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Porphyroblast nucleation, growth and dissolution in regional metamorphic rocks as a function of deformation partitioning during foliation development T. H. BELL & M. J. RUBENACH. Department of Geology, James Cook University. Townsville, Queensland, 481 I , Australia P. D. FLEMING, Department of Geology, La Trobe University, Biindoorri. Victorin. 3083. A iistrciliri Abstract. In regional metamorphic rocks, the partitioning of deformation into progressive shearing and progressive shortening components results in strain and strain-rate gradients across the boundaries between the partitioned zones. These generate dislocation density gradients and hence chemical potential gradients that drive dissolution and solution transfer. Phyllosilicates and graphite are well adapted to accommodating progressive shearing without necessarily building up large dislocation density gradients within a grain, because of their uniquely layered crystal structure. However, most silicates and oxides cannot accommodate strain transitions within grains without associated dislocation density gradients, and hence are susceptible to dissolution and solution transfer. As a consequence. zones of progrcssivc shearing become zones of dissolution of most minerals, and of concentration of phyllosilicates and graphite. Exceptions are mylonites, where strain-rates are commonly high enough for plastic deformation to dominate over diffusion riitcs and therefore over dissolution and solution transfer. Porphyroblastic minerals cannot nucleate and grow in zones of active progressive shearing, as they would be dissolved by the effects of shearing strain on their boundaries. However, they can nucleate and grow in zones of progressive shortening and this is aided by the propensity for microfracturing in these zones, which allows rapid access of fluids carrying the material presumed to be necessary for nucleation and growth. Zones of progessive shortening also have a number of characteristics that help to lower the activation energy barrier for nucleation, this includes a build up of stored strainenergy relative to zones of progressive shearing, in which dissolution is occuring. Porphyroblast growth is generally syndeformational. and previously accepted criteria for

static growth are not valid when the role of deformation partitioning is taken into account. Porphyroblasts in a contact aureole do not grow statically either, as microfracturing, associated with emplacement, allows access of fluids in a fashion that is similar to microfracturing in zones of progressive shortening. The criteria used for porphyroblast timing can be readily accommodated in terms of deformation partitioning, reactivation of deforming foliations, and a general lack of rotation of porphyroblasts, with the spectacular exception of genuinely spiralling garnet porphyroblasts. Key-words: deformation partitioning; fluid infiltration; porphyroblast growth; solution transfer; spiralling garnet porphyroblasts

INTRODUCTION Realization of the significant roles of solution transfer and strain partitioning during deformation has promoted recent advances in understanding the microstructural development of foliations (Durney. 1972; Marlow & Etheridge, 1977; Bell, 1878; Mitra. 1978; Gray, 1979; Gray & Durney. 1979; Bell, 1981). The relationships, however, between progressive deformation and metamorphism during foliation formation have not been defined, even though regional metamorphic rocks commonly contain evidence of these aspects preserved in porphyroblasts (e.g., Bell & Rubenach, 1980, 1983). In particular, the role of deformation partitioning during progressive deformation has received only scant attention (e.g.. Bell, 1981; Williams & Schoneveld, 1981). As a consequence, in this paper we examine conceptually porphyroblast nucleation, growth and dissolution in terms of deformation partitioning, as developed from Bell (1981). but based on our collective microstructural experience in metamorphic terrains which range in age

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from Eocene to Precambrian and in grade from blueschist and greenschist to granulite facies. Some of the more important processes in porphyroblast formation are inferred, by us, to involve: (1) dissolution of material to provide a source of ions; (2) solution transfer of these ions to the site of nucleation and growth; (3) nucleationand(4)growth;(cf. Harker. 1932;Ramberg, 1952). These processes are probably influenced or controlled by microstructural development of the rock (and vice versa), as are a number of related aspects, which include: (1) the creation of suitable nucleation sites for porphyroblasts; (2) location of these sites with respect to micro structural features and, (3) the preservation or dissolution of porphyroblasts during further deformation and metamorphism. Careful studies of the microstructural development of porphyroblasts and their host matrix are critical to a better understanding of the above processes and their controls.

DEFORMATION PARTITIONING

Strain in deformed rocks is extremely heterogneous (Bell, 1981). The heterogeneity arises in several ways and on a range of scales, from ductility contrasts between grains to regional temperature and pressure variation in the crust. Strain heterogeneity can be enhanced by localized strain-softening (White, Burrows, Carreras, Shaw & Humphreys, 1980), the presence of faults, earlier folds and local variations in rock type. The effects of primary heterogeneity during deformation are such that different minerals, rock types, beds and sedimentary structures progressively take up different components of the strain, as described below. Furthermore, once some deformation and/or metamorphism has been imposed, volumes of secondary heterogeneity (e.g., early fold hinges, veins, porphyroblasts and cleavage domains) add to the effects of primary heterogeneity. Both types of heterogeneity result in deformation partitioning on a large variety of scales, whereby, at various stages during a deformation, different minerals, beds, rock types and other portions of rock, take up: (1) no strain; (2) dominantly progressive shortening strain (i.e., progressive coaxial deformation) ; (3) progressive shortening plus-shearing strain (i.e., progressive non-coaxial deformation); (4) progressive shearing-only strain (i.e., progressive non-coaxial deformation); as discussed by Bell (1981). Both 3 and 4 can be coaxial at a

larger scale of consideration of the strain. Consequently, the terms 'coaxial' and 'non-coaxial' are not used in this paper with a deformation partitioning connotation, and are reserved for the description of scale-dependent bulk strain histories or strains. For the purposes of discussion, the foregoing range of possibilities can generally be conveniently reduced to two, because the important factor appears to be the presence, or otherwise, of a significant component of progressive shearing (rotational) strain, as amplified in some detail below. Thus, throughout the remainder of this paper deformation partitioning is discussed in terms of progressive shortening ( 1 and 2 above) or progressive shearing (3 and 4 above) components, except where there is a specific need to describe the effects of, for example, progressive shearing-only strain. An example of these effects is shown for a simple foliation in Fig. Ib, where mica grains wrap around quartz and feldspar grains and take up the progressive shearing component of the strain. whereas quartz and feldspar tend to take up the progressive shortening components. Even for the inhomogeneous shear-dominated end-member of deformataion history, progressive inhomogeneous simple shear (Fig. 2a), similar relationships occur in rocks because of these heterogeneities. For example, consider a felsic gneiss or granite being mylonitized by progressive inhomogeneous simple shear. Feldspar is generally more competent than quartz or mica in this situation and consequently takes up little or none of the progressive shearing strain; any strain is localized at grain or kink-band boundaries. Progressive shearing strain cannot remain constant along any plane parallel to the mylonitic schistosity because of heterogeneities which result from the presence of this feldspar (Fig. 2b). Therefore, the shear planes anastomose around the feldspar grains and, if the shear zone width is to remain constant, the mylonitic foliation must thicken and thin along its length (Fig. 2b). Hence zones of shortening or dilational strain must be developed. as too many feldspar grains are randomly distributed through the rock of laminar flow to be possible. For a more general deformation history, such as progressive bulk inhomogeneous shortening, some narrow zones of developing foliation conceivably undergo no shortening component on t h e scale of the width of the zones. That is, the partitioning of the progressive shearing components of the deformation at that scale is complete and only progressive shear strain is

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Fig. I(a) Strein-tield diagram f o r a geometry develtrped by noncoaxial progressive bulk inhomogeneous shortening. Thrcc zonc typcs of deformation partitioning are indicated. No strain has occurred within the dashed lines (the elliptical areas outlined by these are not strain ellipses). Between the dashed and dotted lines the deformation was dominantly coaxial progressive shortening. Between the dotted lines the deformation involved progressive shortening plus shearing strain. This diagram shows no examples of progressive shearing without a component of shortening strain. I(b) Sketch showing phyllosilicates wrapping around grains such as quartz and feldspar. The phyllosilicates have taken up the progressive shearing component of the deformation. whereas the quartz and feldspar grains have dominantly taken up the progressive shortening component. progressive shearing having occurred only on their rims.

accommodated by this zone, whereas the surrounding rock takes up components of both progressive shortening and progressive shearing strain. An example of such a zone might be a fully differentiated crenulation cleavage which consists of phyllosilicates with no remains of quartz or feldspar (Fig. 2c, cf. Bell, 1985). Folding appears to be a function of the inhomogeneous distribution of the progressive shearing component of the deformation (e.g., Bell, 1986; see later). Several orders of parasitic folds indicate various scales of inhomogeneously distributed progressive shearing strain. Further, inhomogeneity exists on the smaller scale of crenulations and foliation planes. In general, as a fold tightens, more of the hinge is affected by the progressive shearing component of deformation. Decrenulation, accompanied by reactivation of the crenulated foliation (Bell, 1986), is a function of the redistribution of progressive shearing strain in a more pervasive

fashion through that zone. Thus, local zones of progressive shortening become zones dominated by progressive shearing. The reverse also happens, due to the local cessation of deformation, possibly many times during a single deformation event (e.g.. Jones, 1986). This discussion is particularly relevant to porphyroblast formation because, as detailed later, we contend that sites for nucleation and growth of porphyroblasts occur in areas of the rock that deform by progressive shortening, whereas porphyroblast dissolution occurs against, or within, parts of the rock that deform by progressive shearing. Some porphyroblasts have syntectonically overgrown and preserved differentiated crenulations and crenulation cleavage (Bell & Rubenach, 1983). Therefore, zonal patterns of progressive shearing strain must shift through the rock during deformation, if our hypothesis is correct (see later).

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fluid; (2) solute diffusion within a fluid, combined with mass transfer by fluid flow and, (3) removal of the solute from solution at a favourable site by chemical reactions, including phase transformations, precipitation and surface absorption; (Durney, 1972; Bell, 1978; Beach, 1979; Gray & Durney, 1979; Robin, 1979; Raj & Chyung, 1981; Etheridge, Cox, Wall & Vernon, 1984). Dissolution occurs due to chemical potential gradients which are established by factors such as pressure, strain and strain-rate gradients. However, microstructural observations in deformed rocks indicate that dissolution and solution-transfer of material away from specific locations generally have occurred in zones which are demonstrably affected by a large component of shear strain (Fig. 3; see below). As a consequence, the intimate controls on sites of dissolution in non-mylonitic rocks appear to be only indirectly related to pressure

Fig. 2(a-e) Skctchcs showing (a) ii schcmntic gcomctry which results from thc simplest modcl of non-coaxial progressive inhomogcneous simplc shcar (i.c.. thc carddcck rnodcl) and (b) a morc rciilistic skctch which results from bulk progressive inhomogcncous simplc shcar in ii mylonitc zone. whcre shcar displ;iccmcnts anastornose around feldspar porphyroclasts. This results in thickening and thinning of the zones taking up thc shear strain. and consequently deformation partitioning must also occur in such a relative end-memhcr modcl of deformation history. (c) Shows a n cxiirnplc of ii fully diffcrcntiatcd crcnulation clcavage formcd during progressive bulk inhomogcncous shortcning. Prcigrcssivc shortening may not occur in a shearing location such as this. as phyllosilicatcs crystallographically can morc readily accomodate progressive shearing parallel to ((MI)planes than progressive shortening across them.

