When the Continental
Crust Melts
Edward W. Sawyer1, Bernardo Cesare2 and Michael Brown3
1811-5209/11/0007-0229$2.50
P
DOI: 10.2113/gselements.7.4.229
Ga) continental crust appears to be
sl ig ht ly more felsic t ha n
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 modified in order to
b e come cont i nent a l c r ust.
Evidence from modern arcs indicates that more felsic compositions
arise because the mafic magmas
fractionate and because they cause
the crust to partially melt.
Consequently, a layer of mafic
cumulate and residual material
develops at the base of arc crust.
As the arc crust thickens, this
KEYWORDS : continental crust, partial melting, microstructures,
cumulate and residual part at the
metamorphic petrology
base converts to denser material,
detaches (a process called delamiINTRODUCTION
nation) and sinks into the mantle. Thus, the bulk composiThe continental crust is 41.4 km thick on average and
tion of the remaining continental crust becomes more
covers 39% of the Earth’s surface. Information from the
felsic. The residual and cumulate material that returns to
isotopic and trace element composition of >4-billion-yearthe mantle contains, and hence is enriched by, a small
old (Ga) zircon grains and the evolution of mantle isotopic proportion of felsic melt and becomes the Enriched Mantle
reservoirs indicates that 75%, and possibly more, of the I (EMI) isotopic reservoir (Tatsumi 2005).
continental crust was created before 2.5 Ga (Harrison 2009;
Belousova et al. 2010). Thus, the continental crust is much
EVIDENCE THAT THE CONTINENTAL CRUST
longer-lived than oceanic crust and, consequently, has
acquired considerable complexity. This is reflected in the PARTIALLY MELTED
At the beginning of the last century, extensive mapping
petrological and structural characteristics of the rocks
was done in the shield areas of Scandinavia, Canada and
within it.
elsewhere. This pioneering work revealed that large parts
The continental crust began to form in the Hadean, more of the continental crust have been metamorphosed to a
than 4.0 billion years ago, first as the mantle differentiated, higher degree and more strongly deformed than adjacent
then from thickened oceanic crust above “hotspots” and
areas. We now know that the structures in these highly
at shallow levels (~15 km) above convergent margins
deformed regions are similar to those in modern orogens
(Harrison 2009). Since the late Archean (from ca 2.8 Ga),
where continents have collided and that the metamorphic
most new, or juvenile, continental crust has formed in
temperature in these regions was high enough (> 700 oC)
magmatic arcs above subduction zones, but about 10% was
for large areas to partially melt. Some continental crust
formed where mantle magmas were added to existing crust has experienced repeated episodes of modification by
by hotspots or plumes. If new, juvenile continental crust
intense deformation, high-temperature metamorphism and
is formed from mantle magma in magmatic arcs and at partial melting: examples occur in the Grenville Province
hotspots or plumes, then its average composition should
of Canada, in southern West Greenland, in the Western
be mafic. It is not. The average composition of the contiGneisses of Norway and in East Africa. Different terms are
nental crust is broadly andesitic, although Archean (>2.5
used to describe this modification. It is simply called
reworking by petrologists and structural geologists, but from
a geochemist’s perspective, it is intracrustal differentiation.
1 Département des Sciences Appliquées,
The largest and most intensely reworked regions of contiUniversité du Québec à Chicoutimi
nental crust are located where continents collided and
Chicoutimi, Québec G7H 2B1, Canada
E-mail: [email protected]
major mountain chains were formed, for example, the East
African Orogen. Reworking is not restricted to thickened
2 Dipartimento di Geoscienze, Università di Padova
orogens. Mantle melts emplaced into the continental crust
Via Gradenigo 6, I-35131 Padova, Italy
E-mail: [email protected]
at rifts or in large igneous provinces associated with
hotspots can result in high-temperature metamorphism.
3 Department of Geology, University of Maryland
Partial melting in such settings can lead to intense, local
College Park, MD 20742-4211, USA
artial 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, influencing 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.
E-mail: [email protected]
E LEMENTS , V OL . 7,
PP.
229–234
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A UGUS T 2011
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 K 2O relative to the
lower part, which is more mafic 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 intracrustal 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
mafic and residual bulk composition (FIG. 1 AND 2). In addition 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 modification 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 CRUST
Rock types such as metapelite, metagreywacke and granite
may begin to partially melt when the metamorphic temperature 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 H 2O is present as a free fluid in
the pores and grain boundaries of the rock; this is called
H 2O fluid-present melting and takes place at the lowest
temperatures. Melting may also occur when hydrous
minerals (hydrates), such as muscovite, biotite and amphibole, melt incongruently (see glossary); other minerals,
most commonly quartz and feldspar, may also participate
in these melting reactions. Incongruent melting may be
either H 2O fluid-present or, at higher temperature, H 2O
fluid-absent. Crystalline rocks have very low porosity and
so contain very little fluid H 2O; thus the amount of melt
produced from H 2O 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 fluid-absent incongruent melting, except for
instances where large volumes of aqueous fluid were introduced into rocks already at high temperature, as discussed
below.
