1. metamorphic Ceol, 1989, 7, 261-275
Evidence for deformation-induced K-feldspar
replacement by myrmekite
C. SIMPSON
Department of Earth and Planetary Sciences, The johns Hopkins University, Baltimore,
M D 21218, USA
R. P. WINTSCH
Department of Geology, Indiana University, Bloomington, IN 47405, USA
ABSTRACT Several examples of deformation-inducedmyrmekite have been found in two amphibolite facies mylonites
derived from granitic protoliths, namely a muscovite-poor S-C mylonite and a single foliation,
muscovite-poor mylonitic gneiss. Back-scattered SEM and conventional optical microscopy show that in
both rock types, syntectonic myrmekitic intergrowths of oligoclase and quartz formed on the two sides of
&feldspar grains that faced the local inferred incremental shortening direction for the mylonite.
Myrmekite does not occur on the two ends of the grain that faced the incremental stretching direction.
The replacement of K-feldspar by plagioclase and quartz results in a volume decrease and is favoured
on high normal stress sites around the grains. We suggest that the ambient temperature, pressure and
chemical activities were such that the replacement reaction was favoured, but the addition of extra strain
energy along the high-pressure sides of the grains localized the reaction at these sites. This energy could
arise from elastic strain, or strain associated with tangled dislocations or twin boundaries. The relative
roles of stress and strain energy concentrations in driving the replacement reaction are not known, but
both were probably important.
Key w o d x Deformation; mylonite; myrmekite; strainenhanced reaction.
INTRODUCTION
Our understanding of the s i p f i a n c e of the interactions
between deformational and metamorphic processes has
been growing over the last few years with the recognition
that reactions can affect rock strength (White & Knipe,
1978; Rubie, 1983), and that deformation may drive
replacement reactions (Wintsch & Knipe, 1983). The role
of deformation in the replacement of K-feldspar by
mytinekite, a plagioclase-vermicular quartz symplectite,
was considered first by Futterer (1894) who noted a greater
development of mynnekite in deformed plutonic rocks
than in their undeformed equivalents. Eskola (1914),
Sederholm (1916) and Spencer (1945) all remarked on the
tendency for myrmekite to occur in zones of ‘ ‘ ~ h e d )or
’
“granulated” rock. Myrmekite growth in positions of high
stress at grain boundaries has been suggested by Schreyer
(1958), Sarma & Raja (1959), Shelley (1964). Binns
(1966), Parslow (1971), Bhattacharyya (1971) and Phillips
& Carr (1973) among others. Summaries of these early
works are given in Ashworth (1972), Phillips (1974) and
Smith (1974). More recently, a correlation between
mynnekite growth and crystal-plastic deformation has
been suggested by several workers including Hanmer
(1982), Tullis (1983), Watts & Williams (1983), Vernon,
W
i
l
l
i
a
m
s & D’Arcy (1983), S i p n (1983; 1985), Hibbard
(1987),
LaTour (1987) and LaTour & Barnett (1987).
. .
Despite this considerable attention to the problem, a
single model for the relationship between myrmekite
formation and the assoCiated deformation event has
proved elusive. Shelley (1964) considered the quark
vermicules in myrmekite to be recrystallized grains,
inherited from an earlier “cataclastic” event that became
incorporated into recrystalliPng feldspar. Hibbard (1987)
ascribed all myrmekite (including “strained mynnekite
units”) in orthogneisses to recrystallization in the presence
of a late-stage, water-saturated melt during the deformation of an incompletely crystallized magma. Hanmer
(1982), in his study of naturally deformed granites in
Newfoundland, contended that myrmekite formation was
post-deformational on the grounds that the delicate
vermicular textures could not survive appreciable strain.
However, a geometrical correlation between mynnekite
sites on inequant K-feldspar porphyroclasts and the local
k i t e strain axes was noted by Simpson (1985) in mylonitic
granodiorite of the Eastern Peninsular Ranges mylonite
zone, southern California. Progressive rotation of myrmekitic intergrowths into millimetre-scale shear bands was
taken to illustrate syndeformational myrmekite formation,
for which a solid-state difhsion-exsolution model was
proposed (Simpson, 1985).
Recently, several other mylonitic rocks from very
different geological settings have also been found to
contain K-feldspar porphyroclasts in which the position of
m
262 C. SIMPSON & R . P. WINTSCH
mynnekite exhibits a geometrical relationship to the finite
and inferred incremental strain axes. These myrmekites d o
not readily fit into a solid solution model. The best
examples occur in two-foliation mylonites that are Types I
and I1 S-C mylonites of Lister & Snoke (19&Q), where
S = schistosity and C = cisaillement or shear (BerthC,
Choukroune & Jegouzo, 1979). Similar phenomena have
also been observed in mylonitic rocks that contain a single
foliation (e.g. Debat, Soula, Kubin & Vidal, 1978, fig. lb;
Vidal, Kubin, Debat & Soula, 1980, fig. la; Simpson,
1981).
Si@cant modification of the stability fields of minerals
in regions of high strain was first suggested by Harker
(1932) and again by Helgeson, Delany, Nesbitt & Bird
(1978) and Wintsch (1985) but few systematic studies of
the phenomenon have been made until quite recently, and
then only on selected minerals. Nucleation of new grains
may lower the free energy of mica by a combination
of reductions in strain and chemical energy (Etheridge
& Hobbs, 1974; Knipe, 1981). In greenschist facies
mylonites, the relative stability of feldspar with respect
to mica may be shifted by either strain or stress
concentrations (Knipe & Wintsch, 1985). Under very
high-grade conditions, sillimanite may precipitate, and
feldspar and biotite dissolve, at sites of high strain and/or
stress along feldspar grain boundaries (Vernon, 1987;
Wintsch & Andrews, 1988). Quantitative estimates of the
effect of strain energy associated with dislocations on the
relative stabilities of the A l S i 0 5 polymorphs (Kemck,
1986), on the aqueous solubility of quartz (Wintsch &
Dunning, 1985), and of calcite (Bianchetti & Reeder,
1985). indicate that dislocation densities of at least
10’ocm-2 are necessary to perturb a strain-free chemical
equilibrium significantly. Data from recent experiments on
the high-pressure breakdown of plagioclase suggest that
the presence of a high dislocation density may not in itself
promote nucleation and growth of reaction products, but
an enhanced reaction rate can be related to dynamic
recrystallization mechanisms (Yund & Tullis, 1986).
In this paper we describe the geometry and geochemistry of myrmekite rims on K-feldspar porphyroclasts in
two mylonites: the Santa Rosa S C mylonite of
southeastern California and the lineated Ponaganset
gneiss of the Hope Valley shear zone, eastern Connecticut.
The term “porphyroclast” is used to describe K-feldspar
megacrysts, the size of which is reduced by syndeformational recrystallization or by myrmekite replacement and
not by a fracture mechanism. A model is presented for the
syntectonic replacement of K-feldspar by myrmekite as a
result of increasing lattice strain in the K-feldspar host
grain. No attempt is made here to discuss the origin of all
myrmekite occurrences. We limit our discussion to
mynnekite that shows a definite textural asymmetry that
correlates with the surrounding tectonic fabric.
METHODS
Oriented thin sections were cut parallel to the mineral
elongation lineation and perpendicular to the S foliation in
the rocks. All mineral analyses were carried out on an
ARL electron microprobe at Virginia Polytechnic Institute
and State University, and on a Cameca electron
microprobe at the University of Michigan, using natural
standards and ZAF reduction programs. Back-scattered
scanning electron photomicrographs were made on an
Hitachi S570 electron microscope at the University of
Michigan. Activity diagrams were calculated using the
program DIAGRAM and the hydrolysis equilibria
summarized in Bowers, Jackson & Helgeson (1984).
Determination of K-feldspar structural states by X-ray
difkaction would be important information for this
project, but impractical without microstructural control on
the relative positions of the grains within the mylonitic
foliation.
OCCURRENCES
K-feldspar porphyroclasts that show geometrical features
similar to those described here have been found by us in
the following rocks, among others: Santa Catalina
metamorphic core complex mylonites, south-western
Arizona, in an extensional tectonic regime (Keith, Damon,
Reynolds, Shafiqulla, Livington & Pushkar, 1980); Maggia
mylonitic gneiss, Lepontine Alps, Switzerland, in a
compressional tectonic regime (Simpson, 1981; 1982);
mylonitic rocks from the Honey Hill and Tatnic faults,
eastern Connecticut, in a transpressional regime (Wintsch
& Sutter, 1986); Santa Rosa mylonites of the Eastern
Peninsular Ranges mylonite zone (EPRMZ), south-eastern
California, in a compressional tectonic regime (Sharp
1966; 1979; Simpson, 1984; 1985); and lineated Ponaganset
gneiss, Hope Valley Shear Zone, eastern Connecticut, in a
transpressional tectonic regime (O’Hara & Gromet, 1985).
