ELEMENTS, VOL. 7, PP. 229–234 AUGUST 2011229

1811-5209/11/0007-0229$2.50 DOI: 10.2113/gselements.7.4.229

When the Continental Crust Melts

INTRODUCTIONThe continental crust is 41.4 km thick on average and covers 39% of the Earth’s surface. Information from the isotopic and trace element composition of >4-billion-year-old (Ga) zircon grains and the evolution of mantle isotopic reservoirs indicates that 75%, and possibly more, of the continental crust was created before 2.5 Ga (Harrison 2009; Belousova et al. 2010). Thus, the continental crust is much longer-lived than oceanic crust and, consequently, has acquired considerable complexity. This is refl ected in the petrological and structural characteristics of the rocks within it.

The continental crust began to form in the Hadean, more than 4.0 billion years ago, fi rst as the mantle differentiated, then from thickened oceanic crust above “hotspots” and at shallow levels (~15 km) above convergent margins (Harrison 2009). Since the late Archean (from ca 2.8 Ga), most new, or juvenile, continental crust has formed in magmatic arcs above subduction zones, but about 10% was formed where mantle magmas were added to existing crust by hotspots or plumes. If new, juvenile continental crust is formed from mantle magma in magmatic arcs and at hotspots or plumes, then its average composition should be mafi c. It is not. The average composition of the conti-nental crust is broadly andesitic, although Archean (>2.5

Ga) continental crust appears to be sl ightly more felsic than Proterozoic (2.5–0.5 Ga) or Phanerozoic (<0.5 Ga) crust (Rudnick and Gao 2003). Thus, juvenile material added to the crust must be modifi ed in order to become continental crust. Evidence from modern arcs indi-cates that more felsic compositions arise because the mafi c magmas fractionate and because they cause the crust to partially melt. Consequently, a layer of mafi c cumulate and residual material develops at the base of arc crust. As the arc crust thickens, this cumulate and residual part at the base converts to denser material, detaches (a process called delami-

nation) and sinks into the mantle. Thus, the bulk composi-tion of the remaining continental crust becomes more felsic. The residual and cumulate material that returns to the mantle contains, and hence is enriched by, a small proportion of felsic melt and becomes the Enriched Mantle I (EMI) isotopic reservoir (Tatsumi 2005).

EVIDENCE THAT THE CONTINENTAL CRUST PARTIALLY MELTEDAt the beginning of the last century, extensive mapping was done in the shield areas of Scandinavia, Canada and elsewhere. This pioneering work revealed that large parts of the continental crust have been metamorphosed to a higher degree and more strongly deformed than adjacent areas. We now know that the structures in these highly deformed regions are similar to those in modern orogens where continents have collided and that the metamorphic temperature in these regions was high enough (> 700 oC) for large areas to partially melt. Some continental crust has experienced repeated episodes of modifi cation by intense deformation, high-temperature metamorphism and partial melting: examples occur in the Grenville Province of Canada, in southern West Greenland, in the Western Gneisses of Norway and in East Africa. Different terms are used to describe this modifi cation. It is simply called reworking by petrologists and structural geologists, but from a geochemist’s perspective, it is intracrustal differentiation. The largest and most intensely reworked regions of conti-nental crust are located where continents collided and major mountain chains were formed, for example, the East African Orogen. Reworking is not restricted to thickened orogens. Mantle melts emplaced into the continental crust at rifts or in large igneous provinces associated with hotspots can result in high-temperature metamorphism. Partial melting in such settings can lead to intense, local

Partial melting of the continental crust has long been of interest to petrologists as a small-scale phenomenon. Mineral assemblages in the cores of old, eroded mountain chains that formed where continents

collided show that the continental crust was buried deeply enough to have melted extensively. Geochemical, experimental, petrological and geodynamic modelling now show that when the continental crust melts the consequences are crustal-scale. The combination of melting and regional deformation is critical: the presence of melt on grain boundaries weakens rocks, and weak rocks deform faster, infl uencing the way mountain belts grow and how rifts propagate. Tectonic forces also drive the movement of melt out of the lower continental crust, resulting in an irreversible chemical differentiation of the crust.