IMPORTANT PROCESSES IN PORPHYROBLAST FORMATION

Dissolution and solution transfer Solution transfer during deformation and metamorphism involves: (1) dissolution of orieinallv solid material into an intereranular a Y

Fig. 3. Crcnulation cleavagc ( S ? ) anastomosing around a chloritoid porphyroblast and showing dissolution of quartz and feldspar from the crenulation limbs. S, adjacent and nearly parallel to thc porphyrohlaht ih uncrenulated. The shear strain and dissolution intensities in the matrix on the porphyroblast edges. as the chloritoid has taken up none of the strain. Planepolarized linht. Width of base is 3.8 mm. I

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(cf. Durney, 1972; Kerrich, 1977; Gray & Durney, 1979) and more directly related to the amount of shear strain that these sites have undergone (Beach, 1981; Bell. 1981; Bosworth. 1981; Engelder, 1982). High-grade mylonitic rocks appear to have been affected by plastic deformation alone, with no component of solution transfer. During deformation, strain and strainrate gradients, between zones of progressive shortening and progressive shearing (Fig. la), would generate variations in dislocation density, and hence chemical potential gradients across the boundary between these zones. Different deformation mechanisms would probably have operated in the zone of progressive shortening than in the zone of progressive shearing (Bell, 1981), which may enhance diffusion through and away from the latter locations. Thus, dissolution in deformed rocks is most apparent on the limbs of crenulations and along crenulation cleavages (Fig. 3, Marlow & Etheridge, 1977; Gray, 1979; Gray & Durney. 1979; Bell & Rubenach, 1983). presumably because these are zones of high progressive shearing strain. Dissolution, however, is also readily observed at other scales, such as on the limbs of mesoscopic folds (Fig. 4; Hobbs, Means & Williams. 1976, fig. 5.15), and on a mesocopic and macroscopic scale in shear zones, which are commonly depleted in minerals such as quartz and feldspar and enriched in phyllosilicates and carbon (Hammond, 1985). Dissolution has been just as strong, although less striking, in slates, phyllites and schists that show no sign of having developed from a crenulation cleavage (cf., Bell & Rubenach, 1983). For example, Wright & Platt (1982) and Beutner & Charles (1984) have measured volume losses greater than 50% in such rocks. Deformation partitioning-that is, the distribution of progressive shearing and shortening components of deformation-can occur on a grain scale during the formation of a slaty cleavage in the classic sense (Fig. 5a), rather than on the scale of several grains, as occurs during crenulation cleavage development (Fig. 5b). Thus, we infer that progressive shearing strain dominates on the edges of grains parallel to the foliation (Fig. 5c). Since phyllosilicates are common and have extremely well-layered crystal structures, they are generally ideal recipients of the progressive shearing component of deformation by (001) slip (Etheridge & Hobbs. 1974). without the build up of dislocation density gradients across (001) planes (Fig. 5c); graphite has a similar capacity. We assert that crystallization of totally

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Fig. 4. A spcctaculur cxnmplc of dissolution o n the limhs of mesoscopic folds at Mourilyan Hiirhour. North Qucensland. This location is in thc hingc of a macroscopic antiform. and the removal of quartz and feldspar has occurred from the long limhs. with a high component of shear strain on both sides of the mesoscopic fold centred o n the lens ciip. These areiis are topograpically lower by I crn and enhance the spectacular appearance of this outcrop in the field.

new phyllosilicates generally does not occur in active zones of progressive shearing. Instead, old grains commonly straighten, due to the shear strain, and recrystallize by migration of kink or grain boundaries along their lengths (Fig. 6; e.g., Swager, 1985). Recrystallization of these phyllosilicates may involve a slight chemical change. In some cases one composition can be preferentially removed, presumably due to a change in metamorphic conditions (e.g., Marlow & Etheridge, 1977; Swager, 1985). It is possible, however, that some phyllosilicate grains lie in orientations that are not amenable to reorientation into the zone of progressive shearing without extensive internal strain on slip planes other than (001). This could result in the development of large strain

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Fig. 5 . b ) Sketch of a slaty cleavagc of schistosity. in which the phyllosilicates anastornose around ellipsoidal non-

phyllosilicate grains or aggregates of grains. The deformation has partitioned on the scale of grains. such that the ellipsoidal grains take up the progressive shortening components and the phyllosilicates the progressive shearing components. (b) Sketch of a differentiatcd crenulation cleavagc. The deformation has partitioned on the scale of several grains in this type of cleavage. compared t o (a) above. (c) Sketch of an ellipsoidal grain surrounded by phyllosilicates. The shear couple shown is accommodated along the (001) planes of the phyllosilicate grains. These grains are ideal recipients for progressive shearing because of their highly anisotropic crystallographic structure.

Fig. 6. Sketches showing a progression from relatively coaxial crenulation of a foliation to formation of a differentiated crenulation cleavage as the deformation become more non-coaxial (c, d and e). Although the zone of highest strain is the long limb of the crenulation. the stored strain energy within it is decreased. due to dissolution of minerals that are not phyllosilicates and progressive straightening, recovery and recrystallization of those that are. However. the hinges of crenulations retain their stored strain energy.

Porphyroblast nucleation, growth and dissolution gradients across a phyllosilicate at a high angle to (001) and hence cause their dissolution (cf. Bateman, Bell & Rubenach, 1986). Coarsening can occur during recrystallization due to the dissolution of other grains and hence the aggregation of several phyllosilicate grains of similar orientation in the same zone (Fig. 6c,d & e). Minerals which are adjacent to phyllosilicates, such as quartz and feldspar (in pelitic rocks), are also affected on their boundaries by progressive shearing. They do not have the layered crystal structure of a phyllosilicate and. consequently, a gradient in dislocation density builds up across the shear zone on their rims (cf. Figs la and 5c). This sets up a chemical potential gradient, which causes dissolution of quartz, feldspar, carbonate and sometimes garnet grains (see later) against the phyllosilicate, but not, in general, dissolution of the phyllosilicate itself. The effect is shown schematically in Fig. 7. The commonly resulting geometries, such as truncated dustings on quartz grains (Elliott, 1973; Bell, 1978). partly dissolved fossils (Wright & Platt, 1982), partly dissolved klites (Cloos, 1947), embayed cobbles in deformed conglomerates (Mosher, 1981), and truncated zoning (Vernon. 1978; Fig. 6b; Bell & Rubenach, 1983), have led to a resurgence, in the past 15 years, of the pressure-solution concepts of the last century. However, whether or not pressure (stress) is of direct rather than indirect importance, it is apparent that dissolution occurs microstructurally in zones of high progressive shearing strain. This means that syndeformational porphyroblasts cannot grow across active zones of high progressive shearing strain, because they are generally sites of dissolution of all minerals except sheet silicates and graphite. Such zones include active foliation laminae, crenulation limbs, crenulation cleavage and differentiated crenulation cleavage. They also include shear zones, unless deformation partitioning can separate zones of progressive shortening from progressive shearing on a sufficiently large scale within them, or unless the mineral involved has special competency and growth habits, such as garnet (see later). Pre-existing porphyroblasts in these zones or parts of porphyroblasts lying across or next to a zone of progressive shearing tend to dissolve (see below). Conversely, if syndeformational porphyroblasts contain inclusion trails of crenulation cleavage at stages 3 or 4 (fig. 4 in Bell & Rubenach, 1983), we infer from the model, discussed above, that progressive shearing was not operating on crenulation limbs in these loca-

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mica

1

...................

"

1

. .

. .

1; ....................

1

Fig. 7p-e. Schematic diagrams showing shear-controlled dissolution of quartz against mica. The shear strain in the mics (b) is accornmoduted by glide on ((HI) planes. without developing dislocation density variation normal to (001) planes, whereas the quartz is plastically strained on a variety of slip systems against the mica and generates a dislocation density gradient. Quartz is progressively removed during shear of its edges, the end result being shown in (c). Similar relationships would occur if the mica anastomosed around the quartz. Dislocations would be generated within the mica because of its anastornozing character, but would not lead to a dislocation density gradient across (001).

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tions at the time of porphyroblast growth. This implies that the pattern of deformation partitioning was changing during the deformation (e.g., Bell, 1986; Jones, 1985). In general, the source of material for porphyroblast nucleation and growth is provided, or supplemented, from zones of dissolution, such as those dominated by progressive shearing (Bateman, ef al. 1986). Porphyroblast dissolution

The removal .of pre-existing porphyroblasts should be common during pro- and retrograde metamorphism. For example, many high-grade rocks grow andalusite and staurolite during prograde metamorphism, and then these porphyroblasts are removed by subsequent prograde reactions. What takes place microstructurally during porphyroblast removal, and what evidence for these processes is preserved within the rocks? The most obvious microstructural evidence for such processes during both pro- and retrograde metamorphism. consists of porphyroblasts partly or completely pseudomorphed by muscovite, chlorite, biotite or sillimanite (e.g., Vernon, 1976. 1978; Bell & Rubenach, 1983). We can sometimes infer their original nature by the shape and composition of the alteration product (e.g., Kwak, 1974). This implies, however, that there has been much deformation in that locality after the alteration, with the less common exception of porphyroblasts pseudomorphed by other porphyroblastic minerals, such as staurolite replaced by andalusite (Bell & Rubenach, 1983). What has happened to the porphyroblasts for which no pseudomorphs are preserved? We argue above that porphyroblasts cannot overgrow active zones of high progressive shearing and indeed will tend to be dissolved against such zones. Also we argue that, in general, the pattern of deformation partitioning continually changes within a given volume of rock. Porphyroblastic minerals appear to be more competent than the surrounding matrix, as they commonly show no undulose extinction, even though they grew syndeformationally. This may, however, be simply a function of their larger grain size, which would reduce the rate of diffusion and/or fluid flow through them by orders of magnitude relative to the same volume of matrix (Etheridge & Vernon, 1981). because of the presence of grain boundaries in the matrix. For both these reasons, porphyroblasts tend not to deform as readily as the matrix, as shown in Fig. 3. As a consequence,

the deformation partitions around them such that they become zones of progressive shortening, whereas the matrix takes up the progressive shearing component, as shown in Fig. 8. If they become strained on their margins, chemical potential gradients are established and dissolution may commence, especially if the mineral was destabilized by increasing temperature. Bell & Rubenach (1983) described garnet porphyroblasts which had been partially dissolved against mica because they became unstable with increasing grade (Fig. 9). However, we contend that dissolution can also occur without a change in metamorphic conditions,

Fig. 8 Two pliigioclasc porphyrohlasts which contain inclusion trails with a 'millipedc' geomctry (Bcll B Ruhcnach. IYXO). One (top) has a coaxiiil gcomctry and the other has a noncoaxial geometry. Note the cxtensive removal of quartz and feldspar. and consequent concentration of mica. against the porphyroblast at the top of the photograph. This is a function of the porphyroblast taking up little to none of the progressive shearing component of the deformation after its growth and the geometric necessity, therefore. for extremely high shear strain against its rim parallel to the developing Sl. Note also the control of the orientation of S , within the zone of progressive shortening o n the shape of the porphyroblast. N-section. Crossed polars. Base 16.4 mm.