For the sake of brevity we will discuss samples from the
last two of these localities only.
Sang Rosa Mylonite
The Santa Rosa mylonite sample is from a mylonitic granodiorite
section of the EPRMZ (Fig. la); see also Sharp (1966; 1979) and
Sipson (1984; 1985). Theodore (1970). Anderson (1983),
Simpson (1984; 1985) and Erskine & We& (1985) have provided
geological, petrologicaland petrographical details of the mylonitic
rocks of the EPRMZ. Our specimens, especially SR656. were
taken from the Late Cretaceous, west-directed thrust zone
rnylonites, well away from the younger low-angle detachment
faults that occur throughout the region (Engel & Schultejahn,
1984; Schultejahn, 1984; Erskine & Wenk, 1985). Pressuretemperature conditions during deformation are estimated to have
been in the middle amphibolite facies or in the range 450-550°C
and 4-5 kbar by Anderson (1983) and Simpson (1985). The total
finite strain in the sample is thought to be low because the angle
between C- and S-surfaces is approximately 43” and the offset
along individual C-planes does not usually exceed 1.Ocm. The
finite shortening direction is therefore assumed to be close to the
incremental shortening direction for the rock, i.e. subperpendicular to the S-surfaces (Berth6 CI al., 1979).
In thin section, S surfaces are defined by quartz ribbonscontaining 150-250 prn-sized grains, aligned biotite grains (aspect
ratios 50 X loo0 pm in sections perpendicular to (001) plants) and
minor recrystallized muscovite. Tabular morphology and the
absence of undulatory extinction suggcsts that most of the biotite
K-FELDSPAR R E P L A C E M E N T B Y M Y R M E K I T E
N
m
Flg. 1. General locality maps for myrmekite-bearing mylonites discussed in text. Stars mark sample locations;general trend of foliations
and lineationsindicated by strike lines and arrows. (a) Santa Rosa mylonite zone, Palm Canyon, southeastern California (sample
SR656). modified from Erskine & We& (1985); g = Peninsular Ranges granite; mg = mylonitic Peninsular Ranges granite; ms = mylonitic
mctasedimentary rocks; gd = granodiorite; teeth on thrust fault and double ticks on detachment fault both on upper plates. (b) Hope
Valley shear zone, Connecticut (sample B e ) , after Gromet & O'Hara (1985). LCF= Lake Char Fault; HVSZ = Hope Valley shear
zone; HVT (stippled) = Hope Valley subterrane; E-DT = Esmond-Dedham subterrane (of Avalon terrane).
grains are completely recrystallized. Recystalliizcd, 50-100pmsized plagioclase grains of composition An,,, to An, generally
form a polygonal mosaic between quartz ribbons, although
occasional relict porphyroclasts, 2 9 pm to 1.0 mm in diameter,
are present. Elongate K-feldspar megacrysts, 200 pm to 2.0 mm in
diameter, also occur; representative chemical analyses of the
K-feldspar are given in Table 1, columns A and B. Although
difIiculty was experienced in obtaining high quality analyses of
these p i n s , the overall trends were reproducible. The
compositlons of oligoclase grains in the myrmekitic rims of the
K-feldspar megacrysts are within a fairly tight range of An, to
Ass.
Muscovite
occurs with or without d a t e within plagioclase
porphyrodasts and constitutes <0.5% of the total rock. Where
muscovite occurs at the edge of or within plagioclase augen. it
generally has (001) planes aligned subparallel to the adjacent
foliation planes. Within large plagioclase porphyroclasts, muscovite occurs as undeformed. anhedral grains, commonly with their
(001) planes at a high angle to the surrounding foliation.
Muscovite is never seen in close spatial association with
myrmekite, either in the matrix or around augen.
Subhedral danite grains show no deformation features. Where
elongate, the allanite crystals are aligned subparallel to the
adjacent foliation planes. Biotite and/or epidote commonly form
a narrow rim around the allanitc grains; the biotite grains are
parallel to and help define the foliation. Very minor retrogression
of biotite to chlorite occurs along some C-surfaces. Epidote occurs
in isolated stringers of 30pm-sized grains between quartz ribbons.
S i c pre-tectonic garnet grains were observed in two out of ten
thin sections. The garnet grains are fractured, with one main
fracture at a high angle to the S-surfaces. Quartz is the major
precipitate within the fractures. Titanite occurs as a minor
accessory mineral in very small, isolated, subhedral grains; it
shows no evidence of fracturing or other deformation features.
The foliations in this rock form a conspicuous S-C fabric (Fig.
2); also see Simpson (1985. Fig. 3). Ssurfaces are deflected into
C-bands with a consistent sense of curvature, concomitant with a
decrease in rtcrystallized grain-size, and both wrap around the
large K-feldspar augen. The C-surfaces contain nibons of
Iine-grained (30-40pm-sized) polygonal quartz, plagiodase and
K-feldspar, with very line-grained biotite and m i w r muscovite
(aspect ratios 10 x 20-50 pm). Back-scattered electroo imagery
(Fig. 3c), shows nearty monominerallic layers of quartz,
50-2OOpm wide, separating nearly biminerallic K-feldspar+
oligoclase layers of the same thicloess.
Back-scattered electron imagery allows the precise location of
myrmekite around the K-feldspar margins and in the recrystalked
S and C-bands to be determined. Myrmekitic intergrowths of
oligoclase (c. &;Table 1, cdumn~C and D) and quartz
generally occur on the two sides of the K-feldspar augen that
parallel the S-surfaces and face the incremental shortening
direction (Figs 2a, 3a, 4a). Myrmekite also oaufs along some
augen margins that are adjacent to C surfaces, at about 45" to the
incremental shortening direction (Fig. 3a). However, myrmekite
has never been found on the ends of the augen that face the
incremental stretching direction. In these 'tail' regions there are
subgrains and equigranular recrystallized K-feldspar grains (Fig.
4a; Table 1, cohunns E and F). Optical examination with a
264 C. SIMPSON &
R.
-
P. WINTSCH
T.bk 1. Representative electron microprobe analyses of feldspars
from sample SR656. Rim analyses taken within the first 100 pm of
1
myrmekite at porphyroclast margins.
Rim
Core
orthoclase
A
SiO,
MZO,
FeO
CaO
Na,O
K,O
Total
Si
Al
Fe
ca
Na
K
B
oligoclase in
myrmekite
C
D
E
F
14.57
66.74
18.76
0.06
0.04
1
14.52
61.31
25.16
0.03
5.93
7.94
0.19
63.26
24.57
0.01
5.72
8.27
0.12
64.31
20.42
0.04
0.08
1.12
113.89
64.56
20
0.04
0.05
1.09
14.02
99.83
101.14
100.56
101.94
99.82
99.73
2.943
1.101
0.001
2.958
1.08
0.001
0.004
0.002
0.097
65.74
18.53
0.08
0.03
0.88
3.014
1.001
0.003
0.001
0.079
0.852
Cations for 8 oxygens
3.015
2.705
2.748
0.999
1.308
1.257
0.002 0.001
0
0.002 0.28
0.266
0.088 0.679 0.6%
0.837
0.011 0.007
0.099
0.811
0.819
(Ca+ Na + K)
0.932
cationsum
0.927
0.97
0.%9
0.914
0.918
0.15
8.43
91.42
0.19
9.46
90.35
28.89
70.00
1.11
27.46
71.86
0.68
0.43
10.87
88.70
0.26
10.54
89.20
An
Ab
Or
Bi
Tails
recrystallized
microcline
universal stage indicates that the recrystallized K-feldspar grains
immediately adjacent to the host crystal have not greater than 15"
lattice misorientations with their host grain.
