions, which become less felsic and richer in total alkalis and alumina with increasing pressure (Weinberg and Hasalová, 2015; Gao et al., 2016). At relatively low pressures of <~1.2 GPa, HT potassic melts are produced with anorthosite as a restite, which changes to orthopyroxene-plagioclase granulite with increasing pressure. At intermediate pressures between ~1.2 and ~2 GPa, they are replaced by sodic melts with residual clinopyroxene–garnet granulite. At relatively higher pressures of >~2 GPa, alkalic and Al-rich
melts are in equilibrium with HP to UHP eclogite residues. Although the wet solidus temperature may be as low as ~650°C for granitic rocks (Huang and Wyllie, 1981) and ~700C for basaltic rocks (Lambert and Wyllie, 1972), the temperature of ~850°C is required for amphibole breakdown (Aranovich et al., 2014; Gao et al., 2016). Thus, the effective solidus of mafic rocks is close to ~850C (Wyllie and Wolf, 1993; Moyen et al., 2011). At temperatures below ~850°C, nevertheless, muscovite and biotite may break down for dehydration melting. At low pressures of <0.8 GPa, the dP/dT slope of the muscovite breakdown curve is rather flat. Only muscovite-rich metapelitic schists can undergo dehydration-melting
reactions
to
produce
peraluminous
leucogranitic
melts
with
garnet-plagioclase granulite as a restite. Many experiments have been devoted to partial melting of crustal rocks at crustal pressures of 0.5 to 0.8 GPa, but temperatures above 850C for such low-P melting cannot be reached by internal heat sources alone (Clarke et al., 2011); an external heat source is necessary for the HT to UHT metamorphism (Zheng and Chen, 2016). In the deeper lower crust, partial melting of any mafic rocks at temperatures below 850C requires an influx of aqueous solutions. Felsic melts produced in this manner are metaluminous and sodic trondhjemites. If the anatectic reaction of biotite + quartz + K-feldspar + H2O = orthopyroxene + melt is used to define the beginning of crustal anataxis, granulite-facies metamorphism is also marked by the appearance of orthopyroxene in felsic gneisses at lower P–T conditions of ~0.5 GPa and 650-750C (Newton et al., 2014). The bulk of field and experimental evidence favors the fluid-absent model for granulite-facies metamorphism through deep-crustal melting in the absence of active fluids (Vielzeuf et al., 1990; Clemens and Watkins, 2001). As such, granitic magmas, and their supposed residues, the granulites, are generally considered as the common products of fluid-absent melting processes in relatively closed systems. On the other hand, field evidence from various orogens where partial melting is inferred to have taken place, including granulite-facies terranes, favors the view that infiltration of exotic volatiles, including H 2O, was active for crustal anatexis (Aranovich et al., 2013, 2014). Available results from experimental granite-water systems suggest that granitic melts with 4–5 wt.% H2O can be produced by partial melting of the lower continental crust at granulite-facies temperatures of 850-900C (Aranovich et al., 2014). The temperatures would be lowered if there were high rates of H 2O influx in the lower crust to cause rehydration, giving rise to granitic melts with higher H 2O contents. Using large data sets of Ti-in-zircon thermometry of granites and experimental data of zircon saturation temperatures in granitic magmas, Collins et al. (2016) showed that the
temperature scale of granite genesis in accretionary orogens is shifted downward from previous estimates, suggesting the possible presence of greater H2O contents in the granitic melts. Therefore, it is evident that volatile components, especially H2O, were added to the deep crust for granitic magmatism in association with granulite-facies metamorphism. In general, water is necessary for the partial melting of crustal rocks (Campbell and Taylor, 1983; Whitney, 1988). The water is normally released from breakdown of hydrous minerals at elevated temperatures (Johannes and Holtz, 1996; Weinberg and Hasalová, 2015; Gao et al., 2016). Nevertheless, early studies of experimental petrology on crustal anatexis were performed in the presence of excess water (Fig. 13). Later studies utilized hydrous minerals as one of the starting minerals and found that the stability of hydrous minerals is a key to the P-T conditions of partial melting (Fig. 14). Furthermore, it is realized that anatectic melts are produced together with peritectic minerals by consuming reactive minerals through peritectic reactions (Fig. 15). This leads to water-present melting reactions (Fig. 13) and dehydration melting reactions (Fig. 15). Therefore, the partial melting of crustal rocks can be categorized into the following three types according to the mineral paragenesis of parental rocks (Fig. 12). (1) Water-excess anatexis (Fig. 12b). The parental rock is dominated by nominally anhydrous minerals with minor amounts of hydrous minerals. This type of rock characterizes metamorphic zones where fractures favor flow of water