Introduction
The brittle–ductile transition zone (BDTZ) represents the strongest part of
the Earth's crust (Kohlstedt et al., 1995),
the main seismogenic layer
(e.g. Sibson, 1982;
Scholz, 2007), and is a major source and transport region for ore-forming
fluids (e.g. Kolb et al.,
2004). However, it is also the least-understood part of the continental
crust, where the rheological strength estimates and assumptions of rock
deformation mechanisms vary widely.
BDTZ is defined as a transitional layer between the pressure-dependent
brittle rheology of the upper crust and thermally activated viscous creep in
the lower crust (Handy et al., 2007). Highly
localized shear zones control deformation at this depth. The strength of
BDTZ is often estimated using power-law rheology of quartz, which is the
weakest and most abundant phase in granitic assemblages. Experimental data
demonstrate that power-law creep in quartz can be activated at temperatures
as low as 300 ∘C, while feldspar, being another abundant mineral
in granitoids, has a high frictional strength up to temperatures of 500 ∘C (Passchier and Trouw, 2005).
However, in nature, brittle–ductile shear zones are often found to be
altered by syn-tectonic fluids
(e.g.
McCaig et al., 1990; Goncalves et al., 2012; Wintsch and Yeh, 2013). The
chemical and physical consequences of fluid–rock interaction have been
subject of many studies demonstrating the major effects fluid has on rock
rheology. Infiltration of pressurized fluid can cause brittle failure even
at high confining pressures (Byerlee, 1990), allowing
frictional deformation at high temperatures and low differential stresses.
An increasing number of field studies recognize fluid-induced brittle
precursors as the main cause for shear zone nucleation in the middle crust
(Pennacchioni
and Mancktelow, 2007; Fusseis and Handy, 2008; Menegon et al., 2008; Menegon
and Pennacchioni, 2010; Kilian et al., 2011; Brander et al., 2012). The
fluid presence in brittle–ductile shear zones is typically associated with a
variety of weakening mechanisms: (1) chemical breakdown of feldspars into
weak hydrous minerals
(White
and Knipe, 1978; Hippertt, 1998; Oliot et al., 2014); (2)
reaction-controlled grain size reduction by growth of fine-grained
metasomatic assemblages
(White and
Knipe, 1978; Kilian et al., 2011); (3) switch of the dominant deformation
mechanisms from solid state power-law rheology to fluid-facilitated
Newtonian flow
(Menegon et
al., 2008; Wintsch and Yi, 2002; Brander et al., 2012); and (4) hydrolytic
weakening in minerals deforming by crystal-plastic mechanisms
(Tullis and Yund, 1980;
Karato et al., 1986). Based on these observations the strength of the BDTZ
is suspected to be orders of magnitude lower compared to the estimates from
the quartz-based rheology
(Bos
and Spiers, 2002; Mariani et al., 2006; Park et al., 2006; Marsh et al.,
2009). However, the particular role and contribution of each weakening
mechanism is still under debate.
Another incompletely understood problem is the mechanism of the generation
of fluid pathways in middle crust. As pointed out by
Mancktelow (2006), the fact that fluid flows
into and along shear zones, rather than being expelled, requires a brittle
component and local pressure drops during deformation. This conclusion has
led to a concept of a dynamic porosity suspected to exist even at high
lithostatic pressures. The recent work by Fusseis and collaborators
(Fusseis et al., 2009; Menegon et al.,
2015) highlights the importance of
deformation in maintaining fluid pathways in the crust. At the same time,
the study by Billia et al. (2013)
provides potentially direct observations of a significant intergranular
porosity within brittle–ductile shear zone fabrics, suggesting high dynamic
permeability during deformation.
In many cases the difference between fluid-abundant and fluid-limited/absent
conditions is difficult to assess in natural shear zones. Often only one of
the two scenarios is preserved. This study takes advantage of the occurrence
of fluid-facilitated and fluid-restricted deformation within the same
brittle–ductile shear zone, providing the opportunity to discern the
particular microstructural changes and rheological effects caused by
syn-tectonic fluids. Our results demonstrate the microstructural and
rheological difference in each case and reveal the interplay between the
dynamics of fluid flux, deformation mechanism and strain localization within
the middle crust.
Geological setting
Regional geology
The studied shear zone belongs to the Wyangala shear zone system, which is
developed within the Wyangala batholith, situated in the Eastern Lachlan
Fold belt (Fig. 1), one of the three structural regions in the Lachlan
orogeny, SE Australia (Gray, 1997). The Eastern Lachlan
Fold belt consists of voluminous granitic to granodioritic plutons and mafic
volcanics, quartz-rich turbidites, carbonates and shales
(Vandenberg and Stewart, 1992). The regional
tectonic reconstructions suggest that the Lachlan orogeny formed during an
accretion of a volcanic island arc along the eastern margin of Gondwana 450
to 340 Ma in a back-arc or fore-arc basin during a rifting phase in
the Ordovician (Foster et al., 2009).
The Eastern Lachlan subprovince displays characteristic N–S-trending,
eastward-dipping fault-thrust systems
(Glen, 1992; Gray,
1997). The plutons are elongated parallel to these fault systems. Narrow
contact aureoles in the host rock and undeformed margins of the plutons are
considered to support a passive emplacement along pre-existing fault planes
(Paterson and Tobisch, 1992; Paterson
et al., 1990). The crystallization ages of the
plutons in the Eastern Lachlan province are estimated from geochronological
studies and show a range between 435 and 425 Ma
(Lennox et al.,
2005; Squire and Crawford, 2007). The recent zircon Sensitive High-Resolution Ion Microprobe (SHRIMP) U–Pb study by Lennox
et al. (2014) dates the crystallization of the Wyangala
granite as late Silurian (425.2 ± 3.5 Ma).
Smaller-scale shear zones overprinting the major fault system are seen along
the eastern margins of the granitic plutons. These are westward-dipping,
generally N–S-trending and indicate a west-over-east sense of shear
(Paterson et al., 1990). The Wyangala shear
zone system belongs to one such structure, located on the eastern margin
of the Wyangala granite. The main shearing event for these shear zones is
estimated by the Ar–Ar method using recrystallized K-feldspar
(Lennox et al., 2014) to have occurred 375–365 Ma,
corresponding to the late Devonian Tabberabberan deformation event.
Locality map showing the position of Wyangala area in a regional
context (inset) and the location of the studied outcrop on a simplified
geological map (modified after Czarnota, 2002).
Outcrop and general sample description
The study area is located in the vicinity of Wyangala Dam about 25 km SE of
Cowra, NSW (33∘56.855′ S and 148∘57.982′ E; Fig. 1). The
outcrop exposes a weakly to strongly deformed granitic massif. The wall rock
is a weakly foliated monzogranite with up to 8 cm large feldspar
phenocrysts, quartz filling the interstices and biotite marking the
incipient, discontinuous foliation planes (Fig. 2c).
Based on the structural and mineralogical properties we distinguish 2
domains within the studied shear zone: (1) orthogneiss in the shear zone
margins and (2) phyllonite in the central part. The transition from the wall
rock into orthogneiss occurs over a 0.5–1 m distance, characterized by a
gradual decrease in grain sizes and a development of a thin-spaced,
continuous foliation, defined by phyllosilicates. Feldspar grains become
more fragmented towards the shear zone centre, and quartz is more and more
arranged in elongated lenses or layers (Fig. 2a, c).
The boundary between the orthogneiss and phyllonite is sharp and marked by a
sudden strengthening of foliation and disappearance of feldspar grains (Fig. 2a, c). The central domains display centimetre- to millimetre-scale
alteration between two types of phyllonites (Fig. 2b). Phyllonite A is
composed of fine, muscovite-dominated matrix surrounding large, elongated
quartz grains. These large quartz grains in phyllonite A have similar sizes
to quartz in the wall rock. Phyllonite B contains more quartz than
phyllonite A but has a larger matrix mode and exhibits higher phase mixing
(Fig. 2c).
