Textures in Igneous Rocks

The primary factor controlling grain size of rocks is undercooling. Undercoolingtakes place before nucleican growinto crystals under the difference betweenliquidus temperatureand the actual temperature. The general situation with respectto undercooling is mainly controlled by theratio of the nucleation rate (N) to the growth rate (G). At initial stage, both nucleation and growth rise with increasing degrees of undercooling, butgrowth rate increases at a lower rate. After reaching a maximum, both nucleation and growth rate decline withfurther undercooling, because lower temperatures reduce the rates of diffusionof chemical components in the increasingly viscous melt.

At low degrees of undercooling, when N/G ratio is low, large well-formed (polyhedral, euhedral,idiomorphic) crystals develop. At larger degree of undercooling, both Nand G rise, but G tends to decline before N, and a situation of high N/G isreached; at this stage, many small euhedral crystals develop in the ground masses of many volcanic rocks(Fig. 1.2c, d). These are called crystallites (Fig. 1.2c, d),if verysmall, microlites if larger(Fig. 1.2d). At even higher degree of undercooling, both N and Gbecome very small, so that no new nuclei form; consequently, growth is forcedto occur only on existing crystals. At extreme degreesof undercooling, crystallization fails to occur, and volcanic glass is formed.In volcanic rocks, once it is quenched, a range of grain-sizes is produced in the rock as nucleationoccurs during crystallization.

Hot magmas, such as gabbros and ultramafic cumulates, have long, subsoliduscooling periods, even after they have completely solidified. During this phase, changes in the shapes and sizes of grains may occur (Vernon, 2004 and references therein). The process is driven by a tendency to reduce thetotal interfacial energy of the aggregate, so that lower-energy grain-shapes and larger grainsizes are produced by diffusion of chemical componentsfrom high-energy to lower-energy sites. The residual melt (ifpresent) is commonly distributed along solid grain boundaries. Standard grain size ranges are described below:

3 cm

5 mm to 3 cm

1-5 mm

<1 mm

A microcrystalline rock contains small crystals visible only through microscopic examination. The suggested range of crystal sizes that should be regarded as microcrystalline, varies from 1 to 200 microns(Fig. 1.2d, e). Texture of a volcanic rock can be described as microcrystalline.

The term cryptocrystalline defines the texture of a rock consisting of crystals that are too small to be recognized and separately distinguished megascopically oreven under the optical microscope (although crystallinity may be shown by use of the electron microscope); indistinctly crystalline, as evidenced by a confused aggregate effect under polarized light(Fig. 1.1b, d). The upper limit of cryptocrystalline rock is considered to be 0.004 mm.

The most primary factor controlling the shapes (‘habits’) of crystals growingin liquids is the degree of supersaturation, which is the difference between theactual concentration of a chemical component in the liquid and the concentrationat equilibrium (i.e. when the magma is just saturated with the component) at aspecified temperature and pressure (Vernon, 2004). Supersaturation is influenced by the effects of undercooling, changes in pressureand in the concentration of chemical component in the melt. Grain shape is controlled by the diffusion and growth ratio (Vernon, 2004). Euhedral crystals grow at small to moderate degrees of supersaturation at smaller growth/diffusion ratio. In contrast,skeletal and dendritic crystals andspherulitesgrow at greater degrees of supersaturation (Donaldson, 1974) at larger growth/diffusion ratio (diffusion-controlled growth).

Crystals or mineral grains that are well-formed, with sharp, well defined faces are designated as euhedral crystals(Fig. 1.2.a, d). Here the mineral grain is completely bounded by its own coherent crystal faces whose shape is guided by the mineral’s crystal forms and symmetry. Euhedral growth of crystals implies that neithertheir growth during primary crystallization was restrained by or interfered with adjacent grains nor are the facesgreatly affected by post-crystallization physical or chemical processes.

Mineral grain that is bounded partly by its own rational crystal facesand partly by surfaces formed against pre-existing grains as a result of either primarycrystallization or recrystallization. Here, some of the crystal faces relate to the original crystalform and symmetry of the mineralbut some may have been obscured by later physical orchemical erosion(Fig. 1.2b).

Here the mineral grains have irregularprimary faces. This is more common in the interstitial grains that take the shape of the space available between the adjacent grains that might have crystallized earlier.The original shapes of primary faces may also have beenobscured by post crystallization processes, i.e. deformation, metamorphism/alteration, producing anhedral grains.

Crystallites:Incipient or embryonic crystals in volcanic rocks that often lack any recognizable crystallographic form. These are often too small to exhibit optical properties characteristic of the mineral under polarising microscope. They occur when magma congeals so rapidly that crystallization remains incomplete(Fig. 1.1c, 1.2d, e).
Microlites: These are slightly larger forms recognizable as mineral species under polarising microscope(Fig. 1.2c, d, e).

These are crystals which grew as or have been corroded to askeletal framework (Fig. 1.2.5a-c)with a high proportion of internal cavities (Fig. 1.2.5b)that are commonly crystallographically controlled. These cavities are filled with glass or crystalline groundmass material(Fig. 1.2.5d).In volcanic rocks, the high cooling rate often impede growth of good crystal faces or the high temperature of the magma corrodes the pre-existing crystals generating skeletal grains(Fig. 1.2.5e).

A crystal that develops with a typical multi-branching tree-like form. Dendritic crystals typically grow relatively rapidly in response to high growth ratebut can also form in quenched lavas from growth instabilities that occur when the rate of diffusion is slower than the rate of growth. Here, depleted liquid builds up at crystal-liquid interface and crystals reach out in tendrils beyond depletedzone for supply of appropriate elements required for further growth.