Schematic representation of the reworking of continental 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), first 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.
FIGURE 2
Sill and dike network in stromatic metatexite migmatite 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 1
E LEMENTS
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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 fluid-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 o C, 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
mafic rocks quite well. It explains both the volumes of melt
generated and the granulite facies, residual mineral assemblages found deep in the crust that are left behind after
melt has been extracted. However, it is not a good description 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 fluid infi ltrated them and led to what is called
water-fluxed melting at low temperature, around 700 oC.
Such an influx of H2O is now recognised as being responsible for melting of metapelitic, metapsammitic and
metamafic 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, therefore, that many of the places where water-fluxed 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 PROBLEM
The temperature required for H2O fluid–present or waterfluxed 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 orogenesis (FIG. 3). However, large granulite terranes that underwent 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 sufficient radiogenic 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 significant
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.”
E LEMENTS
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 conditions occur in the yellow field to its left, and partial melting can
occur in the pink field. 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 ultrahightemperature (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 flux at
the Moho (30 mW·m -2) are the same for both, but thermal conductivity 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.
FIGURE 3
PETROLOGICAL ASPECTS OF MELTING
THE CONTINENTAL CRUST
The rocks in the continental crust that have partially
melted are called migmatites; the nomenclature specific to
these rocks and the means by which they are identified in
the field 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 melanosome, 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 migmatite becomes attenuated, resulting in a banded or layered
appearance (FIG. 4) typically seen in the deep parts of orogens.
EXPERIMENTS AND PETROGENETIC
MODELLING
The pressure and temperature conditions retrieved from
granulites and migmatites tell us how deep in the continental crust melting occurred and provide minima that
must be achieved by any proposed mechanism of heating.
Basic information for determining the pressure and temperature (P–T) history comes from well-controlled experiments 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
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approaches to better understand the conditions and petrological 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 investigate the combination of parameters governing the development of orogens.
MELTED ROCKS UNDER THE MICROSCOPE
The 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, interfacial or surface energy. The extent to which microstructure
reaches the equilibrium state, often thought of as uniform
grain size and polygonal grain shapes, contains information 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 information 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 information about the crystallization of anatectic melt rather
than the melt-producing reactions. The correct identification of each petrological part of a migmatite is necessary
because each contains information about processes specific
to its origin.
A
Since leucosome cannot be considered as representative of
the initial melt composition, because of crystal fractionation and contamination for example, the chemical composition of quenched glass from melting experiments has
been the principal source of information on the composition of anatectic melts. This situation is changing: micronsized inclusions of glass and “nanogranite” (FIG. 5), believed
to be respectively quenched anatectic melt and its crystallization 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” compositions 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 contribution 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 MELTING
The 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 continental 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.
C
B
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
FIGURE 4
E LEMENTS
migmatite derived from metatonalite partially melted under granulite facies conditions in the Limpopo Mobile Belt, a deeply eroded
orogen. Penknife is 11 cm long. (C) Migmatite in which the garnetbearing 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.
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A UGUS T 2011
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 flow. Over the past two decades, advances in understanding 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 field.
MOVING THE MELT TO DIFFERENTIATE
THE CONTINENTAL CRUST
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 (“negative crystal”; see Cesare et al. 2011) and crystallized into a finegrained aggregate of quartz (Qtz), biotite (Bt), K-feldspar (Kfs),
apatite (Ap) and plagioclase (not visible in this image).
IMAGE COURTESY OF O MAR BARTOLI, U NIVERSITY OF PARMA , ITALY
FIGURE 5
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
REFERENCES
Belousova EA, Kostitsyn YA, Griffi n WL,
Begg GC, O’Reilly SY, Pearson NJ (2010)
The growth of the continental crust:
Constraints from zircon Hf-isotope
data. Lithos 119: 457-466
Berger A, Burri, T, Alt-Epping P, Engi M
(2008) Tectonically controlled fluid flow
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 flow 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
Granites are accumulations of anatectic melt, albeit melt
that has had its composition changed through contamination – by residuum (peritectic phases), wall rocks, or mixing
with different magmas – and through fractional crystallization. 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 concentrated into a more focussed flow 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 flow 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.
ACKNOWLEDGMENTS
Constructive 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 encouragement and help at all stages in the development of this
issue.
Holness MB (2008) Decoding migmatite
microstructures. In: Sawyer EW, Brown
M (eds) Working with Migmatites.