The K-feldspar host grain interface with myrmekite is always
sharply defined (Fig. 3b) and may form rounded, bulbous or
polygonal outlines that are generally not parallel to rational
crystallographic planes in the host grain. The boundary is also
chemically abrupt at the scale of analysis (<5pm electron beam
diameter). Oligoclase grains in the myrmekitic intergrowths, are
50 pm-sizcd at the K-feldspar contact, and decrease to c. 30 pm
in the adjacent polygonally recrystallized material (Fig. 3b). The
quartz vermicules commonly change from roughly cylindroidal,
<5 pm-sized grains near to the K-feldspar host grain, to polygonal
10pm-sizcd grains in the matrix, which suggests that the quartz
grains farthest from the porphyroclast have undergone the most
grain growth by grain boundary migration whereas oligoclase
grains in this position have undergone the greatest grain-size
reduction. Comma-shaped myrmekite grains occasionally occur at
the K-feldspar margin and give a deflection sense consistent with
the sense of shear across the sample (Figs.2b and 4b). Within the
recrystallized material of the S and C-bands there is commonly a
dimensionally preferred orientation of quartz and oligoclase
grains, even where biotite grains are absent. The resultant
secondary foliation is consistent in orientation with the
independently interpreted sense of shear across the sample
(Simpson & Schmid, 1983; Lister & Snoke, 1984; Passchier &
Simpson, 1986).
Along C- and S-surfaces where K-feldspar porphyroclasts are
absent, vermicular intergrowths of oligoclase and quartz may also
occur, particularly where small K-feldspar grains are interspersed.
Myrmekite also occurs along the edges of the larger K-feldspar
'tails' of augen (Fig. 3a). In Fig. 3c, a 0.3 m m diameter oligoclase
porphyroclast between ribbons of quartz contains irregular zones
filled with K-feldspar and a concentration of 10 to 20pm-sized
recrystallized K-feldspar grains that have apparently grown in the
low-pressure regions at the side of the porphyroclast. The
structural state of these K-feldspar grains could not be
Myrmekite
0.2mm
K - Feld/d/Olig
Approx. scale
(b)
K - Feldspar
Q/Olig
Q
L
Q/Olig
0.05mm
Approx. scale
Q. 2.
Thin section sketches from sample SR656 of Santa Rosa
mylonite. (a) K-feldspar (orthoclase) porphyroclast between Sand C- surfaces. adapted from Simpson (1985). Myrmekite only
along K-feldspar grain margins that are parallel to S;recrystallized
K-feldspar in 'tail' regions. (b) Detail of myrmekitic margin to
K-feldspar porphyroclast in (a). Myrmekite morphology is
consistent with independently determined sinistral shear across
the mylonite. Q = quartz; Bi = biotite; Olig = oligoclase.
determined; their small size also inhibits accurate refractive index
measurements so that optically they are easily mistaken for
recrystallized oligoclase grains. The mineralogical differences are
readily apparent, however, in the back-scattered mode of the
SEM (Fig. 3c).
Hope Valley Shear Zone
Sample B4-6 (Fig. lb) is strongly linear, weakly planar,
Ponaganset biotite augen gneiss occurring at stop 6 of Gromet &
O'Hara (1985). c. 2 km east of the dextral Hope Valley shear
zone (O'Hara & Gromet, 1985). The lineation, plunging at 12"
towards 012". is representative of the region (Gromet & O'Hara,
1985). K-feldspar augen are commonly >5cm long, tabular or
ovoid in section, and show welldeveloped 'tails' of recrystallized
K-feldspar that extend parallel to the mineral elongation
Lineation. Many of the augen contain stable cracks subperpendicular to the mineral elongation lineation. O'Hara & Gromet
(1985) suggested deformation temperatures in the range
400-500 "Cfrom the prevalence of evidence of prism ( c ) slip in
K-FELDSPAR R E P L A C E M E N T BY M Y R M E K I T E lf5
Fig. 3. Back-scatted SEM photomicrographs of sample SR656.Bright white (b) = biotite; dull white (kf) = K-feldspar; medium grey
( 0 ) = oligoclase;dark grey (q) = quartz; my = myrmekite. (a) K-feldspar megacryst between S and C-surfaccs. Note absence of
mymekite at short ends of megacryst. i.e. on sides that face incremental stretch direction. (b) Detail of mynnekitic margin in (a) shows
gradual loss of vermicular quark texture and formation of polygonal recrystallized oligoclase grains with distance from mymekite. (c)
Oligodase megacryst between quartz ribbons has 'tail' regions and irregular hctures filled with K-feldspar.
266
C. SIMPSON 81 R.
P. WINTSCH
K - F E L D S P A R R E P L A C E M E N T B Y M Y R M E K I T E 267
TIUC 2. Representative electron
microprobe analyses of feldspars from
simple M.
Rim analyscs taken within
first 100pm of the porphyrodast margin.
core
microdine
A
SiOz
B
C
core
Rim
albitic
IameUae
digodase in
D
E
mymekite
F
I
J
65.61 64.65 64.57
22.52 18.95 18.65
0 0.04 0.05
2.8 0.03 0.04
9.77 0.47 0.57
0.12 15.51 15.35
Total
99.42 99.33 99.63 99.83 99.97 99.99 100.36 100.82 99.62 99.23
Si
Al
Fe
ca
Na
K
Cations for 8 oxygens
2.983 3.002 2.989 2.944 2.996 2.857 2.835
1.032 1.005 1.024 1.074 1.015 1.153 1.171
0
0 0.003
0.001 0.002 0.002 0.002
0 0.033 0.007 0.122 0.165
0.002
0
0.055 0.031 0.m 0.916 0.958 0.847 0.806
0.906 0.933 0.889 0.018 0.002 0.016 0.008
(Ca + Na + K)
cationsum 0.963 0 . W 0.961 0.969 0.%7
An
Ab
or
64.59
22.64
0.01
3.51
9.47
0.14
H
KZO
NazO
FeO
CaO
64.89
22.23
0.07
2.58
9.92
0.29
G
64.46
18.93
0.02
0.05
0.61
15.35
4 0 3
64.75 64.83 67.16 68.7
18.39 18.84 20.79 19.76
0.06 0.04 0.05
0
0.01
0
0.7 0.14
0.34 0.81 10.78 11.33
15.78 15.11 0.35 0.04
Tails
recrystallized
microclillc
2.984
1.031
0.002
0.001
0.827 0.042
0.007 0.913
2.992
1.018
0.96
2.859
1.156
0
0.131
0.002
0.002
0.051
0.907
0.985
0.979
0.965 0.956
0.26 0.05 0.00 3.40 0.68 12.38
5.68 3.17 7.53 94.69 99.09 85.99
94.06 96.78 92.47 1.91 0.23 1.63
16.86
82.34
0.80
13.58 0.15 0.21
85.73 4.40 5.33
0.69 95.45 94.46
quartz and from the presence of recrystallized feldspar and
hornblende; the presence of prism ( c ) slip has, however, been
taken to indicate deformation temperatures as high as 650°C
(Mainprice, Bouchez, Blumenfield & Tubia 1986).
In thin section, the h a t e d Ponaganset gneiss contains narrow
quartz ribbons, Type 3 of Boullier & Boucbez (1978). that contain
350pm to 1.25mm-sizcd grains, rtcrystaUized biotite, and
elongate aggregates of polygonal, 100pm-sized recrystalked
K-feldspar and plagiodase (&
to AI& grains, all of which
define the lineation. "he majority of the K-feldspar porphyradasts in the rock are dose to An,,,Ab,.,Or,.,
in composition
Table 2, columns A to C). Plagiodase lamellae (Table 2, columns
D and E) are scattered throughout augen cores and no systematic
relationship could be detected between the specific location within
the host grain and the lamella composition. In most samples the
boundary between the grain and the myrmekite (Table 2,columns
F to H)is chemically abrupt at the scale of analysis. However, a
few oligodase grains in the myrmekite were seen on
back-scattered SEM images to have a narrow (<5 pm) border of
more aibitic plagioclase between them and the host K-feldspar
(arrowed on Fig. 5a).
Biotite and muscovite are present in roughly equal proportions
and together constitute about 15% of the matrix to the K-feldspar
augen. All biotite grains are SUbparaUel to the foliation plane.
Alteration of biotite to chlorite and titanite, with or without
Fe-xidcs, is patchy; in some areas alteration is almost nil,
whereas in other areas of the same thin &on,
alteration is
almost complete. This alteration is probably post-tectonic, as
tbere is a complete.abscncc of pressure-shadow phenomena
around tbc euhedral to anhedral, 750pm to 1.0mm-sized oxide
grains. Muscovite occu~sas unoriented grains within plagioclase
grain aggngatts and as oriented grains intergrown with biotite.
Myrmekite and muscovite are rarely found in close association.