Porphyroblast nucleation, growth and dissolution

Fig. 9. Garnet porphyroblast with a dodecahedral outline where it is protected by a staurolite porphyroblast from the effects of progressive shearing. However. where it abuts differentiated Sz it has been dissolved. as revealed by the abrupt trucation of the dodecahedral outline. N-section. Plane-polarized light. Base 2.4 mm.

especially where the porphyroblast affected is of low strength, such as chlorite, muscovite or biotite. In such cases, we suggest that there would be no net destruction of the porphyroblastic mineral-just dissolution in unfavourable sites and regrowth in more favourable ones. Redistribution of the pattern of deformation partitioning may generate a zone of progressive shearing through a porphyroblast, causing it to dissolve and/or alter. This is particularly so for porphyroblasts with large length-to-width ratios lying at a high angle to the foliation. Examples are shown in Fig. 10, for biotite, and Fig. 11, for staurolite. This phenomenon would be enhanced by a change in T,P conditions during deformation, such that a porphyroblast became unstable, thus allowing reaction softening to occur (White er a l . , 1980). This could occur in two ways. First, reaction softening on the edges and along a fracture through the porphyroblast would lead to internal strain and the generation

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of kink bands. This in turn would lead to recrystallization andor the rapid access of fluids along kink-band or recrystallized grain boundaries, this would then allow further reaction and the nucleation and growth of matrix minerals. The porphyroblast would then be destroyed by its conversion to matrix schist (Figs 10 & 11). Second, formation of a partial or complete pseudomorph of the porphyroblast by finer grained, more ductile minerals, such as chlorite, muscovite or biotite would allow the pattern of deformation partitioning to redistribute, so that progressive shearing strain could cut through the porphyroblast remains. Nucleation of other minerals on grain boundaries within the deformed pseudomorph would convert it to matrix, e.g. homogenization at stage 6 of crenulation cleavage development, as des(Vernon, 1977); e.g. ornphacite preserved in albite (Bell & Brothers, 1985). Locally. porphyroblastic minerals are replaced a n d o r preserved by incorporation within bigger porphyroblasts of different composition. For example, garnet porphyroblasts are occasionally overgrown by staurolite, and only where they are not fully enclosed are they dissolved against mica (e.g., Fig. 11 in Bell & Rubenach, 1983). Other porphyroblast minerals may react under retrograde conditions and yet be preserved in porphyroblasts that grew partially as a result of that reaction (Vernon, 1977); e.g. omphacite preserved in albite (Bell & Brothers, 1985). NUCLEATION SITES AND GROWTH OF PORPHYROBLASTS Introduction Some minerals preferentially grow as porphyroblasts with a non-uniform distribution, whereas others, such as quartz, feldspar and mica (in general), apparently readily nucleate and grow, in metamorphic rocks, with a roughly uniform distribution and grain size. The latter arrangement, which includes most deformed rocks (as well as the matrix of porphyroblastic ones), is readily explained in terms of grain boundary free energies of like and unlike mineral phases (Bell & Rubenach. 1983). The reason for porphyroblastic growth is not so apparent. It cannot be the availability of constituents, as A12Si05 porphyroblasts consist only of the most abundant ions in metamorphic rocks. It has been suggested that nucleation of a porhyroblastic mineral is difficult (Rast, 1965), and hence, once it has nucleated, it will continue to grow rather than nucleate others. This

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Fig. 10 Biotitc porphyroblast that has been cut after growth by a zone of high strain. duc to redistribution of the deformation partitioning. N-scction. Plancpolarized light. Base 2.6 mm.

Fig. I I Staurolitc porphyroblast that has been cut aftcr growth by two zones of high shear strain. due to rcdistribution of the deformation partitioning. I t has been altered to 'sericite' in these regions. Note the dodecahedra1 garnet porphyroblasts inside the staurolite porphyroblast versus the elliptical ones in the schist matrix. N-section. Crossed polars. Base 10.5 mm.

implies that these minerals generally have a relatively high activation energy barrier for nucleation (Rast, 1965; Vernon, 1976). Therefore, nucleation sites for porphyroblasts may be controlled by microstructural circumstances that lower this energy barrier. Many porphyroblasts from a large number of deformed and metamorphosed terrains appear to have nucleated in specific microstructural sites, such as crenulation hinges (Figs 3,9, 10 & 12). They. also commonly preserve the apparently unaffected to crenulated remains of an earlier foliation, even where no remains of this structure are preserved in the matrix of the rock (Figs 13 & 14; Krige, 1916; Spry, 1963; Williams & Schoneveld. 1981; Bell & Rubenach, 1983; Bell & Brothers, 1985).

and microlithons of crenulated cleavage are zones which are dominated by the progressive shortening component of deformation history during progressive bulk inhomogeneous shortening (compare Fig. l a with Figs 3-5, 8 & 1014; Bell, 1981). As discussed previously, we argue that dissolution occurs in the progressive shearing-dominated zones, such as crenulation limbs, and consequently the only sites available for syndeformational porphyroblast growth in a rock which deforms in this fashion are zones of progressive shortening, which are either unstrained (except for the effects of microfractures, as discussed below) or unstrained in their centres and shortened on their margins, as shown in Fig. la. The irregular distribution of porphyroblasts within compositional layers in which they preferentially nucleate and also within zones of progressive shortening which transect these layers, suggests a very non-penetrative control

Nucleation sites Slightly affected or apparently unaffected remains of earlier foliation, crenulation hinges

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Fig. 12. Biotite porphyroblasts that have overgrown crenulations at stage 2 of their development. The porphyroblasts are prefentially located in the crenulation hinges. They tend to grow up the zones of progressive shortening-partitioned deformation resulting in elongate shapes. which are independent of crystallographic orientation. Plane-polarized light. Base 6.3 mrn.

on nucleation, such as could be expected with irregular microfracturing during deformation. Microfracturing would also provide a means of fluid access and thus a source of ions derived from the adjacent zones of dissolution, a surface for nucleation from a fluid (Fig. 15), and a means of removal of unwanted ions from the replacement process which is associated with porphyroblast growth (see below). Zones of shortening are likely sites for fracture perpendicular to the stretching lineation, or parallel to the stretching lineation but perpendicular to the foliation generated during that deformation event, for the following reasons: 1 . They are zones of extension and contain zones of extreme attenuation (Figs la and 15; cf. Bell, 1981). 2. They commonly contain an earlier anisotropy that lies at a high angle to the x or y axis (or both) of the strain ellipsoid and therefore aids fracture formation (Figs 15, 17 & 18). (3) Most rocks contain a variety of minerals that have a wide range of ductilities at any T,P condition. Relatively strong minerals deform the least, and the deformation consequently partitions around them, such that they become sites of progressive shortening with progressive shearing occurring on their margins. Pressure fringes form on ends that are perpendicular to the stretching lineation in extreme cases of extension (Figs 19 & 20).

Another possibly significant factor is the strain energy that would be expected to accumulate in zones of progressive shortening. The simplest way to envisage this is in terms of crenulation hinges versus limbs. Strain energy is removed on the crenulation limbs by dissolution of the highly strained edges of minerals, such as quartz and feldspar, combined with the recovery and local recrystallization of the micas which occur via kinkband and grain-boundary migration, as discussed above (Fig. 6; cf. Swager, 1985). On the other hand, strain energy accumulates in the crenulation hinges and is microstructurally visible in the form of undulose extinction in quartz, feldspar and mica. We suggest, however, that this stored strain energy can be removed by a reaction associated with the nucleation and growth of a porphyroblast in these zones. Furthermore, this may cause an effective lowering of the activation energy barrier for the nucleation of the porphyroblast, as an overall reduction of free energy will be achieved (Bateman et al., 1986). Figures 3. 8 and 10-14 and 18 show this characteristic localization of porphyroblast growth sites. Potential pressure-fringe and fibre-growth sites, the extreme examples of the sites of attentuation described above, could therefore also be good locations for porphyroblast nucleation, depending on the chemistry of the minerals involved. We have found andalusite grains that are apparently preferentially

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Fig. 13 Largc staurolitc porphyrohlasts that havc ovcrgrown stage 3 of crenulation-cleavagc developmxt with suhscqucnr obliteration of crcnulation (and S,) outside their margins. Quartz has not been dissolved from regions protected from progressive shewing by the staurolites (strain shadows). N-section. Planepolarized light. Base 8.2 mm.