Dendritic crystals(Fig. 1.2.5e)are commonly observed in chilled lavas (Bryan, 1972), meteorites(Fig. 1.2.6e), or in quenched experimental melts (Lofgren, 1980). However, dendritic habits may also occur in intrusive, coarse grained, mafic–ultramafic rocks (Donaldson, 1974) granites (Moore, & Lockwood, 1973; Vernon, 1985), or granite pegmatites (London, 1992). Dendritic crystals in slowly cooled intrusive rocks are explained as a possible result of lower diffusion or delayed nucleation (Vernon, 2004). Skeletal and dendritic crystals are forced to develop by compositional supersaturation. In a quenching silicate melt, if some components that are not required by the growing crystal cannot diffuse away into the liquid fast enough, they become concentrated in a narrow zone adjacent to the interface. Here, projections or spike-shaped crystal-melt interfaces(Fig. 1.2.6e) are favoured than planar interfaces to attain higher interface area and better access to nutrients from the melt. Secondary spikes may be forced to develop on the primary spikes for crystallization to continueleading to a dendritic habit. Skeletal, dendritic and spiky crystals of plagioclase, olivine, pyroxene and chromite/magnetite occur in quenched volcanic rocks such as felsic, intermediate and mafic-ultramafic volcanics, i.e. and komatiites. Skeletal and dendritic crystals of olivine and/or pyroxene (Fig. 1.2.6e)also occur in quenched artificial melts in experiments and slags (Faure et al., 2003), in glassy chondrules in chondritic meteorites, and in shock-melted rocks in meteorite impactsand fault zones (Camacho et al., 1995).

Dendritic crystals commonly grow in felsic lavas such as dendritic overgrowths on crystalsof plagioclase, hornblende and biotite in rhyolite, and microlites of pyroxene, amphibole, biotite, sanidine and magnetite in unalteredfelsic glasses within intermediate volcanic rocks, i.e. andesite.

Swallowtail grains are the results of increasing undercooling (e.g. Lofgren 1974; Faure et al. 2003). Swallow-tail form may commonly occur in plagioclase(Fig. 1.2.6a), olivine, pyroxene and magnetite in quenched volcanic rocks (see Vernon, 2004 and references therein). In rapidly growing crystals, once the growth of the crystal is interface controlled, the crystals grow by forming ‘spikes’ or ‘projections’ (Fig. 1.2.6d-e)to gain better access to the nutrient components from the melt instead of the impurity components (components that are not required in forming the crystal; Vernon, 2004). Often such projections may originate as protuberances on a formerly planar interface or at corners of former polyhedral crystals, which grow preferentially because of lower impurity concentrations at those points. Growth at the corners is facilitated by lower impurity concentrations due to larger ratio of melt volume to crystal surface area at the corners producing ‘swallow-tail’crystals (Vernon, 2004).

A volcanic texture consisting of plate-like crystals of olivine embedded in a fine-grained matrix of dendriticor acicular clinopyroxene, dendritic chromite and altered glass. Plates of olivine crystals are generally millimetres to decimetres long(Fig. 1.2.7a), but can reach lengths exceeding one metre. Length of the spinifex blade or plates is inversely related to the cooling rate of komatiitic magma and thus may vary significantly within a lava flow.For example, many komatiite flows have an upper spinifex-textured layer and a lower olivine-cumulate layer; and other flows grade along strike from layered spinifex-textured portions to massive olivine-phyric units. With the inclusion of the phase about spinifex, the lower olivine-cumulate portions of layered flows or the olivine-phyric units can also be described as komatiite. According to the definition given by Arndt (1994),“spinifex is a texture characterized bylarge, skeletal or dendritic, platy, bladed or acicular grains of olivine or pyroxene, found inthe upper parts of komatiitic flows, or, less commonly, at the margins of sills and dikes”. Thetexture is believed to form during relatively rapid, in situ crystallization of ultramafic or highly mafic liquids. Within the upper parts of komatiite flows, the type of spinifex texture varies systematically. Beneath a thin (1-5 cm) glassy, commonlyporphyritic chill zone, a layer of “random” spinifex texture contains isolated randomlyoriented crystals or centimeter-scale “booklets” of parallel plates of olivine, in a matrix of finegrained clinopyroxene and devitrified or altered glass. Below this, the layer of “platy”olivine spinifex has an organized structure wherein arrays of large bladed olivine crystals(Fig. 1.2.7b), and intermittent pyroxene blades (Fig. 1.2.7c)starting from few decimetre to metre long, are oriented roughly perpendicular to the flow top. In some flows thecrystals form parts of sheaf-like structures that open away from flow tops and serve asreliable facing indicators. The lower parts of spinifex komatiite flows are cumulatescontaining settled solid polyhedral olivine crystals (Arndt et al., 2008). During metamorphism, olivine, pyroxene and chromite are commonly converted into serpentine, chlorite/tremolite and magnetite respectively while glass is commonly devitrified into a fine-grained aggregate of low-temperature secondary minerals(Fig. 1.2.7a).

A texture in which all the component mineral grains are anhedral.

Hypidiomorphic refers to a texture, in which most of the grains are subhedralwhilethose of other mineral species are anhedral(fig 1.2.9).

Panidiomorphic refers to a texture in which, theoretically all the component mineral grains are euhedral(Fig. 1.2a).

Spherulites are radiating arrays of fibrous, needle-like or acicular, crystals that are common in glassy felsic volcanic rocks. Spherulites are typically two-mineral aggregates (mainly quartz and feldspar; Fig. 1.2.11a), formed by initial spherulitic growth of one mineral and later crystallization of a second mineral from the liquid or glass between the fibres. Each fibre has the same crystallographic axis parallel to its length, and each has an orientation slightly different from that of its neighbours. Thus, in contrast to dendrites, spherulites are aggregate of separate crystals, rather than branched single crystals. A dark extinction cross (Fig. 1.2.11b)is common in spherulites observed in crossed polarised light, because many fibres (each with one of the principal optical vibration direction parallel to its length) are parallel or approximately parallel to the vibration direction of the polarizer and analyser of the microscope.

