RIESGOS DE TRABAJO Y ENFERMEDADES PROFESIONALES

6.1 Introduction

Identification of deformation microstructures and mechan-isms in rocks containing melt has several important tec-tonic implications, notably for the problem of melt extraction and in the interpretation of pluton ascent and emplacement mechanisms. However, microstructural criteria to distinguish magmatic, sub-magmatic and non-magmatic deformation mi-crostructures and mechanisms are not well established, and the distinction is best made using a combination of meso-scopic and micromeso-scopic evidence. The mesomeso-scopic evidence is briefly described in Section 6.4.

6.2 Fundamental deformation mech-anisms and microstructures in rocks containing melt

6.2.1 Magmatic flow

Flow of magma (i.e. melt and crystal phases) by transport of rigid crystals is often regarded as the typical deformation mechanism in melt-bearing rocks. Magmatic flow has been defined as flow by displacement of melt and rigid-body rota-tion of crystals without sufficient interacrota-tion to cause crystal plastic deformation (Paterson et al. 1989). As shown be-low, crystal interaction in melts may lead to deformation by cataclasis and diffusive mass transfer (DMT) as well as in-tracrystalline plasticity. A more useful, general definition for magmatic flow is flow of melt and crystals without crystal de-formation; this definition allows for the possibility that crys-tals may be deformed by processes other than crystal plas-ticity, and also describes the flow of a suspension. It leads naturally to a definition of magmatic microstructures as those which indicate melt-present deformation without crystal de-formation.

Some experimental studies suggest that viscosities of melt-laden systems reduce abruptly by orders of magnitude when the proportion of melt increases beyond a value known as the critical melt fraction, CMF (Arzi 1978, Van der Molen and Paterson 1979). The CMF is commonly taken to be 30%, but may be as much as 50% (Vernon et al. 1988) or as little as 10-20% for gabbroic rocks (Nicolas et al. 1988). Experi-ments on two silicate melts reported an increase in viscosity of three orders of magnitude, and a change from Newtonian to non-Newtonian behaviour, as melt fraction increased from

40 to 60% (Lejeune and Richet 1995). The importance of the CMF is further suggested by the observation that the max-imum proportion of phenocrysts in volcanic rocks is 55-65%:

volcanic rocks with larger proportions of phenocrysts may not be able to erupt because their viscosities are too high (Marsh

1981, Wickham l987).

A minimum melt proportion for magmatic flow can also be deduced from the critical packing density of crystals to bring them into a coherent mass. The critical packing density depends on crystal shape, size, size distribution, packing ar-rangement, and amount of compaction. Table 6.1 summarizes porosity at critical packing density, which is equal to the min-imum melt proportion, for some combinations of these factors in geometrical models, and estimates for magmas. Surface energies of the crystal and liquid phases may be important in flow because they can affect melt distribution (e.g. Jarewicz and Watson 1984, 1985), and indeed an explicit relationship between the CMF and surface energy can be formulated (Ri-ley 1990). A more fundamental parameter than melt fraction for determining the rheology of melt-laden systems may be the contiguity (the fraction of grain surface area in contact with other grains), because contiguity affects surface energy and the resistance to shearing at the average surface. A load bearing framework of crystals breaks down at contiguities of less than 0.15-0.2 (Miller et al. 1988), which correspond to variable equilibrium melt fractions depending on surface en-ergies and grain size and shape distributions (German 1985).

These factors may explain the variation in estimates of the CMF.