coming from outside and/or from neighboring dehydration reactions. Free water is added to the rock system and discrete fluid phases are trapped in voids between crystals. Partial melting is caused by infiltration of water into the rock and thus termed as hydration melting. The crustal anatexis takes place at the lowest temperatures at or close to the wet solidus via anatectic reactions, for instance, amphibole + biotite + K-feldspar + plagioclase + quartz + H2O = melt (curve 6 in Fig. 13). There is sufficient water to cause high-degree melting, leaving the lowest amount of residual minerals. The volume of anatectic melts is mostly a function of the quantity of water percolating in the metamorphic zone. This type has been used for a long time in experimental petrology of granite petrogenesis, mimicking the crustal anatexis in the uppermost part above active continental rifts. (2) Water-deficient anatexis (Fig. 12c). The parental rock contains hydrous minerals, but no free water is added to the rock system for partial melting. This type of rock is characteristic of high-grade metamorphic parageneses and shows the transition from amphibolite facies to granulite facies. Water is primarily bound within the hydrous minerals, the most important being micas and amphiboles. Crustal anatexis is associated with the breakdown of hydrous
minerals and thus usually termed as dehydration melting. This process takes place at temperatures above the wet solidus via anatectic reactions, for instance, amphibole + plagioclase = clinopyroxene + garnet + melt in mafic systems (Fig. 12a) but muscovite + plagioclase + quartz = K-feldspar + sillimanite + melt in felsic systems (curves 12 and 2 in Fig. 15). The volume of anatectic melts is a function of not only the abundance of hydrous minerals but also the temperature and duration of dehydration reaction. Although no free water is added to the rock, water released from the breakdown of hydrous minerals may be locally focused to hydrate wallrock, decreasing its solidus for hydration melting (Fig. 13). Nevertheless, there is no sufficient water present to cause high-degree melting. (3) Water-absent anatexis (Fig. 12d). The parental rock is all composed of nominally anhydrous minerals with no hydrous minerals, corresponding to felsic to mafic granulites. Although crustal anatexis takes place at the highest temperatures, it generally produces very low-degree melts that are alkaline, anhydrous and anorogenic granites (A-type). Such high temperatures are only reached at active continental rifts where the orogenic crust was heated to the largest extent and thus termed as heating melting. This process is closely associated with crustal collapse during and subsequent to foundering of the orogenic roots and thus with underplating of the asthenospheric mantle. In summary, the lower crust may experience two kinds of anatectic reaction with increasing temperature. The first is dehydration-melting reactions that occur at temperatures much higher than the wet solidus of crustal rocks. This transfers water from hydrous minerals to the overlying rocks with local focus of free fluid phases to cause hydration melting. The dehydration reactions commonly have positive dP/dT slopes at pressures equivalent to crustal depths. Thus, any pressure decrease (decompression) and/or temperature increase (heating) can result in the breakdown of hydrous minerals for dehydration melting, leaving granulite-facies residues whose mineral components do not dissolve significantly in felsic melts. The second is the hydration melting at temperatures at or close to the wet solidus due to addition of the free water from the underlying metamorphic rocks, leading to large-scale magmatism along the thinned lithosphere at the peak stage of active continental rifting. The results from the above reactions can be outlined by the asthenospheric heating model for the layered dehydration melting of orogenic crust (Zheng and Chen, 2016), which is schematically depicted in Fig. 16. Although the extraction of granitic melts is the dominant process by which continental crust is modified over time, the layered dehydration melting model does adequately account for some important aspects of common granites, migmatites and granulites (Fig. 16). In
addition to the high temperatures, significant amounts of water is also supplied by dehydration of the underlying crustal rocks at temperatures above the wet solidus of these rocks. This provides significant amounts of aqueous solutions and hydrous melts for hydration melting of the overlying rocks. In this regard, water-excess anataxis may have taken place due to influx of the external water from the underlying dehydration of the lowest crust upon the rifting metamorphism. As such, the dehydration-hydration melting model has the capacity to cause large-scale melting of crustal rocks in the thinned orogenic lithosphere. If the infiltration of free water occurs at temperatures below the wet solidus, Barrovian-type amphibolites are produced above the granulite level.