Spherulitic aggregates typically nucleate on existing crystalline material although they are often too small to be seen in thin section. Once a radial growth habit is established, growth continues uniformly in all directions (Harker, 1909) as the crystals grow in a homogeneous material, such as a liquid or glass. Homogeneous growth of fibres from a single point-nucleus (a small crystal or crystal fragment) produces a spherical aggregate or spherulite(Fig. 1.2.11c). Incomplete radiation results in fan, bow-tie (sheaf-like), and plumoseaggregates (Lofgren, 1971a; Fig. 1.2.11d). Axioliticspherulites result from fibres radiating or projecting out from a line or plane, probably owing to water penetratingalong a crack and promoting crystallization of the adjacent glass or viscous melt. Axiolitic aggregates also form by crystallization of glass shards or over phenocrysts in tuffs (Fig. 1.2.11e-h). Spherulites form under conditions of very strong supersaturation. Experimental work on the formation of spherulites in melts (Lofgren 1971, 1974) has confirmed that spherulitic growth is favoured by very low diffusion and growth ratio, even lower than dendritic growth. Low nucleation rates, which are typical of high degrees of supersaturation are essential for both dendritic and spherulitic growth, forcing growth to occur on existing crystals (Fig. 1.2.11d, e, f, g).

Some spherulites consist of micrographic intergrowthsof quartzand alkali feldspar(Fig. 1.2.11i), especially at their margins suggesting simultaneous growth of the two minerals. Here, each poikilitic quartz grain may enclose many feldspar fibres forming ‘snowflake’ structure (Anderson, 1969; Fig. 1.2.11j). Such structure is formed where the nucleation rate forquartz is lower than that for feldspar.

A texture consisting of fine, radiating fibres of plagioclase or pyroxene microlites embedded in a microcrystalline or devitrified glassy matrix. The variolitic texture denotes rapid cooling and the quench growth of the crystals, such as is often found in submarine basalts and chilled margin of shallow-level, basic, igneous intrusions (dykes and sills).

A texture where mineral grains are approximately of the same size is called Equigranular Texture.It suggests that the entire rock probably crystallized under same conditions of coolinge.g. pressure, temperature or rate of cooling).

In igneous rocks, equigranular textures are commonly shown by granites and aplites. In the granitic texture, the constituents are either all coarse or medium grained and the crystals are euhedral to subhedral in nature. In aplites, the rock is equigranular but the grains are mostly microscopic crystals showing perfect outlines.

A texture characteristic of plutonic igneous rocks that consists of crystals that accumulated due to gravitational settling in a magma.

The word cumulate was applied originally to magmatic basic or ultrabasic rocksin which precipitated crystals were inferred to have grown to a relatively largesize before physical accumulation by either floating or sinking. In cumulate texture, crystals accumulate by crystal settling at the bottom or near the margins of the magma chamber. While the cumulate phase accumulates inmutual contact, the intercumulus liquid (melt) fills the interstices and crystallizes as intercumulus phase. If these crystalswere inferred to continue to grow by precipitation from the interstitial liquid(which was inferred to require diffusive exchange of components with themain magma) and so fill space, the resulting rock was called an adcumulate. If the accumulated (cumulus)crystals were inferred to have stopped growing, and new minerals crystallizedfrom the interstitial liquid, the resulting rock was called an orthocumulate.In orthcumulates, intercumulus liquidcrystallizes to form rim or additional growth of the same mineral that initially formed as cumulate phase while there is no exchange between the intercumulus liquid and the main magma chamber. In adcumulates, there is an open-system exchange of components between theintercumulus liquid and the main magma chamber that allows additional components required to crystallize intercumulus phase to escape so that the cumulate phase can further grow to fill most of theavailable space. In mesocumulate texture,intercumulus liquid crystallizes to additional rims on cumulous phases along with other minerals that nucleate poorly and poikiliticallyenveloping the cumulate mineral phase.However,the initial development of single-mineral aggregates inadcumulates in layered mafic–ultramafic igneous complexes isessentially a primary magmatic process, such as size-density sorting (see, forexample, Irvine, 1974) or rhythmic supersaturation (Lofgren & Donaldson, 1975) or flow banding (Fig. 1.3.2e).

Cumulate texture with other minerals occupying the areasinterstitial to cumulate phases(Fig. 1.3.2a, b, c).

Intermediate between ortho- and adcumulate(Fig. 1.3.2f).

Cumulate texture in which the early cumulate minerals grow to fill the interstitial space between cumulate grains.

Porphyritic igneous rocks consist of larger crystals (phenocrysts) in a distinctlyfiner-grained aggregate (groundmass). Thisis regardless of the absolute grainsize, although volcanic and shallow intrusive(Fig. 1.4.1a)rocks are much more commonly porphyritic than deeper, coarser-grainedintrusive rocks. Porphyritic texturein volcanic (Fig. 1.4.1b) and high-level intrusive rocks that has been interpreted as resulting from slow crystallization at depth to form thephenocrysts, followed by rapid cooling (due to extrusion onto Earth’s surfaceor intrusion into colder rocks) to form the groundmass. However, porphyritictexturecan also form in a single, uninterrupted cooling event where the first precipitating (‘liquidus’) mineral must crystallize alone fora certain time(Fig. 1.4.1c).

Porphyritic texture could also be produced by rapid loss of water dissolved in the melt component of a magma crystallizing slowly at depth; this would raisethe equilibrium freezing temperature (‘liquidus’) of the melt(Fig. 1.4.1d). This is because water-bearing melts typically have lower freezingtemperatures than equivalent dry melts.Porphyritic texturein deep intrusive rocks (for example, megacrysts of K-feldsparin granites; Fig. 1.4.1d-e) is very unlikely to be due to two-stage cooling, because theserocks crystallize slowly and at relatively uniform rates in large magma bodies(plutons). It is probably due to different nucleation rates for different minerals. Acommon example of porphyritic texturein intrusive rocks is the occurrence ofvery large euhedral crystals (megacrysts), up to about 20 cm long, of K-feldsparin some granites (Vernon, 1986a). It is explained as these megacrysts developed at conditions ofunusually low nucleation rate compared to growth, presumably at low degrees of undercooling. Such low nucleationrate is connected with local abundance of water, which would tendto break Si–O and Al–O bonds in potential nuclei. The nucleation difficulty is not commonly connected with the major-element chemical composition of the magma (Vernon, 2004).