KEYWORDS: continental crust, partial melting, microstructures, metamorphic petrology

Edward W. Sawyer1, Bernardo Cesare2 and Michael Brown3

1 Département des Sciences Appliquées, Université du Québec à ChicoutimiChicoutimi, Québec G7H 2B1, CanadaE-mail: ewsawyer@uqac.ca

2 Dipartimento di Geoscienze, Università di PadovaVia Gradenigo 6, I-35131 Padova, ItalyE-mail: bernardo.cesare@unipd.it

3 Department of Geology, University of MarylandCollege Park, MD 20742-4211, USAE-mail: mbrown@umd.edu

ELEMENTS AUGUST 2011230

reworking of the continental crust, but such thermal reworking is not generally accompanied by intense deformation.

The deformed and metamorphosed continental crust is not uniform. The upper part is approximately granodioritic in composition and is richer in SiO2 and K2O relative to the lower part, which is more mafi c and richer in Al2O3, FeO, MgO and CaO (Rudnick and Gao 2003). These differences as well as the considerable enrichment in light rare earth elements and the large negative Eu anomaly in the upper crust relative to the lower crust are best explained by partial melting, a process that is also called anatexis. Thus intra-crustal differentiation occurs by partial melting of the lower part of the continental crust and migration of the melt to the upper part, leaving the lower crust with a more mafi c and residual bulk composition (FIG. 1 AND 2). In addi-tion to these geochemical differences, this process imparts a layered structure to the continental crust, which is revealed by an increase in seismic P- and S-wave velocities with depth. Seismic profi les across young continental crust affected by late Paleozoic collision and mountain building in western Europe show the same sub-horizontal Moho and internal velocity structure as old crust in northern Europe that was reworked by mountain building events in the Proterozoic and Archean. Thus, the acquisition of a sub-horizontal layered structured must happen soon after mountains stop growing. This same basic pattern of modi-fi cation to continental crust has been going on since the late Archean, at least.

The geochemical approach has revealed that the large-scale process of intracrustal differentiation occurs by partial melting, but it does not address other concerns, such as

the source of the heat for melting, what happens at the grain scale during anatexis, or how felsic melt moves from the lower to the upper crust. Nor is it concerned with the broader consequences of partial melting, such as its effect on the rheology of the continental crust and how this affects the way mountain chains are built when continents collide. These and other questions are the subject of this issue of Elements on the theme “When the Continental Crust Melts.”

TYPES OF MELTING IN THE CONTINENTAL CRUSTRock types such as metapelite, metagreywacke and granite may begin to partially melt when the metamorphic temper-ature exceeds 650 oC (FIG. 3), and the melt they produce is granitic in composition. Whether they melt or not and the quantity of melt produced depend on the availability of H2O. Melting may occur if H2O is present as a free fl uid in the pores and grain boundaries of the rock; this is called H2O fl uid-present melting and takes place at the lowest temperatures. Melting may also occur when hydrous minerals (hydrates), such as muscovite, biotite and amphi-bole, melt incongruently (see glossary); other minerals, most commonly quartz and feldspar, may also participate in these melting reactions. Incongruent melting may be either H2O fl uid-present or, at higher temperature, H2O fl uid-absent. Crystalline rocks have very low porosity and so contain very little fl uid H2O; thus the amount of melt produced from H2O in the pores is too small to be easily detected. Consequently, the production of large volumes of granitic melt in continental crust is widely thought to occur by fl uid-absent incongruent melting, except for instances where large volumes of aqueous fl uid were intro-duced into rocks already at high temperature, as discussed below.

FIGURE 1 Sill and dike network in stromatic metatexite migma-tite at Maigetter Peak (height 480m) in the Fosdick

Mountains of West Antarctica (76°26’38”S, 146°30’00”W). The image is looking to the SE and was taken from the air (Twin Otter wing tip in upper right). From the aerial perspective and also upon close examination in outcrops, intersecting dikes do not appear to truncate or displace each other; the sills and dikes of granite crosscut foliation but may be continuous with or discordant to leucosomes in the migmatite. The leucosomes contain peritectic garnet and cordierite (see Figure 1 in Brown et al. this issue).

FIGURE 2 Schematic representation of the reworking of conti-nental crust by partial melting. Partial melting occurs

in the lower part of the crust where temperatures exceed the solidus and migmatites are formed (brown). Melt is formed on grain boundaries but segregates from the residual solids along a progressively more focussed pathway (shown in red), fi rst through leucosomes then dykes. The melt collects to form plutons, typically at the transition from ductile middle crust (yellow) to brittle upper crust (green); some felsic lavas may be erupted. It is not yet clear whether melt ascent is uninterrupted or whether melt ponds at intermediate levels, shown by the question marks. The ascent of some melt ends in the middle crust as dyke complexes, without forming plutons.