6. Geodynamic implications Orogeny is characterized by a distinctive relationship between structural deformation, regional metamorphism, basin sedimentation and felsic magmatism. It leads to mountain building with the exposure of metamorphic rocks and magmatic plutons to the surface either during the same cycle or different cycles. Modern orogenic belts are located at convergent plate boundaries where both crustal thickening and uplift are prominent. They occur either above oceanic subduction zones where magmatic arcs are accreting to continental margins or above continental subduction zones where two continental blocks collide. Ancient orogenic belts were also generated by accretionary and collisional orogenies in compressional tectonic settings, but they are generally reworked by later geological processes such as regional metamorphism due to active continental rifting. As a consequence, their previous records of compressional processes are often obscured or even obliterated by rifting metamorphism in extensional regimes. It is the rifting metamorphism that is responsible for continental differentiation through crustal anatexis and transfer of felsic magmas from the lower crust to the upper crust, leaving behind mafic restites. This results in the rifting orogens that are superimposed along previous accretionary and collisional orogens. It is highly possible that rifting orogeny is the rule in all orogens along ancient and modern plate margins, except some young orogens underneath which the subduction structure is still present (e.g., Western Alps). Regional metamorphism is characterized by both a wide areal extent and a regional-scale temperature distribution, which are shown by metamorphic mineral zones, independent of the distribution of individual plutonic masses (Harker, 1932; Dewey, 1988). In the term of modern geology, regional metamorphism is produced by stress and heat transfer from the lithosphere. The HP to UHP metamorphism results from much faster stress transfer than heat
transfer at low geothermal gradients, whereas the HT to UHT metamorphism is caused by much faster heat transfer than stress transfer at high geothermal gradients. While UHP metamorphic rocks unambiguously occur in continental subduction zones that have developed into collisional orogens, it is still uncertain which setting is responsible for the generation of UHT metamorphic rocks. Regional metamorphism associated with crustal extension has been widely recognized in collisional orogens, suggesting the role of extension in the generation of high-T/low-P metamorphic rocks (Peacock, 1991). Lithospheric geodynamics is primarily driven by heat released from the asthenospheric mantle. Heat transport, by conduction, advection or convection, determines the T attained at crustal depths in the lithosphere and, therefore, the type of regional metamorphism, whether partial melting occurs, and the lithospheric rheology. During orogeny, tectonic processes advect heat to shallower levels in association with deformation, metamorphism and magmatism of the lithosphere, and partial melting and melt transport redistribute incompatible elements in the orogenic lithosphere. These processes dictate the evolution and differentiation of continental crust (Rudnick, 1995; Taylor and McLennan, 1995), and produce granulites, migmatites and granites (Vielzeuf et al., 1990; Brown, 2010). There are both positive and negative feedback relationships between heat and anatexis and melt transport and crustal deformation. Orogens are geological systems far from thermodynamic equilibrium that exhibit a coherent arrangement of structural and tectonic features in space and time. As such, tectonic environments are the key to physicochemical mechanisms for the heat transfer from the asthenosphere to the lithosphere. There are generally four active tectonic environments in the lithosphere (Merle, 2011), which are physically defined by the kinematics of tectonic plate boundaries: (1) extensional zones, (2) transform zones, (3) subduction zones, and (4) orogenic belts. All of these tectonic environments correspond to lateral boundaries of tectonic plates, with the common development of subduction zones into orogenic belts. Extensional zones are typified by continental rifts and mid-ocean ridges, which are genetically linked to each other because some continental rifts are ultimately evolved into mid-ocean ridges. On the other hand, the three other active tectonic environments can all be responsible for the formation of a continental rift. These three tectonic environments are the three points of departure for any process of rifting, in the same way as the mid-ocean ridge is viewed as being the point of arrival. While the mid-ocean ridge is a central point toward which all rifts converge, the subduction zone is another central point toward which all plates converge. These two types of tectonism take place regardless of the tectonic setting that initiates them.
Transform zone-based rifts form along conservative plate boundaries, where the simplest tectonic environment occurs to initiate rifting. The extension is achieved by the formation of pull-apart basins along major strike-slip faults, either at releasing bends or at horsetail terminations, and the basins thus formed exhibit many different shapes. Extensional duplexes may also be observed along restraining and releasing bends or stepovers. When pull-apart basins form along major regional transform faults, the extension can continue to completion, with the formation of a new oceanic ridge. The classic relationship between subduction and rift is the generation of backarc basins (e.g., Ziegler and Cloetingh, 2004; Buck, 2015). This may be caused by aging of the subduction zone, to a slowdown of the convergence rate, or to a temporary plate divergence. In either case, rollback of subducting slabs is a primary cause for upwelling of the asthenospheric mantle to result in the formation of backarc basins (e.g., Buck, 2015; Menant et al., 2016). As such, backarc basins are just a kind of active rifts above the mantle wedge overlying subduction zones. Nevertheless, they may move from the normal backarc lithosphere to the thickened arc lithosphere, depending on their tectonic evolution with advance of slab rollback. Some of them develop beneath continental arc terranes, leading