In total 9 samples were collected from all distinguished lithologies,
representing wall rock, orthogneiss and both types of phyllonite.
Methods
Sample selection and preparation
Samples were cut perpendicular to the foliation (yz plane) and parallel to
the stretching lineation (x axis) and polished down to ∼ 30 µm thickness for thin sections. In total 16 thin sections were
prepared for detailed optical and chemical analysis. For electron
backscatter diffraction (EBSD) analysis a colloidal silica–water solution
(mixed in proportion 80 : 20) was used at the final stage of polishing for 3–5 min to reduce surface damage produced by mechanical polishing. All
thin sections were carbon-coated for scanning electron microscopy (SEM)
analysis.
In order to ensure comparability, all quantitative analysis, including X-ray fluorescence (XRF),
EBSD, point counting and image analysis were carried out on the same 4
selected samples, representing each shear zone domain (W17 – wall rock;
W13b – orthogneiss; W21b – phyllonite A; W21c – phyllonite B). Their
locations are given in Fig. 2a, b.
The mineral abbreviations used in the following sections follow the
recommendations of Kretz (1983).
Mineral chemistry and cathodoluminescence (CL) imaging
The chemical composition of the rock-forming minerals was determined using
energy dispersive spectrometry (EDS) on a Carl Zeiss IVO SEM at the
Geochemical Analysis Unit (GAU, Macquarie University), using the AzTec
software from Oxford Instruments. The analytical accuracy of anhydrous
phases, such as feldspars, is 0.1 to 0.2 wt % . The analyses were performed
at high-vacuum conditions with an accelerating voltage of 15–20 kV, with a beam
current of 5.0–10.0 nA at working distances from 12 to 12.5 mm. The same
operating conditions, except for the accelerating voltage which was kept at
15 kV, were used for cathodoluminescence (CL) imaging of quartz
microstructures.
Outcrop and hand specimens. (a) Photo of the studied outcrop
showing the different shear zone domains and collection sites of four studied
representative samples (W17; W13b; W21b; W21c), which were further used for
all the quantitative analysis; (b) fine-scale alternation between
phyllonite A and phyllonite B in the central part of the shear zone; (c) sequence of representative samples collected in a transect across the shear
zone.
Orientation analysis and data processing
Crystallographic orientation data were acquired using HKL NordlysNano
high-sensitivity Electron Backscatter Diffraction (EBSD) detector and indexed
with AzTec analysis software (Oxford Instruments) at the Geochemical
Analysis Unit (GAU, Macquarie University). The analyses were carried out on
a sample tilted to 70∘ angle, in high-vacuum conditions with 20 kV
accelerating voltage and a beam current of 8.2 nA, at working distances from
9 to 13 mm. The typical step sizes ranged from 1 to 5 µm, depending
on the required resolution and the average grain size in the region of
interest. Simultaneously with EBSD data, EDS data were collected to assist
with the phase determination during post-processing of the acquired data.
Channel 5 analysis software from HKL Technology was used for the
post-acquisition processing of the stored EBSD patterns. The obtained EBSD
maps contained from 5 to 26 % of non-indexed points (zero solutions),
mostly resulting from the difficulty to index phyllosilicates. The map
quality was first improved by a “standard” noise reduction following the
procedure of Prior et al. (2002) and Bestmann and
Prior (2003). The second step included the
removal of “fake grains” (all grains with area smaller than a squared
step size), where grains were determined by a minimum grain boundary angle
of 10∘ and subgrains were defined by a boundary angle of
2–10∘ in intra-grain regions. After the processing procedure,
indexing in all maps except for one exceeded 80 % and was close to 100 %
for quartz-dominated areas.
Pole figures were calculated using one point per grain, and plotted on the
upper-hemisphere equal-area projections with stretching lineation parallel
to the x axis and foliation normal to the z axis. All maps and data sets were rotated
consistent with a dextral shear sense as determined by asymmetry of quartz
orientations. In the field it was not possible to determine shear sense
unequivocally.
Modal composition
The modal amounts of mineral phases (Fig. 3a) were determined using
the point-counting method for the four representative thin sections. A minimum of 1000
points for each thin section were counted. Modal amounts of quartz
microstructures (Fig. 3b) were estimated on the same four thin sections
using manually outlined optical micrographs and the imaging software Image J
(http://imagej.nih.gov/ij/index.html). As one quartz
microstructure (Qtz3; details in the following) is very fine grained
and occurs in mixtures with muscovite, it could not be quantified directly
and thus was estimated by subtracting the sum of the other two quartz
microstructure modes (Qtz1 + Qtz2; details in the following)
from the total quartz mode determined by point counting.
Whole-rock analysis
Bulk whole-rock major-, minor- and trace-element concentrations were
measured by XRF analysis for the four chosen
representative samples from each shear zone domain (Fig. 2a). Samples were
processed into fine powders and analysed with a PANalytical PW2400
Sequential WDXRF Spectrometer at the University of New South Wales. Obtained
data sets are presented in Table A1.
(a) Measured modal mineral compositions of each shear zone domain
(Kfs – K-feldspar; Pl – plagioclase; Bt – biotite; Ms – muscovite; Qtz
– quartz; Ep – epidote). (b) The estimated modal amounts of different
quartz microstructures in each shear zone domain (see text for details).
Isocon method
To quantitatively evaluate the mass transfer between the shear zone domains,
we used the isocon method by Grant (1986) based on
the composition-volume equations proposed by Gresens (1967). It allows estimating the absolute change in
the concentration for each individual oxide using the following equation:
ΔMj=CipCiaCjaCjp-1×100,
where ΔM is the mass change in percent; (C) is concentration;
superscripts (p) and (a) stand for protolith and altered sample; subscripts
(i) and (j) stand for immobile and mobile elements.
The slope of the isocon (S=Cia/Cip reflects the
total mass difference between each two analysed samples. The absolute change
in the total mass (in %) thus was estimated using the equation
ΔMtotal=1S-1×100.
For the isocon construction and calculations we assumed the immobility of
aluminium (Al2O3) as (i) it has been shown to be immobile during
the deformation of granitoids in greenschist facies shear zones
(Eilu et al.,
2001; Rolland et al., 2003) and (ii) it is consistent with microstructural
observations in the studied samples.
Results
Optical microstructures, phase abundance and mineral chemistry
Wall rock: weakly foliated monzogranite
The wall rock consists of 34 % quartz, 25 % plagioclase, 21 %
K-feldspar, 14 % biotite, 5 % muscovite and < 1 % accessory
minerals (zoesite, apatite, chlorite, zircons and Fe–Ti oxides) (Fig. 3a).
The most typical quartz microstructure in the wall rock, further referred to as
Qtz1 (Fig. 3b, 4), is characterized by up to 5 mm long grains with
anhedral, equidimensional or slightly elongated shapes having aspect ratios
of ∼ 1 to 1.5. Intracrystalline deformation features such as
undulose extinction and deformation lamellae are common in these grains;
subgrains occur occasionally. Quartz additionally occurs as fine-grained,
largely monocrystalline aggregates associated with Qtz1 grains (Fig. 4a). These microstructures are referred to in the following as Qtz2. In
the wall rock Qtz2 constitutes only a minor amount (Fig. 3b).
Plagioclase and K-feldspar in the wall rock typically occur as 5–8 mm large,
euhedral or subhedral grains, referred to as Pl1 and Kfs1,
respectively (Fig. 4). Pl1 displays normal compositional zoning
characterized by Ca-enriched cores (ab59-81) and Na-rich rims
(ab92-94). The cores are heavily altered to fine-grained muscovite,
K-feldspar and epidote, while the rims rarely contain inclusions. Intergrown
clusters of two or three Pl1 grains are common. Euhedral Pl1 grains occasionally occur in the interiors of larger Kfs1 grains. All
Kfs1 (or96-100) are perthitic, with albite (ab96-98) lamellae
covering about 20 % of the grain area (Fig. 4b, c). These grains often
exhibit fragmentation and pull-apart fractures filled with fibrous or blocky
quartz and biotite. The individual fragments of Kfs1 however are not
significantly displaced with respect to each other.
