JOURNAL OF PETROLOGY
VOLUME 49
NUMBER 7
PAGES 1343^1363
2008
doi:10.1093/petrology/egn028
On the Pseudomorphing of Melt-filled Pores
During the Crystallization of Migmatites
MARIAN B. HOLNESS1* AND EDWARD W. SAWYER2
1
DEPARTMENT OF EARTH SCIENCES, UNIVERSITY OF CAMBRIDGE, DOWNING STREET, CAMBRIDGE CB2 3EQ, UK
SCIENCES DE LA TERRE, DE¤PARTEMENT DES SCIENCES APPLIQUE¤ES, UNIVERSITE¤ DU QUE¤BEC A' CHICOUTIMI,
2
CHICOUTIMI, QUE¤BEC G7H 2B1, CANADA
RECEIVED SEPTEMBER 13, 2007; ACCEPTED MAY 1, 2008
ADVANCE ACCESS PUBLICATION MAY 30, 2008
Whereas it is relatively easy to map melt distribution
in migmatites once the melt has amalgamated into
centimetre-scale pockets or leucosomes (e.g. Brown, 2007),
it is not always straightforward to infer the former presence
of melt in rocks that have either undergone very small
degrees of partial melting or in which the melt has been
almost completely expelled. The grain-scale liquid distribution in melt-poor crustal rocks has commonly been
inferred from the presence of highly cuspate grains with
low dihedral angles (Fig. 1a). Their shape and the low dihedral angles are reminiscent of melt-filled pores in experimental charges (e.g. Jurewicz & Watson, 1985) and it has
been suggested that the cuspate grains nucleated within,
and infilled, melt-filled porosity (Platten, 1982, 1983;
Pattison & Harte, 1988; Harte et al., 1991; Riller et al., 1996;
Holness & Clemens, 1999; Sawyer, 1999, 2001; Rosenberg &
Riller, 2000; Marchildon & Brown, 2002). Such grains have
been termed interserts (Rosenberg & Riller, 2000), pseudomorphs after melt (Marchildon & Brown, 2002) or melt
pseudomorphs (e.g. Harte et al., 1993; Clemens & Holness,
2000; Brown, 2001; Walte et al., 2005). The process of pseudomorphing of melt-filled porosity and the conditions
under which it occurs are not well understood. Why do
some interstitial grains form perfect pseudomorphs of
a melt-filled pore (clinopyroxene in the gabbroic example
shown in Fig. 1b and c), with no overgrowth of the pore
walls, whereas other examples show the (multiply saturated) liquid being replaced by a poly-phase intergrowth
(clinopyroxene and plagioclase in the gabbroic example
shown in Fig. 1d)?
The mineral forming the pseudomorph is commonly different from that of the pore walls (Fig. 1a^c). In quartzofeldspathic crustal migmatites, it is feldspar that forms the
pseudomorphs in quartz-dominated volumes, whereas in
*Corresponding author. E-mail: [email protected]
ß The Author 2008. Published by Oxford University Press. All
rights reserved. For Permissions, please e-mail: journals.permissions@
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Pseudomorphs of melt-filled pores, recognized by their generally cuspate shape, are used as diagnostic for the former presence of partial
melt. They are commonly observed in migmatites from the mid- to
deep crust although they occur in the smaller pores in migmatites
from shallower levels (1^2 kbar). The pseudomorphing of meltfilled pores is controlled by the kinetics of nucleation and is a consequence of the greater supersaturation required for nucleation in a
small pore compared with a larger one. We examine three migmatites
in detail: a contact metamorphosed cherty band from an iron formation; an Archaean regional granulite from an accretionary prism;
and an amphibolite-facies sample from the roots of an Archaean
mountain chain. The greater undercooling required for nucleation in
progressively smaller pores is recorded by the composition of plagioclase pseudomorphs. A study of dihedral angles at the corners of
pseudomorphed pores demonstrates that melt^solid textural equilibrium was probably attained only in the contact aureole.The regional
granulite preserves an almost unmodified reaction-controlled
melt distribution, with little evidence for either melt^solid textural
equilibration or solid^solid re-equilibration, whereas the reactioncontrolled melt distribution in the regional amphibolite-facies example has been modified by a partial approach to solid^solid textural
equilibrium. It is not clear whether the differences in dihedral angle
population are due to differences in uplift and exhumation rates or
due to the presence of H2O on grain boundaries.
KEY WORDS:
migmatite; microstructure; dihedral angle; crystallisation
I N T RO D U C T I O N
JOURNAL OF PETROLOGY
VOLUME 49
NUMBER 7
JULY 2008
Fig. 1. (a) Photomicrograph under crossed polars of a migmatite from the Opatica Subprovince, Quebec. Note the elongate cuspate grains of
microcline on grain boundaries. Scale bar represents 200 mm. (b) Gabbro nodule from Iceland in which glass-filled pores (black) between
plagioclase grains are partially filled with clinopyroxene. (Note how the clinopyroxene inherits the angle at the pore corners.) Scale bar represents 100 mm. (c) Gabbro nodule from Iceland in which thick grain boundary melt films (now glass) are replaced by clinopyroxene. Note the
apparent absence of any plagioclase overgrowth on the pore walls in both this and in (b). Scale bar represents 200 mm. (d) Gabbro nodule from
Iceland, in which large melt-filled pores formed between plagioclase grains are partially filled by plumose intergrowths of clinopyroxene and
plagioclase. Scale bar represents 1mm.
feldspar-dominated volumes it is quartz (e.g. Harte et al.,
1991; Riller et al., 1996; Rosenberg & Riller, 2000). Petrographic studies of migmatites most commonly describe
quartz, alkali and plagioclase feldspar pseudomorphs (e.g.
Sawyer, 2001; Marchildon & Brown, 2002), although other
minerals (such as cordierite and biotite) may also form
pseudomorphs. Uncommonly the pseudomorph is formed
of the same mineral as one or both of the walls, although
it may have a different composition (e.g. plagioclase,
Sawyer, 2001).
The composition of single pseudomorphs is unlikely to
represent the composition of the last liquid in that pore:
even if other components of the liquid solidified on
the walls of the pore itself, some of the liquid must have
solidified elsewhere, forming pseudomorphs themselves or
overgrowths on existing grains. The formation of pseudomorphs therefore necessitates mass transport through the
pore network on at least the grain scale.
If averaged over a centimetre length scale, the composition of all pseudomorphs approximates that of the final
liquid. Table 1 gives four examples of melt compositions
estimated from point counting melt pseudomorphs within
single thin-sections of meta-sedimentary migmatites; these
are all of broadly granitic composition. Overgrowth of restite grains depletes the bulk composition of the pseudomorphs in the phase that forms the dominant restite
phase. The data from the two granulite-facies metagreywackes from Ashuanipi in Table 1 show a low modal
Table 1: Composition of pseudomorphed melt pockets,
determined by point counting 1000 points per section
Duluth aureole pelite
B1-384-20
Ashuanipi metagreywacke
Melt fraction
Quartz
Plagioclase
K-feldspar
(%)
(%)
(%)
(%)
64
39
35
26
251
45
24
31
54
37
22
41
42
37
26
37
BL86-091A
Ashuanipi metagreywacke
BL86-044A
Source: Sawyer (2001).
proportion of plagioclase, probably because plagioclase is
the major residual phase in both these rocks: it is easy to
underestimate the amount of overgrowth on restitic plagioclase, especially if the overgrowths are thin or in optical
continuity with their substrate.
Developing an understanding of the pseudomorphing of
melt-filled porosity necessitates the pseudomorphs being
set into the context of the microstructural development of
migmatites. Migmatite microstructures themselves pose a
considerable interpretative problem, requiring not only an
understanding of the controls on the grain-scale distribution of melt, both during and after reaction, and an
1344
HOLNESS & SAWYER
PSEUDOMORPHING OF MELT-FILLED PORES
appreciation of how and why textures formed during
solidification may vary, but also an understanding of how
these solidification textures may be modified in the subsolidus. We thus first set the scene by providing an overview
of melt distribution in partially melted rocks, together with
a discussion of the role textural equilibration may play
in modifying these. A review of published observations is
then used to demonstrate the joint controls on crystallization textures of cooling rate and pore size: we will show
that pseudomorphs form in the smallest pores in the slowest cooled rocks.
We then present detailed analysis of partially melted rocks
from a contact aureole and two regionally metamorphosed
migmatite terranes using petrographic observations, conventional geochemical analysis and cathodoluminescence
(CL) imaging to reveal successive stages of microstructural evolution during melting and solidification.