Glomeroporphyritic texture refers to a porpyritic texture in which phenocrysts are clustered into aggregates or crystal clots called glomerocrysts(Fig. 1.4.2a-b). Glomeroporphyritic textures are common and often included plagioclase and pyroxenes in basic rocks (Fig. 1.4.2b-c). Glomerocrysts are an important consideration in crystal fractionation by crystal settling since the density of the glomerocryst is an average of that of its constituent phases. Formation of glomerocrysts may in part explain the settling of plagioclase in basic intrusions in which plagioclase crystals are less dense than the surrounding magma. The glomerocrystsplay an important role in the fractionation process, being heavier than the individual crystals of which they consist, they are more prone to sinking in the magma.Hogan (1993) describedglomeroporphyritic aggregates of plagioclase, withinwhich, the grains are random and meet along irregular boundaries as a product of dissolution of crystals owing to changes in temperature, pressure or water activity in the melt followed by joining of the grains by supersaturation and crystallization of the trapped melt at grain contact by dissolving crystals. The dissolution explains embayment of the grains in the aggregate.

The term Synneusis (from the Greek syn: ‘together’ and neusis: ‘swimming’) refers to a texture where oriented attachment and adherence of individual mineralsyield a heterogeneous crystalline association characterized by parallel growth of minerals(Fig. 1.4.2e-i). Vance, (1969) described synneusis as the process of drifting together and mutual attachment of crystals suspended in a meltin specific low-energy conditions and recognized it as a type of ‘glomeroporphyritic aggregate’(Fig. 1.4.2j).The reason for the adhesion in synneusisis interpreted as the low free energy ofspecific faces on which the minerals collide and attach, for example (010) twininterfaces in plagioclase (Vernon, 2004). Also, the magma must be fluidenough to enable movement of the crystals.Vance (1969) suggested that synneusis occurs in the earlier stages of consolidationand is related to turbulence. The centre of concentric zoning in each individual crystal of plagioclase attached to synneusisis offset from themutual boundary. Another example of synneusis is the attachment of small plagioclase(or biotite) crystals to the faces of K-feldspar megacrysts growingin a granitic magma, producing a concentrically zoned (crystallographically controlled) pattern of inclusions in the megacryst.Synneusis of olivine in an olivine-rich volcanic rock from Hawaii has beeninferred by Schwindinger& Anderson (1989) (Fig. 1.4.2h). Synneusis has also been inferred for chromite crystals in some ultramafic rocks (Vogt, 1921; Bastin, 1950).

Poikilitic texture refers to crystals, typically phenocrysts, in an igneous rock which contain numerous small grains of other minerals.Poikilitic texture helps to determine order of crystallization suggesting that the crystallization of the included minerals were followed by the enclosing, host mineral(Fig. 1.4.3a). Also, this texture may be originated as a result of differing nucleation and growth rates, so that a single crystal (like pyroxene) nucleate and grow to a large size (low nucleation rate) in contrast to several other minerals (like feldspars) with higher nucleation rate which necessarily remain relatively small and became successively entrapped in the pyroxene(Fig. 1.4.3b-c). In many cases of poikilitic texture, the enclosed crystals are randomly arranged, in others,they may be concentrated in zone.

The smaller enclosed crystals in poikilitic texture are known as chadacrysts.

The larger crystals are known as oikocrysts.

The most common poikilitic texture involves plagioclase laths enclosed by augite and is known as an ophitic texture (elongate crystals enclosed by another mineral). It is often found in dolerites (Fig. 1.4.4a)and gabbros.

This term is more commonly used once plagioclase laths partially enclosed by augite and is known as an ophitic texture(Fig. 1.4.5 a-e).

Ophiticmicrostructure is the result of more abundant nucleation of plagioclase and lessabundant nucleation of pyroxene (Wager, 1961). Occurrences of either the plagioclaseor the pyroxene as phenocrysts, as well as in intergrowths with the othermineral, implies that the phenocryst mineral grew before the intergrowths, whichcould be due to (1) its presence in excess of eutectic proportions, or (2) delayednucleation of the other mineral (Lofgren, 1980).

This is an igneous texture, in which the wedge-shaped spaces between a network of lath-shaped crystals, such as plagioclase, are filled with smaller crystals of other minerals(Fig. 1.4.2a). Such texture is especially well developed in basalts where the angular interstices between plagioclase grains are occupied by grains of ferromagnesium minerals such as olivine, pyroxene, or iron-titanium oxides(Fig. 1.4.2b-c).

A texture similar to intergranular texture except that the interstices between the network of lath-shaped grains, such as plagioclase, are occupied by glass or cryptocrystalline material(Fig. 1.4.7a-b).The glass may be devitrified or altered to other phases.

Trachytic is a texture of volcanic rocks in which minute tabular crystals (microlites) of K-feldspars (sanidine) occur within groundmass of volcanic glass. The microlites are oriented parallel to each other forming flow lines along the directions of lava flow and around inclusions. Feldspar microlites (typically plagioclase) in groundmass are sometimes arranged subparallelly around larger phenocrysts, reflecting flow in relatively rapidly cooled magma and definepilotaxitic texture (Fig. 1.4.8; 1.4.11d). These textures are typically seen in basalt and andesite.