Biotite (Bt1) is the dominant phyllosilicate in the wall rock
assemblage (Fig. 3a) and occurs as up to 5 mm long grains (Fig. 4a). Rarely,
up to 3 mm long muscovite (Ms1) grains can also be seen in the wall
rock (Fig. 4c). Typically phyllosilicates display well-developed cleavage
and minor kinking and fracturing. Apatite and zircon occur as inclusions in
Bt1 grains.
Typical microstructures and mineral assemblages in the wall rock;
scale bar for (b)–(c) is 200 µm. (a) Optical micrograph (crossed
polarizers with gypsum plate inserted) showing undulose extinction in quartz
(Qtz1), transecting Qtz1 grains, euhedral feldspar grains
(Pl1, Kfs1) and Bt1. Arrows points to developing Qtz2
domains, which often associate with Qtz1 grains; (b) BSE image showing
compositional zonation in plagioclase Pl1 and albite exsolution
lamellae in Kfs1; (c) BSE image showing the microstructure of the
igneous phyllosilicates (Bt1 and Ms1) characterized by large
cleaved and kinked grains.
Shear zone margin: orthogneiss
The orthogneiss is composed of 46 % quartz, 18 % muscovite, 13 %
plagioclase, 11 % K-feldspar, 7 % biotite and 5 % epidote (Fig. 3a).
Up to 70 % of quartz in this domain (Fig. 3b) is represented by the
fine-grained Qtz2 aggregates (Fig. 5a, c, f). The average sizes of
Qtz2 grains are between 15 and 25 µm; the grains have aspect ratios of
about 1.8 and display a shape-preferred orientation (SPO). Qtz1-type
grains are rare (∼ 8 % of all quartz) and commonly
surrounded by mantles of Qtz2 aggregates. Qtz1 grains are
typically up to 1 cm long, with aspect ratios between 3 and 4, and display
undulose extinction, subgrains, slightly serrated boundaries and
intragranular bands and patches of Qtz2 aggregates (Fig. 5a, f).
Another quartz microstructure referred to in the following as Qtz3 can
be distinguished in the orthogneiss (Fig. 3b). It is characterized by very
fine grain sizes and often intermixed with fine-grained muscovite (Fig. 5a, d, e). Here, it is mainly seen in strain shadows of feldspar grains.
Typical microstructures and mineralogy in orthogneiss; scale bar
for (b)–(f) is 200 µm. (a) Optical micrograph (crossed polarizers with
gypsum plate inserted) showing large fractured feldspar grains (Kfs1
and Pl1); undulose extinction in quartz (Qtz1); fine-grained,
monomineralic domains of Qtz2; and thin, muscovite-rich mantles
surrounding Kfs1 and Pl1. (b) Optical micrograph (crossed
polarizers) showing fine-grained muscovite (Ms2) corona around
plagioclase (Pl1). (c) Optical micrograph (crossed polarizers) showing
book-shelf fractures in K-feldspar (Kfs1) and muscovite (Ms2)
aggregates associated with fractures and grain boundaries. (d) EDX- derived
compositional map showing Kfs1 grain with albite exsolution lamellae
and Pl2, Qtz3, Kfs2 and Ms2 in the strain shadow, where
the small plagioclase grains have variable compositions typically having
oligoclase cores and albite rims. (e) BSE image showing Pl1
porphyroclasts with Qtz3 and Kfs2 in the strain shadows and
Ms2 forming strain caps. (b) Optical micrograph (crossed polarizers)
showing Qtz1 grain surrounded by Qtz2 aggregates. Qtz1
displays undulose extinction and subgrains exhibiting similar sizes to
Qtz2 grains.
Both plagioclase and K-feldspar form up to about 3 mm long grains with
aspect ratios of 1.5–1.8. Due to the microstructural and chemical
similarity with feldspars in the wall rock, we refer to them as Pl1 and
Kfs1. Both, Pl1 and Kfs1 occur in association with
fine-grained muscovite (Ms2) mantles (Fig. 5a–e). While Pl1
grains tend to have rounded and irregular shapes, Kfs1 forms clusters
of angular grains and display intragranular fractures, rotated fragments,
bookshelf structures and pull-aparts. The fractures and strain shadows of
feldspar grains are filled with a non-perthitic K-feldspar, further referred
to as Kfs2; fine-grained albitic plagioclase (ab95-99), referred to as
Pl2; fine-grained quartz (Qtz3) and muscovite (Ms2) mixtures
(Fig. 5d, e).
Phyllosilicates in the orthogneiss are mostly represented by Ms2 (Fig. 3a). Biotite and epidote are commonly present in the Ms2-rich bands but
are not in direct contact with feldspar grains.
Shear zone centre: phyllonite A
The mineralogy of phyllonite A is dominated by quartz (62 %),
muscovite (25 %) and epidote (11 %), with minor amounts of plagioclase
(∼ 1.7 %) and biotite (∼ 1.4 %) (Fig. 3a).
Qtz1 grains are larger than in the orthogneiss, approaching sizes
comparable to Qtz1 in the wall rock (Fig. 6a). They are typically
elongated with aspect ratios close to 4 and length up to 8 mm.
Intracrystalline deformation structures – including undulose extinction,
subgrains and deformation lamellae – are common. Fine-grained aggregates of
Ms2, Qtz3, Ep2 and Pl2 form thick layers around
Qtz1 grains. In some parts this matrix can consist of up to 70 % of
very fine Qtz3-type grains with average grain sizes of about 5–8 µm and aspect ratios of 1.8–1.9. The tails of Qtz1 grains commonly
consist of mixtures of fibre-shaped Qtz3–Ms2 (Fig. 6d).
The monomineralic, fine-grained aggregates of Qtz2 in phyllonite A
constitute only 13 % of the total quartz mode, in contrast to the high
abundances in the orthogneiss (Fig. 3b). They show a similar relationship to
Qtz1 grains as in orthogneiss but do not form thick mantles.
Feldspars are represented by up to 1 mm large Na-rich (ab97-100)
plagioclase (Pl2) grains (Fig. 6a, c). They are commonly surrounded by
Ms2-rich matrix and epidote clusters and display well-pronounced,
parallel fracture sets filled by Ms2. Some of the Pl2 grains are
twinned and have undulose extinction.
Typical microstructures and mineralogy in phyllonite A and B;
scale bar for (b)–(g) is 200 µm. (a)–(b) Optical micrographs (crossed
polarizers with gypsum plate inserted), (c)–(f) BSE images and (g) optical
micrograph with crossed polarizers; (a) overview of microstructures in
phyllonite A; (b) overview of microstructures in phyllonite B; (c) fractured
plagioclase grain surrounded by mantle of muscovite, epidote and quartz;
(d) quartz grain having irregular boundaries with muscovite (noted with arrows)
and mixed quartz–muscovite layer in the strain shadow; (e) the extensional
site between two Qtz1 grains filled by mixed quartz–muscovite aggregates;
(f) Qtz1 grains in muscovite–quartz matrix, showing the tendency of
monomineralic muscovite layers to occur parallel to the shear direction and
quartz–muscovite mixtures to occur in the strain shadows; (g) fine-grained
matrix in phyllonite B showing variations of the quartz to muscovite modal
amounts on the millimetre scale.
(a) GB (grain boundary) + IPF (inverse pole figure) map of a
Qtz1 grain from orthogneiss (acquisition location marked in Fig. 5a), scale
bar 200 µm. Subgrain boundaries are marked as grey lines, grain
boundaries are black and Dauphine twin boundaries are red. White pixels are
non-indexed points or other phases. (b) Pole figure plots showing the
crystallographic orientations of all analysed Qtz1 grains (top) and
internal misorientation angle distributions of each pixel in the map
(a) (bottom). (c) Misorientation profiles across two selected regions in
the map (a), where the profiles are marked with white lines and the white circle
represents starting position.