to reworking of the juvenile continental crust for magmatism and metamorphism. In this case, the active rifts have superimposed on accretionary orogens. Orogen-based rifts develop from preexisting subduction zones due to allochthonous terrane amalgamation, arc-continent, or continent-continent collision. In either case, lithospheric thickening was prominent there before rifting due to the presence of orogenic roots (Fig. 11a). Prior to the active continental rifting, orogenic roots are foundered into the underlying asthenospheric mantle due to the asthenospheric erosion (Fig. 11b). As soon as the thickened lithosphere is thinned, high heat flow is introduced from the underlying asthenospheric mantle to the base of the thinned lithosphere, resulting in HT to UHT metamorphism and bimodal magmatism. This tectonism is associated with orogenic collapse, crustal doming and exposure of metamorphic core complexes, with eventual return to its normal thickness (Fig. 11c). If the continental rifting runs into completion, lithospheric rupture results in seafloor spreading. This brings high heat flow to mid-ocean ridges for high-T/low-P metamorphism. Regional metamorphism associated with orogenies provides a mineralogical record that can be inverted to yield geothermal gradients during this process at convergent plate margins. Tectonic settings with low geothermal gradients are characteristic of cold subduction zones whereas those with high geothermal gradients are indicative of hot rifting zones. The duality
of thermal environments at ancient and modern convergent plate margins reflects the dual characters of plate tectonics. Furthermore, the duality of metamorphic belts is the characteristic imprint of plurificial metamorphism due to subducting in the low geothermal gradients at first and then rifting in the high geothermal gradients. Apparent geothermal gradients derived from inversion of age-constrained metamorphic P-T paths can be used to identify the tectonic settings of regional metamorphism and to evaluate the geodynamic regimes of intracontinental orogens. Once a collisional orogen suffered HT to UHT metamorphic overprinting, its P-T path may change from a clockwise one in the early stage of crustal thickening to an anticlockwise one in the late stage of crustal heating subsequent to lithospheric thinning. The orogen-parallel thinning of the lithosphere is the characteristic outcome prior to active continental rifting. Consequent petrotectonic units are generally of linear features and develop along orogenic belts with discrete grabens. They show variable extents of deformation, metamorphism and magmatism. This is primarily attributed to the difference in heat flow into their crustal base during continental rifting. The high heat flow into the lower crust is expected if the bulk mantle lithosphere is foundered into the asthenospheric mantle, whereas the low heat flow occurs if only its lower part is foundered. High heat flow may be only localized for limited high-T/low-P metamorphism on the one hand, it may be of regional scale for extensive anatexis of the lower crust on the other hand. Because the mechanism of continental rifting dictates the extent and geometry of lithospheric thinning, the distinction between active and passive rifts is critical to interpretation of the thermal evolution of thinned petrotectonic units. Axially asymmetric thinning is prominent in HT to UHT metamorphic terranes, suggesting that active continental rifts are more responsible for such metamorphism than passive rifts. Geophysical studies indicate that continental crust, like lithospheric mantle, is normally at the solid state (Stüwe, 2007). However, it does not mean that crustal anatexis cannot happen in the continental lithosphere. A great deal of studies have demonstrated that crustal anataxis is primarily induced by the breakdown of hydrous minerals (Zheng et al., 2011b; Weinberg and Hasalová, 2015; Gao et al., 2016). As such, the thermal structure of lithosphere is a key to the stability of hydrous minerals and thus to the onset of crustal anataxis (Zheng et al., 2016). While the partial melting of crustal rocks can be induced by various mechanisms, lithospheric thinning in response to upwelling of the asthenospheric mantle provides the high heat flow for HT to UHT metamorphism and its contemporaneous magmatism. As a consequence, the crustal anataxis may occur either as a protracted process or as a transient process in modern and ancient
plate margins. In view of the above arguments, we prefer to classify orogens into three types: accretionary, collisional, and rifting. Correspondingly, regional metamorphism is generalized into two categories: subduction and rifting. While the subduction of oceanic plate leads to accretionary orogens at continental margins and in intraoceanic settings, the subduction of continental plate results in collisional orogens at continental margins. Substantially, both accretionary and collisional orogens develop through tectonic compression at convergent plate margins. While the former forms at sites of continuing oceanic subduction without significant thickening of the oceanic crust itself, the latter forms at sites of subsequent continental subduction with significant thickening of the continental crust itself. In either case, the two types of orogens are usually considered as the products of subduction tectonism. However, it does not mean that they are generated in the compressional regime in the early stage of subduction. Substantially, they are induced by transient extension in a late stage of the subduction, which results in either partial melting of the hydrated mantle wedge for arc volcanism above oceanic subduction zones (Zheng et al., 2016) or exhumation of the HP to UHP metamorphic rocks inside continental subduction channels (Zheng et al., 2013). On the other hand, rifting orogens may occur at either divergent or convergent plate margins, and they develop at the same time as the lithospheric extension in response to asthenospheric upwelling subsequent to lithospheric thinning. Although they are initiated in intraplate settings, they may become seafloor spreading centers such as mid-ocean ridges and backarc basins if running into breakup; otherwise they are failed for intracontinental orogeny. Therefore, the rifting metamorphism marks