Large. fntchmd, originally igneous(?) titanite grains have small,
pakr titanite grains around their margins. Rare allanite grains are
rimmed by epidote.
?be margins of a typical tabular K-feldspar augen are illustrated
in Fig. 4c and d. Myrmekite (An,,+ quartz) oczm in quantity on
tbe two long sides of the augen that are parallel to the weak
phnar fabric and the strong mineral elongation lineation (Fig. 4c).
No myrmekite has been found on the two sides of the grain that
face the extension direction (Fig. a),
and only minor, sporadic
occuf~cnceshave been recorded from the two faces that are
perpendicular to the weak foliation and parallel to the Lineation.
Thus, the maximum development of myrmekite is along those
faces that are subperpendicular to the incremental shortening
direction.
Figure 5a illustrates the morphology of the myrmekite in sample
B4-6. As in sample SR656, quartz vermicules 25-50pm in
diameter tend to be subrounded in outline near to the K-feldspar
host crystal (Fig. 5b) but become more polygonal and slightly
larger with distance from the host grain. The K-feldspar crystal
contains numerous inclusions of quartz and white mica, and
exsolution p a t h up to 50pm in diameter. Several of the
K-feldspar augen contain stable cracks at a high angle to the
elongation lineation. These cracks are commonly filled with
indusion-free. polygonal, apparently strain-free K-feldspar or
quartz grains up to 150pm in diameter. Polygonal K-feldspar
grains constitute the major part of the augen 'tails', with quartz
and biotite as secondary components. Small grains of K-feldspar
occur at the boundary between the host crystal and its 'tail'; thcsc
are clear and optically strain-free, and no myrmekite occurs along
them (Fig. 4d).
DISCUSSION
In the previous sections we have shown that quartzplagioclase symplectite (myrmekite) in mylonitic rocks is
highly concentrated along asymmetric K-feldspar megacryst faces that are parallel to the foliation planes in the
rock. In the following sections we discuss those aspects of
the microstructure that a re relevant to interpreting the
driving mechanism of the reaction and its non-random
distribution.
The geometry of the microstructure associated with
these augen suggests that the faces that show myrmekite
268 C. SIMPSON & R . P. WINTSCH
Flg. 5. Back-scattered SEM photomicrographs of sample B4-6. White (kt) = mcrccme; medium grey (0)= oligoclase; dark grey
(q) = quartz; black = holes. (a) Bulbous morphology of myrmekite selvage leaves islands of microcline surrounded by mymekite. (b)
Detail of (a) to show rounded to polygonal outlines of quartz vermicules in oligoclasc and narrow rind of more albitic plagioclase
(mowed) between microcline and oligoclase.
development experienced relatively high normal stress.
This is especially safe to infer in the S-C mylonitic
microstructure. In the usual interpretation of S-C
structure, at relatively low finite strains, where the angle
between S- and C-surfaces is approximately 45", the
Ssurfaces face the maximum incremental shortening
direction and the C-planes are parallel to the bulk rock
plane for simple shear (Berth6 et al., 1979). With
increasing shear strain the S-surfaces become rotated
toward the flow plane and away from the plane of
maximum incremental shortening; i.e. the angle between S
and C decreases. In sample SR656, the average angle
between S - and C-planes is 43", ignoring the small portions
of the S-planes that curve asymptotically into the C-planes,
and thus the main portions of the S-surfaces are still
approximately subperpendicular to the inferred incremental shortening direction. In this orientation the S-planes
would experience a normal compressive stress close to the
maximum and only a very small component of shear stress.
There is thus a very strong positive correlation between
predicted high normal stress sites and the Occurrence of
myrmekite.
Subgrains and recrystallized grains in the 'tail' regions of
K-feldspar augen have lattice orientations very similar to
that of the adjacent host grain, which suggests that the tails
developed by progressive subgrain rotation recrystallization (Porier & Nicolas, 1975). Minor grain-boundary
migration was necessary to produce the observed
polygonal grain shapes. These observations imply the
operation of dislocation glide and climb mechanisms
during the deformation of the K-feldspar megacrysts.
The plastic deformation of K-feldspar requires at least
low amphibolite facies conditions to have prevailed during
the deformation (VoU, 1976; Tullis, 1983). Compositional
information on the feldspars presented here allows
estimates of the temperature of myrmekite development to
be made, provided that K-feldspar and plagioclase were in
exchange equilibrium during the replacement reactions,
and that their compositions have not been modified since
their growth. Except for very thin albite rims on some
grains, neither feldspar is systematically zoned. Thus, in
spite of the obvious replacement textures, exchange
equilibrium may have been obtained. Minimum and
maximum temperatures of replacement were calculated
using the analytical expressions of Stormer (1975) and
Whitney & Stormer (1977), which treat maximum disorder
and maximum order, respectively. Average minimum and
maximum temperatures calculated for the Santa Rosa
rocks are 495 "C and 540 "C, in good agreement with the
estimates of Anderson (1983). The temperature range
calculated for the Hope Valley shear zone samples is
440-445"C, which is consistent with the estimates of
O'Hara & Gromet (1985).
The two main hypotheses for myrmekite development
are those of replacement reactions (Becke, 1908) and
solid-state exsolution (Schwantke, 1909). These ideas and
K-FELDSPAR R E P L A C E M E N T B Y M Y R M E K I T E 2c9
variations thereon have been summarized and discussed in
&tail by Sturt (1970), Widenfalk (1972), Smith (1974) and
Phillips (1974; 1980). In Schwantke's lattice diffusion
model, a change in solid composition m u m in response to
elastic strain gradients, but requires a relatively long time
to accomplish (see also Bayly, 1987). The development of
subgrains and recrystallized grains provides strong
evidence that dislocation glide and climb mechanisms were
operating during myrmekite development, so that elastic
strain gradients would have been continually created but
also destroyed during the plastic deformation. In the
solid-state ditrusion model of Simpson (1985), a concentration of Na' ions is built up along the high-pressure sides of
K-feldspar porphyroclasts. Na+ has such a small ionic
radius that it may diffuse through the K-feldspar lattice
either by a vacancy drag effect (e.g. Hanneman &
Anthony, 1969; Petrovit, 1973) or by direct interstitial
diffusion (Lin & Yund, 1972; Foland, 1974). However, this
model does not adequately account for the calcium
component of oligoclase/quartz myrmekite and, moreover,
a compositional change from alkali to plagioclase feldspar
accomplished solely by a lattice diffusion mechanism
requires the diffusion of all four major cations, which is
unlkely to have occurred at these relatively low
temperatures. We therefore rule out the lattice diffusion
models of Schwantke (1909) and Simpson (1985) as
responsible for the examples described here.
The Becke (1908) replacement of K-feldspar by
myrmekite can be written in the form:
1.25KA1Si308+ 0.75Na'
( 1 3 7 mole-')
~
+ 0.25Caz+
= Na,,7sCa,, UAll.,Si2.7s08
( l a k c mole-')
.
+ 1.25K+ + SiO,.
(1)
( 2 k c mole-')
This reaction is fully consistent with the textures of
sample SR656. The reaction direction can be inferred both
by textural and surface area considerations. Mushroomshaped lobes of mynnekite embay the margins of the
K-feldspar grains and the ratio of plagioclase to quartz is
approximately 4:1, as predicted by reaction (1). The
replacement of myrmekite by K-feldspar, were it to occur,
would occur selectively along quartz-plagioclase interfaces, where the local bulk AljSi ratio is closest to that of
the K-feldspar. This has not been observed. O n the
contrary. the lobes of myrmekite are K-feldspar-free and
have the lowest possible surface area, whereas the surface
area of the K-feldspar grain is much increased by the
replacement. Exactly the same arguments hold for the
sample from the Hope Valley shear zone, but in this rock,
reaction (1) must be modified for the appropriate
composition of plagioclase, c. An,,. The plagioclase to
quartz ratio is then 6.7:l which is consistent with observed
ratios of about 8:l in thin sections. Reaction (1) thus
describes the texture in Fig. 5a very well, when the
reaction proceeds to the right.