Shear zone centre: phyllonite B
Samples from phyllonite B are mainly composed of quartz (83 %) and
muscovite (15 %), with small amounts of epidote (∼ 1.5 %),
biotite (∼ 0.3 %) and minor plagioclase (∼ 0.1 %) (Fig. 3a).
The microstructure in phyllonite B consists of centimetre- to millimetre-scale alternating
bands of mixed, fine-grained aggregates of Qtz3–Ms2. Variations in
the modal amounts of the two phases characterize the layering (Fig. 6b, g).
Ms2 grains are typically 10–20 µm in length, aligned subparallel
to the main foliation and homogenously mixed with Qtz3. Muscovite-poor
layers display typical Qtz1–Qtz2 associations, similar to the ones
in orthogneiss, with elongated up to 5 mm long Qtz1 grains in the
central parts, surrounded by mantles of the fine-grained Qtz2. In
muscovite-rich domains, very fine grained Qtz3 is homogenously
intermixed with fine-grained Ms2. The grain size of Qtz3 depends
on the Ms2 amount present (Fig. 6g). It varies from 8 µm in
domains with 30 % muscovite to 20 µm in domains with 5 % Ms2.
(a) GB (grain boundary) + IPF (inverse pole figure) map of
Qtz2 microstructures (acquisition locations marked in Figs. 5a, 6b),
scale bars 200 µm. (b) Pole figures of the crystallographic orientation
data for map 1 and map 2, plotted as one point per grain using equal-area
projection and the upper hemisphere. The white dot in the pole figure for
map 1 marks the orientation of the Qtz1 grain, which occurs at the
central part of the Qtz2 domain. (c) Misorientation profiles across two
selected regions in map 2, where the profiles are marked by black lines
and the circle represents starting position.
(a) GB (grain boundary) + IPF (inverse pole figure) map of
Qtz3 microstructures in phyllonites (acquisition locations marked in
Figs. 5a, 6b). Map 2 is separated in two areas for pole figure plots. Area
(i) represents quartz-rich part (15 % muscovite), while area (ii) has
large muscovite mode (30 % muscovite). Map 3 is an enlarged image from map
2b, showing subgrain-scale microstructures; scale bars are 200 µm long.
(b) Pole figures of the crystallographic orientation data for map 2(i) and map
2(ii), plotted as one point per grain using equal-area projection and the
upper hemisphere. The white dot in the pole figure for map 1 marks the
orientation of the Qtz1 grain, which occurs at the central part of
Qtz2 domain. (c) Misorientation profiles across two selected regions in
map 1 and map 3, where the profiles are marked by black lines and the circle
represents starting position.
Isocon diagrams (Grant, 1986) showing
major element concentrations for (a) orthogneiss,
(b) phyllonite A and (c) phyllonite B
plotted against the wall rock composition. The dashed line
represents a situation of a zero mass change. The isocon (solid line) is
constructed assuming immobile Al2O3. Elements enriched in the
altered domains lie above the isocon; the elements which are depleted lie
below. Scaling factors (shown in front of each element) have been applied to
avoid data clustering. The detection error is smaller than the diameter of
the data points (cf. Table A2).
Crystallographic orientation analysis
Crystallographic preferred orientation (CPO) analysis focuses on quartz
microstructure. No EBSD analysis is necessary for muscovite, where c axis is
perpendicular to the elongation of the grain, and the majority of the grains
display well-pronounced elongation parallel to the general foliation.
Feldspars show dominantly brittle deformation; thus no specific
crystallographic orientation data were sought for these. In the following, we
characterize in detail the different quartz types as identified from optical
analysis.
Qtz1 porphyroclasts
Figure 7b shows a combined pole plot for six Qtz1-type grains from all
deformed shear zone domains. The bulk orientation of these grains is
non-random, with the poles of c axis clustering 5–45∘ clockwise
from z direction. This arrangement is consistent with the general shear
sense in the analysed samples. In the second pole figure (Fig. 7b), a CPO
pattern of a single strongly deformed Qtz1 grain is presented, showing
the range of internal deformation in the crystal lattice. Two c axis maxima
appear next to the opposite poles of z direction, lying on the same great
circle.
Internally Qtz1 grains show a high degree of lattice distortions, with
high abundance of subgrain boundaries, most of them oriented approximately
at 45∘ angles to the foliation plane (Fig. 7a). The variations of
internal misorientation across a single grain can reach more than
40∘ (Fig. 7c; profile 1). Dauphine twin boundaries, identified by
the 60 ± 2∘ lattice rotation around the c axis, are common
(Fig. 7a). They typically occur as irregular patches in highly deformed
internal parts or along highly stressed edges of the grains.
Qtz2 aggregates
Grain orientations in the fine-grained Qtz2 domains always display a
strong CPO. In the areas surrounding Qtz1 porphyroclasts, the CPO of
Qtz2 grains tends to cluster close to the porphyroclast orientation, with
a slight rotation towards the centre of the pole figure (Fig. 8b; map 1). In
the homogenous Qtz2 domains, devoid of Qtz1 grains, the CPO of
Qtz2 is arranged in an asymmetric Type Ia crossed-girdle pattern
(Passchier and Trouw, 2005), subparallel to z direction and
synthetic with the shear direction (Fig. 8b; map 2).
Subgrain and Dauphine twin boundaries are common (Fig. 8a). The variations
in internal misorientation angles across single grains typically range from
2 to 4∘ but can reach up to 9∘
misorientation in the largest grains (Fig. 8c).
Qtz3 in muscovite–quartz mixtures
Two associations of Qtz3 grains were analysed in phyllonite A and
phyllonite B: one from a tail of a Qtz1-type grain, another from an
interior of a fine-grained quartz–muscovite band (Figs. 6a, 9a). The
crystallographic orientation of Qtz3 grains at the tails of Qtz1
typically cluster close to the orientation of the Qtz1 grain (Fig. 9b; map 1). Often the Qtz3 grains in these tails form string-like
aggregates. Misorientation profiles across individual grains parallel to the
elongation of these “strings” do not exceed 15∘ between adjacent
grains (Fig. 9c; profile 1).
In contrast, Qtz3 grains in the matrix-forming Qtz–Ms bands display
close-to-random CPO (Fig. 9b; map 2ii). Domains with smaller muscovite
content may have similar, or slightly larger, grain sizes but show much
stronger CPO than muscovite-rich domains (Fig. 9b, map 2i vs. map 2ii).
Despite the small grain size (average 8 µm) the individual Qtz3
grains display subgrains, Dauphine twinning and lattice bending, typically
ranging from 1 to 3∘, but angles up to 13∘
are not uncommon (Fig. 9c; profile 2).
CL, EBSD and BSE images of quartz domains in orthogneiss and
phyllonite B; scale bar is 100 µm. (a) CL patterns from a Qtz2
domain showing a structure consisting of lighter polygons surrounded by
darker rims; (b) EBSD image of (a) showing the grain and subgrain boundary
network; (c) CL pattern from a Qtz3–Ms2 domain showing a structure
dominated by thin intragranular bright lines or polygons, bordering or
surrounding darker polygons; (d) area (c) with some of the polygonal
structure traced; (e) BSE image of area (c) showing mineralogy and grain
boundaries.
Cathodoluminescence (CL) microstructures
CL pattern imaging of quartz is a commonly used, powerful technique for
distinguishing multiple growth events and fluid/melt influx events in
igneous (e.g. D'Lemos et al., 1997), metamorphic
(e.g. Bergman and Piazolo, 2012)
and sedimentary rocks (e.g. Demars et
al., 1996). Accordingly, we performed CL imaging to investigate the
potential difference in patterns developed in Qtz2 and Qtz3
microstructure, present in the orthogneiss and phyllonite, respectively. As
shown in Fig. 11, several differences can be recognized.