termination of the compressional regime, providing a complement to the subduction metamorphism. As such, the rifting orogeny is independent of both accretionary and collisional orogenies. Nevertheless, it may take place due to slab breakoff and thus have a rapid transition in regional stress regime from compression to extension at mantle depths. Otherwise it happens at lower crust depths subsequent to the lithospheric thinning due to either the lithospheric delamination or the asthenospheric erosion in the post-subduction stage, postdating both accretionary and collisional orogenies at variable timescales. In either case, it may have superposed on the preexisting accretionary and collisional orogens. As such, the compressional tectonics in HT to UHT metamorphic terranes is inherited from the previous subduction tectonism, whereas the extensional tectonics is caused by underplating of the asthenopsheric mantle (active rifting). Because of the low geothermal gradients during subduction (Zheng et al., 2016), partial melting of the subducting crust rarely takes place in compressional regimes in the early stage of
accretionary and collisional orogenies. Instead, the crustal anatexis would occur considerably in extensional regimes. This may happen during exhumation, particularly when rollback of the subducting slab leads to the conductive heat transfer from laterally convective asthenosphere-derived material into the base of the accretionary wedge. Whereas the synexhumation anatexis occurs in a later stage of the accretionary and collisional orogenies, postorogenic anatexis also takes place in the extensional regime where active continental rifting develops along the previously thickened orogenic lithosphere subsequent to its thinning to the largest extent. In this case, high heat flow is conductively transferred into the base of the thinned lithosphere through underplating of the asthenospheric mantle, leading to the high geothermal gradients for crustal anatexis. In this regard, the UHT metamorphism is associated with the extensional regime rather than the compressional regime. As such, it marks rifting of the assembled continental block along the previous suture rather than amalgamating of the continental blocks through collisional orogeny. The first occurrence of Buchan-type HT to UHT metamorphic rocks is in the Archean (Fig. 6), signifying the geodynamic regime of plate tectonics in the early stage of the Earth’s evolution. The UHT metamorphism is dominantly a Precambrian phenomenon and thus inferred to have developed in active rift settings where intraoceanic subduction zones underwent lithospheric thinning via active rifting. In addition, HT metamorphism is also common in the Paleozoic Caledonides and Variscides of Europe, recording the rifting orogeny that developed on preexisting accretionary to collisional orogens. In comparison, HP to UHP metamorphic belts primarily occur in the Phanerozoic (Fig. 5), signifying a change in the geodynamic regime of plate tectonics. Although plate tectonics may have started from subduction of oceanic lithosphere in high geothermal gradients, subduction in low geothermal gradients possibly begins as early as the Neoproterozoic (Stern, 2005). This change registers the beginning of a modern plate tectonics regime that is characterized by cold apparent geothermal gradients and deep subduction of continental crust. In either case, both ancient and modern orogens were switched into active continental rifts in response to upwelling of the asthenospheric mantle, which take place subsequent to thinning of the subduction-thickened lithosphere. Similar to the Buchan-type HT to UHT metamorphic rocks, Barrovian-type HT to UHT metamorphic rocks have also occurred since the Archean (Figs. 5 and 6). Such uniform age distributions may record intraoceanic subduction and continental growth through accretionary orogeny in the geological history on the one hand and rapture of continental lithosphere into orogen-parallel rifts via active rifting on the other hand. In this regard, they indicate the
operation of plate tectonics as early as the Mesoarchean. In general, the P-T records of Barrovian facies series are associated with Buchan facies series in rifting orogens, resulting in segmented P-T paths with near-isothermal decompression and near-isobaric heating legs, paths that involve heating during decompression, and paths with looping structure (Fig. 7). In either case, rapid cooling is common for the three facies series of high-grade metamorphic rocks in Phanerozoic orogens, but it is less common for those in Mesoarchean to Mesoproterozoic orogens. This may be related to the difference in the thermal gradient of orogenic belts, which remains to be resolved by further studies. Nevertheless, the all HT to UHT metamorphic rocks postdate the Alpine facies series, resulting in partial inheritance of the former from the latter. However, Alpine-type HP to UHP metamorphic rocks primarily occur in the Phanerozoic (Fig. 5). The difference in age distribution between Alpine and Barrovian/Buchan facies series is prominent (Figs. 5 and 6). This may be caused by the difference in the thermal structure of subduction zones. It is possible that the ancient plate tectonics in the Archean may be dominated by crustal subduction to the forearc depths for HT/HP metamorphism at warm subduction zones, whereas the modern plate tectonics in the Phanerozoic is dominated by crustal subduction to the subarc depths for UHP metamorphism at cold subduction zones. The Neoproterozoic is a transitional period that records the change in the thermal regime from warm to cold ones. Although Archean subduction is characterized by high geothermal gradients, the HT to UHT metamorphic rocks at ancient convergent plate boundaries are not the product of subduction orogeny. Instead, they are the product of rifting orogeny. On the other hand, Phanerozoic subduction is dominated by low geothermal gradients, resulting in HP to UHP metamorphic rocks at collided continental margins. In either case, active continental rifts occur at convergent plate margins since the Archean. Therefore, the HT to UHT metamorphic rocks provide a petrological record of the rifting metamorphism and thus rifting orogeny subsequent to the lithospheric thinning in precedingly accretionary and collisional orogens.