The Phillips (1980) modification of the Becke model
requires the introduction of H,O into the reactions, so that
muscovite forms at the expense of the existing K-feldspar
and plagioclase, and K,O is removed from the local
environment. The Santa Rosa mylonites (sample SR656)
were almost certainly deformed in the presence of an
aqueous phase (Anderson, 1983). The alteration of biotite
to chlorite in the Ponaganset gneiss indicates that it too
had access to an aqueous fluid, but this CouId have
occurred at a later stage of metamorphism. In order to
investigate the possible effects of changing the activities of
potassium, sodium and hydrogen ions on the system
K-feldspar-muscovite-albite,
we illustrate solid-fluid
equilibria for this system at 400 "C and 450 "C, 300 MPa
(Fig. 6). Reaction (2) in Fig. 6 can be represented, using
ideal formulae, by:
Musc Qtz + 2K' = 3Ksp + 2H'
(2)
+
and reaction (3) by:
Musc + Qtz + 3Na' = 3Ab + K+ + 2H'.
(3)
Figure 6 shows that with increasing temperature, the
stability field of albite increases with respect to K-feldspar
and muscovite, and the K-feldspar stability field increases
with respect to muscovite. Increasing pressure enlarges the
stability field of muscovite relative to both feldspars, and
enlarges the plagioclase field relative to K-feldspar. The
diagram thus shows that plagioclase is stabilized relative to
K-feldspar by increasing temperature or pressure, either of
which could theoretically lead to the development of albite
replacement. ,
The production of myrmekite with muscovite during
retrograde metamorphism (Phillips, Ransom & Vernon,
1972) is also elucidated by Fig. 6. Consider a Kfeldspar-plagioclase-quartz-bearing granite at 450 "C in
equilibrium with a fluid (point X in Fig. 6). With cooling,
the muscovite stability field grows and by 400°C totally
encompasses point X, whereupon muscovite would
precipitate from the fluid. Such precipitation would lower
the local activity of aqueous K+, thereby destroying the
equilibrium of reaction (1). K-feldspar would then dissolve
and be replaced by plagioclase plus quartz. This reaction is
predicted for all rocks, but the modal significance is
directly proportional to the concentration of ions in the
fluid and to the amount of fluid present in the rock. The
amount of mynnekite formed during cooling would thus be
directly proportional to porosity. Under these conditions a
close textural relationship would be expected between
muscovite and myrmekite. Although a small quantity of
muscovite may have been produced during my-nnekite
formation in mylonitic rocks of the Hope Valley Shear
zone, the general lack of association between myrmekite
and muscovite in both sample studied here, and the
sparsity of muscovite in the Santa Rosa samples, argue
against reactions (2) and (3) and in favour of reaction (1).
The equilibrium constant, log K1,for activity balance of
reaction (1) is given by:
+ 1.25 (a,+) + log (sP)
log K,= log (a&
- 1.25 log (aw) - 0.75 log (a,,+) 0.25 log (+.z+).
-
(4)
The forward reaction (1) can therefore be driven by an
4.70.
3
4.80.
0" 4.90
5.00
5.1 0
-2A
MUSCOVITE
I
I
I
I
I
4:4
'/
QTZ SAT.
300 MPa
1
/
/
h
3.8
4.80
4.90
5.0C
MUSCOVITE
/3
ALBITE
300 MPa
+ QTZ
I
I
unstrained
K-feldspar
4.4
Fig. 7. Calculated effect of adding 100 cal(418 J) to the free energy of K-feldspar on
decreasing the size of the stability field of K-feldspar at 450 "C, 300 MPa at quartz
saturation. Solid lines show stability fields for strain-free crystals; dashed line shows
reduced size of stability field of strained K-feldspar. For further discussion, see text.
pb. 6. Calculated effect of decreasing temperature from 450 "C to 400"Con the stability
fields of muscovite, albite, and 'equilibrium K-feldspar' (Helgeson ei ul., 1978) at 300 MPa
(3 kbar) with quartz saturation, in t e r n of the log aqueous activity ratios Na+/H+ and
K+/H*. Note that with decreasing temperature, the K-feldspar field expands relative to
albite, and the muscovite field expands relative to both feldspars. Similar changes occur
for calcium-bearing plagioclase. For further discussion see text. Numbers 1,2 and 3 refer
to equations (l), (2), (3) in text.
w
0
-
a.
Z
t
5.1(
K-FELDSPAR R E P L A C E M E N T BY M Y R M E K I T E
increase in the activities of K-feldspar or of aqueous Na+
and G2+.
Assuming that the coupled diffusion of
and silicon is relatively slow in comparison with
that of aqueous Na' and Ca2', the volume change Of the
solid phases during the reaction of K-feldspar to form one
mole of the plagioclase product from reaction (1) is
14 cc mole-' plagioclase, a volume decrease of more than
10%. The forward reaction (1) will thus also be favoured
by an increase in the normal stress on the K-feldspar grah
and is more likely to occur on the edges of the K-feldspar
grains that faced the incremental shortening direction. The
reverse reaction (1) will be favoured in the lower pressure,
'W regions of the augen.
The effect of an increase in applied differential stress on
the K-feldspar grains during steady-state deformation
would be to increase the dislocation density and decrease
the rotation-recrystaked grain-size (Twiss, 1977) especially on the sides of the augen that faced the incremental
shortening direction. The addition of strain energy
associated with dislocations could be sufficient to shift the
K-feldspar stability field boundaries on Fig. 6, as shown on
Fig. 7. Material close to the boundary of reaction (1)
would be stable as K-feldspar at low values of stored strain
energy, but would form plagioclase (i.e. reaction (1) would
proceed to the right) at higher energy values. Point H on
the plagioclase-K-feldspar stability field boundary (Fig. 7)
describes the composition of a fluid in equilibrium with
unstrained solids under a hydrostatic stress state, i.e. the
composition of the fluid buffered by the rock prior to
deformation. Tbis equilibrium strain-free state is compared
to a deformed state (dashed m e on Fig. 7), calculated by
raising the free energy of K-feldspar by 100 cal (418 J); the
energy could be associated with tangled dislocations. The
composition of the fluid H is unchanged by the
deformation, but it is no longer in equilibrium with the
deformed K-feldspar grains; such a fluid would have a
composition along the dashed curve. Fluid H is now
entirely within the plagioclase stability field, and to change
its composition to be in local equilibrium with the strained
K-feldspar, the activity of aqueous Na' must drop, and
that of aqueous K' must rise. This is readily accomplished
by the replacement of the strained regions of K-feldspar by
plagioclase.
During high-temperature steady-state deformation by
recovery-accommodated dislocation creep, much of the
internal strain energy of the lattice is reduced by the climb
of dislocations to form low-energy tilt walls (e.g. TuUs &
Yund, 1985). However, at the lower temperatures of
deformation in the Santa Rosa and Hope Valley
mylonites, regions close to the grain boundaries of large
K-feldspar augen probably contained a high density of
tangled dislocations, particularly along the edges of the
grains that faced the incremental shortening direction, and
a dislocation density of 1010cu-2 is readily obtained in
such tangles (Knipe & Wintsch, 1985; Wintsch & Dunning,
1985).
To determine whether or not dislocation densities
reached such high values during the deformation of the
K-feldspar rims, a transmission electron microscope study
M
of the myrmekite margins of the augen would be of great
interest. Unfortunately, neither the present dislocation
density nor the recrystallized grain structure in the augen is
necessarily the one that was present prior to myrmekite
formation. Our examples contain clear evidence for
syntectonic myrmekite development and most of the
myrmekitic aggregates have undergone deformation. A
higher dislocation density now present in the vicinity of the
myrmekite could reflect post-formation strain, whereas a
lower dislocation density in that region could simply
indicate that grain boundary migration had occurred, or
that reaction (1) had led to the replacement of all the
strained feldspar. In addition, the original size of
rotation-recrystallized new grains would be modified by
grain boundary migration.
Etheridge & Hobbs (1974), Knipe (1981) and Lee,
Peacor, Lewis & Wintsch (1986) noted that chemical
changes accompany recrystallization in micas; this is to be
expected if the pressure and temperature conditions of
&formation are different from those at the time of
formation of the protolith. In their examples, however, the
free energy drop associated solely with the chemical
metastability of the micas was insufficient to cause
recrystallization of micas of a different composition; strain
energy must have been added. The mica examples parallel
the mymekite discussed here, with the important
differences that our samples require a net change in the
mineral assemblage, and that the reaction mechanism was
one of ion exchange.