The quartz in the orthogneiss was analysed from the monocrystalline
Qtz2 domains, which are seen in the tails of the recrystallizing
porphyroclasts. The CL image displays a polygonal structure where the
lighter central areas are enclosed by darker rims (Fig. 11a). Comparison of
the CL image and EBSD map, acquired from the same area, shows that the dark
rims in the CL pattern coincide mainly with grain, and in some cases
subgrain, boundaries (Fig. 11b).
In phyllonite B, quartz was analysed from the fine-grained, mixed
Qtz3–Ms2 domains. In general, the CL pattern for Qtz3 is more
intricate than that of Qtz2. It consists of thin, lighter lines,
cross-cutting darker quartz grains (Fig. 11c). These linear features do not
represent any visible cracks or grain boundaries. Lighter and darker
polygons are also seen adjacent to each other, and the variations in CL
shading do not correspond to grain boundaries. In summary, the variations in
the CL patterns of Qtz3, in contrast to the patterns in Qtz2, are
clearly intragranular in nature (Fig. 11d, e).
Whole-rock geochemistry
Figure 10a shows that the chemical composition of the orthogneiss closely
resembles the composition of the wall rock. The slope of the isocon
approximates 1 and, all the major elements lie near to the isocon. Slight
enrichment in CaO and depletion in Mn3O4, MgO, P2O5 and
K2O may be related to the heterogeneity of the analysed samples.
The element concentrations in phyllonite A deviate from the isocon more
significantly (Fig. 10b; Table A1). SiO2 and FeO are enriched,
reflecting the increased quartz and epidote modes in phyllonite A (Fig. 3a). All the other major elements, especially Na2O, are depleted. The
major loss in Na2O (-79 %) coincides with the decrease in feldspar
mode from 46 % in the wall rock to 1.7 % in phyllonite A (Fig. 3a).
Furthermore, the isocon itself lies below the “constant mass” line (dashed
line in Fig. 10), indicating a total mass increase of 12.2 % compared to
the wall rock domain.
Phyllonite B maintains the chemical trends displayed by phyllonite A but
shows more extreme deviations from the wall rock composition (Fig. 10c;
Table A1). The isocon slope of 0.29 equates to a total mass gain of 239 %.
Large gains in SiO2 (+332 %) concentration correspond to the
increase in quartz mode from 34 % in the wall rock to 83 % in
phyllonite B (Fig. 3a). Na2O again shows depletion; the near-zero
concentration reflects the loss of feldspars in the shear zone centre.
Discussion
Orthogneiss and phyllonite: same protolith, or not?
The studied shear zone displays two mineralogically and structurally
different domains: (1) marginal domains containing orthogneiss, similar in
composition to the wall rock, and (2) central domains with muscovite–quartz
phyllonite exhibiting a significant change in bulk rock chemistry. An
important question regarding the interpretation of the shear zone
microstructures is whether both of these domains originated from the same
protolith, or whether they represent pre-existing heterogeneities in the wall rock, such
as dykes, veins or compositional layering.
The mineralogical and chemical similarity between the orthogneissand the
wall rock, as well as the gradual microstructural and mineralogical change
over a broad transition zone clearly implies a wall-rock-type protolith for
the orthogneiss domains. In contrast, the origin of the central phyllonites
with markedly different chemical and microstructural characteristics is more
ambiguous to interpret. Two scenarios can be envisaged: (i) local alteration
of the wall-rock-type protolith by a metasomatic fluid or (ii) initial
heterogeneity in the wall-rock-type protolith, such as a quartz-rich vein
overprinted by the mylonitic deformation. Although the quartz vein scenario
cannot be completely ruled out, no direct evidence was found to support it.
Phyllonite A, which is still relatively similar to the wall rock
composition, does not display a sharp boundary with the quartz-enriched
phyllonite B (Fig. 2b). Within the wall rock, there is no evidence of
pre-mylonitic quartz veinlets. On the other hand, asymmetric reaction rims,
metasomatic minerals in pressure shadows and dilational sites and
syn-tectonic fracturing of porphyroclasts with associated localized but
discontinuous quartz growth in extensional sites indicate the presence of a
metasomatic fluid during the deformation. This, together with the consistent
trends in chemical and mineralogical changes towards the shear zone centre
(Fig. 10), suggests a metasomatic, rather than a vein-derived, origin for
phyllonites.
Mineral reactions and mass transfer
The metasomatic system in Wyangala shear zones bears many similarities to
the alteration patterns in other fluid-affected greenschist facies shear
zones, where the feldspar breakdown to phyllosilicates and albitization of
feldspars is a consistent trend
(Ramberg,
1949; Kerrich et al., 1980; Pryer and Robin, 1995; Hippertt, 1998; Wintsch
and Yeh, 2013). The general reaction can be summarized as follows:
plagioclase + K-feldspar + biotite + H2O = muscovite + quartz + albite + epidote.
Further below, the particular reactions are interpreted from the specific
mineral relationships observed in orthogneiss and phyllonites. The chemical
formulas for the minerals as used in the reactions are given in Table 1.
Orthogneiss: reactions and local mass transfer in a closed, fluid-limited
system
The mineral composition of orthogneiss differs from the wall rock by
increased modes of muscovite, quartz and epidote, and decreased amounts of
feldspar and biotite (Fig. 3a).
The new metasomatic muscovite (Ms2) occurs as fine-grained aggregates,
aligned subparallel to the foliation fabric. Muscovite-rich bands often
surround or cross-cut the igneous feldspars (Pl1 and Kfs1) and
biotite (Bt1) (Fig. 5b–e). Plagioclase, more often than K-feldspar, is
surrounded by fine-grained muscovite–epidote mantles. Reactions responsible
for these fabrics can be written as
3Kfs1+2H2O→Ms2+6Qtz3+2K++2OH-
and
4Pl1+0.8H2O+0.6K++0.2Fe++2.6Qtz3→3.2Pl2+0.6Ms2+0.4Ep2,
where small amounts of pore fluid, probably provided by
deformation-facilitated breakdown of igneous biotite (Bt1) and
muscovite (Ms1) interact with the metastable igneous feldspars. As a
result Kfs1 (Reaction R1) breaks down to form muscovite (Ms2),
releasing silica and K+ in the solution, both of which together with
small amounts of Fe+, derived from the biotite breakdown, can be
further incorporated in Reaction (R2), where igneous plagioclase of
andesitic composition (Pl1) breaks down to form albitic feldspar
(Pl2), muscovite (Ms2) and epidote (Ep2).
Pryer and Robin (1995) described
a similar multi-stage reaction sequence involved in the albitization of
K-feldspar to form flame perthite.
Abbreviations and chemical formulaes for the minerals as used in the reaction
equations.
Abbreviation
Mineral
Formula
Ep2
epidote
Ca2Fe0.5Al2.5(SiO4)3(OH)
Kfs1, Kfs2
K-feldspar
KAlSi3O8
Ms2
muscovite
KAl3Si3O10(OH)2
Pl1
oligoclase
Ca0.2Na0.8Al1.5Si2.5O8
Pl2
albite
NaAlSi3O8
Qtz3
quartz
SiO2
The presence of intergranular fluids is
necessary for the transport of the reaction components, as the parent
minerals and reaction products often do not occur in a direct contact
relationship. The growth of Kfs2 in pressure shadows and extensional
sites (Fig. 5d, e) especially suggests that material transport is controlled
by a pressure gradient, where the nucleation of new phases preferably occurs
in low-pressure dilational sites. The lack of perthite lamellae in the
metasomatic Kfs2 indicates the immiscibility of K+ and Na+
in the conditions of alteration. Ca content in Pl1 promotes the
formation of epidote (Ep2). The excess Fe+ necessary for Ep2
production is possibly released from Bt1, which is also seen to break
down to Ms2. Quartz released in Reaction (R1) is not seen to intermix
with muscovite but rather precipitate in strain shadows and fractures (Fig. 5e), suggesting high mobility of silica in the circulating fluid.