We demonstrate how the details of pseudomorph geometry can provide information on the balance of reaction
and textural equilibration during anatexis, and qualitative constraints on the rates of cooling after the metamorphic peak.
A N A LY T I C A L T E C H N I Q U E S
A N D C HOIC E OF SAM PLES
Plagioclase compositions and the titanium content of
quartz were determined using wavelength-dispersive electron microprobe spectrometry at the Department of Earth
Sciences, University of Cambridge, using a Cameca SX100
electron probe microanalyser. For the plagioclase analyses,
operating conditions were an accelerating potential of
15 keV, with a nominal beam current of 10 nA for Na, Al,
Si and Ca, and 100 nA for Mg, Sr, Ti, Mn and Fe, and
beam diameter of 5 mm. The quartz analyses used an accelerating voltage of 25 keV, with a nominal beam current of
300 nA and a beam diameter of 5 mm. The detection limit
was 7 ppm Ti. The PAP corrections procedure was used for
data processing and was run using Cameca Peak Sight
software.
Recent work coupling trace element concentration with
CL in quartz has demonstrated a strong link between concentration of Ti and brightness (Wark & Spear, 2005),
which can further be exploited using the TitaniQ thermometer (Wark & Watson, 2006; Wiebe et al., 2007). CL and
back-scatter electron (BSE) images were obtained using a
Jeol JSM-820 scanning microscope fitted with a Gatan
MonoCL. Polished thin-sections were scanned at a large
working distance (37 mm) with an accelerating voltage of
15 kV, and a beam current of 10 nA. The BSE aperture was
centred while the CL detector was off-centreçthis permitted BSE and CL images to be obtained over a wide
area (3 mm 3 mm) with minimal disruption. The presence of the CL detector created some astigmatism of the
Fig. 2. CL spectra obtained for quartz from the Biwabik iron formation sample ELY-6B. The locations of the spots from which these
spectra were obtained are shown in Fig. 10.
beam but the effect on the final images at low magnification was not significant.
CL spectra were obtained from four quartz generations
discernible by variations in brightness, using a beam current of 30 nA and a dwell time of 2 s. Spectra were
recorded with the Gatan MonoCL spectrograph, with a
paraboloidal mirror for light collection. The spectra are
shown in Fig. 2, with the detector background omitted for
clarity. The darkest, most weakly luminescent, quartz has
four clearly defined peaks at 320, 430, 580 and 660 nm.
All four peaks are present in more strongly luminescent
quartz but the 430 nm peak is dominant. The intensity of
the 430 nm peak, and thus the brightness of the quartz
under CL, correlates with Ti concentration, consistent
with the conclusions of Wark & Spear (2005) and Wiebe
et al. (2007).
In a texturally equilibrated poly-mineralic rock in which
grains have no preferred crystallographic orientation,
dihedral angles form a narrow population with a spread
about the mean determined by the extent of anisotropy of
interfacial energies (Herring, 1951; Laporte & Provost,
2000; Holness, 2006). When the microstructure is out of
equilibrium, the dihedral angle population may no longer
be unimodal, and the standard deviation is likely to be
large (Holness et al., 2005b). Although it is possible to constrain the extent of textural disequilibrium from an analysis of apparent angles measured in two dimensions using a
conventional microscope stage (Jurewicz & Jurewicz, 1986;
Elliott et al., 1997), a direct picture of the population of
true, three-dimensional (3-D), dihedral angles can be
gained using a universal stage mounted on an optical
microscope (Vernon, 1997). As the details of the population
of true angles are essential to interpreting disequilibrium
microstructures, we measured dihedral angles using a universal stage mounted on a J. Swift monocular optical
microscope at a magnification of 320. The error on each
measurement is of the order of 28, and populations of
90^162 angles were measured for each sample. The universal stage was also used to identify and obtain a qualitative
1345
JOURNAL OF PETROLOGY
VOLUME 49
picture of the distribution of grain boundary melt films,
which become difficult to see if they lie at a low angle to
the plane of a conventional microscope stage.
M E LT D I S T R I B U T I O N O N
THE GRAIN SCALE
Because of the long time scales involved it is impossible to
find a regionally metamorphosed anatectic rock preserving
unmodified microstructures dating from the early stages
of melting: primary textures are invariably overprinted.
Access to the earliest stages of reaction, during which the
melt distribution and geometry are totally controlled by
the kinetics of reaction, is generally possible only via
laboratory experiments and by examination of natural
examples of contact metamorphism in which time scales
of the metamorphic event were short. These show that
melting results in the development of progressively thicker
melt films separating reactant grains (Fig. 3a; Butler, 1961;
Mehnert et al., 1973; Bu«sch et al., 1974; Paquet & Franc ois,
1980; Platten, 1982, 1983; Maddock, 1986; Holness,
1999; Cesare, 2000; Holness & Watt, 2002; Holness et al.,
2005a; Acosta-Vigil et al., 2006), which may contain
euhedral grains of peritectic phases (Fig. 3b; Kenah
NUMBER 7
JULY 2008
& Hollister, 1983; Vernon & Collins, 1988; Grant &
Frost, 1990; Brown, 2001; Marchildon & Brown, 2002;
Vernon, 2004).
The volume increase associated with some fluid-absent
melting reactions may generate over-pressure, which can
result in hydrofracture (Fig. 3c and d; Clemens & Mawer,
1992; Connolly et al., 1997; Rushmer, 2001; Holness & Watt,
2002; Holness et al., 2005b), although micro-cracking
appears to be confined to experimental charges and pyrometamorphic environments. This is perhaps because only
in such extreme environments is the overstep sufficient to
overcome the natural tendency of systems to buffer along
reactions in P^T space.
If the melt production rate is sufficiently slow (with
reaction and deformation rates commensurate with that
of grain-boundary mass transport) melt-bearing systems
can attain textural equilibrium. Melt geometry becomes
a function of porosity or melt fraction (f), the equilibrium melt^solid dihedral angle () and the extent of
anisotropy of interfacial energies. Equilibrium dihedral
angles are controlled by the balancing of the two interfacial energies involved: the grain boundary energy
between the two grains of the same phase, gb, and
the energy of the interface between the two different
Fig. 3. (a, b) Brown glassy melt rims in partially melted psammitic rocks from the pyrometamorphic aureole of the Glenmore gabbro plug,
Ardnamurchan, Scotland, imaged in plane-polarized light. Feldspar is partially melted, forming a sieve-like texture. Biotite is entirely replaced
by spinel, new biotite and mullite. Scale bars represent 200 mm. Note the parallelism of the grain boundary melt films, indicating no relative
movement of liquid and restite (a), and in (b) the acicular orthopyroxene grains, which grew consequent to biotite breakdown. (c) CL image of
psammite collected 44 cm from the contact of the Traigh Bhan na Sgurra sill, Isle of Mull. The melt present at the peak of metamorphism was
derived from reaction between quartz and feldspar. Each feldspar grain in this image is surrounded by a brightly luminescing granophyric rim,
from which inferred melt-filled cracks emanate, forming a completely interconnected network. Scale bar represents 500 mm. (d) Reacted muscovite grains in a psammite collected 80 cm from the contact with the Traigh Bhan na Sgurra sill, Isle of Mull, imaged in plane-polarized light.
The muscovite is now replaced by mullite, biotite, K-feldspar and spinel, with some prismatic grains of cordierite (now entirely replaced by
yellow sheet silicates). The sites of the reacted muscovite grains are linked by melt-filled fractures, which are filled with brown granophyre. Most
of the fractures in this image contain melt but it is only apparent in BSE or CL. Scale bar represents 200 mm.
1346
HOLNESS & SAWYER
PSEUDOMORPHING OF MELT-FILLED PORES
phases, gi (Fig. 4a). For isotropic materials satisfies the
relation
gb ¼ 2gi cos
2
(Smith, 1948), but because most materials of interest to
geologists have significant anisotropy of interfacial energy
(e.g. Kretz, 1966; Vernon, 1968; Laporte et al., 1997), the
equation above is modified to incorporate torque forces,
which act to rotate interfaces to lower energy orientations
(Herring, 1951; Hoffman & Cahn, 1972; Cahn & Hoffman,
1974). This stabilizes planar interfaces at pore corners (e.g.
Kretz, 1966; Vernon, 1968; Laporte & Watson, 1995; Cm|¤ ral
et al., 1997; Lupulescu & Watson, 1999; Laporte & Provost,
2000), and a range of possible equilibrium angles, depending on the relative orientation of the solid grains (Herring,
1951; Laporte & Provost, 2000).