Fig. 14 Gamct porphyrohlasts that havc ovcrgrown crcnulations at stage 1/2 of their development. The porphyrohlasts havc a non-coaxial ‘millipede‘ geometry. Note the strain-shadows where the amount of quartz removed is much less because this portion of the foliation was protected from the shear strain by the presence of the porphyrohlasts. These zones are not prcssurcshadows or fringes. and locally. immediately adjacent to the porphyroblasts, a few rclics of S, arc preserved. N-section. Crossed Polars. Base 26.3 m m .

nucleated in strain-shadow regions adjacent to earlier cordierite porphyroblasts (Fig. 19). Other examples are biotite porphyroblasts that nucleated against staurolite porphyroblasts (Fig. 20) and between boudinaged garnet porphyroblasts (Fig. 16).

through the rock once deformation has commenced, because of competency contrasts and strain softening effects (Figs la & 21). Grains, and the rocks they lie within, are constrained by the grains and rocks around them and hence cannot deform without considerable interplay and interaction with their surroundings. This enhances the shifting of zones of progressive shearing and progressive shortening relative to one another (Bell, 1986; Jones, 1986). Figure la shows a possible strain field in a rock that has been deformed by progressive bulk inhogeneous shortening. Neglecting fracturing, strain rates would have been slow to negligible in the zones of rock within this strain

Nucleation sites and apparent rate of growth Many porphyroblasts appear to have grown rapidly relative to the strain rate (Bell & Rubenach, 1980, 1983). However, alternative interpretations follow. We mention above that zones of progressive shearing and progressive shortening move

Porphyroblast nucleation, growth and dissolution

Fig. 15. Sketch showing a zone of progressive shortening bounded by zones of progressive shearing. in which dissolution has occurred. A strain ellipsoid for the crenulation event is shown. with the X-axis nonorthogonal to the earlier foliation preserved in the zone of progressive shortening. Fractures can readily form in this zone (shown in solid black) because the anisotropy due to the earlier foliation lies at a high angle to the extension direction (the X-axis). This allows access of fluid associated with dissolution which occurs in and immediately adjacent to the zone of progressive shearing. to the zones of progressive shortening.

field that underwent progressive shortening, particularly in those zones surrounded by dashed lines. Thus, as long as the pattern of deformation partitioning did not shift for some period, a porphyroblast could have grown relatively slowly in a zone of progressive shortening, yet give the appearance of having grown relatively fast, because the inclusion trails may not reveal any geometric effects associated with the synchronous deformation (Figs 22 & 23). Subsequently, the strain heterogeneity pattern might shift, or the porphyroblast might grow to the limits of a shortening-dominated strain (Figs 22 & 24). Further growth in that location would be precluded because progressive shearing was concentrated on the porphyroblast margins which cause dissolution (Figs 24 & 25). In this respect, garnet (in particular),

49

Fig. 16 A f - v x t i o n , in which garnct porphyrobliists

have boudinaged. duc to high extensiond strains parallel to the stretching lineation. The garnct boundarics mesh back together in the left-hand porphyroblasts. which indicates that it has been pulled apart and not simply replaced by biotite. Biotite porphyroblasts have grown in the pressure-shadows. This gives microstructural confirmation of the highly localized progressive extensional strains which are possible within progressive shortening-partitioned zones (cf. Fig. 19). Plane-polarized light. Base 5.3 mm.

chloritoid, cordierite, plagioclase, staurolite and andalusite commonly show straight inclusion trails, especially in their centres-evidence of very early syndeformation growth, interpreted by many petrologists as ‘pre’ deformation growth. Commonly, only an examination of numerous thin sections from the one specimen or locality, at very high magnification, reveals a slight curving (strain) of inclusion trails on some of the porphyroblast margins (e.g., Figs 24 & 25). The rims that contain bent inclusion trails are generally very narrow, because localization of progressive shearing against the porphyroblast may cause dissolution of these edges (Figs 9 & 11). Thus, a lack of curvature of inclusion trails on the edge of a porphyroblast does not preclude syndeformation growth. Porphyroblasts commonly appear to predate

103 ,her\ IOU paau s a w q)moi% iselqoJdqdJod ‘asuanbasuos e sv .d!qsuopelaJ qimor% uo!) -ewopp-isod iuaiedde ue u! pasnpoid isnl uoy -ego3 aqi M O J ~ J ~ Apue O S ~ U O Zpauo!i!ved-%u! -UauOqS U! MOJ% PInOM S)SE1qOJdydJOd ‘U!eJ)S JeaqS q%!g30 SaUOZ MOUEU ‘paseds d[ap!M aJOW qi!m am!i ieqi ie pasodm! u!eJis leioi aqi saiep -0mmo33e dllesuiarnoaS y s o ~aqi pue sdoJp aiei u!eJis aqi se ‘uo!ieuuopp 30 pua aqi i e a1eJado osle I[!m euamouaqd asaqL . ( ~ puei s % g ) uoyeauy Su!qsiaJis mau agio1 sal%ueq%!q 01 aieiapom ie %u$ dluowmos uo!ie!loj %u! -is!xa-aid e 01 anp ‘AdoJios!ue ue Suole Ued 01 u!eiis %u!Jeaqs qS!q 30 sauoz uaaMiaq I ~ O JJOJ dsuapuai aqi dq pamequa aq osle plnom uop -eapnu asuaq pue p!nu 30 ssassv .uo!ieuuopp 30 iuauodmos pauo!i!iied-%uueaqs aqi dn uaqei aAeq ieqi S ~ ~ Oiua3eFpe J moij uo!iem~opp e u! dpea alqepme Appeal ale qimoi%iselqoidqd -Jod ioj siuauodmos aqi ieqi iiasse a~ q!eJi uo!snpu! iq%!eiisu!eiuos pue S!q a i p b aq ppo3 UUOJ )eqi siselqoidqdJod aqL w a h a s!uolsai e %upnpdpea JO aio3aq U M O ~ %aAeq 01 ieadde ieq] ‘sist?[qoJdqdlod JO JaqWnU aA!lelaJ ale -iaSPexa pinom s!qL .uo!ieuuojap isry e suunp ueqi paqesol a ~ o mdlqeqoJd aJe ‘(aldwexa 103) sploj Jayiea 30 asuasaid aqi se qsns ‘sa!iaua%oJaiaq hemud-uou 01 anp ‘sisa33a Bu!uayos 3!JIaUIOa% ‘uo!leluJopp 30 saseis dpea aqi u] %~!mp iuaiedde s!qi asuequa sJol -3e3 30 Jaqmnu v woyeuuojap iaiel JO puosas

Porphyroblast nucleation, growth and dissolution

51

Fig. 19 Andalusite grains (arrowcd) that have apparently preferentially nucleated i n the quartz-rich pressureshadows which originate on cordierite porphyroblasts (c). The cordierite has been completely pinitized. The crenulations post-date both porphyroblast types and fold the pinitized cordierite. The sections are cut perpendicular to the crenulations and nearly parallel to the stretching lineation in the crenulatcd foliation. Scale bar i s I mm. Sketches traced from a camera lucida image of a specimen from north-east Victoria. Australia.

always so. The replacement of strained phyllosilicates by porphyroblastic material in zones of shortening-partitioned deformation may lower the overall free-energy more than the replacement of quartz and feldspar grains. This is because the latter minerals can recover by subgrain formation, whereas phyllosilicates, in general, cannot (Etheridge et al.. 1984; Vernon, 1977; Bell, 1979), although they can kink, which allows dislocation densities to decrease in a somewhat similar manner to that which occurs during recovery. However, phyllosilicates within zones of progressive shortening, such as crenulation hinges, in general do not kink but bend uniformly, and consequently retain all the strain energy associated with that curvature. They may, thus, be more important in providing the energy contribution needed to overcome the, apparently, relatively high activation energy barrier for nucleation of porphyroblasts. Phyllosilicates in these zones of progressive shortening would also readily fracture along (001) or their boundaries with other minerals, providing easy access for the fluids needed to nucleate and grow porphyroblasts, and they or adjacent minerals, such as quartz and feldspar, may even act as a substrate for nucleation. During replacement, the phyllosilicates presumably react with these fluids as the porphyroblast nucleates. A possible situation which

Fig. 20 A staurolite porphyroblast that has been partially replaced by andalusite (light) immediately adjacent to garnet (at extinction). This is a f-section that is cut parallel to the stretching lineation. and biotite beards and porphyroblasts have grown off the, staurolite and garnet porphyroblasts i n a number o f locations. Crossed polars. Base 10.5 mm.

shows a phyllosilicate/quartz grain boundary at the stage of nucleation of a porphyroblastic mineral, is drawn in Fig. 26. This boundary is depicted with an island structure (Raj & Chyung, 1981) which would allow ready fluid access, even after the microfracture (which enabled rapid fluid and ion access for porphyroblast nucleation) had closed. We speculate that during replacement, mica islands dissolve and porphyroblast islands enlarge (Fig. 26) in a continuous and progressive fashion, until the phyllosilicate is consumed by the porphyroblast. The fluid along the grain boundary must be continually enriched in ions that are more abundant in the phyllosilicate than in the porphyroblast, and continually depleted in ions that are more abundant in the porphyroblast. Thus, ions must diffuse andor flow into and away from the site of replacement, presumably

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T. H.Bell et al.

Fig. 21 Schematic diagram that illustrates variation in the concentration of zones of dissolution which are generated from partitioning of the deformation into progressive shearing and progressive shortening components, the former anastomosing around the latter (cf. Fig. la). The amount of shear on each zone of dissolution is approximately the same. and the effects of an increase in concentration of these zones on an earlier layering is also known. Redistribution of deformation partitioning, due to further deformation. would generate more zones of progressive shearing through the less strained region, tightening the fold. Other zones of progressive shearing may cease operating as a result of fold formation and become zones of progressive shorteningpartitioned deformation in a strain-field diagram applicable to that time.

under the control of chemical potential gradients. The effective net consequence is the sweeping of a high-angle grain boundary across the phyllosilicate,converting it to the porphyroblastic mineral. Nucleation sites and porphyroblast shape A possible shape control on the amount of rotation of porphyroblasts has been suggested by several authors (e.g., Ferguson & Harte, 1975; Ward, 1984). However, for apparent rotations of the internal inclusion trails of less than or equal to 180",one cannot be certain that a porphyroblast has rotated. 'Rotation' of the internal inclusion trails within a porphyroblast, relative to the matrix, can be due to overgrowth of zones of shortening strain, such as crenula-

tions, andor reactivation of the external foliation due to subsequent deformations (Bell, 1985, 1986). The geometry of deformation partitioning, and its effects on previously formed foliations, can considerably affect the length-to-width ratios of porphyroblasts relative to the orientation of their internal inclusion trails. This is shown schematically in Fig. 17, where porphyroblasts that overgrow an earlier foliation, that has been rotated more towards the axialplane crenulation cleavage, tend to have the greatest length-to-width ratios. These ratios appear to be a function of the old foliation control on the geometry of microfractures in zones of shortening strain, combined with growth by replacement from fluids that enter these fractures, as explained above (Fig. 26).