ELEMENTS AUGUST 2011231

Pelitic rocks contain a large amount of muscovite and biotite – 30 to 50 vol% is not unusual – and will produce melt progressively as the temperature rises above the temperatures of the incongruent melting reactions involving these minerals, typically ~720 oC and ~820 oC, respectively. Other rock types also undergo fl uid-absent incongruent melting. Metagreywackes and meta-andesites begin to melt between 750 oC and 800 oC. Amphibolites follow at about 850 oC, but they produce melt of tonalitic composition. Fluid-absent incongruent melting of micas in metapelites and metagreywackes can produce as much as 50 vol% melt. After all the mica is consumed at about 925 oC, the rate of melt production decreases, and the composition of the melt is no longer granitic.

Fluid-absent incongruent melting of micas and amphibole describes the melting of metapelite, metagreywacke and mafi c rocks quite well. It explains both the volumes of melt generated and the granulite facies, residual mineral assem-blages found deep in the crust that are left behind after melt has been extracted. However, it is not a good descrip-tion of melting in hydrate-poor quartzofeldspathic rocks, such as leucocratic granites, trondhjemites and tonalites. Recent studies in metamorphic terranes, ranging in age from Archean to Phanerozoic, show far higher degrees of partial melting in granitic rocks than can be accounted for by H2O in pores or by the breakdown of their mica and amphibole. Melting in these rocks occurred because an aqueous fl uid infi ltrated them and led to what is called water-fl uxed melting at low temperature, around 700 oC. Such an infl ux of H2O is now recognised as being respon-sible for melting of metapelitic, metapsammitic and metamafi c rocks in some anatectic terranes (Ward et al. 2008; Berger et al. 2008). Oxygen stable isotope studies reveal diverse sources for this H2O. In some terranes it came from dehydration reactions in nearby metapelites or from crystallizing plutons, whereas in others it originated as deeply penetrating seawater or meteoric water, and in yet others it came from the mantle. It is not surprising, there-fore, that many of the places where water-fl uxed melting has occurred in the continental crust are adjacent to major crustal-scale shear zones that provided the pathways for the H2O to infi ltrate the continental crust (Sawyer 2010).

THE HEAT PROBLEMThe temperature required for H2O fl uid–present or water-fl uxed melting (700 oC) might be reached as a result of mantle heat entering the base of the crust and radiogenic heat generated in a continental crust thickened by orogen-esis (FIG. 3). However, large granulite terranes that under-went melting at temperatures well above 850 oC and appear to have lost substantial volumes (>600,000 km3 for the Ashuanipi subprovince in Quebec; Guernina and Sawyer 2003) of granitic melt as determined from the composition of their residual rocks are problematic in that they required a great deal of heat. The average continental crust does not contain enough K, Th and U to produce suffi cient radio-genic heat to sustain this degree of melting on the required timescale. Other sources of heat are required. The mantle is an obvious source, and strain heating may be signifi cant in some circumstances. New measurements (Whittington et al. 2009) indicate that the thermal diffusivity of rocks at high temperature is low; consequently, the middle and lower crust may retain heat better than previously thought. Identifying the source of heat and the combination of parameters or circumstances required to focus the heat into thickening crust and produce a high degree of partial melting remains a major problem. Hence, the article by Clark et al. (2011 this issue) is the starting point for “When the Continental Crust Melts.”

PETROLOGICAL ASPECTS OF MELTING THE CONTINENTAL CRUSTThe rocks in the continental crust that have partially melted are called migmatites; the nomenclature specifi c to these rocks and the means by which they are identifi ed in the fi eld are outlined by Sawyer (2008a, b). Migmatites are basically simple rocks with two components. One, which is partially melted, is called neosome, and consists of the crystallized products from the melt and the complementary residual material. The second, called paleosome, consists of rock that did not melt. In most cases, however, the melt and residual solid have segregated from each other, although not completely. The neosome then consists of two petrologically different parts, one derived from the melt and called leucosome, and the other derived from the residual solid material and, if dark coloured, called melano-some, otherwise simply residue. In most cases this simple petrological framework is made morphologically complex by deformation during the melting process. Deformation results in the translation, rotation and distortion of the constituents parts. If the strain is high enough, the migma-tite becomes attenuated, resulting in a banded or layered appearance (FIG. 4) typically seen in the deep parts of orogens.