For geologically relevant systems the equilibrium melt^
solid dihedral angle is less than 608 [see Holness (2006)
for a recent review] with a standard deviation of 128
(Laporte & Provost, 2000; Holness, 2006) so the equilibrium distribution is one of interconnected channels along
three-grain junctions (Fig. 4b and c; Smith, 1964; Beere¤,
1975; von Bargen & Waff, 1986). For isotropic materials
these channels remain open and interconnected even for
vanishingly low porosities but, because all minerals are
anisotropic to some extent, there is a threshold level for
interconnectivity of f less than a few volume per cent
(Wark & Watson, 1998; Maumus et al., 2004; Price et al.,
2006; Yoshino et al., 2006).
In a system in which liquid can move freely, complete
equilibration will entail melt flow (e.g. Jurewicz &
Watson, 1985) until it reaches the minimum energy porosity, which ranges from 0 vol. % at 4608 to 23 vol. %,
or the porosity required for complete disaggregation, for
¼ 08 (Park & Yoon, 1985; Cheadle, 1989). For the likely
range of equilibrium melt^solid dihedral angles in geological systems (20^408, Holness, 2006), the minimum energy
Fig. 4. (a) Rounded pigeonite glomerocrysts in textural equilibrium with the andesitic fine-grained matrix, Mull, Scotland (from Holness,
2006). The direction of the force exerted by the energy of the grain boundary is labelled gb, and that of the interface between melt and solid is
shown as gi. The dihedral angle, , is a function of the balance between these two energies. Scale bar represents 500 mm. (b) Schematic illustration showing the difference in equilibrium melt topology for dihedral angles above and below the critical value of 608 for melt interconnectivity.
After Watson & Brenan (1987). (c) Scanning electron microscope image of an experimental charge in which quartz was equilibrated with brine
at 94 kbar and 8008C: the dihedral angle under these conditions is less than 608 (Holness,1992). (Note the channels on all three-grain junctions.)
The small acicular crystals scattered on the quartz surface formed during the quench. Scale bar represents 20 mm.
1347
JOURNAL OF PETROLOGY
VOLUME 49
porosity is 10^18 vol. % (Cheadle, 1989). The corollary of
this is that textural equilibration in a melt-present system
in which the liquid is free to move will result in the infiltration of melt along previously dry three-grain junctions
(Watson, 1982; Daines & Kohlstedt, 1993; Hammouda &
Laporte, 2000).
The extent of textural equilibration depends on the rate
and mechanism of deformation during anatexis. Although
diffusion creep does not change the melt distribution, dislocation creep can result in a significantly different melt
distribution compared with the static case (Ave’Lallemant
& Carter, 1970; Urai, 1983; Jin et al., 1994; Hirth &
Kohlstedt, 1995; Kohlstedt & Zimmerman, 1996; Daines &
Kohlstedt, 1997). Similarly, if the rates of reaction, whether
solidification in igneous rocks or melting in migmatites,
are faster than that of textural equilibration, melt geometries during reaction will be controlled by reaction
kinetics.
Much of our understanding of the process of textural
equilibration has been obtained from studies of plutonic
igneous rocks (Hunter, 1987; Holness et al., 2005b; Higgins,
2006). These show that textural equilibration, or maturation, takes place in a series of stages, the earliest of which
is the rotation of grain boundaries in the vicinity of threegrain junctions to achieve the equilibrium dihedral angle
(Holness et al., 2005b). As a consequence of angle change,
the grain boundaries in the vicinity of the pore corner
develop a step-change in curvature. Because changes in
mean curvature increase the energy of the interface
(Beere¤, 1975), the pore wall must move to attain constant
mean curvature, at least in isotropic systems (e.g. Holness
& Siklos, 2000). There is thus a complex interplay between
the gradually changing angle at the melt^solid^solid junction and the propagation of a change in grain boundary
orientation and curvature further from the junction.
Because continuous films of liquid on grain boundaries
are stable only for equilibrium melt^solid dihedral angles
of 08 (Smith, 1964), and equilibrium dihedral angles are
always 408 for silicate systems, reaction-generated grain
boundary melt films will neck down to form strings of
melt-filled lenses of a shape controlled by the dihedral
angle. Melt-filled, intra-grain cracks formed by hydrofracture will heal, leading to arrays of melt-filled inclusions
with a shape controlled by the interfacial energies of the
host crystal.
The details of establishment of the equilibrium dihedral
angle at pore corners depends on the geometry of the
initial reaction-controlled melt geometry and on whether
the equilibration takes place in the super- or sub-solidus
(or both). Thin grain boundary melt films, generated
either by reaction between adjacent phases or by intergrain fracturing, in a rock with an initially wellequilibrated solid-state microstructure will initially generate dihedral angles of 1208 (Fig. 5). As the walls melt back
NUMBER 7
JULY 2008
and the melt rim becomes thicker, the dihedral angles will
tend towards 1808, with a low standard deviation. Melt
films propagating down two-grain junctions are akin to
propagating fractures and likely to have very low dihedral
angles at their tips. Melt^solid^solid dihedral angles in a
melting rock in which melt distribution is dominated
by grain boundary films will therefore form as many as
three separate peaks (Fig. 5c). The rate at which pore junctions will progress towards melt-present equilibrium will
depend on their geometry. The tips of propagating films
are close to melt^solid^solid equilibrium and will not
change significantly. The junctions of grain boundaries
with thick melt films are farthest from equilibrium, will
have the greatest driving force for microstructural change,
and will therefore undergo rapid adjustment of the pore
walls. The peak at 1208 will migrate relatively slowly
towards lower angles.
M I C RO S T RU C T U R E S F O R M E D
D U R I N G C RY S TA L L I Z AT I O N
Cooling rate exerts a primary control on the microstructures formed during crystallization. At one extreme we
find supercooled melt (glass), whereas more slowly cooled
melt crystallizes as increasingly coarse polycrystalline
aggregates as the cooling rate is decreased. This can be
illustrated by a consideration of quartzo-feldspathic rocks
from a series of contact aureoles.
Pyrometamorphism at 120 bar around the 50 m diameter gabbro plug at Glenmore, Ardnamurchan, resulted
in glassy melt rims separating quartz and feldspar (Fig. 3a
and b; Butler, 1961; Holness et al., 2005a). Solidification of
melt rims in the 150 bar aureole of the gabbro plug in
Kinloch, Isle of Rum, and in the 600 bar aureole of the
Traigh Bhan na Sgurra sill (Holness & Watt, 2002)
resulted in a fine granophyric intergrowth (Fig. 6a).
Solidification of melt rims in amphibolitic gneiss metamorphosed by the Rum Igneous Complex at 100^200 bar
formed a coarser granophyric intergrowth in which
quartz paramorphs after tridymite are visible (Fig. 6b and
c; Holness & Isherwood, 2003).
A further reduction in cooling rate results in melt rims
on quartz^feldspar grain boundaries being replaced
by either an aggregate of equigranular, equant grains of
feldspar and quartz (termed the string of beads texture;
Holness & Isherwood, 2003) with randomly oriented
grains, or by a highly irregular grain boundary (Fig. 6e
and f). The irregularity of the latter is due to the formation
of extremely coarse granophyric intergrowths, with each
phase having nucleated and grown on its respective sidewall of the melt rim. Which of these two textures forms is
probably a consequence of the ease of nucleation during
solidification.
1348
HOLNESS & SAWYER
PSEUDOMORPHING OF MELT-FILLED PORES
As the cooling rate decreases further, for larger or
deeper contact aureoles and for regional metamorphism,
the extended time scale permits the onset of melt migration, resulting in the grain-scale draining of melt from its
localized source, the destruction of thick, parallel-sided,
melt films such as those depicted in Fig. 6b, and segregation of melt into large pockets. This permits the identification of the second major control of crystallization
microstructures: the size of the melt pocket.
In the deeper parts of the contact aureole associated
with the Rum Layered Suite (Holness & Isherwood,
2003), and in that associated with the 3 kbar, 10 km diameter, Ballachulish Igneous Complex (Pattison & Harte,
1988), the largest melt pockets are filled with granophyre
(Fig. 7a) or by oikocrysts with abundant inclusions
(Fig. 7b; Grant & Frost, 1990; Harte et al., 1991). Thick melt
films on quartz^feldspar grain boundaries are replaced by
highly cuspate intergrowths or a string of beads (Fig. 6e
and f), whereas the smallest pores (the thinnest melt rims,
and melt pockets bounded by 3^4 grains) are generally
pseudomorphed by single crystals of the volumetrically
minor phase (Fig. 7c).