THE MODEL
From the foregoing discussion, a conceptual model for the
myrmekite replacement reaction can be developed and this
is presented in Fig. 8. The model requires only a
pre-existing K-feldspar megacryst in a matrix of quartz,
plagioclase and K-feldspar. A non-hydrostatic stress state
develops in the granitic rock and the large K-feldspar
augen act as stress risers, such that normal stresses are
highest along those margins facing the incremental
shortening direction (open arrows on Fig. 8). Under the
low amphibolite faaes conditions of our samples, the
imposed stress causes concentrations of elastic strain
energy, and strain hardening (Taylor, 1934), associated
with relatively immobile dislocation tangles, is likely to
occur along the K-feldspar grain boundaries (e.g. T u b &
Yund, 19?7; 1987). The composition of the ambient
aqueous fluid present would be buffered in the unstrained
grain boundaries by the K-feldspar-plagioclase-quartz
assemblage (Fig. 6). However, along the sfrained margins
of the K-feldspar, the equilibrium is destroyed and a
replacement reaction ensues (Fig. 7). This replacement is
accompanied by a 10% volume loss, so that the high
imposed stresses on the grain margins facing the
incremental shortening direction are relieved. The
replacement reaction is thus a form of incongruent
pressure solution (Beach, 1982).
Reaction (1) is a replacement reaction and so sources
and sinks of aqueous components must be considered. The
272 C . SIMPSON &
R. P. WINTSCH
plag l a y e r
qtz
1a y e r
(Ca++, .Na+>
/'
(Ca++. Na+>
augen
1ayer
Ksp. replacement
by plag.
gc
s
poss b l e plag. rep acement
by Ksp.
'VQ"
Fig. 8. Conceptual model for development of deformation-induced myrmekite. High normal stress at some K-feldspar megacryst
boundaries favours replacement of K-feldspar by plagioclase + quartz. Sources of plagioclase components and sinks for K+ may be the
reverse reaction in the 'tails'. Open arrows indicate sites of high normal stress. For further discussion, see text.
replacement of K-feldspar occurred at approximately
constant Al and Si and can be modelled as a cation
exchange reaction. Aqueous K' is large relative to Na+
and Ca2' (Helgeson & Kirkham, 1976) and is thus more
likely to be concentrated in the lower normal stress 'tail'
regions of the augen. In fact, K-feldspar is concentrated in
the augen 'tails' almost at the expense of plagioclase, and
it also occurs in fractures, unlike plagioclase. 'The
migration of K' to the lower normal stress regions would
increase the K+/Na' activity ratio, such that the fluid
precipitates K-feldspar. Such an interpretation is required
by the filling of fractures in plagiwlase by K-feldspar (Fig.
3c). Potassium may thus be conserved on the scale of a
thin d o n . The possibility exists that muscovite may also
act as a sink for some of the K+ released during myrmekite
formation. However, the paucity of muscovite in the Santa
Rosa samples argues against it acting as the predominant
potassium sink in these rocks.
The sodium-calcium budget is more difficult to deal
with, because calcium sources are less obvious; the
replacement of K-feldspar clearly offers an important sink.
One answer to the source problem may again be found in
Fig. 2.In the 'tail' region labelled kf, K-feldspar is Seen to.
surround and embay plagioclase grains - a texture that
may be evidence of plagioclase replacement by K-feldspar.
Another possible local source of calcium could be provided
by the dynamic recrystallization of plagioclase to a lower
temperature, more sodic composition. It is, therefore,
quite possible that the strainenhanced myrmekitic
replacement of K-feldspar may occur with complete
elemental conservation, i.e. a closed system on the scale of
centimetres. The possibility of a totally open system, in
which the composition of the fluid remains fixed by
equilibria in an adjacent rock, may be discounted for our
examples because such a fluid would invade all grain
boundaries, not just the highly stressed ones. Rather, the
K-FELDSPAR R E P L A C E M E N T B Y M Y R M E K I T E 273
concentration of the reaction at selected microstructurd
sites demands gradients in fluid composition on the scale of
mm, controlled by reactions in the local (cm-scale)
environment.
We do not, however, concur with the contention of
Hibbud (1979; 1987) that textures consistent with 'mobile
feldspar' necessarily indicate deformation in the presence
of H,O-bearing magmatic fluids, i.e. deformation of an
incompletely crystallized magma. Hibbard (1987) considered that the presence of myrmekite in mylonitic granite
gneisses reflects relocation of late-stage magmatic fluids
and is not the result of deformation of a completely
crystallized rock. The examples that we have discussed
here contain myrmekitic intergrowths that are clearly on
the high-strain sides of the K-feldspars, in direct contrast
to the low-pressure sites that are required by Hibbard's
hypothesis. The high temperatures (>450°C) that are
apparently necessary for the dynamic recrystallization of
all feldspars at natural strain rates (e.g. Voll, 1976; Tullis,
1983) certainly favour the production of a mylonite gneiss
if deformation follows immediately upon, or at a very late
stage in, the crystallization of a pluton, but this is surely
not the only way to produce amphibolite facies
metamorphism in a tectonic environment. Deformation of
a pluton buried more than about 15km beneath thrust
nappes could .result in the formation of mylonitic gneisses
at that depth, regardless of the age gap between pluton
crystallization and deformation. Indeed, the myrmekitebearing gneisses of the Maggia Nappe in Switzerland
formed during Alpine deformation of Hercynian-aged
plutonic rocks (Simpson, 1983). Similarly, deformation of
plutonic rocks of the lowermost upper crust (c. 15 km and
below) during large-scale extension, as in the metamorphic
cure complexes of Nevada and Arizona, can produce
mylonitic gneisses that contain strain-related myrmekites,
e.g. the Santa Catalina mylonites in southern Arizona. In
addition, myrmekite is also found in retrograde mylonitic
rocks of the Willyama Supergroup, Broken Hill, Australia
(Phillips et d.,1972) and in metamorphosed sedimentary
rocks (e.g. Nold, 1984). We suggest, therefore, that a
significant number of myrmekite-bearing mylonitic
gneisses may have formed some considerable time after
crystallization of the magma was complete and that the
presence of myrmekite, or of apparently mobile
K-feldspar, should not be taken a priori to indicate
deformation before crystallization was complete. Finally,
the 'late-stage magmatic' argument is unsubstantiated for
the samples we describe because the temperatures of
formation of the rnyrmekite are entirely subsolidus.
As Ashworth (1972). Smith (1974) and Phillips (1980)
pointed out, there may be many different geneses for
myrmekitic intergrowths. We have discussed here only
myrmekite that is clearly related to the strain state in the
rock. Other Occurrences of mynnekite that might bear
closer examination include its widespread development in
undeformed granites. Although we do not rule out the
possibility of myrmekite formation in tectonically undeformed granites in the absence of a local differential stress
field, nevertheless most granites contain quartz that shows
evidence of lattice strain in the form of undulatory
extinction and subgrain formation. This lattice strain
presumably occurs in response to differential volume
changes that result from Merences in the thermal
conductivity and elastic anisotropy of adjacent minerals in
the cooling rock. Kirby & Etheridge (1981) suggested that
deformation-aided exsolution of Ca-clinopyroxene from
orthopyroxene could occur due to thermal conductivity
differences among adjacent minerals in a cooling rock. We
suggest that myrmekite could also form due to local stress
differences on K-feldspar grain boundaries during cooling
of a completely crystallized granitic rock.
CONCLUSIONS
A geometric relationship exists in many granitic mylonites
between myrmekite development on K-feldspar augen and
the finite, and inferred incremental, strain state of the
rock. A magmatic origin for this mynnekite is ruled out by
its geometry and the subsolidus temperatures of formation.
The absence of an association between muscovite and
myrmekite in the rocks that we examined suggests to us
that replacement reaction (1) predominated. Reaction (1)
involves a net volume decrease and is thus favoured at
sites of high normal stress on K-feldspar grain boundaries.
In the samples we have discussed, myrmekite is syntectonic and found only on the high-pressure sides of
the augen; its presence cannot be attributed solely to
chemical processes. The presence of small, dynamically
recrystallized new grain boundaries in the augen 'tails'
demonstrates that in spite of the relatively low
temperatures, deformation by crystal plastic mechanisms
did OCCUT in these K-feldspar megacrysts. Elastic strain
energy or energy associated with dislocations localized the
replacement reactions to sites of high normal stress,
causing the observed asymmetrical distribution of
mynnekite.
ACKNOWLEDGEMENTS
Financial support was provided in part by NSF grants
EAR-8121438, EAR-8305846 and EAR-8606120 to CS and
EAR-8313807 and EAR-8618305 to RPW. Helpful
discussions with B. Bayly, G. W. Faher, B. E. Hobbs, J.
A. Tullis and R. A. Yund are gratefully acknowledged.
The comments of R. H. Vernon and an anonymous
reviewer helped us to clarify parts of the text. I. D.