EXPERIMENTS AND PETROGENETIC MODELLINGThe pressure and temperature conditions retrieved from granulites and migmatites tell us how deep in the conti-nental crust melting occurred and provide minima that must be achieved by any proposed mechanism of heating. Basic information for determining the pressure and temper-ature (P–T) history comes from well-controlled experi-ments on the partial melting of rocks such as pelite, greywacke and amphibolite. Phase equilibria modelling using internally consistent thermodynamic datasets derived from experiments has now been added to the set of tools available for understanding the P–T conditions for partial melting in the continental crust. The article by White et al. (2011 this issue) compares the results from both

FIGURE 3 Types of melting in P–T space for continental crust thickened to 71 km. The base of average (41.4 km)

crust is shown by the blue dashed line. The red curve is the H2O-present solidus in the haplogranite system; subsolidus condi-tions occur in the yellow fi eld to its left, and partial melting can occur in the pink fi eld. Fields for melting by hydrate breakdown are shown: blue for muscovite (Ms), brown for biotite (Bt) and green for amphibole (Amp). The purple line marks the start of ultrahigh-temperature (UHT) metamorphism. Two equilibrium geotherms for crust of normal thickness are shown as dotted black lines. Crustal radiogenic heat production (0.61 µW·m-3) and a mantle heat fl ux at the Moho (30 mW·m-2) are the same for both, but thermal conduc-tivity is 3.0 W·m-1·K-1 for geotherm A and 2.0 for B; hence geotherm B is hotter but still does not reach UT conditions.

ELEMENTS AUGUST 2011232

approaches to better understand the conditions and petro-logical processes that occur when the continental crust melts.

Dating the time of formation of metamorphic minerals and adding this time constraint to P–T information results in a P–T–t trajectory, which charts the movement of rocks through the continental crust. These trajectories provide a powerful tool for testing numerical models that investi-gate the combination of parameters governing the develop-ment of orogens.

MELTED ROCKS UNDER THE MICROSCOPEThe microstructure in rocks continually readjusts to changes in conditions. Minerals disappear, new ones grow, and grain boundaries move, driven by the need to reduce energy (e.g. Holness 2008), whether that is lattice, inter-facial or surface energy. The extent to which microstructure reaches the equilibrium state, often thought of as uniform grain size and polygonal grain shapes, contains informa-tion on driving forces and the kinetics of grain-boundary migration. These factors could be related to such diverse and interesting parameters as the cooling and deformation histories of the rocks. The type of microstructural informa-tion sought must be matched to the rock sampled. It is fruitless, for example, to attempt to understand the melting reactions or mineral–melt equilibration microstructures by examining the paleosome, since it did not melt. Similarly, the microstructure of a leucosome contains infor-mation about the crystallization of anatectic melt rather than the melt-producing reactions. The correct identifi ca-tion of each petrological part of a migmatite is necessary because each contains information about processes specifi c to its origin.

Since leucosome cannot be considered as representative of the initial melt composition, because of crystal fraction-ation and contamination for example, the chemical compo-sition of quenched glass from melting experiments has been the principal source of information on the composi-tion of anatectic melts. This situation is changing: micron-sized inclusions of glass and “nanogranite” (FIG. 5), believed to be respectively quenched anatectic melt and its crystal-lization products, have been found in minerals from migmatite terranes (Cesare et al. 2011). These inclusions could provide the major, trace and isotopic compositions of natural anatectic melts; such “starting-point” composi-tions are required to understand what changes occur to anatectic melts in the crust. How can anatectic melt remain as glass in slowly cooled rocks from deep in the continental crust? This and other questions are addressed in the contri-bution by Holness et al. (2011 this issue), which outlines what recent studies of the microstructure in partially melted rocks tell us about the processes that occur when the continental crust melts and subsequently cools.

TECTONIC AND GEODYNAMIC IMPLICATIONS OF PARTIAL MELTINGThe onset of partial melting has a profound effect on the continental crust. The types of structures that form change and strain rates increase when the temperature of the conti-nental crust passes the solidus temperature. Because anatectic melt is less dense and less viscous than either the protolith or the solid residue, it is more mobile than the solid fraction and will separate from it. Buoyancy is a driving force, but differential stress acting on an inevitably anisotropic crust induces pressure gradients, and these constitute another, locally stronger, driving force for the movement of melt. Differential stress in anisotropic rocks results in the formation of many different types of dilatant structures, the space between boudins being one well-known example. Melt migrates to and collects in these structures.