Porphyroblast nucleation, growth and dissolution

shearing'

shortening

DEFORMATION

53

shearing

PARTIONING

Fig. 22 Schematic diagram that shows porphyroblast nucleation as a result of deformation partitioning in a zone of progressive shortening. Dissolution in zones of high progressive shearing provides ions and possibly fluids from the reactions that occur there. These ions and fluids are transported to zones of progressive shortening due to microfracturing. which preferentially develops in the latter zones because they are sites of extension with a strong anisotropy generally lying at an angle to the stretching lineation. Porphyroblasts nucleate from the ions transported into these zones and generally grow by replacing the phyllosilicates. They can grow steadily larger until they reach a zone of progressive shearing. at which point they stop enlarging, because of the effects on their rims of dissolution associated with these zones.

Porphyroblasts that grow in these locations tend to become elongate up the zone of shortening, as shown schematically in Fig. 27. This characteristic can often be used mesoscopically as well as microscopically, to time the growth of porphyroblasts relative to formation and deformation of the surrounding foliations. For example, the biotite porphyroblasts in Fig. 12 tend to be elongate up the shortening zones defined by the weakly overprinting crenulations. Porphyroblast growth versus matrix growth

Porphyroblastic and matrix minerals tend to be chemically dissimilar, except for biotite, muscovite and feldspar, which grow as both matrix and porphyroblastic minerals. Why are matrix minerals generally non-porphyroblastic? Microstructural studies of strain transitions, across

which a new foliation has developed, indicate that the original microstructure of the rock can be totally destroyed if the deformation is sufficiently intense, and, as a result, the grain size and distribution of matrix minerals become fairly homogeneous at a different scale (or even the same scale). Examples of this are phyllonite development in the centres of major mylonite zones and stage 6 of crenulation cleavage development (Bell & Rubenach, 1983). This relatively homogeneous matrix microstructure occurs at much lower strains in slates, because bedding fissilities which result from diagenesis are generally weaker than tectonically generated foliations. This phenomenon has been attributed to like-phase mineral boundaries that have free energies equal to or higher than unlike ones (Smith, 1952; Vernon, 1968). Nucleation of matrix minerals may involve preformed nuclei, such as partially dissolved,

54

T. H.Bell e t al. to be precursors within the matrix, we suspect that porphyroblasts of these minerals have a different composition from similar matrix grains when they grow, and that these minerals have higher grain-boundary free energies against other minerals. However, phyllosilicates, apparently, readily chemically reequilibrate (Vernon, 1977) and hence it would be difficult to demonstrate that porphyroblastic phases were chemically different than matrix ones, although one such example for biotite has been described by Kamineni & Carrara (1973). This should be easier to demonstrate for plagioclase porphyroblasts versus matrix grains, as the tendency for this mineral to re-equilibrate should be much less. TIMING OF PORPHYROBLAST GROWTH Syndeformation growth

Fig. 23 'Millipcdc' plagioclase porphyrohlast that nucleated in a zone of progressive shortening-partitioned deformation. hut has not grown to the limiting boundary for growth caused by the zone of progressive shearing on the right. The deformation history in this example is very close to coaxial. Cross p l a n . Base 13. I mm.

disaggregated or complexed species. The predominance of matrix minerals such as phyllosilicates, quartz and feldspar in rocks from very low or moderately high metamorphic grades, may mean that numerous nuclei are available and, consequently, porphyroblasts tend not to develop. This is not true for minerals like andalusite, garnet, cordierite and staurolite which, in general, have no structural or chemical precursors in lower grade rocks. Hence, multiple nuclei of these minerals are much less likely and this results in porphyroblastic growth. As an alternative, grainboundary free energies for porphyroblastic minerals against other minerals may be much higher than against the same mineral; i.e., the reverse of the norm (Smith, 1952). This favours continued growth at the one site, rather than nucleation throughout on other mineral boundaries (e.g., Bell & Rubenach, 1983). Concerning muscovite, biotite and plagioclase porphyroblasts. for which there do appear

Deformation and concurrent metamorphism can cause dissolution and provide fluids (especially during prograde metamorphism) which contain a large range of the components essential to the nucleation and growth of porphyroblasts. Deformation provides sites for nucleation and growth, and a means of overcoming the energy barrier for nucleation, in the form of the energy contribution available from stored strain energy removal from, for example, crenulation hinges. Deformation and concurrent high fluid pressure (Etheridge et a!. , 1984) can also provide the fractures for local fluid access, subsequent reaction and replacement. Why should porphyroblasts ever nucleate, andor grow in a static situation, during which all these favourable conditions are removed? Even porphyroblasts in contact metamorphic rocks are not really growing in a static situation (Vernon & Powell, 1976). The intrusion of granite not only provides considerable heat and fluid circulation, but also causes fracturing and, at least in some places, ductile deformation of the surrounding country rock, no matter what the intrusion mechanism. This will allow ready and rapid access of fluid through the surrounding country rock, enabling porphyroblast development. Many rocks have been macroscopically folded several times, uplifted and eroded, and yet retain stored strain energy generated at the time of porphyroblast growth three or four deformations previously. This is discernable as regularly varying extinction about the relics of crenulation hinges formed in that even in

Porphyroblast nucleation, growth and dissolution

55

Fig. 24 Chloritoid porphyroblast that nuclcatcd in ii zonc of progrcssivc shortcning and stopped growing whcn it rcachcd the zones of progressive shcuring o n thc tims o f this zonc. Finc ixlusion trails that arc difficult to ohscrvc at this magnification. arc just heginning to curve on thc ends of thc porphyrohlast (hcst scen on thc right-hand end). The ovcrall crenulation gcomctry across thc porphyrohlast p~irallclto the crcnulntion clcawgc (S2).is that of antiforms in S I changing to synforms (espccially obvious on thc left). This indicatcr that the dcformation was closc t o coaxial in the earlier stages and hccamc morc non-coaxial suhscqucnt to thc porphyrohlnst growth. Planepoliirizcd light. Basc S . 0 mm

minerals such as quartz, feldspar and phyllosilicates; e.g., the Robertson River Metamorphics (Black, Bell, Rubenach & Withnall, 1979) and the Juntala Schists (Duncan, 1983) in north-east Australia. Yet during all that subsequent history, not one porphyroblast may have grown, the only new minerals occurring in zones of more intense crenulations associated with the subsequent deformations (e.g., Bell & Rubenach, 1983). Other rocks do not preserve this strain energy in crenulation hinges, and yet retain the crenulation geometry (e.g., Marlow & Etheridge, 1975). the minerals having recrystallized subsequently, though syndeformationally (e.g., Bell & Brothers, 1985). Therefore, it appears to us that mineral growth in metamorphic rocks in general, and porphyroblast growth in particular, is syndeformational, the activation energy barrier for nucleation and growth of these minerals being overcome by the energy input from concurrent deformation. This appears to be supported by a critical re-examination of the published literature on porphyroblast timing, as follows. Literature reassessment

Zwart (1960a. b, 1962) published a set of micro structural criteria based on internal inclusion

trails (Si) versus external foliation (S,) to-infer porphyroblast growth timing relative to deformation in multiply deformed tectonites. Many investigators have attempted to use these and other microstructural criteria. However, as the number of workers has increased, so has the amount of contention about the interpretation of a number of porphyroblast/matrix relationships. This particularly applies to the interpretation of relationships between inclusion trails in porphyroblasts and the foliation trends in the adjacent matrix. Discussions about alternative explanations for some of these relationships include Ferguson & Harte (1975). Dixon (1976), Vernon & Powell (1979). Vernon & Flood (1979), and Bell & Rubenach (1983). With these problems in mind, and noting that many workers using microstructural criteria arrive at porphyroblast growth vs deformation histories that are rather complex, we have attempted to resolve variable interpretations of a number of common microstructural relationships. This has been done to arrive at histories that are as simple as possible (yet consistent with the microstructural evidence), and that also seem to better fit histories predicted on the basis of what we regard as the more important metamorphic and deformation processes

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T. H. Bell et al.

Fig. 25 'Millipede' andalusite porphyroblast, from the aureole of the Cannibal Creek diapiric granite. Queensland. Australia. that has grown in a zone of progressive shortening at an early coaxial stage and stopped growing when it reached the zone of progressive shearing. as can be seen by the millipede geometry revealed on its rims. Photograph courtesy of Bateman (1985). Crossed polan. Base 3.2 mm.

involved in porphyroblast formation, as discussed later. A result of this approach is that porphyroblasthatrix microstructural relationships, with apparently different timing implications (on the basis of current interpretations), can be reconciled. For example, Bell & Rubenach (1983) explained all porphyroblast-matrix relationships (including ones that would be conventionally rejected as conflicting) as indicating synD2 growth of all the porphyroblasts. Abundance of syndeformation [ D z ( + ~ ) / porphyroblasts

In the assessment of rocksfrom avariety of areas, we have been impressed by the large proportion of porphyroblasthatrix microstructural relationships that indicate (or at least areconsistent with) growth during crenulation (commonly the D2event) of an earlier foliation, other than a bedding fissility. Furthermore, many published

examples can be fitted into one of three categories: (1) those described by the authors as indicating growth during a crenulation D2(+") event; (2) those that can be reasonably reinterpreted as having formed at such a time and, (3) those that need not be interpreted in this way, but nevertheless show microstructural relationships consistent with syn D2(+n) growth. Porphyroblasthatrix microstructural relationships that indicate growth during a Dz(+,,) crenulation or crenulation cleavage event may take a number of forms. The literature contains a few convincing examples of porphyroblasts that grew and rotated during D2. and that formed 'snowball' structures (e.g., Rosenfeld, 1968; Schoneveld, 1977; Powell & Vernon, 1979). Another form, where the porphyroblast, from core to rim, has successively inherited crenulations of increasing amplitude and decreasing wavelength, is one of the diagnostic cases listed by Zwart (1962 fig. 1; 1960, fig. 4.4), but these are fairly rare. Forms where the porphyroblast preserves an earlier stage of crenulation cleavage development than that seen in the adjacent matrix are much more common, and are diagnostic of growth during that deformation episode. An example from the Mt Lofty Ranges, South Australia, is shown in the central staurolite porphyroblast of Fig. 28. Here the matrix has a dominant foliation (S2) that is a highly evolved crenulation cleavage (stage 516 of Bell & Rubenach, 1983). The inclusion trails in the central porphyroblast reveal crenulations (with axial planes parallel to S2) that preserve an earlier stage of crenulation development; see also Zwart (1960, fig. 4.5) and Bell & Rubenach (1983, figs 6 and 7). The 'millipede' microstructures of Bell & Rubenach (1980) are a special variation of this form. The second category of relationships, mentioned above, are those that singly, are not necessarily diagnostic, but which may nevertheless, singly or considered with neighbours, permit an interpretation of growth during a crenulation or crenulation cleavage event. Numerous examples exist in the literature, often with quite different interpretations. A few are discussed below. Figure 29a is traced from a photograph in Spry (1963, plate 1 and 2). The garnet porphyroblast was interpreted by Spry (1963, p. 203) as having 'a snowball centre (syntectonic) with a massive idioblastic rim (posttectonic) . . . Micas outlining S2 wrap around the garnet which is thus pretectonic to S2.. By implication this would make the 'snowball' core syn-SI. The inclusion trails do not have the appropriate geometry for a 'snowball garnet'

Porphyroblast nucleation, growth and dissolution

57

Fig. 26 Schematic diagram of a grain boundary between quartz (clear) and a phyllosilicate (dotted) at the initial stage of nucleation of a porphyroblastic mineral. The boundary, drawn with an island structure, has parted. allowing fluid access along the resulting microfracture. in a fashion similar to that which could occur in zones of progressive shortening. A porphyroblastic mineral (black) has nucleated between two phyllosilicate islands from ions or ionic complexes brought in with the fluid. Replacement of the phyllosilicate grain boundary islands by the porphyroblastic mineral (striped) has begun. It is common for the phyllosilicate to be preferentially replaced and hence the striped region would expand along the quartzlphyllosilicate grain boundary and into the phyllosilicate as the porphyrblastic mineral grows.