Indications of Na+ mobility in fluid come from the observations of
albitic rims and veinlets associated with Kfs1 and can be described by
a simple exchange reaction:
Kfs1+Na+→Pl2+K+,
possibly governed by interface-coupled dissolution–precipitation processes
(Putnis, 2009), as a fluid phase must participate in
the reaction as a transport medium for Na+ and K+, as they are not
always available from the surrounding phases (Fig. 5d).
In summary, the microstructures in orthogneiss suggest metastability of the
wall rock assemblage during the deformation. The major reactions are
feldspar-to-muscovite and biotite-to-muscovite breakdown; albitization of
feldspars; and growth of metasomatic K-feldspar (Kfs2), albite
(Pl2), epidote (Ep2) and quartz (Qtz3), indicating a local
mobility of K+, Na+, Ca+, Fe+ and Si+. However, the
bulk rock composition of orthogneiss is highly similar to the wall rock
composition (Fig. 10a) excluding the possibility of a large-scale mass
transfer during alteration. A minor amount of free fluid is required to
enable Reactions (R1) and (R2) to occur and to facilitate the transport of
the components in all described reactions. Thus, the reactions in
orthogneiss are best explained by limited fluid influx and dominance of
local element “recycling” in “closed-system” conditions, where mass
transfer did not exceed distances of millimetres to centimetres.
Phyllonite A and B: reactions and mass transfer in an open, fluid-abundant
system
Both phyllonites display dramatic changes in mineral and chemical
composition with respect to the wall rock (Fig. 10b, c). Mineralogically,
the differences include almost complete loss of K-feldspar, plagioclase and
biotite, as well as increase in muscovite, quartz and epidote (Fig. 3a)
content.
No igneous K-feldspar (Kfs1) or plagioclase (Pl1) remnants were
observed in phyllonites, but the rare albite grains (Pl2) typically
associated with muscovite–quartz–epidote mantles, can either be products of
the breakdown reactions of Pl1 – with Pl2, Ms2, Ep2 and
Qtz3 as reaction products (Reaction R2) – or a result of an
arrested/incomplete breakdown of Pl2:
3Pl2+2H++K+→Ms2+6Qtz3+3Na+,
resulting in the production of muscovite (Ms2) and quartz (Qtz3)
and the release of Na+ in the pore fluid.
Although in some cases Reaction (R2) may be operating, Reaction (R4) is necessary
to explain the eventual disappearance of albite and the general trend of
Na+ depletion indicated by the XRF data (Fig. 10b, c).
Both domains in the central part of the shear zone indicate net mass gains,
mainly due to the increased Si+ concentration. The modal analysis also
shows the increase in the quartz content in both domains, suggesting
external influx of Si+ fluids. No loss of Si+ is observed in the
adjacent orthogneiss, excluding the possibility of a lateral transport.
In summary, the chemical reactions described in this section requires
“open-system” conditions with an influx of external Si+-rich hydrous fluid,
undersaturated in Na+.
Deformation mechanisms and strain localization
Shear zone margins: orthogneiss
Deformation in shear zone margins occurs mainly via crystal-plastic and
cataclastic processes. Ductile deformation is concentrated in quartz, while
feldspar deforms in a brittle manner.
The larger Qtz1 grains, representing porphyroclastic remnants of the
igneous quartz, show abundance of undulose extinction, deformation lamellae
and subgrain boundaries, indicating the operation of dislocation creep
processes (White, 1977;
Poirier, 1980; Urai et al., 1986). The finer-grained Qtz2, often
surrounding porphyroclasts (Qtz1), has similar sizes and shapes to the
subgrains, suggesting formation by subgrain rotation (SGR) recrystallization.
The presence of serrated grain boundaries suggests a component of bulging
(BLG) recrystallization involving low-temperature grain boundary migration
(Hirth and Tullis, 1992), where bulging may have been
partly facilitated by intergranular fluids
(Mancktelow and Pennacchioni,
2004). CL patterns from orthogneiss domains also support the origin of Qtz2
by dynamic recrystallization. The newly developed grain boundaries show a
distinctly different CL signature (Fig. 11), consistent with grain boundary
generation within fluid and trace-element conditions different to that of
the original igneous Qtz1. The fact that Qtz2, further interpreted
as the recrystallized fraction in the orthogneiss, represents close to
70 % of the orthogneiss and forms interconnected, fine-grained layers
suggests relatively low rheological strength of quartz in the particular
assemblage, highlighting the significance of crystal-plastic processes in
the strain accommodation.
Combined EBSD data for analysed Qtz1 grains shows c axes oriented
synthetically oblique to the shear direction. This pattern has been
previously documented to occur in mid-crustal mylonites and interpreted as a
selective preservation of only those grains, which are optimally oriented
for a slip on basal planes (Menegon et
al., 2011).
The CPO of the recrystallized Qtz2 grains either clusters around the
orientation of the adjacent porphyroclast (Fig. 8b; map 1), reflecting a
“parent”-controlled misorientation relationship during a progressive
subgrain rotation
(Kruse et al., 2001; Stünitz et al., 2003), or, in cases when the recrystallized
fraction is higher, displays a hybrid between an asymmetric single-girdle
and crossed-girdle patterns (Fig. 8b; map 2), indicating slip on prism < a >and rhomb < a >crystal systems
(Schmid and Casey, 1986; Law
et al., 1990).
Porphyroclastic remnants of the igneous feldspar phenocrysts (Pl1 and
Kfs1) display clusters of angular, internally fractured and slightly
displaced grains (Fig. 5a, c), indicating deformation via cataclasis. The
relatively small displacement between the individual fragments suggests that
fracturing itself plays a minor role in the accommodation of the finite
strain. The distribution of low-solubility minerals (Ms2) at high-stress
sites between porphyroclasts and perpendicular to the shortening
direction in contrast to the distribution of more soluble phases (Pl2,
Qtz3, Kfs2) in strain shadows (Fig. 5e) suggests that mass
transfer occurred dominantly by pressure solution, involving a fluid phase
(Rutter, 1983; Wintsch and Yi,
2002). Accordingly, elements with higher mobility – such as Si+,
Ca+, Na+ and K+ – were transported by fluid and precipitated in
the locally occurring low-pressure sites, while the elements of lower
mobility, in this case Al+, were immediately incorporated in the
Ms2 structure, which was the new, more stable mineral phase after the
dissolution of Kfs1 and Pl1. It is important to note that the
largest displacement between feldspar fragments occurs where they are
surrounded by the thickest mantles of the reaction products (Fig. 5a). This
observation indicates strain partitioning from feldspars to the newly
created reaction products as soon as a thick-enough mantle is formed.
Shear zone centre: phyllonite
Phyllonite A and phyllonite B in the shear zone core structurally display
many similarities in terms of Qtz1 and Qtz2 microstructures.
Qtz1 grains in both domains are internally deformed and shows
development of subgrain boundaries, while Qtz2 grains have a strong CPO
indicating rotation from porphyroclast orientation towards rhomb < a> and prism < a > slip systems (Fig. 8b; map 1; Schmid and Casey, 1986) with an increasing degree of
recrystallization. The relatively minor fraction of the recrystallized
quartz (Qtz2) in phyllonite A suggests strong strain partitioning
into the surrounding muscovite-rich matrix.
The two main differences between phyllonite A and phyllonite B domains
are (1) quartz-to-muscovite ratio and (2) matrix-to-clast ratio, both of
which are higher in phyllonite B (Figs. 3a, b; 6a, b).
All muscovite in both phyllonites is of metasomatic origin (Ms2),
formed either by a breakdown of igneous biotite or in feldspar reactions
(Reactions R1, R2) and distributed by pressure solution processes, as Qtz3
tends to be associated with extensional sites in strain shadows and
fractures, while muscovite is concentrated in high-stress sites around
porphyroclasts, subparallel to the foliation. The consistent SPO of the
muscovite grains with elongation subparallel to the shear direction thus may
be controlled by syn-tectonic growth in the direction of maximum elongation.