For migmatites in the mid- to deep crust, the same
pattern of solidification microstructures is observed.
Smaller melt pockets (typically those bounded by fewer
than four grains) are partially or wholly pseudomorphed
by single grains. In contrast, large melt pockets, or leucosomes, crystallize as poly-mineralic aggregates (although
granophyric intergrowths are apparently confined to
shallow to mid-crustal environments). We suggest that this
microstructural progression, from poly-mineralic aggregates in the larger pores to monocrystalline pseudomorphs
of the smaller pores, which can be observed within a single
rock (and therefore at a constant cooling rate), reflects an
increasing barrier to nucleation as the pore size becomes
smaller.
The role of pore size in nucleation
Because an atom on the surface of a small crystal suspended in a solution is in a highly energetic condition compared with an atom of the same material on a planar
surface (i.e. on the surface of an infinitely large, spherical,
crystal) in contact with the same solution, it has a stronger
Fig. 5. (a) Image of Ashuanipi granulite-facies migmatite in which
the melt phase has been pseudomorphed by K-feldspar. The sketch in
(b) shows grain outlines, with melt as white, and the quartz and plagioclase as light and dark grey, respectively. Field of view is 09 mm
wide. The different types of pore^solid^solid junctions should be
noted: those corresponding to the tips of grain boundary melt films
have dihedral angles 208, whereas those at the junction of
grain boundaries intersecting melt films at a high angle have
dihedral angles of 1808. Those formed at erstwhile solid^solid junctions of 1208 have an initial dihedral angle of 1208. (c) During
melt-present textural equilibration, the low angles at film tips do not
need to increase much to attain equilibrium dihedral angles of 308,
whereas those at high-angle, grain boundary intersections need to
decrease by 1508çthe driving force is greatest for these high
angles. (d) During sub-solidus textural equilibration, the endpoint
for dihedral angles is now 1208, and it is the low angles that now
have the greatest departure from, and hence driving force towards,
textural equilibrium.
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JOURNAL OF PETROLOGY
VOLUME 49
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Fig. 6. A progression of solidification microstructures developed in grain boundary melt films in quartzo-feldspathic rocks as the cooling rate
decreases. (a) Fine-grained granophyre replacing a feldspar grain in the 600 bar aureole of the Traigh Bhan na Sgurra sill, Isle of Mull,
Scotland. Scale bar represents 200 mm. (b) Continuous and parallel-sided granophyric rims separating quartz and feldspar in gneiss from
the 100^200 bar aureole of the Rum Igneous Complex, Scotland. Scale bar represents 1mm. (c) Close-up view of the granophyric rim from
(b), showing the elongate quartz paramorphs after tridymite. Scale bar represents 200 mm. (d) The string-of-beads texture developed in more
slowly cooled, deeper parts of the Rum aureole. Quartz paramorphs after tridymite are no longer present. Scale bar represents 200 mm.
(e) Highly cuspate quartz^feldspar grain boundaries are the manifestation of extremely coarse-grained intergrowths of quartz and feldspar,
in which the two phases nucleate on the walls of the melt rim. Gneiss from the deeper parts of the Rum aureole. Scale bar represents 1mm.
(f) Close-up of (e) showing the highly cuspate grain boundaries. Scale bar represents 200 mm.
tendency to leave the crystal (i.e. to dissolve). To maintain
chemical equilibrium, to permit a correspondingly higher
rate of movement of atoms from the solution onto the crystal, the small crystal needs to be in contact with a stronger
solution than does a larger crystal. The concentration of
solute, C, in equilibrium with a spherical crystal of radius
r is given by
2g RT
C
¼
ln
r
V
C0
(the Gibbs^Thompson, or Ostwald^Freundlich equation), where g is the energy of the solid^liquid
interface, R the gas constant, T the temperature, V
the molar volume of the solid phase, and C0 is the solubility of a liquid in equilibrium with an infinitely large
spherical crystal or a planar interface (Cahn, 1980;
Adamson, 1990).
This means that the thermodynamics of crystallization
in pores is different from that in a free fluid: most importantly, a confined fluid can become more supersaturated
before crystallization begins compared with an unconfined
fluid of the same composition (e.g. Bigg, 1953; Melia &
Moffitt, 1964; Cahn, 1980; Scherer, 1999; Putnis & Mauthe,
2001). For liquid-bearing rock with a range of pore sizes the
point at which crystals (be they interstitial grains in a plutonic igneous rock or cement in a sedimentary rock) nucleate will be determined by pore size. The first observation of
this effect in rocks was reported by Putnis & Mauthe
(2001), who showed that small pores in sandstone remain
empty whereas nearby larger pores are filled with halite
cement. A similar effect is probably implicated in the preservation of the thick grain boundary films of melt in
almost completely solidified crystalline mafic nodules
described by Holness et al. (2007).
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HOLNESS & SAWYER
PSEUDOMORPHING OF MELT-FILLED PORES
sluggish. It is rare to find completely equilibrated solidstate microstructures in polymineralic rocks.
Typical median angles for solid-state textural equilibrium fall in the range 100^1408 (e.g. Kretz, 1966; Vernon,
1968, 1970), with standard deviations of 10^208 (Vernon,
1968, 1970). Therefore it is the pseudomorphed propagating
film tips that are furthest from equilibrium and likely to
migrate the fastest (Fig. 5d). The 1808 angles on thick melt
films will migrate to lower angles relatively slowly.
C O N TAC T- M E TA M O R P H O S E D
Q UA RT Z O - F E L D S PAT H I C
RO C K S AT 2 k b a r
Fig. 7. (a) Granophyric intergrowths of quartz and feldspar, nucleating on euhedral crystals of plagioclase in Lewisian gneiss from the
aureole of the Rum Igneous Complex. Scale bar represents 1mm;
crossed polars. (b) Melt pool defined by oikocrystic quartz enclosing
euhedral plagioclase crystals that grew directly from the liquid.
Contact metamorphosed Lewisian gneiss, Isle of Rum, Scotland.
Scale bar represents 200 mm; crossed polars. (c) Contact metamorphosed arkose from the aureole of the Ballachulish Complex,
Scottish Highlands, under crossed polars. This sample was collected
1m from the contact and contains rounded quartz grains with highly
cuspate feldspar grains with low feldspar^quartz^quartz dihedral
angles. The feldspar grain shapes are suggestive of melt pseudomorphing. Scale bar represents 200 mm.
T E X T U R A L M O D I F I C AT I O N I N
T H E SU B - SOL I DUS
Further microstructural adjustment towards lower internal
energies may occur in the sub-solidus (Fig. 5d). This process is much slower than the melt-present case because
mass transfer is achieved by grain boundary diffusion
rather than mass transport through a liquid. Furthermore,
whereas sub-solidus textural equilibration of microstructures in monomineralic domains is rapid, as it necessitates
only mass movement across grain boundaries (Hunter,
1987), equilibration in polymineralic domains requires
mass transport along grain boundaries and is therefore
The Lower Proterozoic Biwabik Iron Formation in
Minnesota comprises a cherty and slaty banded iron formation (Morey, 1992) and is part of the Animike Group,
which was metamorphosed at 2 kbar by the Duluth
Igneous Complex, an arcuate mass of troctolitic and
anorthositic intrusions of Middle Proterozoic age located
in the mid-continent rift of North America (Van Schmus
& Hinze, 1985). Samples of cherty horizons of the Biwabik
Formation were obtained from a drill core through the
intrusion and into the footwall.
The cherty horizons are dominated by coarse-grained
quartz, feldspar, clinopyroxene (with well-developed exsolution lamellae of Ca-poor pyroxene), magnetite and ilmenite (Fig. 8a). Coarse-grained granophyric intergrowths are
common and mark centimetre-sized pockets of melt
(Fig. 8b). Quartz-rich parts of the rock are dominated by
rounded quartz grains separated by a network of thin
single crystals of plagioclase (with minor clinopyroxene),
which outline the grain boundaries and form cuspate
grains at three-grain junctions. Each plagioclase grain is
approximately one, or at most two, quartz grains in
length. This network occupies 10 vol. % of the rock.
A similar pattern is observed in relatively quartz-poor
regions, in which the Fe-oxides and clinopyroxene form
the network of interserts (Fig. 8a).
A hundred randomly chosen plagioclase^quartz^quartz
dihedral angles form a unimodal population with a
median of 688, and a standard deviation of 1738 (Fig. 9b).