7. Concluding remarks The plate tectonics provides a paradigm to understand the global tectonics of the lithosphere – the strong outer layer of Earth overlying the softer asthenosphere. The thermal structure of continental lithosphere determines the physicochemical behavior of regional metamorphism at convergent plate margins. Extreme metamorphism at a regional scale is usually referred to tectonic processes at UHP and UHT conditions, respectively. We have
extended their products to two facies series at low and high geothermal gradients, respectively, in order to decipher their genesis within the framework of plate tectonics. Alpine-type HP to UHP facies series is produced during lithospheric subducting at the low geothermal gradients, and the metamorphic products of subduction orogeny are characterized by clockwise P–T paths. In contrast, Barrovian/Buchan-type HT to UHT facies series is produced during active continental rifting at the high geothermal gradients, and the metamorphic products of rifting orogeny often show anticlockwise P-T paths. The abnormal P-T values are reached by extreme tectonism in specific settings due either to the rapid subduction of crustal rocks to mantle depths or the slow heating of the base of the thinned lithosphere by the upwelling asthenosphere. Continental rifts may be built on the basis of subduction zones, transform faults and orogenic belts on the one hand, and they develop into seafloor spreading such as mid-ocean ridges and backarc basins on the other hand. While continental rifting is a basic process for hot to ultrahot underplating of the asthenospheric mantle on the thinned orogenic lithosphere, it is also the initial and fundamental process by which the separation of a single continent into two tectonic plates takes place along preexisting orogens. As such, the tectonic switch between subduction and rifting has great bearing on the major plate tectonic processes at work on Earth. In terms of the tectonic evolution in continental rifts, the active rift is caused by upwelling of the asthenospheric mantle. This process may occur at sites where the lithosphere undergoes thinning, which may occur at backarc sites due to rollback of the subducting slab, or along orogenic belts due to slab breakoff, lithospheric delamination or asthenospheric erosion. In either case, active continental rifting in response to the asthenospheric upwelling leads to the HT to UHT metamorphism at lithospherically thinned sites. This results in lithospheric extension in its upper part but compression in its lower part, with extensive magmatism along the thinned orogens. The HT to UHT metamorphic processes in continental crust have been operative since the Archean. There, crustal anatexis occurs at active continental rifts, involving a series of dehydration melting reactions in the water-deficient environment of orogenic crust at different depths. They give rise to an abundant supply of aqueous solutions to the overlying rocks, leading to Barrovian-type amphibolite-facies metamorphism at temperatures below the wet solidus but hydration melting for migmatization at temperatures above the wet solidus. It is expected that the hydration melting is abundant in the middle continental crust of accretionary and collisional orogens, which have evolved into rifting orogens in intracontinental settings. The water added is primarily derived from dehydration of the deeper crustal rocks upon heating below from the upwelling asthenosphere. The quantity of water was sufficiently high and the
P-T conditions were suitable for extensive anatexis of the lower crust for hydration metamorphism and bimodal magmatism during active continental rifting. Regional metamorphism associated with different types of orogeny provides petrological records that may be inverted to yield apparent thermal gradients for different metamorphic belts, which in turn may be used to infer geodynamic regimes. While UHT metamorphic rocks have occurred since the Archean, UHP metamorphic rocks primarily occur in the Phanerozoic. The temporal difference between the two extreme metamorphic rocks may indicate a switch in the thermal regime of plate tectonics in the Neoproterozoic. The occurrence of UHT metamorphic rocks since the Archean suggests that the lithospheric thinning at the high geothermal gradients has operated very early in the geological history. In contrast, the occurrence of UHP metamorphic rocks mainly in the Phanerozoic indicates that the thermal regime of many subduction zones has changed to the low geothermal gradients since the Neoproterozoic for the cold subduction of oceanic and continental slabs. In order to unravel the tectonic history of an orogenic belt it is necessary to have sufficient knowledge on the change of pressure and temperature with time in regional metamorphic rocks, and on the relationship of these to the generation of metamorphic mineral assemblages. Information that enables us to address this issue potentially include the following eight aspects: (1) data on the temperature and pressure of prograde, peak and retrograde metamorphic conditions; (2) data on the nature and age of protolith lithologies; (3) data on the timing of prograde, peak and retrograde metamorphic conditions; (4) data on the P-T-t evolution of the metamorphic rocks during orogeny; (5) data on the temporal and spatial relationships between metamorphic mineral growth and fluid action, including the overprint of a previous metamorphic event by a later metamorphic event (HP granulite facies by UHP eclogite facies or vice visa); (6) data on the temporal and spatial relationships between crustal thickening and thinning to metamorphic conditions; (7) data on the timing of pre-, syn- and post-orogenic magmatic events; and (8) data on the ascent of metamorphic core complexes. Although a complete understanding of the evolutionary history of an orogenic belt requires much or all of this information, studies to date are rarely so complete. Nevertheless, our understanding of the relationship between metamorphism and tectonics has increased dramatically during the past two decades through the combination of different types of geological, geochemical, and geodynamic studies.