Local variations in muscovite abundance throughout the phyllonites affect
the deformation behaviour of quartz. In the fine-grained muscovite-rich
layers (Ms2=90-30 %) quartz occurs as Qtz3-type grains and
displays a lack of a clear CPO pattern (Fig. 9b; map 2). Qtz3 CPO
becomes increasingly weaker with increased amounts of muscovite (Fig. 9b; map 2i and map 2ii). The lack of CPO coupled with the small grain sizes and
phase abundance is consistent with GBS as the main
deformation mechanism
(Kruse
and Stünitz, 1999; Ree et al., 2005; Svahnberg and Piazolo, 2010). The
elongated grain shapes in these domains and the abundance of
intracrystalline deformation features (Fig. 9a, c) suggest that grain
boundary sliding was accommodated by dislocation glide
(Rybacki
et al., 2010; Svahnberg and Piazolo, 2010).
A special case of Qtz3 microstructures can be seen on the tails of
larger porphyroclasts (Fig. 9a; map 1) where CPO clusters around the
orientation of the adjacent porphyroclasts (Fig. 9b; map 1). This is
consistent with quartz precipitation from pore fluids at low-pressure sites,
where the nucleation of the new grains is host-controlled. Their elongated
shape subparallel to the main foliation suggests anisotropic growth of
minerals commonly reported to occur during pressure-solution processes
(Passchier and Trouw, 2005).
Feldspars are rare in the phyllonite and thus play a minor role in the bulk
rock rheology. The prominent fracture sets filled by fine-grained muscovite
suggest mechano-chemical breakdown processes. The albitic composition of
Pl2 is interpreted as a result of interface-coupled
dissolution–precipitation reactions (Putnis, 2009) as indicated
by arrested replacement structures in the less-altered orthogneiss, where
asymmetric albite rims occur around Kfs1 and Pl1. Simultaneously,
albite grains may also form by neo-nucleation processes from pore fluid
where a chemically different composition is obtained in response to the
disequilibrium state between the deformation conditions/fluid composition
and igneous feldspars. This scenario has been previously suggested by
Stünitz and Fitz Gerald (1993) for
similar shear zones in the Wyangala area. Both of these options indicate
fluid-accompanied mass transfer processes in the shear zone.
Deformation conditions
The deformation conditions within the studied shear zone are estimated from
(1) syn-tectonic mineral assemblages, (2) recrystallization microstructures
in quartz and (3) CPO patterns in the dynamically recrystallized quartz
domains.
The observed reactions (Sect. 5.2) – such as the
breakdown of feldspars and biotite to muscovite, albitization of K-feldspar
and plagioclase, and formation of epidote – commonly occur in many middle
crustal shear zones, deformed at greenschist facies conditions
(Hippertt,
1998; e.g. Kerrich et al., 1980; Park et al., 2006). In combination with the
brittle–ductile rheology expected at middle crustal depths, this gives the
first indication of the deformation conditions.
More accurate temperature ranges can be obtained by analysing deformation
microstructures in quartz. Microstructural features – such as undulose
extinction, subgrain development, bulged grain boundaries and CPO patterns
in the quartz-rich domains – indicate the operation of crystal-plastic
processes (Hirth and
Tullis, 1992; Stipp et al., 2002). Crystal plasticity in quartz occurs by
one of the three mechanisms: bulging (BLG), subgrain rotation (SGR) or
high-temperature grain boundary migration. Which one of these is the most
dominant largely depends on temperature. In our samples, SGR is the
dominant recrystallization regime in quartz. The lobate grain boundaries
observed for the recrystallized quartz grains indicate a minor component of
grain boundary migration typical of the BLG recrystallization regime. In
nature, the transition from BLG to SGR is found to occur at temperatures
between 350 and 400 ∘C
(Stipp et al., 2002). However,
bulging-type microstructures in quartz have been reported to form as a
result of post-deformational activity of grain boundary fluids
(Mancktelow and Pennacchioni,
2004). This imposes some uncertainty on the lower temperature constraint.
Quartz CPO pattern therefore is a better indicator for deformation
temperatures, as it does not strongly depend on the absence or presence of
fluid (Mancktelow and
Pennacchioni, 2004). Our samples show CPO patterns characterized by a
combination of basal and prism slip (Fig. 8b; corresponding to
temperatures between 300 and 400 ∘C; Schmid and Casey,
1986). Other authors interpret similar patterns as being
indicative of deformation at temperatures close to 400 ∘C
(Stipp
et al., 2002; Lee et al., 2012). The latter temperature range is supported
by the observations of frequent Dauphine twins, which in quartz form in
temperatures between 300 and 400 ∘C
(Wenk et al., 2007; Menegon
et al., 2011).
There are no direct indicators for the confining pressure; however, an
approximate estimate based on the obtained temperature range (300 to 400 ∘C), can be drawn using an assumption of a normal geothermal
gradient (30∘ km-1). Assuming the average density of continental
crust (2.8 g cm-3), depths of 11 and 13 km and pressures between 3.02 and 3.57 kbar can be obtained.
(a–c) Schematic illustration of the temporal (stages) and
spatial (cases) fabric development of the studied shear zone; (d) the
proposed mechanism for Qtz3 incorporation in the shear zone fabrics.
Refer to the text for details.
Model for shear zone development
In this section we discuss a conceptual model for the development of the
studied brittle–ductile shear zone. It is important to note that, although
feldspar-to-muscovite reactions in the shear zone centre indicate softening
and thus shear zone narrowing, the strain gradient in the studied shear zone
does not represent a progressive fabric evolution with time as would be
expected for an ideal Type II shear zone (Means, 1995).
As we further discuss, the shear zone margins and shear zone centre
experienced vastly different geological and fluid history due to a brittle
precursor. A hypothetical example of such a case has been theoretically
investigated by Means (1995), demonstrating how the
deformation can proceed by different scenarios inside and outside the
initial zone of cataclasis even when the late fabrics across the strain
gradient show an apparently progressive evolution.
Here we propose a three-stage temporal development of the studied shear zone,
distinguishing between two cases, representing different domains, namely case
I – the shear zone margins – and case 2: the central domains (Fig. 11).
Stage 1: localized cataclasis
We suggest that the first stage of the deformation was marked by high
effective stresses and fluid pressures leading to brittle failure and
formation of a cataclastic fracture zone in the granitic protolith (Fig. 12a). This interpretation is based on the observation of the abundance of
the large, only weakly deformed quartz porphyroclasts, in phyllonite A (Fig. 6a). If shear localization had initiated through crystal plasticity, then
quartz would be highly recrystallized as seen in the orthogneiss (Fig. 5a).
The preservation of the Qtz1 grains can only be possible if the weak
matrix of reaction products in the central parts of the shear zone is
created early in the deformation history and partitions most of the strain
before any significant recrystallization of quartz takes place.
Stage 2: ductile deformation
Stage 2 is marked by ductile deformation in response to the initial
cataclastic stage. Brittle failure (stage 1) resulted in a localized zone of
interconnected fault-fracture mesh as described by
Sibson (1996). The opening in brittle fractures and
dilatancy-related pressure fluctuations around the fractures facilitated and
localized the infiltration of external fluids. Subsequent microstructural
development and rheology in the shear zone was mainly controlled by its
proximity to this initial fracture zone. The sharp boundary separating the
orthogneiss and phyllonites is interpreted to represent the extent of the
initial cataclastic damage zone affected by high fluid fluxes, and the
largely intact crystalline wall rock with restricted pathways for fluid
infiltration (Fig. 12a, b). This distinction is also reflected in
differences in phase distribution (Fig. 3a), microstructures (Figs. 3b, 4, 5)
and whole-rock chemistry (Fig. 10).
Case I in Fig. 12b represents deformation at fluid-limited conditions in
the shear zone margins, outside the initial cataclastic zone (Fig. 12b).