Fitzgerald, T. E. LaTour and R. A. Yund are thanked for
their incisive comments on a very early (1985) version by
CS. Assistance with electron microprobe and SEM
analyses was ably provided by T. N. Solberg and C.
Henderson.
REFERENCES
Anderson, J. R., 1983. Petrology of a portion of the Eastern
Peninsular Ranges mylonite zone, southern California.
Contributions to Mineralogy and Petrology, 84,253-271.
Ashworth, J. R., 1972. Myrmekites of exsolution and replacement
origins. Ceobgical Magazine, 189,45-62.
n4 C . SIMPSON 81 R. P. WINTSCH
Bayly, B., 1987. Nonhydrostatic thermodynamics in deforming
rocks. CMOdian J o d Of Earth sciences, 24,572-579.
Beach, A., 1982. Deformation mechanisms in some cover thrust
sheets from the external French Alps. J o d of Sturcwal
Geology, 4, 137-150.
Becke, F., 1908. Ueber Mymekite. Schweizerische Miwalogbche und Petrogrophische Minedungen, 27,377-390.
Berth&,D., Choukroune, P. & Jegoum, P., 1979. Orthognciss,
mylonite and non-coaxial deformation of granites: the example
of the South Armorican shear zone. J o d of Structural
Geology, 1,3142.
Bhattacharyya, C.. 1971. Mynnekite from the charnockitic rocks
of the Eastern Ghats, India. Geological Magazine, 108,
433-438.
Bianchetti, S. P. & Reeder, R. J., 1985. Variable dissolution rates
of deformed and undeformed calcite. Geological Society of
Anuricn Abstracts wilh Programs, 17,7, 522.
Binns, R. A., 1%. Granitic intrusions and regional metamorphic
rocks of Permian age fmm the Wongwibinda district,
northeastern New South Wales. Proceedings of the Royal
Society of New South Wales, 99,5-36.
Boullier, A. M.& Bouchez, JCL., 1W8. Le quartz en rubans dans
les mylonites. Bulktin Societie Gtologique de France, 7,
253-262.
Bowers, T. S., Jackson, K. J. & Helgeson, H. C., 1984.
Equilibrium Activity Dipgrams for Coexbfing Minerah and
Aqueous Solutions at Pressures and Temperatures to 5 kb and
600 "C.Springer-Verlag, New York.
Debat, P., Soula. J. C., Kubin, L. & Vidal, J. L., 1978. Optical
studies of natural deformation microstructures in feldspars
(gneisscf and pegmatites from Occitanie, southern France).
L a , ll,133-146.
Engel, A. E. J. & Schultejahn, P. A., 1984. Late Mesozoic and
Cenozoic tectonic history of south central California. Tectonics,
3, 659-676.
Erskhe, B. G. & W e d , H.-R., 1985. Evidence for Late
Cretaceous crustal thinning in the Santa Rosa mylonite zone,
southern California. Geology, W,274-277.
Eskola. P., 1914. On the petrology of the Orijarvi region in
southwestern Fiiand. Bulletin de la Conunisswn G6ologque de
Finfande, 40,285 pp.
Etheridge, M. A. & Hobbs, B. E., 1974. Chemical and
deformational controls on recrystallization of mica. Contributions to Mineralogy and Petrology, 43, 111-124.
Foland, K. A.. 1974. Alkali diffusion in orthoclase. In:
Geochemical Transport and Kinetics (eds Hoffman, A. W. et
al.), pp. 77-98, Carnegie Institute, Washington.
Futterer, K., 1894. Ueber granitporphyr von der Greisscharte in
den Zillerthaler Alpen. News Johrbuch jZr Geologic und
P&ntologie, 9,509-553.
Gromet, L. P. & O'Hara, K., 1985. The Hope Valley shear zone
- a major late Paleozoic ductile shear zone in southeastern New
England. In: NEIGC 77th. Annual Meeting, Yale Univeniry.
Connecticut, Guidebook for Fwldtrips, 6,277-295.
Hanmer, S. K., 1982. Microstructure and geochemistry of
plagiodase and microcline in naturally deformed granite.
Journal of Structural Geology, 4, 197-214.
Hanneman, R. E. & Anthony, T. R.. 1969. Effects of
non-cquilibrium segregation on near-surface diihsion. Acta
Metalhgica, 17, 1133-1140.
Harker, A., 1932. Metamorphism. Methuen, London.
Helgeson, H.C. & Kirkham, D. H.. 1976. Theoretical prediction
of the thermodynamic behaviour of aqueous electrolytes at high
pressures and temperatures. HI.Equations of state for aqueous
species at infinite dilution. American Journal of Science, 276,
97-240.
Helgeson, H. C., Delany, J. M.,Nesbitt, H. W. & Bud, D. K.,
1978. Summary and critique of the thermodynamic properties of
rock forming minerals. American Journal of Science, 278A,
1-229.
Hibbard, M. J., 1979. Mymekite as a marker between
preaqueous and postaqueous phase saturation in granitic
systems. Geological Society of
AIWI~CM BJlehir,
90,
1047-1062.
Hibbard, M. J., 1987. Deformation of incompletely
magma systems: granitic gneisses and their tectonic implications. J o d of Geology, 95,543-561.
Keith, S. B.. Damon, P. E., Reynolds, S. J., Shafiqulla, M.,
Livingston, D. E. & Pushkar, P. D.. 1980. Evidence for
multiple intrusion and deformation within the Santa Catalina&con-Tortolita metamorphic core complex. In: Cordillrrm
Metamorphic Core Complexes (eds Crittcnden, M. D., Jr.,
Coney, P. J., & Davis, G . H.), Geological Society of Americo
Memoir, l53,217-267.
Kemck, D. M.,1986. Dislocation strain energy in the A g i o ,
polymorphs. Physics and Chemistry of Minerals, W, 221-226.
Kirby, S. H. & Etheridge, M. A., 1981. Exsolution of
Ca-clinopyroxene from orthopyroxene aided by deformation.
PhySicr Md Chemistry Of Miner&. 7,105-109.
Knipe, R. J., 1981. The interaction of deformation and metamorphism in slates. Tectonophysics, 78,249-272.
Knipe, R. J. & Wintsch, R. P. 1985. Heterogeneous deformation,
foliation development, and metamorphic processes in a polyphase mylonite. In: Advane- in Physical Geochenuitry, 4,
Metamorphic Reactiotw (eds Thompson, A. B. & Rubie, D.
C.), pp. 180-210, Springer-Verlag, New York.
LaTour, T. E., 1981. Geochemicd model for the symplectic
formation of mynnekite during amphibolite grade progressive
mylonitintion of granite. Geological Society of America
Abstracts wifh Programs,19,741.
LaTour, T. E. & Barnett, R. L., 1987. Mineralogical changes
accompanying mylonitintion in the Bitterroot dome of the
Idaho batholith: implications for timing of deformation.
Geological Society of Amcrico Bulletin, 98,356-363.
Lee, J. H., Peacor, D. R., Lewis, D. D. & Wmtsch. R. P. 1986.
Evidence for syntectonic crystallization for the mudstone to
slate transition at Lehigh Gap, Pennsylvania. Journal of
Structural Geology, 8,767-780.
Lin, T.-H & Yund. R. A., 1972. Potassium and sodium
selfdif€usion in alkali feldspar. Coruriburiotw to Mineralogy and
Petrology, 34,177-184.
Lister, G . S . & Snoke, A. W.,1984. S-C mylonites. Journal of
Structural Geology, 6, 617-638.
Mainprice, D., Bouchez, J.-L., Blumenfeld, P. & Tubia, J. M.,
1986. Dominant c slip in naturally deformed quartz:
implications for dramatic plastic softening at high temperature.
Geology, 14,819-822.
Nold, J. L., 1984. Myrmekite in Belt Supergroup metasedimentary rocks-northeast border zone of the Idaho batholith.
American Minrralogist, 69,1050-1052.
O'Hara, K. & Gromet. L. P., 1985. Two distinct Precambrian
(Avalonian) terranes in southeastern New England and their
late Paleozoic juxtaposition. American Journal of Science, 285,
673-709.
Parslow, G. R., 1971. Variations in mineralogy and major
elements in the Caimsmore of Fleet granite, SW Scotland.
Lithos, 4,43-55.
Passchier, C. W. & Simpson, C., 1986. Porphyroclast systems as
kinematic indicators. J o d of Smccnual Geology, 8,831-843.