The transfer of heat in the continental crust is largely by the slow process of conduction, so the deep parts of the crust are slow to heat up and slow to cool. Consequently, metamorphic temperatures can remain above the solidus (650 oC) for times as long as 30 million years, e.g. in the Himalayan–Tibetan system. In that period melt can move from one set of dilatant structures to the next as the crust progressively deforms, crystallizing partially in each and creating a complex network of leucosomes.

FIGURE 4 Examples of partially melted rocks. (A) Migmatite derived from pelite and psammite protoliths,

Nemiscau subprovince, Quebec. The lightest-coloured parts are leucosome and the darkest parts, rich in biotite and conspicuous red garnet, are residual material; together these are the neosome. The medium-grey-coloured part is a psammite that did not partially melt; it is paleosome. Scale is 15 cm long. (B) Highly strained

migmatite derived from metatonalite partially melted under granu-lite facies conditions in the Limpopo Mobile Belt, a deeply eroded orogen. Penknife is 11 cm long. (C) Migmatite in which the garnet-bearing neosomes have been highly strained, creating a banded or layered structure typical of shear zones developed in melt-bearing rocks. Scale is 15 cm long.

A

B C

ELEMENTS AUGUST 2011233

Approximately 80% of grain boundaries have melt on them when the melt reaches ~7 vol%, and this results in a loss of about 80% of the pre-melting strength of the protolith (Rosenberg and Handy 2005). Rocks become very weak long before melting advances enough (~26 vol%) to turn them into magma, i.e. a suspension of crystals in melt. The onset of melting and the weakening it causes have a profound effect on the rheology of the continental crust, on the way it deforms and on how orogens develop. The location of weak rocks is controlled by where the heat source is and by the rate at which hot rocks and cold rocks are moved to advect heat and mass. These factors are controlled in part by isostasy, by the development of a ductile root at the bottom of the continental crust and by erosion at the top of it. A weak region in the crust is

produced when and where rocks become hot and melt. Strain and advected heat may be focussed into a narrow zone between a reverse-sense shear zone at the bottom and a normal-sense one at the top, in a phenomenon called channel fl ow. Over the past two decades, advances in under-standing these topics and other tectonic and geodynamic consequences of “When the Continental Crust Melts” have occurred through the use of highly sophisticated numerical models, and the article by Jamieson et al. (2011 this issue) presents the state of the art in this critical fi eld.

MOVING THE MELT TO DIFFERENTIATETHE CONTINENTAL CRUSTGranites are accumulations of anatectic melt, albeit melt that has had its composition changed through contamina-tion – by residuum (peritectic phases), wall rocks, or mixing with different magmas – and through fractional crystalliza-tion. Melting takes place deep (>25 km) in the continental crust. However, most plutons of granite are emplaced in its upper part, mostly at depths of 12 to 15 km where the transition from ductile to brittle rheology occurs (FIG. 2). To accomplish the differentiation of the continental crust, anatectic melt must migrate from the grain boundaries where it was formed and become progressively concen-trated into a more focussed fl ow pattern. Thus, the melt is able to traverse rocks that are at subsolidus temperatures in the middle crust without freezing as dykes. In other words the fl ow of granite melt must become organised. How this happens “When the Continental Crust Melts” is discussed by Brown et al. (2011 this issue) in the fi nal article.

ACKNOWLEDGMENTSConstructive reviews and comments by principal editor Hap McSween and reviewers Tracy Rushmer, Nick Petford and Gary Stevens have greatly improved this contribution. On behalf of all the contributors we would like to express our collective thanks to Pierrette Tremblay for her encour-agement and help at all stages in the development of this issue.

FIGURE 5 Backscattered electron image of a “nanogranite” derived from a small (6 µm) inclusion of granitic melt

trapped in a garnet (Grt) crystal from a migmatite at Ronda (Spain). The melt inclusion has a typical polyhedric shape (“nega-tive crystal”; see Cesare et al. 2011) and crystallized into a fi ne-grained aggregate of quartz (Qtz), biotite (Bt), K-feldspar (Kfs), apatite (Ap) and plagioclase (not visible in this image). IMAGE COURTESY OF OMAR BARTOLI, UNIVERSITY OF PARMA, ITALY

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Berger A, Burri, T, Alt-Epping P, Engi M (2008) Tectonically controlled fl uid fl ow and water-assisted melting in the middle crust: An example from the Central Alps. Lithos 102: 598-615