Viscosities of suspensions depend on fluid (melt) com-position, pressure, temperature, and proportion, size and shape distribution of solids. The viscosity of melts contain-ing spherical crystals is often approximated by the Einstein-Roscoe equation (Einstein-Roscoe 1952):

where is the viscosity of the pure melt, and is the crystal fraction for coherent packing. Values of 0.6 for seem to represent silicate systems well. n has a theoretical value of 2.5 for variably-sized spheres, which is also a good fit to data for silicate systems, but the viscosity-melt fraction relationship has a slight temperature dependence which can be allowed for by letting n vary between 2.0 and 2.5 with temperature (Lejeune and Richet 1995). Another relationship between viscosity and crystal fraction includes the effect of grain size (Sherman 1968). At lower melt fractions, the rhe-59

60 CHAPTER 6. MAGMATIC AND SUB-MAGMATIC DEFORMATION

ology of the magma may be at least partly controlled by the mechanical properties of the solid phases or by the factors af-fecting diffusive mass transfer (DMT) through the melt phase (see below).

While some experiments and observations suggest a signi-ficant mechanical change at the CMF, other experiments cast doubt on the existence of a large change in viscosity over a narrow range of melt fraction (e.g. Rushmer 1995), and it has been argued that the abrupt change in viscosity observed in some previous experiments was due to a combination of temperature effects and increase in water contents of melts (Rutter and Neumann 1995). There may be a continuous de-crease in magma viscosity as melt fraction inde-creases, and no CMF, if the change in water content is allowed for. Yet other experiments suggest that the change in strength at the CMF is due to dilatancy hardening (cf. Brace and Martin 1968) at low melt fractions (Renner et al. 1999).

6.2.2 Sub-magmatic flow

Sub-magmatic flow can be defined as deformation involving flow of melt and crystals with crystal deformation. Sub-magmatic microstructures are accordingly those indicating melt-present deformation with crystal deformation, and cor-respond to the pre-full crystallization fabric of Hutton (1988).

Crystal deformation by DMT through the melt phase is prob-ably the dominant deformation mechanism at low melt frac-tions and lower strain rates. The thermodynamic consider-ations for DMT through melt are the similar to those de-scribed for DMT through solution in Section 3.2. Experi-ments show that strength decreases by an order of magnitude, and intracrystalline plasticity changes to melt-enhanced diffu-sion creep, when melt proportion increases from zero to only 3-5% (Cooper and Kohlstedt 1984, Dell’Angelo and Tullis 1988, Dell’Angelo et al. 1987). The weakening occurs due to enhanced grain boundary diffusion though the melt. Ex-periments on rock analogues by Park and Means (1996) also demonstrate the importance of this process in some systems at low melt fractions. Surface energy considerations may be im-portant in sub-magmatic deformation. Most diffusion creep models assume isotropic surface energy, but measurements suggest an order of magnitude anisotropy in olivine, for ex-ample (Cooper and Kohlstedt 1982). This may affect strength by allowing a continuous film of melt along two-grain bound-aries instead of limiting melt to three or more grain junctions, as usually assumed (Hirth and Kohlstedt 1995).

6.2.3 Magmatic and sub-magmatic flow and rheology

It appears that two fundamental changes in deformation mechanisms occur as melt fraction is increased from the solid to liquid states, and these correspond to reductions in strength or viscosity. An order of magnitude decrease in strength occurs as melt fraction rises from zero to only a few percent melt, and solid-state deformation mechanisms change to melt-enhanced diffusion creep, with other deform-ation mechanisms (cataclasis, grain boundary sliding, and in-tracrystalline plasticity) also possibly accommodating crys-tal deformation. As melt fraction increases still further, sus-pension flow can occur, reducing strength by three orders of magnitude, and crystal deformation ceases. Figure 6.1 shows these two transitions in mechanism schematically, but it should be emphasized that the values and range of melt frac-tions over which the transifrac-tions occur, and the magnitudes of the changes in viscosity, vary with magma composition, and are presently the subject of some debate.

6.3 Mesoscopic evidence for magmatic and sub-magmatic flow

It is important to use mesoscopic evidence in conjunction with microscopic observations to distinguish magmatic, sub-magmatic and non-sub-magmatic flow. Fabrics in sub-magmatic rocks can be defined by phenocryst alignment, matrix mineral shape fabrics, compositional banding, enclave or schlieren shape fabrics, or magnetic anisotropy. Phenocrysts may be tiled or imbricated (e.g. Blumenfeld 1983, Blumenfeld and Bouchez 1988). The fabrics can have any shape from prolate to ob-late. The mere existence of any of these fabrics is not dia-gnostic of melt-present flow, but detailed examination on the mesoscopic scale may provide some strong indicators to dis-tinguish magmatic/sub-magmatic from non-magmatic flow.