The median is significantly higher, and the standard deviation slightly higher, than expected for melt^quartz equilibrium (208 and 108, respectively; Holness, 2006). For
comparison with solid-state equilibrium, plagioclase^
quartz^quartz angles were measured in texturally wellequilibrated regions of an upper amphibolite-facies rock
(sample EL180, which is described at length below). The
median and standard deviation of a population of 100
angles are 1128 and 1608, respectively (Fig. 9a), comparable
with a mean angle of 1058 (and standard deviation of
1248) measured in granulites from Broken Hill (Vernon,
1968), and also with a median of 1178 measured by
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Fig. 8. Contact metamorphosed banded iron formation of the Biwabik Formation, from the aureole of the Duluth Igneous Complex,
Minnesota. (a) The rock comprises quartz, plagioclase, Fe-oxides and clinopyroxene. Note the rounded shape of the quartz grains and the
generally cuspate nature of the others. Plane-polarized light. Scale bar represents 1mm. (b) Melt is replaced by granophyre in the large pockets
in quartz-rich regions, whereas the smaller pockets are pseudomorphed by plagioclase. Note the almost continuous grain boundary network of
plagioclase grains, with low plag^quartz^quartz dihedral angles where the network is discontinuous. Cross polars. Scale bar represents 1mm.
(c) Back-scattered electron image showing the contrast between the granophyre-filled large pores and the pseudomorphed thin grain-boundary
and grain-edge channels. The change from granophyre to pseudomorphs occurs in melt pockets 250 mm wideça transitional pore is arrowed.
Scale bar represents 1mm. (d) Plane-polarized light view of quartz-rich regions showing the network of plagioclase-filled pores. Scale bar
represents 1mm.
transmission electron microscopy in pelitic schists (Hiraga
et al., 2002). These angles are significantly higher than the
median measured in the Biwabik migmatite.
CL imaging reveals four distinct quartz generations
(Fig. 10). Dark, weakly luminescing lobate cores are overgrown by a slightly more luminescent quartz generation,
which does not completely surround the cores in many of
the grains, signifying a period of quartz dissolution before
the growth of the third quartz generation. The latter forms
brightly luminescent, discontinuous rims around the partially resorbed grains, and is rare in the quartz-rich zones.
The final quartz generation is slightly less luminescent
than the brightest generation and forms continuous rims
around the quartz grains. This fourth generation is that of
the granophyre in the large melt pockets, and forms wide
margins and irregular extensions in optical continuity
with the quartz grains bounding the granophyric pockets.
It forms only thin and discontinuous rims in the quartzrich regions.
Having established that luminescence intensity is correlated with Ti content (Fig. 2), temperatures corresponding
to each of the four quartz generations along four profiles
across quartz grains were determined using the TitaniQ
geothermometer of Wark & Watson (2006). Because
rutile is absent, the temperatures (shown in Fig. 10) are
minimum estimates, although the relative differences
between the generations are accurate if TiO2 activity
remained constant (Wark & Watson, 2006). This assumption depends on the mineral assemblage remaining the
same throughout the metamorphic history. Although a
complete assessment of the TiO2 activity would necessitate
examination of lower-grade samples from the aureole,
ilmenite is common in these rocks and we believe the
TiO2 activity was most probably close to unity for the
duration of quartz recrystallization and growth. With
these caveats, the dark, resorbed cores give a temperature
of 630 108C. That of the first overgrowth is variable, with
a monotonic increase to 730^8008C. The bright, discontinuous overgrowths have temperatures between 870 and
9908C, whereas the outermost rims and granophyric intergrowths record temperatures of 820^8308C.
Plagioclase compositions were determined across a
single thin-section (sample ELY-6B), with randomly
chosen analyses within the large granophyric pocket and
the thin melt pseudomorphs. The composition is shown as
a function of the width of the plagioclase interserts (or the
width of the granophyric pockets) in Fig. 11. For transitional regions, in which the walls of the pocket are lobate,
the width is taken to be that at the widest part. The plagioclase composition varies by a few molar per cent about an
average of Ab51 until the width of the pockets decreases
below 40 mm. For smaller pockets the plagioclase is more
albitic, reaching Ab88 for a significant proportion of the
thinnest grain boundary films.
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HOLNESS & SAWYER
PSEUDOMORPHING OF MELT-FILLED PORES
Fig. 9. Populations of true, 3-D dihedral angles, measured with the universal stage. The number of measurements for each population is given
by n. (a) Equilibrium populations, showing the spread of angles as a result of anisotropy of interfacial energies. The melt^plag^plag population
is taken from Holness (2006), whereas the plag^quartz^quartz angles were measured on sample EL180 from regions in which melt-derived
textures were absent. (b) Plag^plag^quartz angles from the Biwabik iron formation. (c) All K-feldspar angles from the Ashuanipi granulitefacies migmatite. The bimodal population comprises a peak at low angles corresponding to the tips of melt films on (predominantly)
plagioclase^quartz grain boundaries, whereas the higher peak corresponds to that on quartz^quartz, or plagioclase^plagioclase, boundaries
intersecting melt films at high angles. The K-feldspar grain is outlined for clarity. (d) All microcline dihedral angles from the upper amphibolite-facies migmatite from the Opatica Subprovince.
R E G I O N A L Q UA RT Z O F E L D S PAT H I C G R A N U L I T E S
Ashuanipi Subprovince, Quebec
The Ashuanipi Subprovince, lying NE of the Opatica
Subprovince, forms the eastern end of a 2000 km long
meta-sedimentary belt that crosses the entire width of
the Superior Province (Card & Ciesielski, 1986; Percival
et al., 1992) and is believed to be the remains of a single,
late Archaean, accretionary prism (Percival, 1989).
Sample DL96-1006A [previously described by Sawyer
(2001)] is a meta-greywacke that experienced a prolonged
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Fig. 10. Main image shows a composite of CL images of the Biwabik sample ELY-6B from the region shown in Fig. 6c, with single quartz grains
outlined in white (the quartz component of the granophyre is outlined only where it forms an overgrowth on restitic grains). The four distinct
quartz generations, distinguishable by variations in luminescence, should be noted. Profiles of Ti concentration measured along the four traverses shown as white lines in the CL image were converted to minimum temperature estimates using the thermometer of Wark & Watson
(2006) and are shown as plots to the right of the CL image, each colour-coded for the four generations of quartz. Spot analyses of plagioclase
are shown, calculated to molar per cent albite, on the CL image. The positions of the CL spectra shown in Fig. 1 are shown as spots labelled
x, y and z, located both on the CL image and in the temperature traverses (a and b).
(40^50 Myr, Guernina & Sawyer, 2003) metamorphic
event that peaked in the granulite facies at about 8508C
and 6^65 kbar (Guernina & Sawyer, 2003; Percival, 1991).
It is a foliated, medium-grained, quartzo-feldspathic
rock containing abundant plagioclase, biotite and orthopyroxene, and with a melt-depleted bulk composition
(Guernina & Sawyer, 2003).
Melt, pseudomorphed predominantly by non-twinned
K-feldspar (which appears only as pseudomorphs), with
subordinate amounts of quartz and plagioclase, is concentrated in films and pools around reactant biotite grains
(Fig. 12a). Quartz and plagioclase pseudomorphs tend
to be more compact than the K-feldspar, with high dihedral angles against the adjacent quartz and plagioclase
grains. In contrast, K-feldspar pseudomorphs form
extended, cuspate films several grain diameters in length.
Melt films occur mainly on biotite^quartz, biotite^
plagioclase, and plagioclase^quartz grain boundaries
and are associated with rounded, corroded grains of
biotite, plagioclase and/or quartz (Fig. 12a and b), suggesting the reaction
biotite þ plagioclase þ quartz ¼ orthopyroxene
þ melt ilmenite
(Sawyer, 2001). Melt films on plagioclase^plagioclase and
quartz^quartz boundaries are uncommon, are generally
in close proximity to biotite grains, and are always a continuation of a thick melt film on a plagioclase^quartz or
biotite^quartz grain boundary. The thinnest films terminate along an inter-phase boundary (Fig. 12c and d) and
isolated thin lenses may be present. Slightly thicker films
are usually only a single grain width long and terminate
at a three-grain junction. The thickest films are several
grain diameters in length, with several melt-free grain
junctions terminating at the film wall (Fig. 5a and b).