Fig. 27 Schematic diagram. modified from Fig. 17, which shows the tendency

for porphyroblasts to eventually become elongate up the zones of progressive shortening. Compare this with Figs 8 & 12.

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T. H.Bell et al.

Fig. 28 Camcra lucida skctch of stiiurolitc porphyrohlasts in an andalusite-staurolite schist, eastcrn Mt Lofty Ranges. South Australia (after Fleming, 1971). The central porphyroblast contains complete remains of crenulations. the axial planes of which are parallel to the external foliation. The axial plane of the trails in the other porphyroblast (shown approximately by dashed lines) is not fully determinable. as the sigmoid preserved does not crenulate the earlier foliation through more than 9(P. In our experience, inclusion trails which curve through less than Yoo give apparent axial trends that are variable, whereas those that curve through more than YO" (and less than IXOo) in the same thin section have constantly oriented axial planes.

(Schoneveld, 1977)-unless the sense of rotation changed during growth-but are much more easily interpreted as preserving Dz crenulations. Figure 28b shows our interpretation of the internal microstructure. The axial traces of the included crenulations curve, but this is easily explained if the garnet grew during Dz, and successive growth stages inherited inactive crenulations, the axial traces of which bowed around the growing garnet while Dz continued. Some of the other 'snowball' porphyroblasts which have been interpreted by Spry (1963) as syn-D, are also consistent with syn-D2 growth, especially where enough matrix is illustrated to show that the axial traces of sinusoidal inclusion trails parallel the axial traces of crenulations (Spry, 1963, fig. 3111, redrawn as Fig. 30 of this paper). A similar example from Spry (1963. fig. 2E, redrawn as Fig. 31 of this paper) depicts an albite porphyroblast considered by him to record post F,-pre F2 growth, because of the helicitic inclusion trails. The external foliation

is a crenulation cleavage (S2). with residual F2 fold hinges, the trends and shape of which are close to the curves in the inclusion trails. These relationships are consistent with syn-D2 growth of the porphyroblast along a zone of shortening. An albite porphyroblast from Takagi & Hara (1979, fig. 4; redrawn as fig. 32 of this paper) is more difficult to interpret. These authors reported that all albite porphyroblasts in the host amphibolitic schist show a similar relationship, with the Si of the core at a uniform 25" to the main matrix 1 schistosity (Se), They reasonably concluded that the curved part of the Si near the albite rims indicates albite growth during a second deformation. However, they ascribed the more planar Si in the cores to post DI-pre Dz growth. This is possible, but considering the later history of the porphyroblast we suggest that the cores have grown early syn-D2 when the progressive shearing component of the deformation, in this location, was less penetratively developed, or even later than this, in parts of the rock where Dz deformation partitioning had left (or temporarily left) broad domains of progressive shortening surrounded by widely spaced anastomosing zones of progressive shearing (Figs 1 & 17). A final point is that in many rocks, careful observation of porphyroblasts with approximately planar Si trails reveals that the t r a h show slight and/or distinct curvature just near the edge of the porphyroblast (e.g., Figs 8, 10, 14, 24,25), this implies that the porphyroblasts were syntectonic and stopped growing when they reached zones that were accommodating progressive shearing at the time of their growth. Lack of syn DI growth of porphyroblasts Porphyroblasts have only relatively rarely been interpreted as having grown syn-DI in regional metamorphic rocks (e.g., Fleming & Offler, 1968). The reason for this may lie in the scale of deformation partitioning in such rocks. The great bulk of psammitic and pelitic sediments have fine- to medium-grain size. They may or may not be fully lithified prior to the first deformation, but in general have more pore space than sediments that have been deformed and metamorphosed. Hence, during a first deformation, a large amount of layer-parallel shortening can be expected before the nucleation of folds, as pore space is removed and because of a general lack of penetrative anisotropy of similar intensity to a deformation-produced foliation. As a consequence, strain must be relatively

65

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T. H . Bell et al.

s 1' Fig. 30 Re-sketched from Spry (1963. fig. 3111). Spry interpreted the dodecahedra1,garnet porphyroblast in the upper centre of the sketch as syn-FI and the other porphyroblasts (albite) as post-S1and pre-SZ. We suggest that the garnet may have grown early during D2 at stage 2 of the crenulation cleavage development. The large and small albite porphyroblasts also grew early during D2. The straight trails in the large porphyroblasts versus the curved trails in the small porphyroblasts are simply products of the different scales of deformation partitioning at these two localities, prior to albite growth.

Fig. 32 Trends of S,and S, in. and adjacent to. an albite porphyroblast in an amphibole schist from the Sambagawa Schist, Central Shikoku. Japan. traced from a photo in Takagi & Hara (1979). who referred to S,as 'S'and S, as 'Sl*. The inclusions are mostly amphibole with epidote and opaque grains. 'S' in the centre of this porphyroblast. as well as others in the They concluded that the rock. is oriented at 2.5" to 51'. inner core grew post-5'. pre-'S,'. the outer core grew syn-'SI'. and the mantle area late-'SI'. We infer that the porphyroblast grew syn-DZ and equate Dz with their DI. ~~

1m m Fig. 31 A re-sketching typical example of an albite porphyroblast in Spry (1963. figs 2b. 3 and plate I . No. 4). Spry interpreted the albite porphyroblast as post-tectonic to FI, but pre-tectonic to Fz. We interpret the albite as having overgrown slightly crenulated S, early during D2, with some reactivation of the SI foliation after the porphyroblast had formed. Same scale as for Fig. 30.

homogeneously distributed in the early stages of deformation without localization of zones of

intense shear, and this may control the continuing development of strain as folds nucleate. Prior to deformation (to discount the effects of reactivation of So), bedding fissilities are only very strong in fine-grained sediments like shales and mudstones. Hence, when a crenulation cleavage does develop in these rocks during the first deformation (Weber, 1976), it has a fine spacing, because the wavelength of the initial crenulation is partially a function of the very fine-grain size of the clay or shale. If the rock

Porphyroblast nucleation, growth and dissolution deforms without crenulating, the partitioning of deformation takes place on the even scale of individual grains, as discussed previously (Figs l b and 5). Porphyroblasts apparently cannot grow while dissolution is operating on the scale of individual matrix grains or crenulation limbs only 10-50 microns apart. Deformation and metamorphism generally generate a much stronger and coarser-grained foliation than a bedding fissility in a far greater range of rock types and grain sizes. Thus, crenulation wavelengths are considerably larger in a second or later deformation event, causing coarser deformation partitioning and thus providing sites for porphyroblast growth. Crenulations are also very commonly developed during a second deformation within many different rock types, and the contribution to overcoming the activation energy of porphyroblast nucleation by removal of stored energy in crenulation hinges may be one of the controlling factors in the large number of D2-generated porphyroblasts. For syn-D, growth of typically porphyroblastic minerals, the fine scale of deformation partitioning inhibits sizeable growth of grains that otherwise might form porphyroblasts. Only small grains of these minerals may thus be preserved. These may be numerous but are not big enough to inherit, in the form of inclusion trails, a large enough portion of the matrix fabric to allow later diagnosis of the syn-DI signature. Furthermore, small index mineral grains, if preserved, are liable to be overgrown or enveloped by porphyroblast growth of the same mineral during D2 growth, and so the preD2 growth of the mineral would not be obvious.

PORPHYROBLAST ROTATION (OR LACK THEREOF) AND FOLIATION REACTIVATION These topics have been dealt with extensively by Bell (1985, 1986) and so are only briefly discussed here. Rotation of porphyroblasts during ductile deformation of the rocks in which they have grown or are growing has generally been accepted as typical of metamorphic rocks (Krige, 1916; Read 1957; Cox, 1969; Spry, 1969; Dixon, 1967; Powell & Treagus, 1970; Williams & Schoneveld, 1981; Olesen, 1982) possibly mainly because of the prominence, in the literature, of articles on unambiguously spiralling garnet porphyroblasts (Rosenfeld, 1968, 1970; Schoneveld, 1977; Powell & Vernon, 1979). The porphyroblasts described

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in the earlier references, however, can all be interpreted as never having rotated (Bell, 1985; Figs 28-33 of this paper), the apparent rotation being due to overgrowth of crenulations and/or reactivation of foliations (Bell, 1986). Bell (1985) has suggested that porphyroblasts that do not deform can only rotate in a special structural environment, in which the deformation involves progressive simple shear and cannot partition. This explains the relative scarcity or porphyroblasts that contain definitely spiralling inclusion trails, although it does not explain why they appear to be restricted to garnet (e.g., Rosenfeld, 1968; Schoneveld, 1977; Powell & Vernon, 1979; Bell & Brothers, 1985). We have demonstrated that porphyroblasts grow in specific locations relative to the deformation partitioning and cannot grow across active zones of progressive shearing. How, then, can garnet porphyroblasts grow in rocks undergoing progressive simple shear where deformation partitioning cannot occur? Garnet is a very competent mineral that rarely deforms plastically. When garnet does dissolve against an actively shearing foliation (e.g., Figs 9 & 11) the metamorphic conditions have changed such that the rims of the garnet are no longer stable (e.g., Bell & Rubenach. 1983). Hence, reaction softening of the porphyroblast rims can occur and allow the development of strain gradients between the core and rim, which also cause dissolution (e.g.. Fig. 7). However, if metamorphic conditions do not change, so that reaction softening does not occur when deformation partitioning brings zones of progressive shearing against the rims of garnet porphyroblasts then they do not dissolve but maintain their original shape. Hence, if a garnet grain nucleates in a zone of rock that is undergoing progressive simple shear, in which the deformation cannot partition, it will not dissolve on those rims that abut the actively shearing foliation (e.g.. fig. 4c in Bell, 1985). The foliation planes in this location the zone of crenulation that must form adjacent to those boundaries which lie perpendicular to the actively shearing foliation (e.g., Fig. 4c in Bell, 1985). The foliation planes in this location appear to part commonly allowing precipitation of 'saddle-reef'-like quartz (Bell, 1985). This allows access of fluids, and hence ions, in a manner very similar to that shown in Figs 15 and 26, except that the crenulations lie almost normal to the developing foliation. Therefore, as the garnet grain is forced to rotate, a porphyroblast can grow because a new surface is continually being brought into contact with the