Mineralogical Association of Canada,
Short Course Volume 38, pp 57-76
Holness MB, Cesare B, Sawyer EW (2011)
Melted rocks under the microscope:
Microstructures and their interpretation. Elements 7: 247-252
Jamieson RA, Unsworth MJ, Harris NBW,
Rosenberg CL, Schulmann K (2011)
Crustal melting and the flow of mountains. Elements 7: 253-260
Jessup MJ, Cottle JM, Searle MP, Law RD,
Newell DL, Tracy RJ, Waters DJ (2008)
P-T-t-D paths of Everest Series schist,
Nepal. Journal of Metamorphic Geology
26: 717-739
Clark C, Fitzsimons ICW, Healy D, Harley
SL (2011) How does the continental
crust get really hot? Elements 7:
235-240
Rosenberg CL, Handy MR (2005)
Experimental deformation of partially
melted granite revisited: implications
for the continental crust. Journal of
Metamorphic Geology 23: 19-28
Guernina S, Sawyer EW (2003) Largescale melt-depletion in granulite
terranes: an example from the Archean
Ashuanipi Subprovince of Quebec.
Journal of Metamorphic Geology 21:
181-201
Rudnick RL, Gao S (2003) The composition of the continental crust. In:
Rudnick RL (ed) The Crust. Treatise on
Geochemistry 3, Elsevier-Pergamon,
Oxford, pp 1-64
Harrison TM (2009) The Hadean crust:
Evidence from >4 Ga zircons. Annual
Review of Earth and Planetary Sciences
37: 479-505
E LEMENTS
Sawyer EW (2008a) Atlas of Migmatites.
The Canadian Mineralogist Special
Publication 9, NRC Research Press,
Ottawa, Ontario, Canada, 371 pp
233
Sawyer EW (2008b) Identifying the parts
of migmatites in the field. In: Sawyer
EW, Brown M (eds) Working with
Migmatites. Mineralogical Association
of Canada, Short Course Volume 38, pp
29-36
Sawyer EW (2010) Migmatites formed by
water-fluxed partial melting of a leucogranodiorite protolith: Microstructures
in the residual rocks and source of the
fluid. Lithos 116: 273-286.
Tatsumi Y (2005) The subduction factory:
How it operates in the evolving Earth.
GSA Today 15: 4-10
Ward R, Stevens G, Kisters A (2008) Fluid
and deformation induced partial
melting and melt volumes in lowtemperature granulite-facies metasediments, Damara Belt, Namibia. Lithos
105: 253-271
White RW, Stevens G, Johnson TE (2011)
Is the crucible reproducible?
Reconciling melting experiments with
thermodynamic calculations. Elements
7: 241-246
Whittington AG, Hofmeister AM, Nabelek
PI (2009) Temperature-dependent
thermal diffusivity of the Earth’s crust
and implications for magmatism.
Nature 458: 319-321
A UGUS T 2011
GLOSSARY
processes such as fractional crystallization and
contamination may have modified its composition.
Anatectic front – The surface marking the beginning
of partial melting in the continental crust. It corresponds to the fi rst occurrence of neosome in the
direction of increasing metamorphic grade.
Melanosome – A type of residuum composed predomi-
nantly of dark-colored minerals, such as biotite,
garnet, cordierite, amphibole or pyroxene
Anatectic melt – A melt, generally granitic in composi-
tion, produced by anatexis
Metatexite – A type of migmatite in which coherent
pre–partial melting structures, such as bedding, foliation and folds, are preserved
Anatexis – Partial melting of the continental crust, irre-
spective of the degree of partial melting
Brittle–elastic fracturing – Open-mode fracturing by
Migmatite – A metamorphic rock formed by partial
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.
melting. At the outcrop scale migmatites are heterogeneous. In addition to two petrogenetically related
parts called leucosome and residuum, migmatites can
also contain rocks, called paleosome, which did not
melt.
Constrictional strain – Deformation resulting in
prolate fabrics in which linear structures dominate
over planar structures
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 undergone segregation.
Diatexite – A migmatite in which neosome dominates
and pre–partial melting structures (bedding, foliation, folds) have been destroyed and commonly
replaced by syn-anatectic flow structures
Orogenesis – The process of forming a mountain chain
in the Earth’s continental crust due to the convergence and collision of tectonic plates
Ductile fracturing – Fracturing due to creep and
growth of microscale voids—fi lled with either fluid
or melt in rock—that become interconnected leading
to rupture.
Paleosome – The non-neosome part of a migmatite that
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.
Peritectic mineral(s) – A new mineral (or minerals)
produced in addition to melt during incongruent partial
melting of a rock, mineral or mineral assemblage
Flattening strain – A deformation resulting in oblate
Protolith or parent rock – The rock from which the
fabrics in which planar structures dominate over
linear structures
Pseudosection – A map of phase assemblages for two
was not affected by partial melting because of its bulk
composition
neosome in a migmatite was derived
specified intensive and or/extensive variables (for
example, pressure and temperature) and a specified
bulk composition
Haplogranite system – A simplification 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.
Residuum – The solid fraction left in a migmatite after
partial melting and the extraction of some or all of
the melt
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-
Segregation – The overall process in which anatectic
melt is separated from the residuum in a migmatite
Solidus – The boundary separating the solid (± fluid)
phase assemblage fields (generally at lower temperature) from the melt-bearing phase fields (generally at
higher temperature) in a P–T phase diagram
gated partial melt. Leucosome does not necessarily
have the composition of an anatectic melt because
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
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