Preferred orientation of elongate K-feldspar phenocrysts is a common feature in felsic or syenitic plutons with porphyritic texture. Many of them also exhibit other primary magmatic structures such as grain-sizegradation, subparallel, planar, or weakly concave-up layering. Magmatic flow promotes the crystals to rotate passively into alignment without them deforming internally once sufficient melt is available for the crystals (Paterson et al., 1989; Vernon, 2000a). This causes parallel or subparallel alignment of elongate euhedral crystals commonly feldspar, hornblende or olivine, that are not internally (plastically) deformed, implying rotation of the crystals in a relatively weak medium, such as melt. Layering in igneous rocks, especially from felsic plutonshave been reported from around the world associated with wide range of tectonic settings (Gilbert, 1906; Barbeyet al., 2008).Gravitational settling, the sinking of denseminerals in the less dense felsic melt that contains them is the key process thatmight have been responsible for layering and preferred alignment of elongated minerals in these rocks(Clarke & Clarke, 1998). The other possible mechanisms include density-driven separation of minerals due to convection within magma chamber (Irvine, 1987) or hydrodynamic sorting (i.e. by size and shapeof particles) in density currents (Barbey et al., 2008).Solgadi and Sawyer (2008) marked a strong similarity in the orientation of minerals, such as feldspar phenocrysts and their size gradation with primary sedimentary structures derived from hyper-concentrated and concentrated sediment gravity flows.The K-feldspar crystals, having lesser density than the melt are held in the melt in suspension and are oriented parallel to the direction of flow.The feldspar crystals havea strong preferred orientation consistent with movementby a traction current.

Exsolutionisa process through which an initially homogeneous solid solution mineral separates into at least two different crystalline minerals without the addition or removal of any materials. In most cases, it occurs upon cooling below the temperature of mutual solubility or stability of the solution. The separated phases occur as fine lamellae (mostly microscopic but may be also visible under naked eye) whose orientations are crystallographically controlled within the original grain(Fig. 1.5.1a, b). Exsolution takes place by the solid state diffusion of ionsthrough the structure, it is very slow and so the two phases generallyexist on a microscopic scale(Fig. 1.5.1c).

Schiller effects are the result of thin microscopic inclusions within a translucent mineral, usually as exsolution lamellae, which refract and reflect incident light. Labradorite, a Na-rich anorthite (a plagioclase feldspar), sometimes exhibits a very pronounced iridescent schiller (Fig. 1.5.1d), also termed labradorescence, in particular where it contains appreciable K. This is due to the presence of a miscibility gap, causing exsolution of orthoclase from the plagioclase host during cooling. K-feldspars (probably between microcline and orthoclase in structure) with well-developed schiller due to microperthite (exsolution of albite), produce semi-precious gemstones usually known as moonstones. Adularia, a low temperature K-feldspar of variable structure, also can exhibit this milky schiller, also termed adularescence. Sunstone or aventurine feldspar is a variety of feldspar (microcline or oligoclase) that has schiller with an orange/brown background color, and containing small hematite crystals that give it an additional sparkle.

Perthite defines an intergrowth texture of two feldspars due to exsolution where the host grain is potassium-rich alkali feldspar (near K-feldspar, KAlSi3O8, in composition) that includes exsolved lamellae or irregular intergrowths of albite-rich (NaAlSi3O8) plagioclasefeldspar(Fig. 1.5.2 a-d).Typically, the host grain is orthoclase or microcline, and the lamellae are albite. Perthite forms by exsolution due to cooling of a grain of alkali feldspar with a composition intermediate between K-feldspar and albite. Albite and K-feldspar maintain complete solid solution at ~ 700 °C in surficial pressures while a miscibility gap is present at lower temperatures. Therefore, an alkali feldspar grain with an intermediate composition cools slowly enough, K-rich and more Na-rich feldspar domains separate from one another in form of exsolution lamellae.

An antiperthite is an intergrowth arising due to exsolution where potassic feldspar is present as blebs or lamellae within a sodic (albite-rich) feldspar(Fig. 1.5.3). The term mesoperthite is used when sodic and potassic feldspars are in broadly equal abundance. Perthite that can only be observed with the aid of a microscope is known as microperthite.

Minerals belonging to a solid-solution series (e.g. plagioclase) continuously reacts with the surrounding liquid as it grows. The attainment of equilibriumbetween a complete growing crystal and the melt during cooling should producea compositionally uniform crystal. However, diffusion and ion exchange in manyminerals is commonly not sufficiently fast (or it rather falls with falling temperature as crystallization progresses) for the crystal to adjust its compositionto changing conditions (e.g. falling temperature), so that only the rims canequilibrate with the liquid leading to compositional zoning. The most common form of compositional zoning is concentric zoning, in which the zones are parallel to the advancing crystal faces. The other forms of zoning, are patchy zoning and sector zoning. Concentric zoning patterns reflectthe growth history of the crystal. For example, in plagioclasein igneous rocks, the compositional variationis guided by the substitution of Ca+Al↔Na+Si substitution involving breaking of strong O–Al and O–Si bonds. This explains why chemicalequilibrium is rarely achieved through the whole crystal (Vernon, 2004). Concentriczoning in plagioclase may be marked by zones rich in inclusions representing marked changes in growth rate, or resorption due to partial melting or chemical reaction.

The normal zoning pattern is characterized by a high temperature composition of the crystal gradually followed by the low temperature composition towards the rim of a growing crystal within magma. Plagioclase cores commonly reflect crystallization at relatively high temperatures,so that they are more anorthitic, viz. relatively rich in Ca and Al, and successive outer zones are progressivelyalbitic, viz. richer in Na and Si, in response to falling temperature. The resultant zoning pattern is called normal zoning (Lofgren, 1974). This feature is very common in slowly cooled felsic plutonic rocks, i.e. granite.

Reverse zoning, although uncommon, occurs in igneous rocks which exhibit opposite pattern to the normal zoning, viz. a lower temperature composition of the core gradually followed by a higher temperature composition towards the rim. This commonly happens during magma mixing when a relatively hotter magma with more primitive composition mixes with the magma that hosts the growing crystal. For example, reverse zoning in plagioclase(Fig. 1.4.11a) may occur by sudden increase of magma temperature by introduction of fresh batch of magma, rapidmovement of magma in a chamber or conduit,or magma mixing.