The following criteria are proposed as diagnostic of magmatic or sub-magmatic flow because they suggest extremely large, non-systematic strain gradients in an homogeneous rock on a meter scale, which are mechanically unlikely in the non-magmatic state. Microscopic studies should be used to back up these lines of evidence.

1.

2.

3.

Abrupt and non-systematic changes in fabric orientation.

Abrupt and non-systematic changes in fabric intensity.

Irregular, polyclinal, disharmonic or rootless folds in

massive rocks that have no sign of mechanical aniso-tropy or shear surfaces (e.g. McLellan 1984).

The relationship between igneous layering and fabrics can also be used to suggest magmatic/sub-magmatic flow. For example, Pons et al. (1995) describe centimetre-scale cycles with sharp bases consisting of a layer of fine-grained am-phiboles to a medium grained amphibole-plagioclase layer to a slightly porphyritic plagioclase-K-feldspar-quartz layer in alkali granites. The layers are parallel to the preferred orient-ation in any of these minerals, and the fabric never cross-cuts the layering. The layering and fabric can be explained by trapping of a mafic cumulate layer from a slowly convecting magma, leaving an increasingly felsic rich magma to crys-tallize as the upper layers. By contrast, Paterson and Ver-non (1995) discuss several examples of an apparently mag-matic foliation that cross-cuts both gradational compositional changes and contacts between phases of different magmatic

compositions. These foliations may have formed in a short interval of cooling after the melt has been emplaced but be-fore final solidification.

Phenocryst density and size distributions may be affected by magmatic flow. Interaction between phenocrysts in con-centrations greater than 8% in a moving fluid creates a “grain dispersive pressure” which is proportional to the velocity gradient in the magma and is therefore greatest at the margins of an intrusion (Bagnold 1954, Komar 1972a, b). Phenocrysts are concentrated into the centre of the intrusion, and may also coarsen in this direction. The velocity distribution and pheno-cryst concentration should be plug-shaped with a central re-gion of high density and an abrupt decrease in density towards the sides of an intrusion, due to the effect of phenocryst con-centration on viscosity, and this is indeed observed in many natural examples (e.g. Ryan 1995).

Discordance between fabrics in a xenolith and a surround-ing igneous rock is often interpreted as evidence for

mag-62 CHAPTER 6. MAGMATIC AND SUB-MAGMATIC DEFORMATION

matic flow, but this criteria is not diagnostic: fabrics may be preserved in xenoliths during non-magmatic deformation due to competence contrasts between the xenolith and sub-solidus matrix. Sub- or non-magmatic fabrics on this scale are demonstrated by deformation of individual crystals such as feldspar megacrysts. However, concordance between fab-rics within enclaves and the host rock can be evidence for magmatic/sub-magmatic flow if the enclaves can be inter-preted as partially molten during deformation (e.g. Vernon and Paterson 1993).

Shear zones may develop during intrusion by magmatic, sub-magmatic and non-magmatic mechanisms (e.g. Guine-berteau et al. 1987, Pons et al. 1995). Magmatic shear zones can be identified by fabrics defined by unstrained ig-neous minerals (e.g. Miller and Paterson 1994) or reorient-ation of a magmatic fabric in the shear zones (e.g. Pons et al. 1995). Non-magmatic shear zones in the latter study had pressure shadows filled by epidote around megacrysts, and ribbon grains formed by intracrystalline plasticity in quartz.

Magmatic shear zones were identified in the experiments of Park and Means (1996) by analyzing movements of solid in-clusions, but there was very little direct microstructural evid-ence to distinguish the shear zones from the unsheared walls.