- -

- -- Fig. 33 Re-sketched from Olesen (1982). Garnet porphyroblasts show orientation of internal inclusion trails versus the external foliation in a P-section (i.e., one cut perpendicular to the foliation but parallel to the stretching lineation). The garnet porphyroblasts are slightly elongate WSW-ESE across the diagram. This could be due to growth up shortening zones. as shown in Fig. 27. or dissolution of the garnet rims, as shown in Figs 9 and I I. The internal inclusion trails are similar in orientation from one porphyroblast to another. the range from the average orientation being less than +20" for the majority of the porphyroblasts. We interpret the external foliation as a product of rotation and reactivation of t h e foliation which is preserved within the garnet cores. and not a product of rotation of the porphyroblasts. as in Olesen's interpretation.

growth-zone abutting the crenulation, against which no shearing is occurring. As the porphyroblast rotates, the newly grown surface will remove into a zone of shearing strain. It will not, however, dissolve because of the high competency contrast between it and the matrix, and hence the lack of plastic deformation of the rim of the garnet grain, as discussed above. A consequence of the interpretation that porphyroblasts d o not typically rotate is that porphyroblast timing relationships become more easier to decipher. However, it is critical that thin sections are cut in quite specific orientations relative to structures in the rocks that are relevant to the timing of growth of the porphyroblasts. For example, if the porphyroblasts grew during D2,which locally or generally has little effect on the geometry of S I ,then thin sections cut normal to the stretching lineation Lt may give no information on porphyroblast timing relationships, because they may lie at a very low angle to the SI/S2 intersection lineation, LA, and so reveal no Fi folds (Bell & Rubenach, 1983). Thus, it is usually necessary to cut several thin sections before those containing the critical relationships are found.

The goemetrical relationships of stretching lineations, from the event that produced the porphyroblast to that which generated the previous foliation, are critical as to whether or not readily decipherable inclusion trails are preserved and can be used. For example, Olesen (1982) interpreted all the garnet porphyroblasts, resketched in Fig. 33, as having rotated. However, this section was cut parallel to the stretching lineation and this inclusion trail geometry is typical of thin sections with this orientation, where the external foliation has formed by reactivation of the foliation preserved within the porphyroblast (Bell, 1986) due to a subsequent deformation. The external foliation is rotated and reactivated about the porphyroblast, which does not rotate (Bell, 1985; Fig. 34). Olesen also cut a thin-section normal to the stretching lineation, but considered that the inclusion trails were random or uninterpretable. A possible reason for this was that the stretching lineation for the earlier deformation was approximately parallel to that for the subsequent one. Foliations commonly appear better developed in a plane parallel to the stretching lineation, due to the greater UW

Porphyroblast nucleation, growth and dissolution

Fig. 34 Sketch showing a porphyroblast surrounded by anastornosing foliation. A single layer of inclusion trails within the porphyroblast is shown. The dotted region in the centre is planar. To the east and west of this (with north pointing to the top of the figure) the curvature is produced by crenulation of the original foliation. To the southwest and north-east of this the curvature is produced by rotation of this crenulation in the original foliation. during subsequent reactivation. towards the stretching lineation for the deformation which causes the reactivation. The porphyroblast has grown during this deformation.

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ratio of grains in this sectidn and, as a result, the inclusions define much clearer trails. We infer that none of the garnet porphyroblasts had rotated (Bell, 1985) and that they have approximately the same orientation throughout the slide (the inclusion trails average 44" from the external schistosity, with a standard deviation of 13.5". because they grew early during the deformation that produced the external foliation (cf. Fig. 34). I t is possible that if more sections had been cut perpendicular to the stretching lineation, in this example, a clearer picture would have emerged as to the inclusion trail geometry in this plane, due to original inhomogeneities of the strain perpendicular to this plane. Another problem arises from rotation of the crenulationlintersection lineation associated with the event that produced the porphyroblast (e.g., the intersection of S, and S 2 ) towards the stretching lineation, due to continued deformation subsequent to the porphyroblast's growth. This is brought out in Fig. 34. Here, the crenulation/intersection lineation has been rotated into near parallelism with the stretching lineation. If the sections are cut perpendicular to the stretching lineation, they are not oriented perpendicular to the S l / S r intersection lineation at the time of porphyroblast growth, even though they are nearly perpendicular to the external intersection lineation. As a consequence, they will contain very little information about the timing relationships compared with the section cut perpendicular to the intersection lineation which is preserved within the porphyroblast. Hence, it is good practice to initially cut thin sections at 45", as well as parallel and perpendicular to the stretching lineation (and perpendicular to the foliation) in situations where this may have occurred. SUMMARY

The partitioning of deformation into progressive shearing and progressive shortening components causes strain and strain-rate gradients across the boundaries between the partitioned zones. These result in dislocation density gradients that generate chemical potential gradients and drive dissolution and solution transfer. Phyllosilicates and graphite can readily accomodate progressive shearing and hence, in general, are not affected by these strain gradients. Because of this they are concentrated in zones of progressive shearing by the dissolution and removal of other silicates and oxides. Exceptions to this are more

strongly deformed mylonitic rocks, in which strain rates are so high that plastic deformation dominates over diffusion in accomodating the deformation. Therefore, porphyroblasts cannot nucleate and grow in zones of active progressive shearing and the only syntectonic sites which are available to them are zones of progressive shortening. These zones are ideally suited for the nucleation of porphyroblasts. because they commonly contain earlier foliations that enablc microfracturing at a high angle to the maximum extension direction. This allows rapid access of fluids which carry some of the necessary components for porphyroblast nucleation from the adjacent zones of progressive shearing and dissolution. Porphyroblasts are commonly dissolved when repartitioning of the deformation brings a zone of progressive shearing against their margins or causes such a zone to cut through them as. for example, when the grade increases and reaction softening occurs. However, porphyroblasts are generally more difficult to deform than the surrounding matrix, due to their larger grain size and commonly more competent character. Hence, they control the distribution of deformation partitioning, such that progressive shearing is accommodated in the matrix on their margins. As a result, their adjacent strain shadows usually preserve evidence for the composition of the rock at the time they grew, demonstrating that volume losses (often after porphyroblast growth), which are associated with zones of progressive shearing. are very large at moderately high as well as low metamorphic grades. Thc criteria which are used by many geologists to indicate static growth of porphyroblasts between deformation events are not valid when considered in terms of the effects of deformation partitioning. I t appears that most, if not all porphyroblasts, have grown syntectonically. aided by the role of deformation in fluid generation and movement during metamorphism. Even porphyroblasts which grow in contact metamorphic aureoles essentially grow syntectonically, as the emplacement of the pluton provided the fluids and the fractures for their access, in a similar fashion to deformation during regional metamorphism. Porphyroblasts preserve various stages of the microstructural development of a rock. As a consequence, they reveal the extensive role of reactivation of pre-existing foliations during deformation. This is because they are not strained during redistribution of deformation partition-

Porphyroblast nucleation, growth and dissolution ing about their margins. This characteristic led geologists to regard porphyroblasts as objects that generally rotate during deformation. However, in general they do not rotate, with the relatively uncommon exception of genuinely spiralling garnet porphyroblasts.

ACKNOWLEDGEMENTS

Bell acknowledges financial support from the Australian Research Grants Scheme and N.S.F. EAR-8406572. We thank Roger Bateman and Ron Vernon for reading and criticizing various versions of the manuscript. REFERENCES Bateman. R. J.. 1985. Aureole deformation by flattening around a diapir during in situ ballooning: the Cannibal Creek Granite. Journal of Geology, 93, 293-301. Bateman, R. J . , Bell, T. H. & Rubenach. M. J., 1986. The effect of deformation partitioning, strain energy and solution transfer on the stability relations of metamorphic reactions. Journal of Metamorphic Geology, (submitted). Beach, A,. 1979. Pressure solution as a metamorphic process in deformed terrigeneous sedimentary rocks. Lithos, 12, 51-58. Bell, T. H.. 1978. The development of slaty cleavage across the Nackara Arc of the Adelaide Geosyncline. Tectonophysics, 51, 171-201. Bell, T. H.. 1979. The deformation and recrystallization of biotite in Woodroffe Thrust Mylonite Zone. Tectonophysics. 58, 139-158. Bell. T. H . , 1981. Foliation development: the contribution. geometry and significance of progressive bulk inhomogenous shortening. Tectonophysics. 75, 27% 296. Bell, T. H.. 1985. Deformation partitioning and porphyroblast rotation in metamorphic rocks: a radical reinterpretation. Journal of Metamorphic Geology, 3, 109-1 18.

Bell.-T. H . 1986. Foliation development and reactivation in metamorphic rocks: the reactivation of earlier foliations and decrenulation due to shifting patterns of deformation partitioning. Journal of Metamorphic Geology, (in press). Bell, T. H. & Brothers, R. N.. 1985. Development of PT prograde and P-retrogradelT-prograde isogradic surfaces during blueschist to eclogite regional metamorphism in New Caledonia as indicated by progressively developed porphyroblast microstructure. Journal of Metamorphic Geology. 3, 59-78. Bell. T. H. & Rubenach. M. J., 1980. Crenulation cleavage development-evidence for progressive bulk inhomogeneous shortening from 'millipede' microstructures in the Robertson River Metamorphin. Tectonophysics, 68, T9-TI5. Bell, T. H. & Rubenach. M. J . . 1983. Sequential porphyroblast growth and crenulation cleavage development during progressive deformation. Tectonophysics, 92, 171-194. Beutner, E. C. & Charles, E. G..1984. Large volume loss during cleavage formation in the Hamburg

sequence. Pennsylvania. Geological Society America, Abstracts with Programmes, 16, 445.