This texture is commonly observed in granites where coarse crystals of K-feldspar is rimmed (mantled) by smaller crystals of plagioclase (typically oligoclase). The existingmodels for their genesis include both magmatic or subsolvusprocesses. The magmatic models include changes infeldspar stabilities during crystallization through variationsin pressure, temperature, activity of H2O, or magma composition.Many of the magmatic models involve mixingof silicic and basic magmas as a general mechanism forthe formation of rapakivi feldspars. Others invoke sub-isothermaldecompression of low-volatile crystal-saturatedsilicic magmas. In the magma mixing models, themingling of an alkali feldspar-saturated felsicmagmawith a more mafic magma changes the magma compositionand increases its temperature, leading to resorptionof alkali feldspar megacrysts and precipitation of plagioclase around them(Fig. 1.6.3a-c). However, recent study also suggested that rapakivi feldspars canform by subsolidus fluid-induced dissolution of feldsparmegacrysts and pseudomorphic replacement by oligoclaseand albite (Mondal et al., 2017). Similarlysecondary overgrowth of different mineral (i.e. monazite) may occur over pre-existing crystal (i.e. apatite; Fig. 1.10c).

Oscillatory zoning involves repeated small composition changes that occur duringgrowth of a crystal, and is particularly common in plagioclase, though it also occurs in other minerals. Oscillatory zoning has beenproduced experimentally at constant temperature, indicating that it is caused bylocal compositional variations in the melt immediately surrounding the growingcrystal, rather than by temperature variations (cf. Vernon, 2004).Oscillatory zoning in plagioclase, involving alternating calcic and sodic zones(Fig. 1.6.4 a-c) with small compositional differences, has been explained by diffusion-controlled, recurrent supersaturation of the melt in anorthite and then albite components adjacent to the growing crystal (i.e. Harloff, 1927; Vance, 1962). Oscillatory zoning may also occur in other minerals in igneous rocks, such as clinopyroxene, alkali feldspar, olivine, hornblende and tourmaline. For example, in zoned titanium-rich clinopyroxene, the composition oscillates between augite and subcalcic augite.

Patchy zoning is common in many igneous rocks (especially calc-alkaline, plagioclase-rich volcanic andplutonic rocks), especially in plagioclase containing irregular corroded calcic cores, surrounded in crystallographic continuity by more sodic plagioclase. Such microstructure has been interpreted (Vance, 1965) as being due to initial crystallization of relatively calcicplagioclase in a water-undersaturated magma at depth, followed by a decrease inconfining pressure, causing resorption, owing to the fact that the melting pointdecreases with falling pressure in most water-deficient systems. The resorptionappears to be followed by new crystallization of more sodic plagioclase that is stable under the lower pressure conditions (Vance, 1965), as rims on the coresand as fillings of cavities in the cores, forming pseudo-inclusions of sodic in morecalcic plagioclase. Small inclusions (e.g. of pyroxene) that commonly occur inthe patches of more sodic plagioclase may have crystallized from melt trappedin the corroded cores (Vance, 1962). Patchy zoning in plagioclase has also been attributed to magma mixing where early formed plagioclase cores react with the changed magma composition resulting in overgrowth of plagioclase with different composition in equilibrium with the hybridized magma the early crystals are retained as the core with resorbed margin.

Crystal twinning occurs when two separate crystals share same crystal lattice points in a symmetrical manner forming an intergrowth of two separate crystals in a variety of specific configurations. The surface along which the lattice points are shared in twinned crystals is called a composition surface or twin plane. Twinning may occur during the growth of a crystal, or in post crystallization stages if the crystal is subjected to stress or temperature/pressure conditions different from those under which it originally formed. Twinning is important to recognize, because when it occurs, it is often one of the most diagnostic features enabling identification of the mineral. The nature of twinning is guided by twin laws. Twining can be defined by the symmetry operations based upon plane of symmetry (twin plane), axis of symmetry (twin axis) and centre of symmetry (twin center). Twinning can originate in three different ways, as growth twins, transformation twins, and glide or deformation twins.

New crystal may be added to the face of an already existing crystal when regular growth is impeded by accidents during crystal growth causing occurrence of twinning. Here the new crystal shares lattice points on the face of the existing crystal but has an orientation different from the original crystal. Such growth twins have either planar (contact twins) or irregular (penetration twins) compositional surfaces between each other.

Transformation twinning occurs when a pre-existing crystal undergoes a transformation due to a change in pressure or temperature. This commonly occurs in minerals that have different crystal structures and different symmetry at different temperatures or pressures. When the temperature or pressure is changed to that where a new crystal structure and symmetry is stable, different parts of the crystal become arranged in different symmetrical orientations, and thus form an intergrowth of one or more crystals.

During deformation, if atoms are pushed out of their original position in the lattice with a symmetrical arrangement, deformation twins are formed. Such twining is common is calcite and plagioclase which show tapered twin lamellae.

Simple twins represent twinning following a single twin law or if the twin plane associates with only two distinct halves(Fig. 1.6.5 a-c). For example, plagioclase (NaAlSi3O8 - CaAl2Si2O8) very commonly shows albite polysynthetic twinning following the albite Law - {010} which indicates that the twining occurs perpendicular to the b crystallographic axis. Albite twinning is a diagnostic property for identification of plagioclase.The most common twins in the monoclinic system occur on the planes {100} and {001}.

The combination twins represent overlapping of twinning following more than one twin laws. For example, the pericline twinning usually occurs in combination with albite twinning in microcline, producing a cross-hatched pattern(Fig. 1.6.5d), called tartan or crossed hatched twinning.Thistwinning pattern is one of the most characteristic diagnostic properties for the identification of microcline.

Under some conditions, minerals form regular intergrowths in igneous rocks. The term ‘intergrowth’ expresses interlocking of grains of two different minerals. Intergrowth texture results due tosimultaneous crystallisation of two mineral components of the magma at aparticular temperature for example, eutectic crystallisation.