Acknowledgments This study was supported by funds from the Chinese Ministry of Science and Technology (2015CB856100) and the National Science Foundation of China (41590620). Thanks are due to Chunjing Wei, Wenjiao Xiao, Guochun Zhao and anonymous reviewer for their comments, which greatly helped the improvement of the presentation.
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Figure captions Figure 1. Phase diagram for regional metamorphic rocks in three facies series. The P-T conditions for the polymorph transition of SiO2 are after Bose and Ganguly (1995) and Al2SiO5 are after Pattison (1992), and those for the mineral reaction of Ab = Jd + Qz are after Holland (1980). Green lines denote the geothermal gradients at 5°C/km, 10C/km and 30°C/km, respectively.
Figure 2. Metamorphic facies according to the interests of petrology (a) and geodynamics (b). Shown is the P-T region of relevance to the lithosphere and fields are inserted semi-quantitatively (modified after Stüwe, 2007). The axes are the same in both diagrams but labeled in GPa for petrology (a) and in km for geodynamics (b). Bands delineating fields in (a) indicate that boundaries vary according to bulk composition of metamorphic lithologies. Red curve denotes the granite wet solidus. The normal geotherm in (b) is drawn as to pass about 500°C at 35 km depth and reach 1200°C in 120 km depth. Metamorphic facies abbreviations: LG = low grade; GS = greenschist facies; BS = blueschist facies; EC = eclogite facies; UHP = ultrahigh-pressure facies; GP = garnet pyroxenite facies; HGR = high-P granulite facies; GR = granulite facies; PA = plagioclase amphibolite facies; UHT = ultrahigh-temperature facies. Green lines denote the geothermal gradients at 5°C/km, 10C/km and 30°C/km, respectively.
Figure 3. Metamorphic facies and stability fields of hydrous minerals under subduction-zone conditions. The dark-green area denotes the UHP metamorphism above the coesite/quartz transition line, whereas the light-green area denotes the HP metamorphism below the coesite/quartz transition line. Dashed lines denote the geothermal gradients at 5°C/km, 10C/km and 30°C/km, respectively.
Figure 4. Plot of peak metamorphic P-T conditions for different metamorphic facies series (modified after Brown, 2014). Red circles denote the Buchan facies series, pink diamonds denote the Barrovian facies series, and blue squares denote the Alpine facies series. Also shown in (b) are geothermal gradients at 5°C/km, 10C/km, 30C/km and 50°C/km, respectively.
Figure 5. Plot of peak metamorphic pressures versus age for regional metamorphic belts (revised after Brown, 2014). Blue squares denote the Alpine facies series,.red circles denote the Buchan facies series, and pink diamonds denote the Barrovian facies series.
Figure 6. Plot of peak metamorphic temperature versus age for regional metamorphic belts (revised after Brown, 2014). Red circles denote the Buchan facies series, pink diamonds denote the Barrovian facies series, pale blue squares denote the Alpine HP facies and dark blue squares denote the Alpine UHP facies.
Figure 7. Schematic diagram illustrating the characteristic P-T-t paths of Alpine-type metamorphic slices in collisional orogens (revised after Zheng et al., 2011a). Note that the P-T-t paths follow the initial geotherm only during the early phase of crustal subduction for prograde metamorphism, but they become oblique to the geotherm during exhumation to form convex rightward loops. As a result, peak pressures are attained at t1 with metamorphic temperatures lower than peak temperatures at t2 and decreased pressures during the exhumation.
Figure 8. Pressure–temperature paths of Alpine-type UHP terranes in continental subduction zones (revised after Hermann and Rubatto, 2014). Small UHP terranes show decompressional cooling during exhumation, whereas large UHP terranes tend to exhibit near-isothermal decompression during exhumation. If UHP terrains suffered UHT overprinting at the maximum depth, they have the capacity to record both highest P and T conditions. Abbreviations for UHP terranes: LC, Lago di Cignana; WG, Western Gneiss Region; DB, Dabie; DM, Dora Maira; SL, Sulu; KK, Kokchetav; EG, Erzgebirge; GL, Eastern Greenland.