Syn-tectonic retrograde reactions of igneous biotite (Bt1) and
muscovite (Ms1) provide pore fluid which migrates along fractures and
grain boundaries, facilitating microcracking, pressure-solution processes
and mineral reactions to weak fine-grained polycrystalline phases (Fig. 5).
The majority of igneous quartz grains behave as a weak phase, developing
subgrain and new grain boundaries through a progressive distortion of the
crystal lattice. Up to 70 % of the igneous quartz recrystallizes into
fine-grained interconnected aggregates, accommodating most of the strain by
dislocation creep processes, a typical deformation behaviour for
monomineralic quartz domains (Obee and White, 1985;
Fliervoet and White, 1995). Simultaneously, the igneous feldspar phenocrysts
undergo fragmentation by brittle failure (Fig. 5). CL patterns seen in the
dynamically recrystallized Qtz2 in the orthogneiss (Fig. 11a, b) are
consistent with the presented scenario. The dark grain and subgrain
boundaries indicate that recrystallization occurred at fluid conditions
different to the original igneous conditions. The fact that the grain
boundaries are dark is consistent with a syn-tectonic fluid containing CL
signal suppressing trace elements such as Ti or Al
(Rusk et al., 2008; Bestmann and
Pennacchioni, 2015). However the limited amount of fluid available from the
local reservoirs, and possibly the slow rates of its release, do not permit
rapid weakening to take place. Consequently, strain localization is limited
and deformation occurs throughout the rock.
Case II (Fig. 12b) represents deformation at fluid-abundant conditions in
the shear zone centre. Here we assume that fabrics in phyllonite A is an
early stage in the evolution of the central shear zone domains. This
assumption is based on the fact that mineralogically and chemically
phyllonite A is closer to the wall rock composition than to phyllonite B
(Fig. 10b, c). In the shear zone centre, the cataclastic protolith provides
easy fluid pathways, and the dilatational low-pressure sites facilitate the
infiltration of external fluids, leading to widespread and rapid mineral
reactions of the metastable igneous wall rock assemblage. As a result, the
load-bearing feldspar framework is rapidly destroyed and the newly produced
matrix of fine-grained reaction products localizes the strain. The
deformation is mainly accommodated by pressure solution processes and
reaction creep in the fine-grained matrix of the reaction products. As
proposed by Wintsch and Yi (2002),
pressure-solution-controlled reactions can be an efficient deformation
mechanism in middle crust due to the low activation energy required, leading
to significant shape and volume changes even at low differential stresses
(Wintsch and Yi, 2002). Consequently, the
fraction of dynamically recrystallized quartz in phyllonite A is small and
the remnant quartz porphyroclasts preserve sizes similar to the igneous
grains in the wall rock (Fig. 6a).
Stage 3: Continued fluid influx and alteration
Phyllonite B represents phyllonite A fabric, modified by a prolonged fluid
percolation (Fig. 12c). The chemical and mineral evidence of SiO2 gains
in the phyllonites (Figs. 3a, 10c) indicates that a significant amount of
quartz is incorporated during the deformation. We suggest that this “new”
quartz has been added in the phyllonite B microstructure through local
precipitation from a pore fluid in transient dilatational sites, similar to
the creep cavitation process described by Fusseis et al. (2009) and Menegon et al. (2015).
The first stages of the silicification may be related to stress and strain
heterogeneities in the vicinity of porphyroclasts. As seen in phyllonite A,
the initial matrix, resulting from the feldspar and biotite breakdown
reactions, contains little quartz. However, the mixed fine-grained
Qtz3–Ms2 aggregates typically occur in the strain shadows of
Qtz1 and feldspar porphyroclasts, where the rheologic contrast between
the rigid clasts and the soft matrix is high, promoting dilation
(Passchier and Trouw, 2005;
Mamtani et al., 2011). Low pressures in the dilational voids attract pore
fluids and enable the precipitation of material from the supersaturated
fluids (Fig. 12di). Initially the growth of the new material occurs on the
surface of Qtz1 porphyroclasts, as seen by the related CPO patterns of
the adjacent Qtz3 grains (Fig. 9a, b; map 1; Fig. 12dii). With
increasing length of these “tails” and “necks”, Qtz3 grains become
disconnected from the porphyroclasts (Fig. 12diii) and, due to the fine
grain sizes, transit toward the deformation by GBS, as is indicated by the
random CPO patterns (Fliervoet
et al., 1997) (Fig. 9a, b; map2ii). The further fluid flow (Závada et
al., 2007; Rybacki et al., 2008) and silicification is then localized in
these fine-grained quartz–muscovite layers by cavitation creep mechanisms.
When a cavity is created next to an existing quartz grain, the nucleation of
new silica will occur syntaxially due to kinetic advantages in inheriting
the crystallographic orientation (Fig. 12div). Some of the subgrain
boundaries observed in Qtz3 grains (Fig. 9a; map 3) thus may reflect
separate events of silica precipitation, rather than dynamic
recrystallization in Qtz3. As a result the individual grains will
increase in size and the interparticle spacing of the second phases
(Ms2) will also increase, leading to the development of quartz- instead
of muscovite-dominated polyphase layers. CL signature of Qtz3 grains in
these polyphase layers is distinctively different from the dynamically
recrystallized Qtz2 in the orthogneiss (Fig. 11), often displaying
intricate internal patterns consisting of lighter and darker domains and
lines. As the CL signal in quartz is shown to correlate with Ti content and
other trace elements (Rusk et al., 2008), we
suggest that these patterns reflect continuous precipitation of silica from
a slightly changing intergranular fluid.
The Qtz3 grain coarsening may continue until the increasing grain sizes
and quartz mode lead to the switch to dislocation creep mechanisms within
the fine-grained material. The resulting drop in strain rates by switching
from GBS deformation to dislocation creep will arrest or greatly decrease
the porosity generation (i) within the fine-grained material itself and (ii)
at porphyroclast–matrix boundaries, leading to the hardening of the
phyllonite B domain. In this case, the resulting microstructure displays CPO
and resembles Qtz2 layers, attributed to the origin by dynamic
recrystallization. This interpretation is supported by the fact that the
“reaction” quartz (Qtz3) cannot account for all the increase in
quartz mode in phyllonite B compared to the wall rock (Fig. 3b), and thus at
least part of the Qtz2 microstructure must also have an external
origin.
We suggest that the shear zone may be abandoned completely once it becomes
rheologically hard, leading to the activation of a new fracture in the
adjacent host rock. This could happen when the rock deforms dominantly by
dislocation creep and not cavitation creep due to fluid flux and
cavitation-creep-related increase in quartz grain size and abundance.
Conclusions
The studied shear zone represents an example of a brittle–ductile
deformation in the middle crust accompanied by circulating syn-deformational
fluids. The fluid flow was highly localized in narrow central parts of the
shear zone due to a cataclastic precursor and rapid metasomatic reactions,
which created necessary porosity for fluid infiltration. As a result two
structurally and chemically different domains developed across the shear
zone, depending on the availability of the fluid. The “fluid-limited”
shear zone margins experienced little effect of chemical processes during
the deformation, preserving the granitic assemblages similar to the wall
rock and deforming mainly by crystal-plastic processes in quartz. In
contrast, the “wet” shear zone centre experienced extreme chemical
alteration, rapid reduction in grain sizes by chemical breakdown reactions
and a development of interconnected foliation partitioning the strain. The
fine grain sizes further enabled fluid infiltration and creep cavitation
coupled with grain boundary sliding in the highly anisotropic reaction
fabric. Consequently, circulating pore fluids led to further alteration of
the shear zone fabric and rheology by mineral reactions, mass transfer and
incorporation of significant amounts of “external” quartz.
In summary, the existence of an initial brittle fracture system facilitated
enhanced fluid flow, rapid reaction and subsequent phase mixing. This then
enabled strain localization and focussed fluid flow due to creep cavitation
and associated fluid pumping.