Criteria which are sometimes misused to demonstrate mag-matic flow include imbrication, which may occur in mylon-ites (Section 7.11.3), magnetic anisotropy, which is well known in metamorphic rocks, and S-C fabrics, which may form in either sub-magmatic or non-magmatic deformation (Blumenfeld and Bouchez 1988).

6.4 Magmatic microstructures

6.4.1 Grain shape fabrics

The typical microstructure of magmatic flow consists of a preferred orientation of euhedral phenocrysts, in an isotropic matrix that shows igneous textures (Plate 36). Grain shape fabrics are commonly defined by feldspars and micas in felsic rocks, and may be defined by feldspar, olivine and pyroxene in mafic rocks (e.g. Benn and Allard 1989). This is diagnostic of magmatic flow if there is no sign of any other deforma-tion mechanism. Since intracrystalline plastic deformadeforma-tion microstructures are readily visible in quartz after only small strains, the lack of deformation in quartz (no undulose ex-tinction, subgrains, or kink bands) is a reliable criterion for magmatic flow. Lack of sub-solidus deformation can also be checked from inclusions (e.g. rutile) in quartz: the inclusions should be undeformed themselves and randomly orientated (e.g. Mitra 1976, Stel 1991). Other primary, undeformed igneous features such as strongly euhedral crystals, igneous (e.g. idiomorphic) zoning, growth twins, and ophitic texture can be used to demonstrate the absence of intracrystalline de-formation. A fabric in igneous rocks can only be confirmed as a magmatic microstructure by lack of any evidence for other deformation microstructures or mechanisms, including any of the microstructures described in previous Chapters.

It may be difficult to distinguish magmatic deformation from non-magmatic deformation followed by static recrystal-lization. Post-deformational annealing can be revealed by

de-formed inclusions in annealed minerals, and by a granoblastic texture in mineral aggregates (especially quartz) which them-selves define a shape fabric. Primary igneous fabrics can also be recognized by random spatial relationships between dif-ferent phases. By contrast, solid state deformation produces higher frequencies of contacts between like grains because new grains preferentially nucleate on the same phase (e.g.

Ashworth and McLellan 1985).

6.4.2 Crystallographic fabrics

Magmatic crystallographic fabrics are due to orientation of euhedral, inequant phenocrysts during flow. In mafic rocks, (010) crystal faces are parallel to the magmatic foliation in olivines, pyroxenes and feldspars (Benn and Allard 1989).

The [001] direction of olivine and clinopyroxene is parallel to the magmatic lineation, and [100] has been observed parallel to the magmatic lineation in feldspars within gabbros and ton-alites (Benn and Allard 1989). Such crystallographic fabrics can be distinguished from fabrics due to intracrystalline plas-ticity by complete lack of evidence for intracrystalline plastic microstructures (Chapter 4), and in the case of olivine, by the fact that high-temperature non-magmatic deformation leads to [100] parallel to the lineation rather than [001], as observed for magmatic fabrics (Benn and Allard 1989).

6.5 Sub-magmatic microstructures

6.5.1 Grain shape fabrics

Grain shape fabrics may be expected to form in sub-magmatic flow by both rigid body rotation and crystal de-formation by DMT or intracrystalline plasticity. The crystal deformation combined with evidence of melt such as matrix with igneous textures is diagnostic of sub-magmatic flow (e.g.

Quick et al. 1992). Localization of melt in structures that are coeval with the foliation constitutes good evidence for sub-magmatic deformation: this sort of evidence is more readily seen at outcrop scale. As for magmatic flow, static recrystal-lization may give the matrix the false appearance of a primary igneous texture.