Quartz is generally featureless under CL, apart
from a slightly brighter grain core compared with
the rim, and the presence of numerous, late-stage,
1354
HOLNESS & SAWYER
PSEUDOMORPHING OF MELT-FILLED PORES
geometry of the melt film. The terminations of the thinnest
melt films have low dihedral angles, forming a well-defined
peak at about 208 (examples are labelled ‘a’ in Fig. 14). The
slightly thicker melt films, which terminate at three-grain
junctions, have angles generally below about 608, whereas
the junctions of melt-free grain boundaries, which terminate at the thickest melt films, form a well-defined peak
at 1008 (examples of which are labelled ‘b’ in Fig. 14).
Although this does correlate with mineralogy, as lowangle, terminating melt films almost invariably occur on
plagioclase^quartz grain boundaries whereas it is predominantly quartz^quartz grain boundaries that terminate at
the thickest melt films (Figs 12 and 14), we do not think
that these variations are caused by interphase variations
in interfacial energy. This is because equilibrium dihedral
angles for melt^quartz compared with melt^plagioclase,
and quartz^plagioclase^plagioclase compared with
plagioclase^quartz^quartz systems are very similar
(Vernon, 1968; Holness, 2006).
Opatica Subprovince, Quebec
Fig. 11. (a) Plagioclase composition as a function of width of the
plagioclase pocket in sample ELY-6B from the Biwabik Formation
in the aureole of the Duluth Igneous Complex. (b) Detail of (a) for
pockets smaller than 200 mm. The grey area shows the range of compositions observed in the larger pockets with the average (An508)
shown by the horizontal line. (c) BSE image showing the position of
each analysis together with the measured plagioclase composition.
The most albitic plagioclase is found in the thinnest part of the films
and at the film tips.
dark fractures (Fig. 13). The absence of the kind of detail
observed in the Biwabik samples is probably because Ti
variations are able to decay by diffusion during extended
high-temperature episodes.
Dihedral angles at the junctions of the K-feldspar melt
pseudomorphs with either plagioclase or quartz were measured in sample DL96-1006A, which is typical of the terrane (Sawyer, 2001). The angle population is strongly
bimodal (Fig. 9c). The bimodality is controlled by the
The Opatica Subprovince is a 200 km wide belt of amphibolite-facies plutonic gneisses from the southeastern part of
the Superior Province, Quebec. It is interpreted as the
deeply eroded interior of an Archaean mountain chain
and represents a section of reworked Archaean crust without major tectonic disruption (Sawyer & Benn, 1993). The
highest-grade, and later exhumed, migmatites are in the
centre of the Subprovince and record maximum temperatures of 760^8108C at a pressure of 6^7 kbar (Sawyer,1998).
Sample EL180, collected from approximately 508330 N
and 778330 W, is a typical residual grey gneiss containing
plagioclase, green hornblende and biotite, with accessory
apatite and zircon [see Sawyer (1998) for a bulk composition analysis]. Titanite is abundant and forms strings of
small grains within aggregates of biotite and amphibole.
Orthopyroxene is absent. Biotite grains are commonly partially chloritized, and are deformed by the adjacent growth
of lenses of untwinned K-feldspar, many of which are partially albitized. Both chloritization and lens growth are
indicators of low-temperature migration of hydrous fluids
(Holness, 2003), which therefore significantly postdate formation of the migmatites.
Solidified melt is pseudomorphed by microcline, plagioclase and quartz. Extended, fine-grained, poly-mineralic
aggregates commonly occur along boundaries between
larger grains (Fig. 15a), and we interpret these as solidified
melt rims. The quartz and plagioclase grains tend to be
equant, with 1208 dihedral angles against adjacent quartz
and plagioclase grains (Fig. 15b). Only the microcline
grains are cuspate and elongate, with dihedral angles
51208, reminiscent of the interserts described elsewhere
(Fig. 15b and c), but even the microcline pseudomorphs
are relatively rounded and equant compared with those in
the Ashuanipi granulite (Fig. 12).
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Fig. 12. Microstructures in Ashuanipi granulite-facies migmatite, DL96-1006A. (a) Back-scattered electron image, showing the distribution of
K-feldspar in thin grain boundary films associated with biotite (elongate pale grey grain). Note the termination of these films part-way along
plagioclase^quartz grain boundaries, and the rounded and resorbed nature of plagioclase and quartz within K-feldspar pools. Scale bar represents 50 mm. (b) Photomicrograph under crossed polars. The K-feldspar-filled melt pools surround corroded biotite grains, with much of the
melt forming elongate pools parallel to the margins of the biotite grains. Scale bar represents 200 mm. (c) Back-scattered electron image demonstrating the general absence of melt films along plagioclase^plagioclase and quartz^quartz grain boundaries. Scale bar represents 50 mm.
(d) Photomicrograph under crossed polars. Note the tapering melt films ending part-way along plagioclase^quartz grain boundaries and the
corroded nature of the quartz and plagioclase. Scale bar represents 200 mm.
Melt is not concentrated adjacent to biotite or amphibole
but is evenly distributed, forming thick grain boundary
films on plagioclase^plagioclase, plagioclase^quartz,
plagioclase^biotite and quartz^biotite grain boundaries,
together with compact, triangular pools at quartz or feldspar three-grain junctions: these are rounded with blunt
apices. Thin melt films are uncommon, and intra-grain
cracks are rare (only one example was observed).
The melting reaction is not as clear-cut as in the
Ashuanipi rocks. However, the dearth of melt around biotite grains, the absence of peritectic phases, and the concentration of melt rims on plagioclase^quartz grain
boundaries, suggest that it may have been
Quartz þ Plagioclase þ K-feldspar þ H2 O ¼ Melt:
The confinement of K-feldspar to melt pseudomorphs suggests not only that all of the original K-feldspar grains have
been melted, but also that significant melt loss has
occurred (see Sawyer, 1998). The source of the hydrous
fluid that drove melting is uncertain.
In a similar fashion to the Ashuanipi samples, the
Opatica rocks do not reveal any detailed information in
CL images. Once again, the study concentrated on
dihedral angles formed at the corners of melt pores pseudomorphed by K-feldspar. The dihedral angle population
is bimodal with peaks at 708 and 1008 (Fig. 9d). Angles
below 508 are rare. In a similar manner to DL96-1006A,
the lower peak corresponds to the tips of grain boundary
melt films and the corners of pores at three-grain junctions, whereas the higher peak corresponds to grain
boundaries truncated by melt films.
T H E DEVELOPM EN T OF POR E
PSEU DOMORPHS
The best information about the development of pore pseudomorphs is preserved in the cherty bands in the Biwabik
Formation. The four generations of quartz shown in Fig. 10
record metamorphism, anatexis and solidification of the
protolith. The last episode of quartz growth, to form the
granophyre, occurred at a temperature consistent with
melt^solid chemical equilibrium in the Qtz^Ab^H2O
system at 2 kbar (Johannes, 1978; although he presented
data only for 5 kbar). The earlier growth of highly luminescent quartz may have occurred during a period of
1356
HOLNESS & SAWYER
PSEUDOMORPHING OF MELT-FILLED PORES
Fig. 14. Photomicrographs of Ashuanipi migmatite, DL96-1006A
showing examples of low dihedral angles preserved at the tips of
grain boundary melt films (examples are marked ‘a’ in both images),
whereas high angle junctions of grain boundaries between nonreactants (i.e. intra-phase boundaries such as quartz^quartz) with
melt films have high dihedral angles of 120^1808 (examples are
marked ‘b’ in both images). Scale bars in both images represent 200 mm.
Fig. 13. (a) Back-scattered electron image of Ashuanipi granulitefacies migmatite DL96-1006A with corresponding region imaged
under CL (b). Note the absence of fine detail, the generally more
luminescent cores to the quartz grains and the fine tracery of nonluminescing fractures associated with late-stage hydrous fluids.
Scale bar represents 50 mm.
lower H2O activity (and higher melt temperatures);
the details of this remain unclear.
The quartz overgrowths on the restitic grains bounding
the largest pores show that at the metamorphic peak the
melt was saturated only in quartz. This melt occupied
20 vol. % of the quartz-rich regions (of which about
half now comprises feldspar) and also formed much
larger, scattered pockets (Fig. 10). During cooling, overgrowths formed on the quartz restite until the liquid
moved onto the quartz^feldspar cotectic. At this point feldspar nucleated, and quartz and feldspar grew simultaneously to form granophyric intergrowths in the largest
pores. CL images of the largest granophyre-filled pores
demonstrate that undulose quartz^feldspar boundaries signify simultaneous crystallization of quartz and feldspar. In
contrast, the smooth quartz^feldspar grain boundaries in
the smaller pores denote sequential growth, with feldspar
(or clinopyroxene or Fe-oxides in other parts of the rock;
Fig. 8a) crystallizing alone, following quartz overgrowth
on the pore walls. The critical pore width at which this
change occurs is 250 mm (Fig. 8c). It should be noted
that in some pores quartz continued crystallizing until the
pore was completely filled (Fig. 10).