65 of

Black. L. P.. Bell, T. H.. Rubenach, M.J. & Withnall, I. N., 1979. Geochronology of discrete structuralmetamorphic events in a multiply deformed Precambrian terrain. Tectonophysics, 54, 103-138. Bosworth. W., 1981. Strain induced preferential dissolution of halite. Tectonophysics, 78, 5W526. Cox, F. C.. 1%9. Inclusions in garnets: discussion and suggested mechanism of growth for syntcctonic garnets. Geological Magazine, 106, 57-62. Dixon. J . M . , 1976. Apparent 'double rotation' of porphyroblasts during a single progressive deformation. Tectonophysics,34, 101-1 16. Duncan. A. C., 1983. On geometric analysis. Unpubl. PhD thesis. James Cook University, Townsville. Queensland.

Durney. D. W., 1972. Solution transfer, an important geological deformation mechanism. Nature, US.315317. Elliot, D.. 1973. Diffusion flow laws in metamorphic rocks. Geological Society of America Bulletin, 84, 2645-2664.

Engelder, T., 1982. A natural example of the simultaneous operation of free-face dissolution and pressure solution. Geochimica cosmochimica Acta. 46, 69-74. Etheridge. M. A.. Cox, S. F.. Wall, V. F. & Vernon, R. H., 1984. High fluid pressures during regional metamorphism and deformation-implications for mass transport and deformation mechanisms. Journal of Geophysical Research, 89. 43444358. Etheridge, M. A. & Hobbs, B. E., 1974. Chemical and deformational controls on recrystallization of mica. Contributions to Minerology and Petrology, 43, 1 11124. Etheridge. M. A. & Vernon, R. H.. 1981. A deformed polymictic conglomerate-the influence of grain size and composition on the mechanism and rate of deformation. Tectonophysics. 79, 237-254. Ferguson. C. C. & Harte. B.. 1975. Textural patterns at porphyroblast margins and their use in determining the time relations of deformation and crystallization. Geological Magazine, 1I t , 2 467-480. Fleming, P. D.. 1971. Metamorphism and folding in the Mt Lofty Ranges. South Australia, with particular reference to the Dawesley Kanmantoo Area. Unpubl. PhD thesis, University of Adelaide. Fleming. P. D. & Offler. R.. 1%8. Pre-tectonic metamorphic crystallization in the Mt Lofty Ranges. South Australia. Geological Magazine, 105, 356-359. Gray. D. R.. 1979. Microstructure of crenulation cleavage differentiation: implications of solution-deposition processes. Journal of Structural Geology. I , 73-80. Gray. D. R . , 1979. Microstructure of crenulation cleavages: an indicator of cleavage origin. American Journal of Science, 279,97-128. Gray, D. R. & Durney. D. W., 1979. Investigations of the mechanical significance of crenulation cleavage. Tectonophysics. 58, 35-79. Hammond. R. L.. 1985. Melange microstructures: the roles of shear strain, dissolution and deformation partitioning. Submitted to Geological Society of America, Bulletin.

Harker, A,. 1932. Metamorphism. Methuen. London. Henley. K. J.. 1970. The structural and metamorphic history of the Sulitjelma Region;Nonvay, with special reference to the Nappe Hypothesis. Norsk Geologisk Tidsskrifr, 50, 97-136.

66

T. H.Bell et al.

Hobbs, B. E., Means, W. D. &Williams, P. F.. 1976. An Outline of Structural Geology. John Wiley & Sons. Inc., New York. Jones, P. A,, 1986. The microstructural development of concentric ellipsoidal compositional shell (CECS) around nodules in metamorphic rocks. Submitted to Journal of Metamorphic Geology. Kamineni. D. C. and Carrara, A.. 1973. Comparison of the composition of porphyroblastic and fabric-forming biotite in two metamorphic rocks. Canadian Journalof Earth Sciences, 10, 948-953. Kerrich. R., 1977. An historical review and synthesis of research on pressure solution. Zentralblatt fur Geol. Paldontologic, I, 5 12-550. Krige. L. J.. 1916. Petrografixhe Untersuchungen uber das Verhaltnis der Schieferung zur Faltung under Berucksichtigung des Stockwerkproblems. Eclogae Geologicae Helvetiae. 14, 519-654. Kwak. T. A. P., 1974. Natural staurolite breakdown reactions at moderate to high pressures. Contributions to Minerology and Petrology, 44, 57-80. Marlow. P. C. & Etheridge, M. A,, 1977. Development of a layered crenulation cleavage in mica schists of the Kanmantw Group near Macclesfield. South Australia, Geological Society of America, Bulletin. 88, 87-2. Meneilly. A. W., 1983. Development of early composite cleavage in pelites from West Donegal. Journal of Structural Geology, 5, 83-97. Mitra, S.. 1978. Microscopic deformation mechanisms and flow lavas in quartzites within the South Mountain anticline. Journal of Geology. 86, 129-152. Olesen. N. 0..1978. Distinguishing between interkinematic and syn-kinematic porphyroblastesis. Geological Rundschau, 67, 278-287. Olesen. N. 0.. 1982. Heterogeneous strain of a phyllite as revealed by porphyroblast-matrix relationships. Journal of Structural Geology. 4, 48 1490. Powell, D. & Traegus. J. E.. 1970. Rotational fabrics in metamorphic minerals. Mineralogical Magazine. 37, 801-814.

Powell. C. McA. & Vernon. R. H., 1979. Growth and rotation history of garnet porphyroblasts with inclusion spirals. Tectonophysics. 54. 2543. Raj. R. & Chyung. C. K., 1981. Solution-precipitation creep in glass ceramics. Acta metallurgica, 2a, 159-166. Ramberg, H., 1952. The Origin of Metamorphic and Metasomatic Rocks. University of Chicago Press., Chicago. Ramsay. J. G.. 1962. The geometry and mechanics of formation of 'similar'-type folds. Journal of Geology, 70, 309-327.

Rast. N.. 1%5. Nucleation and growth of metamorphic minerals. In: Controls of Metamorphism (eds Pitcher, W.S . & Flinn, C. W.), pp. 73-102. Read. H . H.. 1957. The Granite Controversy. Thomas Murby & Co.. London. Robin, P. Y. F.. 1979. Theory of metamorphic segregation and related processes. Geochimica cosmochimica Arta. 43, 1587-1600. Rosenfeld. J. L., 1968. Garnet rotations due to the major Paleozoic deformations in south-east Vermont. In: Studies of Appalachian Geology (eds Zen, E.. et al.), pp. 185-202. Wiley Interscience. New York. Rosenfeld. J . L., 1970. Rotated garnets in metamorphic rocks. Special Paper of the Geological Society of America, 129.

Schoneveld. C.. 1977. A study of some typical inclusion patterns in strongly paracrystalline rotated garnets. Tectonophysics, 39, 453-471. Smith, C. S., 1952. lnterphase interfaces. In: Imprrfections in Nearly Perfect Crystals (ed. Shockley. W.) pp. 377401. Wiley. New York. Spry. A.. 1%3. The chronological analysis of crystallization and deformation of some Tasmanian Precambrian rocks. Journal of the Geological Society of Australia, 10, 193-208.

Spry, A., 1%9. Metamorphic Textures. Pergamon. Oxford. Swager, N.. 1985. Solution transfer versus mechanical rotation and kink band boundary migration during crenulation cleavage development. Journal of Structural Geology, 1, 421429. Takagi. K. & Hara, I . . 1979. Relationship between growth of albite porphyroblasts and deformation in a Sambagawa schist, Central Shikoku, Japan. Tectonophysics. 58. 113-126. Vernon. R. H.. 1968. Microstructure of high-grade metamorphic rocks at Broken Hill. Australia. Journal of Petrology, 9, 1-22. Vernon, R. H.. 1976. Metamorphic Processes. p. 247. Wiley, New York. Vernon, R. H., 1977. Relationships between microstructure and metamorphic assemblages. Tectonophysics, 39, 4 3 M 5 2 . Vernon. R. H., 1977. Microfabric of mica aggregates in partly recrystallized biotite. Contributions to Minerology and Petrology. 61, 175-185. Vernon, R. H., 1978. Porphyroblast-matrix microstructural relationships in deformed metamorphic rocks. Geologische Rundschau. 61. 288-305. Vernon. R. H. & Flood, R. H., 1979. Microstructural evidence of time-relationships between metamorphism and deformation in the metasedimentary sequence of the northern Hill End Trough, New South Wales. Australia. Tectonophysics. 58, 127-137. Vernon. R. H. & Powell, C.McA.. 1976. Porphyroblastesis and displacement: some new textural criteria from pelitic hornfels-a comment. Mineralogical Magazine, 40,787-788. Ward, C. M.. 1984. Geology of the Dusky Sound area. Fiordland. with emphasis on the structural-metamorphic development of some porphyroblastic staurolite. Unpubl. PhD thesis. University of Otago, Dunedin, New Zealand. Weber. K.. 1976. Gefugeuntersuchungen an transversal geschieferted Gesteinen aus dem ostliched Rheinischen Schiefergebirge. Geologische Jahrbuch, IS. 398.

Williams, P. J. & Schoneveld. C.. 1981. Garnet rotation and the development of axial plane crenulation cleavage. Tectonophysics. 10, 307-334. Wilson. M. R., 1971. On syntectonic porphyroblast growth. Tectonophysics, 11. 239-260. White. S. H.. Burrows. S . E.. Carreras. 1.. Shaw. N. D. & Humphreys. F. J.. 1980. On mylonites in ductile shear zones. Journal of Structural Geology, 2, 175-188. Wright, T. P. & Platt. L. B.. 1982. Pressure solution and cleavage in the Martinsburg Shale. American Journal of Science, 282, 122-135. Zwart. H. J.. 196Oa. Relations between folding and metamorphism in the Central Pyrenees, and the chronological succession. Geol. Mijnb., 22, 162-180.

Porphyroblast nucleation, growth and dissolution Zwart. H . J . . 196Ob. The chronological succession of folding and metamorphism in the Central Pyrenees. Geologische Rundschau, 50. 203-218. Zwart. H. J . , 1962. On the determination of polymetamorphic mineral associations and its application to the Bosost area (Central Pyrenees). Geologische Rundschau. 52, 38-65,

Received 24 December 1984; revision accepted 6 August 1985

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