Graphic intergrowth is represented by intergrowth of quartz and alkali feldspar in granites and pegmatites andin the interstices of some mafic rocks, i.e. gabbro.Graphic intergrowths are so called because many of them resemble ancient forms of writing. Here, the quartz blebs arealigned parallel to a well-defined crystallographic orientation giving rise to theeffect of cuneiform writing on a background of K-feldspar (host; Fig. 1.7.1a-e). Quartz is disposed in the form of prismatic, wedge-shaped crystals intersecting at an angle of about 60°. In graphic texture, all the quartz intergrowths have the samecrystallographic orientation, and the host K-feldspar also has a uniform optical orientation.This indicates that both minerals are parts of large single grains respectively, implying a lownucleation rate for each.

The microscopic equivalent of graphic texture is called micrographic (term used fortexture observed under microscope) and the rock is called granophyre. In other words, the term ‘granophyric’ representsmicrographic intergrowths of quartz and alkalifeldspar(Fig. 1.7.2 a-c).Sometimes such type of an intergrowth is ultra-small and can be seen underthe scanning electron microscope under high resolution. In some cases, the quartz : alkali feldspar ratio and chemical composition of granophyricintergrowths approximate the composition of the ‘ternary minimum’ (cotectic) in the system Or–Ab–Qtz (Vogt, 1921; Dunham, 1965; Barker, 1970; Hughes, 1971). Occurrence of granophyric intergrowth in alternatively explained by Lentz & Fowler (1992). They inferred that slow diffusion of Al to a growing feldspar interface causes increment of Si concentration saturating the local melt in quartz. This depletes the adjacent meltin Si, causing feldspar saturation and rhythmic precipitation of both minerals.

In graphic and similar intergrowths, both minerals growas single crystals forming rods or lamellae with shapes at least approximatelycontrolled by crystal structure. The quartzin graphic intergrowths occurs as angular rods with an irregular inner surfaceand a planar outer interface against the alkali feldspar. Graphic quartz–feldspar intergrowths can also be generated experimentally that undercooling of felsic melts (London, 1999). Some granites have euhedral phenocrysts of quartz, with interstitial micrographicintergrowths of quartz and alkali feldspar. This is caused by supercooling once the magma moves to shallower crustal levels, or by loss of water. Nucleation of granophyric intergrowthson quartz phenocrysts typically results in optical continuity of the quartz ofthe phenocryst and the quartz in the adjacent intergrowth (Hughes, 1972).

Myrmekite is a common symplectitic growth consisting of vermicular intergrowthof quartz and sodic plagioclase(Fig. 1.7.3a). Myrmekite typically occurs as lobes projecting into K-feldspar fromadjacent grains of plagioclase(Fig. 1.7.3b). In general, myrmekite is attributed to solid-statereplacement of K-feldspar accompanying deformation. Common occurrence of myrmekitein deformed granitoid rocks,such as augen gneisses and felsic mylonites (Binns, 1966) suggests that the development of myrmekite isconnected with deformation (Vernon et al., 1983). Myrmekite also occurs in metapelitic gneisses (Vernon et al., 1990), where it replaces K-feldsparduring late deformation.Most commonly, myrmekite occurs as lobes or colonies projecting intoK-feldspar (typically microcline) from the margins. The myrmekiteforms at the contact of K-feldspar and plagioclase, replacing the former. The plagioclase component typically nucleates on adjacent primaryplagioclase grains, so that the plagioclase of the myrmekite has the same crystallographicorientation as that of the primary plagioclase (Hubbard, 1966).

The myrmekitic intergrowth involves a chemical reaction where K-feldspar is replaced by plagioclase producing excess silica in form of quartz. The proportion of quartz in the intergrowth increases progressively withincreasing anorthite content of the plagioclase (Phillips & Ransom, 1968), probably as more calcic plagioclase contains less silica (Vernon, 2004). This reaction can also be fluid driven where Ca and Na may be released from plagioclase and transferred to myrmekite replacing K-feldspar, from which K could be expelled to form muscovite growing in plagioclase (Simpson &Wintsch, 1989).Myrmekite of this type may occur along fractures in the K-feldspar, indicating the role of fluid access (Vernon et al., 1983). Some myrmekite develops with muscovite by hydrous replacementof K-feldspar (Ashworth, 1972), during retrograde metamorphism or the deformation of granite (Vernon et al., 1983).

In mylonites, myrmekitenucleates on relatively weakly deformed or undeformedK-feldspar, while being deformed and recrystallized at the rear of the growinglobes (Vernon et al., 1983). Myrmekitetypically grows at ranges of 450–500 ◦C (Tribe &D’Lemos, 1996) to 500-670 ◦C (Wirth &Voll, 1987).

Magmatic enclaves are volumes of rock (or aggregates of mineral) surrounded by emplaced host rock of related but distinct composition and of separated genesis.

Vernon (2004) categorized different types of enclaves as (1) fragments of rock (xenoliths) orindividual grains (xenocrysts) that are broken off thewalls of the magma chamberand incorporated into the ascending magma, the process being known as ‘magmaticstoping’; (2) mafic microgranular and microgranitoid/felsic microgranularenclaves (MME or FME),which are globules of other magma bodies mingled with the host magma(Fig. 1.9.1a, b, d, e); (3) concentrations of fine-grained, early-precipitating aggregates thathave been broken up and incorporated into the same or a later magma body and (4) refractory residual grains and aggregates, which may be carried in themagma from the source area(Fig. 1.9.1f). The last category may be difficult to detect due to partial or near complete reactionwith the host magma and re-equilibration to lower temperatures and pressures during ascent. However, the most common categories of enclave are microgranular enclaves and xenoliths.