There is no kinetic barrier to overgrowth of pore walls,
but crystallization of the phases not forming the pore
walls requires homogeneous nucleation. Although this
becomes easier as overgrowth on the walls proceeds
(which increases the concentration of the other phases in
the remaining liquid), the smaller the pore, the greater
the supersaturation required for nucleation. Nucleation
of the minor phase therefore cannot occur until the temperature has decreased to create sufficient undercooling to
overcome the barrier to nucleation. Once nucleation does
occur, the small pores are pseudomorphed by a few grains,
which spread out in the porosity over 1^2 restitic grain diameters. In the case of the Biwabik migmatite, this plagioclase growth must have been accompanied by quartz
overgrowth elsewhere, as the rejected quartz component
migrated through the remaining porosity away from the
growing feldspar.
Pore-size-controlled nucleation delay of the plagioclase
in the Biwabik sample, ELY-6B, is consistent with the general increase in Na content of plagioclase in progressively
smaller pores (Fig. 11). Within any single pore, the variation
in Na content shows that growth occurred first in the
widest region and subsequently propagated into the smallest extremities (Fig. 11c). When all the data are taken as a
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Fig. 15. Photomicrographs of Opatica amphibolite-facies migmatite
EL180. (a) An annealed and recrystallized string-of-beads texture
developed between reacting grains of plagioclase and quartz. The
larger grains are interpreted as restite and are marked qtz and plag,
whereas the crystals solidified from thick grain boundary melt films
are marked m (microcline), q (quartz) and p (plagioclase). Note how
the microcline within the thinner melt film (bottom left-hand corner of
image) is elongate and cuspate, in contrast to the quartz and plagioclase grains within the thicker melt film. The plagioclase restite grain
at the top centre of the image has an extension with a slightly different
extinction position separating the two white quartz restite grains. This
extension nucleated on the plagioclase restite grain during solidification of the thick melt film. Scale bar represents 200 mm. (b) String of
quartz grains forming a poly-crystalline pseudomorph of a grain
boundary melt film. The asterisk marks the spot where four quartz
grains meet, two of which are inferred to have crystallized from melt.
The two three-grain junctions have 1208 angles, demonstrating subsolidus textural equilibration. The microcline grains are cuspate.
Scale bar represents 200 mm. (c) Microcline pseudomorphs after melt
films. Note the cuspate shape denoting limited sub-solidus textural
equilibration, although the dihedral angles at the corners have
increased somewhat in the sub-solidus. Scale bar represents 200 mm.
whole it appears that significantly delayed nucleation
occurs in melt films below about 40 mm thickness. For
ELY-6B, reduced solidification rates associated with
delayed feldspar nucleation in the small pores permitted
NUMBER 7
JULY 2008
attainment of smoothly curved pore walls and (perhaps)
also the equilibrium melt^quartz^quartz dihedral angle
(208, Holness, 2006). The change from Ab50 to Ab90
corresponds to a solidus temperature decrease of 308C in
the Qtz^Ab^An^H2O system at 5 kbar (Johannes, 1978).
Although the cooling rate of the aureole of the Duluth
Complex is not known, it is likely that a decrease in
temperature of several tens of degrees would take longer
than the few hundreds of years thought to be required to
equilibrate a melt-bearing rock with a 1mm grain size
(Cheadle, 1989; Holness & Siklos, 2000).
A critical issue to address is the preservation of melt^
solid dihedral angles in pseudomorphed pores. In the
Biwabik rock, nucleation delay appears to have promoted
melt^solid textural equilibration, and it is these equilibrium angles that are retained by the plagioclase porefilling. However, in the Ashuanipi granulite, and perhaps
also in the Opatica migmatite, dihedral angles recorded
by the K-feldspar pseudomorphs are far from textural
equilibrium. This implies both that any nucleation delay
was insufficient to promote melt^solid textural equilibration before the melt was replaced by K-feldspar, and
also that overgrowth on the walls of K-feldspar-filled
pores in the vicinity of the pore corners was minimal.
To interpret the preserved angle population in terms of
melt distribution and subsequent cooling history it is
essential to be certain that this population accurately
reflects melt^solid angles, rather than being an artefact
of overgrowth on the pore walls. Although a mechanism
for pseudomorphing based on the kinetics of nucleation
in small pores can be used to explain the observed microstructure in the Biwabik migmatites, there is no corresponding information on the progressive stages of
solidification of either the Ashuanipi or the Opatica
migmatites.
The Opatica sample described here contains a pseudomorph population that can plausibly account for a granitic
liquid (i.e. quartz þ plagioclase þ microcline). All three
phases crystallized together in larger pools and the thicker
melt rims, whereas isolated single grains infill the smaller
pores and thinner melt rims. Restitic plagioclase grains
commonly have either overgrowths or elongate extensions
into adjacent inferred melt pools, discernible by a variation
in extinction position (Fig. 15a). In contrast, the Ashuanipi
granulite is dominated by K-feldspar pseudomorphs.
The absence of preserved CL information means that we
cannot be sure about the quartz pore walls, but plagioclase
pore walls bounding K-feldspar pseudomorphs appear
optically to be compositionally uniform, suggesting that
at least the plagioclase has not been overgrown. It is
not clear why some melt-filled pores in the Ashuanipi
migmatite were apparently perfectly pseudomorphed by
K-feldspar, with the plagioclase and quartz components of
the melt crystallizing elsewhere.
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HOLNESS & SAWYER
PSEUDOMORPHING OF MELT-FILLED PORES
I N T E R P R E TAT I O N O F
PSEU DOMORPH ED PORE SH A PES
Melt^solid textural equilibrium is most likely to have been
attained in the Biwabik sample, ELY-6B, which contained
about 20 vol. % melt within the quartz-rich regions
(Fig. 10). This is close to both the disaggregation porosity
of 23% and the minimum energy porosity of 185% for a
dihedral angle of 208 (Cheadle, 1989). Distinction between
the possibilities that the quartz-rich parts of the rock
formed a crystal slurry at the metamorphic peak, or that
their microstructure records complete textural equilibrium, with excess melt expelled to form large pockets
(such as that shown in Fig. 10; Jurewicz & Watson, 1985) is
difficult, as the melt percentage is variable on the millimetre scale.
If textural equilibrium were achieved in rocks containing 5 vol. % melt, inferred for the Ashuanipi and
Opatica samples, melt^solid dihedral angles of 208
would lead to most of the melt residing in tubular channels
on three-grain junctions (Smith, 1964; Beere¤, 1975). A randomly oriented population of such junctions will be predominantly sectioned at high angles to their length, so
tubular melt pockets would generally appear as cuspate
pockets at three-grain junctions. The observed dominance
of elongate pockets on two-grain junctions in both EL180
and DL96-1006A cannot be accounted for by random
cross-sectioning (see Wark et al., 2003) but must reflect a
dominance of grain-boundary melt films, suggesting textural disequilibrium on the grain scale (even if equilibrium
dihedral angles are established on a smaller scale). Melt
films are most common on grain boundaries between reactant phases, suggesting that the grain-scale melt distribution is controlled predominantly by reaction kinetics.
Reported examples of dihedral angles associated with
pore pseudomorphs (Holness & Clemens, 1999; Rosenberg
& Riller, 2000; Marchildon & Brown, 2002) only rarely
form populations consistent with either melt^solid (e.g.
Rosenberg & Riller, 2000) or solid^solid textural equilibrium. This is in agreement with the results presented
here for the contact metamorphosed rock from the
Biwabik Formation, in which the unimodal plagioclase^
quartz^quartz dihedral angle population (Fig. 9b) is intermediate between melt^solid and solid^solid textural
equilibrium (Fig. 9a). A comparison with partially
re-equilibrated angle populations in olivine cumulates
(Holness et al., 2005b), in which a similar, intermediate,
unimodal population occurs in clinopyroxene oikocrysts,
suggests that the present dihedral angle population in the
Biwabik Iron Formation sample represents the partial
approach to solid-state textural equilibrium of an inherited
fully equilibrated population of melt^solid dihedral angles,
potentially providing a measure of the cooling rate in the
sub-solidus.