6.5.2 Intracrystalline plasticity

Intracrystalline plasticity in quartz may be expected in the presence of melt under appropriate differential stresses (>1 MPa) and temperatures (700 to 800°C) on the basis of granite and quartz flow laws (Rutter and Neumann 1995). The very common appearance of undulose extinction in quartz within granites suggests that intracrystalline plasticity occurs during sub-magmatic deformation (Plate 37). Subgrain formation or recrystallization of quartz has been taken as an indicator of sub-magmatic deformation where an overall igneous texture is preserved and no other non-magmatic deformation event is known (e.g. Bouchez and Gleizes 1995). Deformation twins and bent twins in feldspar may also hint at sub-magmatic intracrystalline plasticity (Plate 38). However, these micro-structures on their own do not demonstrate that melt was present during deformation.

6.5.3 Diffusive mass transfer

Despite the experimental evidence for the importance of melt-aided diffusion creep, microstructural evidence has remained elusive, probably because melt films have almost no poten-tial for preservation in the geological environment. However, experiments suggest some features that could be used: trun-cated, embayed, scalloped or overgrown grain boundaries in igneous rocks may be analogous to some of the features dis-cussed in Chapter 3 that indicate DMT through a fluid phase (e.g. Dell’Angelo et al. 1987, Park and Means 1996). The potential importance of surface energy anisotropy in melt dis-tribution needs to be investigated and may have some micro-scopic expression in the form of differences between crystal faces.

6.5.4 Cataclasis

Paradoxically, some of the clearest sub-magmatic microstruc-tures involve cataclasis of the crystals, such as microfracmicrostruc-tures which are healed by melt (e.g. Hibbard 1987, Bouchez et al 1992, Karlstrom et al. 1993). Microfractures in plagioclase can be demonstrated to have been filled by melt from the fol-lowing criteria:

1.

2.

3.

4.

The microfractures are intragranular. This allows that the plagioclase crystals were in contact with melt.

The microfracture filling is compositionally and crystal-lographically continuous with the same phase in the ig-neous groundmass of the rock.

The composition of the microfracture filling is compat-ible with the later stages of the igneous petrographic his-tory of the rock. Plagioclase microfracture fillings may have lower anorthite contents than their host crystals, consistent with progressive evolution towards a min-imum melt composition (Bouchez et al. 1992). The rela-tion between quartz and plagioclase in the microfracture fillings also suggests a residual melt: feldspars are on the walls or tips of the microfractures.

Early crystals (e.g. biotite, sphene) are trapped within the microfracture fillings.

Cataclastic microstructures observed in experiments con-firm the potential importance of cataclasis in sub-magmatic deformation in the experiments of Rutter and Neumann (1995). Up to 10% melt, axially-orientated cracks formed and filled with melt, and the sample was faulted. Between 10 and 45% melt, cataclastic flow occurred with pore col-lapse. Axially orientated microcracks formed by high melt pressures have been observed in granitic aggregates contain-ing 2-15% melt (Dell’Angelo and Tullis 1988). Connolly et al. (1997) have demonstrated that microcracking caused by volume increase during melting is a viable way to create permeable fracture and melt-pool networks in a muscovite-bearing quartzite. The syn-kinematic experiments by Park and Means (1996) also recorded fracture, localized along a kink band boundary.

On a larger scale, several cataclastic features are commonly associated with melt in migmatites. Metatextites often consist

of a competent body sub-divided by fractures that are filled by melt. Melt may form in pressure shadows at the ends of boud-ins, and fill faults (e.g. Quick et al. 1992). Quartz-feldspar neosomes have been described accumulated under imper-meable refractory layers such as amphibolite sheets which are boudinaged, allowing the neosomes to rise into boudin necks, and forming a geopetal structure (Burg 1991). Segregations in shear zones and along axial planes of folds in migmatites are common. Localization of the melt in these structures in-dicates that the melt was syntectonic. This sort of evidence has great relevance to the problem of extracting melt to form plutonic bodies (e.g. Wickham 1987).

6.6 Other microstructures

Park and Means (1996) introduced the term “contact melting”

to describe melting at contact points between grains observed in their experiments, and suggested that indented boundaries

to describe melting at contact points between grains observed in their experiments, and suggested that indented boundaries