Brown M, Korhonen FJ, Siddoway CS (2011) Organizing melt fl ow through the crust. Elements 7: 261-266

Cesare B, Acosta-Vigil A, Ferrero S, Bartoli O (2011) Melt inclusions in migmatites and granulites. Journal of the Virtual Explorer 40: paper 2

Clark C, Fitzsimons ICW, Healy D, Harley SL (2011) How does the continental crust get really hot? Elements 7: 235-240

Guernina S, Sawyer EW (2003) Large-scale melt-depletion in granulite terranes: an example from the Archean Ashuanipi Subprovince of Quebec. Journal of Metamorphic Geology 21: 181-201

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GLOSSARY

Anatectic front – The surface marking the beginning of partial melting in the continental crust. It corre-sponds to the fi rst occurrence of neosome in the direction of increasing metamorphic grade.

Anatectic melt – A melt, generally granitic in composi-tion, produced by anatexis

Anatexis – Partial melting of the continental crust, irre-spective of the degree of partial melting

Brittle–elastic fracturing – Open-mode fracturing by crack propagation normal to the direction of minimum compression. It occurs when stresses at the crack tips equal fracture toughness, or when reduced stresses lead to subcritical crack growth.

Constrictional strain – Deformation resulting in prolate fabrics in which linear structures dominate over planar structures

Diatexite – A migmatite in which neosome dominates and pre–partial melting structures (bedding, folia-tion, folds) have been destroyed and commonly replaced by syn-anatectic fl ow structures

Ductile fracturing – Fracturing due to creep and growth of microscale voids—fi lled with either fl uid or melt in rock—that become interconnected leading to rupture.

Ductile-to-brittle transition zone – The depth in the Earth’s crust where the brittle strength equals the ductile strength. It occurs in the range of 12 to 18 km.

Flattening strain – A deformation resulting in oblate fabrics in which planar structures dominate over linear structures

Haplogranite system – A simplifi cation of the composi-tion of granite to just albite + orthoclase + quartz + H2O components (the Ab–Or–Qz system). Adding an anorthite component creates the haplogranodiorite system.

Incongruent melting – The process by which partial melting of a rock, mineral or mineral assemblage produces one or more new (peritectic) minerals, in addition to melt

Leucosome – The part of a migmatite derived from segre-gated partial melt. Leucosome does not necessarily have the composition of an anatectic melt because

processes such as fractional crystallization and contamination may have modifi ed its composition.

Melanosome – A type of residuum composed predomi-nantly of dark-colored minerals, such as biotite, garnet, cordierite, amphibole or pyroxene

Metatexite – A type of migmatite in which coherent pre–partial melting structures, such as bedding, folia-tion and folds, are preserved

Migmatite – A metamorphic rock formed by partial melting. At the outcrop scale migmatites are hetero-geneous. In addition to two petrogenetically related parts called leucosome and residuum, migmatites can also contain rocks, called paleosome, which did not melt.

Neosome – The part of a migmatite formed by partial melting and consisting of melt-derived and residual fractions. The neosome may, or may not, have under-gone segregation.

Orogenesis – The process of forming a mountain chain in the Earth’s continental crust due to the conver-gence and collision of tectonic plates

Paleosome – The non-neosome part of a migmatite that was not affected by partial melting because of its bulk composition

Peritectic mineral(s) – A new mineral (or minerals) produced in addition to melt during incongruent partial melting of a rock, mineral or mineral assemblage

Protolith or parent rock – The rock from which the neosome in a migmatite was derived

Pseudosection – A map of phase assemblages for two specifi ed intensive and or/extensive variables (for example, pressure and temperature) and a specifi ed bulk composition

Residuum – The solid fraction left in a migmatite after partial melting and the extraction of some or all of the melt

Segregation – The overall process in which anatectic melt is separated from the residuum in a migmatite

Solidus – The boundary separating the solid (± fl uid) phase assemblage fi elds (generally at lower tempera-ture) from the melt-bearing phase fi elds (generally at higher temperature) in a P–T phase diagram

Stromatic migmatite – A type of metatexite migmatite in which the leucosome and melanosome, or just the leucosome, occur as laterally continuous, parallel layers called stroma, which are commonly oriented along the compositional layering or the foliation

Supercontinent – A large continental landmass created from the collision of several continental cores or cratons

Ultrahigh-temperature (UHT) metamorphism – Metamorphism that occurred at temperatures above 900 oC and pressures compatible with the stability of sillimanite