PetroviC, R., 1973. The effect of coherency stress on the
mechanism of the reaction albite + K+ = K-feldspar + Na' and
on the mechanical state of the resulting feldspar. Contributions
to Mineralogy and Petrology, 41, 151-170.
Phillips. E. R., 1974. Mymekite-one hundred years later. Litlux,
7, 181-194.
Phillips, E. R., 1980. On polygenetic myrmekite. Geologml
Magazine, 17,29-36.
Phillips, E. R. & Carr, G. R., 1973. Myrmekite associated with
alkali feldspar megacrysts in felsic rocks from New South
Wales. L W , 6,245-260.
Phillips, E. R., Ransom, D. M. & Vernon, R. H., 1972.
Myrmekite and muscovite developed by retrograde metamorphism at Broken Hill, New South Wales. Minerologiccrl
Magazine, 38,570-578.
K-FELDSPAR R E P L A C E M E N T BY M Y R M E K I T E
Poirier, J. P. & Nicolas, A., 1975. Deformation-induced
rtcrystallization due to progressive misorientation of subgrains
with special reference to mantle peridotites. Journal of
Geology, 83,707-720.
Rubie, D. C. 1983. Reaction enhanced ductility: The role of
solid-solid univariant reactions in deformation of the crust and
mantle. Tectotwphysics, 96,331-352.
Sarma, S. R. & Raja, N., 1959. On myrmetite. Quarterly Journal
of the Geolorical, M i n e r a l o u and Metalluraical Society of
I&,
31,
Schrever. W..1958. Die Quarz-FeldsDat-GeNge der miamatischin desteh von Vilshofen an der Donau. N & a Jahrbkh f i r
Mineralogie, 92,147-170.
SchultejaJm, P. A., 1984. The Yaqui Ridge antiform and
detachment fault: Mid-Cenozoic extensional terrane west of the
San Andreas fault. TectoniCr, 3, 677-691.
Schwantke, A., 1909. Die beimischung von Ca in Kalifeldspat und
die Mymekitbildung. Zentralblati fiir Geologic und
P & ~ l o g i e , 311-316.
Sederholm, J. J., 1916. On synantetic minerals and related
phenomena (reactions rims, corona minerals, kelyphite,
myrmekite, etc.). Bullerin de la Commission Gkologque de
Ehlkande,48,146 pp.
Sharp, R. V., 1966. Ancient mylonite zone and fault
displacements in the Peninsular Ranges of southern California.
Geological Society of America Special Paper, 101, 333.
Sharp, R. V., 1979. Some characteristics of the Eastern
Peninsular Ranges mylonite zone. United States Geological
Survey Open F i i Report, 79-1239,258-267.
Shelley, D., 1964. On myrmekite. American Mineralogist, 49,
41-52.
Simpson, C., 1981. Ductile shear zones: a mechanism of rock
deformation in the orthogoekes of the Maggia Nappe, Ticino,
Switzerland. Unpubl. PhD Thesis, ETH-Zuirch, Switzerland.
Simpson, c . , 1982. The structure of the northern lobe of the
Maggia Nappe, Ticino, Switzerland. Eclogae Geologicm
Helveripc, 75,495-516.
Simpson, C., 1983. Strain and shape fabric variations associated
with ductile shear zones. Jountal of Smctural Geology, 5,
61-72.
Simpson, C., 1984. Borrego S p M g - Santa Rosa mylonite zone:
a late Cretaceous west-directed thrust in southern California.
Geology, l2,8-11.
Simpson, C., 1985. Deformation of granitic racks a c r m the
brittle-dude transition. Journal of Smr~hual Geology, 7 ,
503-511.
Simpson, C. & Schmid, S. M.,1983. An evaluation of criteria to
deduce the sense of movement in sheared rocks. Geological
society of America Bullcrin, 94,1281-1288.
Smith, J. V., 1974. Feldspar Minerals 2: Chemical and texhvol
properties. Springer-Verlag. New York.
Spencer, E., 1945. Myrmekite in graphic granite and in vein
perthite. Mineralogical Magazine, 27,79-98.
Stormer, J. C., Jr.. 1975. A practical two-feldspar geothermometer. American Mineralogist, 60,661-674.
Shut, B. A., 1970. Exsolution during metamomhism with
particular reference to feldspar solid &lutions. Mineralogical
Magazine, 37,815-832.
Taylor, G. I., 1934. The mechanism of plastic deformation of
crystals: Part I - Theoretical. Proceedings of rhe Royal Society
of London, Serics A , 145,362-387.
l k o d o r e , T. G., 1970. Petrogenesis of mylonites of high
metamorphic grade in the Peninsular Ranges of southern
California. Geological Society of America BuUetin, 81,435-450,
I
~
-
- .
Zn
Tullis, J. A., 1983. Deformation of feldspars. In: Feldspar
Mineralogy (ed. by Ribbe, P. H.),Mineralogical Society of
America short c o m e notes, 2 (2nd. edn), 297-323.
TuUs, J. A. & Yund, R. A., 1977. Experimental deformation of
dry Westerly granite. Journal of Geophysical R u e d , 62,
5707-5718.
Tullis, J. A. & Yund, R. A., 1385. D y n d c recrys&z&on of
feldspar: a mechanism for ductile shear zone formation.
Geology, U,238-241.
T W , J. A. & Yund, R. A., 1987. Transition from catadastic
h w to dislocation creep of feldspar: Mechanisms and
microstructures. Geology, 15,606409.
Twiss, R. J., 1977. Theory and applicability of recrystallized grain
size paleopiewmeter. R u e and Applied Geophysics, 115,
227-244.
Vernon, R. H.,1987. Oriented growth of sillimanite in andahsite,
PLaatas-Juan Tabo area, New Mexico, U.S.A. CPMdian
Journal of Earth ScicnCe, 24,58%590.
Vernon, R. H., Williams, v. A. & D'Arcy, W. F., 1983.
Grain-size reduction and foliation development in a deformed
granitoid batholith. Tectotwphysics, 92, 123-145.
Vidal, J-L., Kubin, L., Debat, P. & Soula, J. C., 1980.
Deformation and dynamic recrystallization of K-feldspar augen
in orthogneiss from Montagne Noire, Occitania, southern
France. Lithos, U,241-255.
VoU, G., 1976. Recrystallization of quartz, biotite, and feldspars
from Erstfeld to the Leventina Nappe, Swiss Alps. and its
geological significance. Schweizerische MincrrJogirchc und
Petrographische Miaeilungen, 56,641-647.
Watts, M. J. & Williams, G. D., 1983. Strain geometry,
microstructure and mineral chemistry in metagabbro shear
zones: a study of softening mechanisms during progressive
mylonitization. Journal of Structural Geology, 5,507-518.
White, S . & Kaipe, R. J., 1978. Transformation and reaction
e n h d ductility in rocks. Journal of the Geological Society of
London, WS, 513-516.
Whitney, J. A. & Stormer, J. C., Jr., 1977. The distribution of
NaAISiO, between coexisting microcline and plagioclase and its
effect on the geothermometric calculations. AmcricM
Mineralogist, 62,689-891.
Widenfalk, L., 1972. Mymekite-Like intergrowths in larvikite
feldspars. L k ,5,255-267.
Wintsch, R. P., 1985. The possible effects of deformation on
chemical proccssts in metamorphic fault zones. In: Adwnccs in
physical geochemistry, 4, Metamorphic Reactions (eds Thomp
son, A. B. & Rubie. D. C.). pp. 251-268, Springer-Verlag,
New York.
Wmtsch, R. P. & Andrews, M. S., 1988. Deformation-induced
growth of sillimanite: stress minerals revisited. J o d of
Geology (in prtss).
Wintsch, R. P. & Dunning, J., 1985. 'Ihe effect of dislocation
density on the aqueous solubility of quartz and some geologic
implications. Journal of Geophysical Research, W,3649-3657.
Wintsch, R. P. & Knipe, R. J. 1983. Growth of a zoned
plagioclase porphyroblast in a mylonite. Geology, 11,360-363.
Wintsch, R. P. & Sutter, J. F.,1986. A tectonic model for the Late
Paleozoic of southeastern New England. Journal of Geology,
94,459-472.
Yund, R. A. & Tullis, I. A., 1986. The effect of ductile
deformation on plagioclase breakdown kinetics. Geological
Society of A m e h Abstracts with Programr, 18,799.
Received 25 January 1988; revision accepted 27Jwu 1988