Figure 9. Schematic cartoons showing three types of orogeny with different types of regional metamorphism. (a) Accretionary orogeny due to subduction of oceanic crust, accompanied by arc volcanism and emplacement of accretionary wedge. (b) Collisional orogeny due to subduction of continental crust, accompanied by Alpine-type HP to UHP metamorphism. (c) Rifting orogeny due to lithospheric thinning in response to asthenospheric upwelling at either convergent or divergent plate margins. While HT to UHT metamorphism and bimodal magmatism take place at convergent plate margins, mid-ocean ridges or backarc basin mafic magmatism occurs at divergent plate margins.
Figure 10. Schematic diagram illustrating two end-member models for initiating mechanisms of continental rifting (modified after Olsen and Morgan, 1995).
Figure 11. Schematic diagram illustrating the convective removal of the mantle lithosphere in thickening orogens for active continental rifting (revised after Zheng and Chen, 2016). (a) Thickening of the lithosphere during accretionary or collisional orogeny (crust is light gray, mantle lithosphere is light green). (b) Foundering and thinning of the lithosphere by convective erosion of the asthenospheric mantle. (c) Balancing of the orogenic lithosphere with adjacent lithosphere subsequent to the thinning. The arrows indicate the direction of forces in different stages.
Figure 12. Different types of anatectic reactions in the crustal rocks of mafic composition (revised after Moyen, 2011). (a) The anatectic reactions are shown as lines, which are actually multivariant fields for most of them. Thick grey lines: water-excess reactions. Dot–dash lines: water-absent reactions. Other lines: other reactions (boundaries of the stability field of key minerals). Finer lines with reference: epidote reaction lines with SPD02 for Skjerlie and Patiño-Douce (2002) and VS01 for Vielzeuf and Schmidt (2001). The effect of water behavior (ignoring epidote) on the solidus position is further illustrated by three sketches of the P–T diagram in bottom: (b) water excess, with a solidus shifted to high temperatures above the plagioclase/amphibole outline; (c) water deficient, but in closed system such that the system becomes water excess when the amphibole breaks down; (d) water absent, in an open system where water is lost when hydrous phases break down. A refractory eclogite forms at a high pressure and does not melt below 1000°C or more.
Figure 13. Water-present melting reactions of crustal rocks as determined by experimental petrology (revised after Weinberg and Hasalová, 2015). Reaction curves: 1. H2O saturated Mus-granite solidus; 2. H2O saturated solidus in Qz+Or+Ab+H2O system, with the activity of H2O in melt to be 1 (2a), 0.7 (2b), 0.5 (2c), 0.3 (2d) and 0.1 (2e); 3. melting reaction Qz + Pl + Kfs + H2O = melt, with the activity of H2O in melt to be 1 (3a) and 0.1 (3b); 4. melting reaction Bt+Qtz+Kfs+H2O = melt; 5. H2O-saturated melting of tonalite Bt+Pl+Qtz+H2O = Hbl + melt; 6. tonalite H2O-saturated solidus; 7. wet basalt solidus; 8. H2O-saturated melting reaction Qtz+Kfs+Ab+H2O = melt; 9. H2O-saturated granite solidus (Qtz+Ab+Or+H2O = melt); 10. melting reaction (Bt+Pl+Qtz+Kfs+H2O = Hbl+Grt+ melt) of gneiss with 4 wt.% added H2O.
Figure 14. Schematic diagram illustrating comparison between water-present and absent melting reactions of crustal rocks as determined by experimental petrology (revised after Weinberg and Hasalová, 2015). Note different slopes and solidus for different types of anatectic reactions.
Figure 15. Dehydration melting reactions of crustal rocks as determined by experimental petrology (revised after Weinberg and Hasalová, 2015). Reaction curves: 1. muscovite dehydration melting of muscovite–biotite schist (1a) and muscovite schist (1b); 2. muscovite dehydration melting; 3. biotite dehydration melting of metagreywacke; 4. biotite dehydration melting of metagreywacke (a) and gneiss (b); 5. biotite dehydration melting in the MASH system; 6. Biotite dehydration melting of gneiss; 7. biotite dehydration melting of metapelite; 8. amphibole dehydration melting of amphibolite; 9. amphibole dehydration melting of basalt.
Figure 16. Schematic diagram illustrating the asthenospheric heating model for layered dehydration melting of orogenic crust at active continental rifts (abstracted from Zheng and Chen, 2016).
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Graphical abstract
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Highlights
Alpine-type HP to UHP metamorphism is associated with crustal subduction at low geothermal gradients;
Barrovian/Buchan-type HT to UHT metamorphism is associated lithospheric thinning at high geothermal gradients;
Metamorphic dehydration of crustal rocks is dominant in cold to ultracold subduction zones;
Hydration and dehydration melting of metamorphic rocks is significant in hot to ultrahot orogens;
Active continental rifting develops subsequent to thinning of subduction-thickened lithosphere.