On the assumption that the angles preserved in the
Ashuanipi granulite reflect those inherited from the melt
with no modification by pore-wall overgrowth, their
strongly bimodal population is also clearly out of any kind
of textural equilibrium, supportive of the wider disequilibrium recorded by the dominance of grain boundary reaction rims. As the films on plagioclase^quartz boundaries
are commonly adjacent to reacting biotite grains, these
films probably represent a propagating reaction front. The
angle at propagating fronts is likely to be low and so melt^
solid equilibration, at least in the immediate vicinity of the
tip, would require relatively little adjustment. The absence
of isolated K-feldspar lenses on plagioclase^quartz grain
boundaries (see Fig. 1b) demonstrates that equilibration of
the films, either melt^solid or solid^solid, did not occur.
The 908 peak, which represents the evolution of
reaction-controlled geometries with an initial 1808 angle,
is below that expected for solid^solid textural equilibrium
(with a median of 1208), suggesting that these angles
were adjusting towards melt^solid equilibrium. The angle
population comprises a significant proportion of angles
between the two main peaks and these generally correspond to melt film geometries that appear to follow original grain boundaries, where the initial angle was likely to
be in the range 1208 551808 (Fig. 5).
Melt-filled pores pseudomorphed by K-feldspar in the
Ashuanipi granulite thus appear to record a snapshot of
melt geometries preserved at or close to the metamorphic
peak, with only minor modifications in the sub-solidus.
Melting occurred on grain boundaries between reacting
phases, concentrated near biotite grains and propagating
out along plagioclase^quartz grain boundaries. Reaction
occurred sufficiently fast that equilibration of the microstructures was not possible apart from perhaps establishment of the equilibrium dihedral angle in the immediate
vicinity of pore corners that already had low angles.
The geometry of the pseudomorphed melt in the
Opatica amphibolite-facies migmatite is very different
from that in both the Biwabik chert and the Ashuanipi
granulite. Although thin cuspate interserts are common,
much of the melt has been replaced by poly-mineralic
aggregates on grain boundaries (Fig. 15a and b). These are
highly reminiscent of the string-of-beads texture shown in
Fig. 6d, and we suggest that their coarser grain size and
more equilibrated texture (i.e. straight grain boundaries
and 1208 triple junctions) are due to annealing and equilibration in the sub-solidus. The melt films in the Opatica
migmatite that solidified as poly-crystalline aggregates are
significantly thicker than those replaced by microcline
alone (Fig. 15a and b), reflecting the greater ease of nucleation within the larger pores.
The population of angles preserved at the corners of
microcline-filled melt pores in the Opatica amphibolitefacies migmatite shows the same peak at 908 seen in the
1359
JOURNAL OF PETROLOGY
VOLUME 49
Ashuanipi sample, but the lower peak has a median of 608,
instead of 208. The large driving force for grain boundary
migration near the tips of pseudomorphed reactiongenerated melt films permitted substantial approach to
solid^solid textural equilibrium, whereas the inherited
angles at the terminations of intra-phase boundaries at
the walls of thick melt films did not evolve as much,
because of their generally higher angles.
The relatively uncommon, isolated, pseudomorphs filled
by quartz and plagioclase have very different shapes from
those filled by microcline, with angles close to, or in subsolidus textural equilibrium (Fig. 15b). This is probably
because the adjacent grains are quartz and plagioclase
and so microstructural adjustment towards lower energy
configurations requires only mass movement across the
grain boundary (Hunter, 1987). For the microcline (and
the K-feldspar in the Ashuanipi granulite) there is no corresponding K-feldspar in the pore walls and so sub-solidus
textural equilibration is much slower. The important corollary of this is that recrystallization and sub-solidus equilibration of inherited melt microstructures is dependent on
the minerals forming the pseudomorphs. They are likely
to retain a shape that is recognizable as a pseudomorph of
a melt-filled pore only if they are formed of the minor
component.
The three samples thus show a range of histories
recorded by the geometry of the melt-filled pores. The
Biwabik rock is likely to have attained complete melt^
solid textural equilibrium, not only with equilibrated
melt^solid dihedral angles but also possibly with an
equilibrated pore geometry on the grain scale. This was
followed by a partial approach to sub-solidus textural equilibrium during cooling. The Ashuanipi granulite records
an almost unmodified reaction texture in which only partial approach to melt^solid equilibrium occurred, and the
Opatica amphibolite records a reaction texture that has
been significantly modified in the sub-solidus.
It is somewhat surprising that textural equilibrium
was not attained, even on the scale of the pore corners,
while melt was present in the granulite-facies regional
(Ashuanipi) migmatite, whereas the Biwabik sample
demonstrates the attainment of melt^solid textural equilibrium during contact metamorphism. The contrast may be
due to the slightly higher temperatures in the Biwabik
sample (49008C compared with 820^9008C) or perhaps
to the relative mineralogical simplicity of the Biwabik migmatite. Additionally, the quartz grains in the Biwabik rock
record a complex thermal history and this may have contributed to a prolonged period of (fairly) constant high
temperature during which reaction was slowed and melt^
solid equilibration was facilitated.
The two regional terranes studied here record very
different extents of sub-solidus textural equilibrium. This
difference may have been a function of thermal history,
NUMBER 7
JULY 2008
with the Ashuanipi granulite cooling much more rapidly
than the Opatica rocks as a consequence of their very different tectonic settings. This is consistent with the presence
of microcline in the latter, and untwinned K-feldspar in
the former. The Opatica Subprovince is thought to be the
deep roots of a mountain chain whereas the Ashuanipi
Subprovince is an ancient accretionary prism. It is highly
likely that the Ashuanipi samples experienced rapid exhumation as a result of tectonism and active faulting in a
terrane that remained at, or close to, a cratonic margin,
whereas the Opatica rocks could have been exhumed only
by erosion.
However, a further possibility is that the difference in
sub-solidus textural equilibration was caused by the difference in melting reactions (with the Ashuanipi terrane
undergoing dehydration melting whereas melting of the
Opatica rocks was probably fluxed by externally derived
H2O). The presence of H2O is known to facilitate textural
equilibration in the sub-solidus (Holness et al., 1991) and it
is possible that sufficient free H2O was present after melting in the Opatica rocks, whereas the grain boundaries in
the Ashaunipi rocks remained dry. A much slower rate of
sub-solidus textural maturation for the Ashuanipi rocks
would avoid the necessity for an abrupt decrease in temperature coinciding with the almost complete depletion of
K-feldspar that is required to preserve reaction textures in
a residual gneiss. Given the potential of microstructural
analysis for the decoding of migmatite history it is important that this question is addressed in future work.
S U M M A RY A N D C O N C L U S I O N S
Comparison of reported microstructures in migmatites
from a variety of crustal environments shows that the
time scale of metamorphism exerts a crucial control on
melt distribution and on the microstructures formed
during solidification. Pseudomorphing of pores to form
cuspate grains and pockets on grain boundaries generally
necessitates temperature^time histories sufficient to permit
melt migration. This is because pore pseudomorphing is
dependent on there being a barrier to nucleationçsuch a
barrier occurs in a small pore.
Observation of the distribution and shape of pseudomorphed pores in migmatites from a range of crustal
environments demonstrates that the balance between textural equilibration, melting and deformation is only rarely
in favour of textural equilibrium in melt-bearing rocks,
even in regionally metamorphosed granulites. Observed
low dihedral angles only rarely form populations consistent with melt^solid equilibrium (e.g. Rosenberg & Riller,
2000), and even where melt^solid equilibrium is achieved
at pore corners, the grain-scale distribution may still be
dominated by grain boundary melt films rather than channels on three-grain junctions.
1360
HOLNESS & SAWYER
PSEUDOMORPHING OF MELT-FILLED PORES
Because the smallest pores are those that require the
greatest supersaturation, these will be the last to solidify.
The minerals filling the smallest pores will thus have
more evolved compositions than grains of the same
mineral in larger pores. A further consequence of the
later in-filling of the smallest pores is that these will show
the lowest solid-state dihedral angles. This is because they
will have had a shorter time in which to undergo solid-state
textural maturation compared with the earlier-formed
interstitial grains in the larger pores (see Holness et al.,
2005b). The control of pore size on solidification kinetics
means that, given the same rate of cooling, a migmatite in
which the melt is distributed in many small pores will take
longer to solidify than one in which the same amount of
melt is contained in fewer, larger, pores. This will affect
the rheological response to cooling.
AC K N O W L E D G E M E N T S
We are grateful to Stephen Reed for assistance with the
CL and for suggesting that we examine Ti in quartz.
Chris Hayward is thanked for help with microprobe analyses. Helpful comments from Jon Blundy, Jean-Louis
Vigneresse, Richard White and Gary Stevens improved
and clarified the manuscript.
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