Most xenoliths and xenocrysts have angular or irregular shapes(Fig. 1.8.1a, d). Xenocrysts are foreign igneous crystals (not crystallized from the melt) that has been introduced into the melt from an external source, e.g. the surrounding country rocks (Fig. 1.9.1i)or a previously crystallized part of the same igneous body. Xenocrysts, which are usually in chemical and/or thermal disequilibrium with the melt, may become rounded due to reaction or part melting in effect of the host magma(Fig. 1.9.1c, e, i). In submarinebasalts, olivine grains inferred to be xenocrysts are anhedral, and may showfracturing and optical evidence of plastic deformation. If xenocryst or xenoliths are unstable in the magma, they show reactionrims or partly melted/resorbed rims and/or interiors (Fig. 1.9.1i-j).Some xenoliths may be enclosed by microgranitoid enclaves forming a variety of ‘double enclave’. This forms by incorporation of a rock fragmentin a magma that subsequently becomes mingled with the present hostgranite. Other xenoliths or xenocrystsmay act as nucleus around which spectacular, compositionallylayered rims develop formingorbs in orbicular granites.

Microgranular enclaves (ME) are very common in granitoid rocks (Didier &Barbarin, 1991; Hibbard, 1981; Reid et al., 1983). They are commonly felsic to intermediate in composition (rarely mafic), are rounded, scalloped or lenticular in shape, and are finer-grained and typically contain more mafic minerals than the host granite. They have igneous microstructures, commonly with euhedral phenocrysts, oscillatory zoning in plagioclase, and may have mineral alignment reflecting magmatic flow. Some exhibit chilled margins against the host granite. These features indicate that theenclaves were originally magma globules that flowed and quenched to finergrainedsolid enclaves in the host magma (Reid et al., 1983; Vernon, 1983, 1984). Although some early workers favoured a solid-state (xenolith) origin of these enclaves (i.e. Nockolds, 1933; Grout, 1937).However, recent work has produced overwhelming evidence of magmatic origin referring to the role of magma mingling/mixing after formation of microgranular enclaves (Vernon, 1996b).Microgranitoid enclaves commonly show evidence that they were formed bymagma mixing or hybridisationbefore they were incorporated into the host granitic magma, a process called ‘magma mingling’ (mechanical interaction; e.g. Didier and Barberin, 1991; Vernon. 1984). Magma mingling involves the intermingling of twoor more magmas without pervasive mixing of their melt components, whereasmagma mixing involves homogenization of the melts (chemical mixing) and either the conversionof any pre-existing crystals to minerals stable in the hybrid (mixed) melt or their armouring by stable minerals.Generally, the enclaves are too small and isolated for mixing to take place in situ. Mixing only takes place at the margins. Thefollowing microstructures in microgranular enclaves indicate magma mixing:

(1) Quartzxenocrysts incorporated in the more mafic magma from the more felsic magma may exhibit glassy rims on corroded boundaries. The latent heat of crystallization required todissolve the quartz is taken from the immediately adjacent magma, and thiseffectively undercools the magma at the margin. Similar process promotes fine-grained crystallizationof minerals around the xenocrysts in which the host magma is presently saturated. As a result, fine grained mineral aggregate such as hornblende- or orthopyroxene rich mantles around quartz xenocrysts in metaluminous or peraluminousgranites and volcanic rocks are commonly found. These mantled quartz xenocrysts are usually called ‘ocelli’ (Vernon, 1983, 1990a, 1991a).

(2) Alkali or predominantly K-feldspar megacrysts in microgranular enclaves are quite common (Reid et al., 1983; Vernon, 1983,1984; 1986a, 1990a, 1991a; Cox et al., 1996). These crystals are commonly identical to K-feldspar megacrysts in the adjacent host suggesting that they are xenocrystic in origin and the mixing involved the actual host magma. These xenocrysts typically show oscillatory zoning, or overgrowth owing to the change of chemical composition of the host magma due to mixing. Theymay also exhibit partly dissolved and rounded margins being in disequilibrium with the new, hybridised host magma. Many K-feldsparxenocrystsare rimmed with plagioclase as the hybridized new magma, which is more mafic in composition than the previous, stabilizes plagioclase (Vernon, 1990a). Resorbed K-feldspar xenocrysts also occur inmagmatic enclaves in rhyolite (Bacon & Metz, 1984) and trachyte (Cantagrel etal., 1984). Often, the enclaveminerals show alignment against the megacryst, indicating that the enclave magmaflowed after the overgrowth was formed and further suggesting that the enclave magma was in molten stage, i.e. magma globule.

(3) Corrosion, overgrowths and sharp zoning discontinuities (compositionalspikes) are common features of plagioclase in microgranitoid enclaves and in thehost granite owing to crystal growth under compositionally changing (more mafic) magma composition. Hibbard (1981) described corroded grains of plagioclase (from the more felsic magma) with dendriticovergrowths of more calcic plagioclase (precipitated from the hybrid melt), the dendritic habit resulting from the strong undercooling and compositionalchange of the melt caused by the magma mixing.

Phenocrysts, especially quartz and olivine, in volcanic rocks exhibit embayed grain margins. These are interpreted as the result of magmatic corrosion (i.e. resorption;Fig. 1.8.1, 1.9.1i,dissolution), resulting from a change of conditions, i.e. pressure or a change in chemical composition of the melt caused by mixing of magmas. This causes a previously stable crystal to become unstable with respect to the liquid that results in dissolving of the crystal from the margin. Embayment is characterized by rounded corners of the crystals. Often, compositional zoning is truncated by embayment. Embayment can often be related to fracture in the crystal as dissolution tends to begin with them. Often new minerals, i.e. neoblasts grow on the surfaces of embayed crystals as product of the reaction with the melt(Fig. 1.9.1j). In this instance, the embayed crystal may be foreign to the magma or may be a phenocryst that has reacted in response to changing conditions. Although embayments are typically rounded(Fig. 1.9.1g), some relatively planar, crystallographically controlled embayments may be present (M¨uller et al., 2000).

Orbicular rocks are igneous rocks that contain concentric shells of different texture and/or mineralogy about a central core. Orbicular rocks, which are worldwide in occurrence, have been described in more than one hundred localities. Orbicules are found in igneous, metamorphic, and migmatitic terrains and are not restricted to unusual or limited compositions (Leveson, 1966). Orbicule structure is composed of concentric shells of contrasted texture and mineralogy about a central core, which is often characterized by radially and/or tangentially oriented minerals. Hypotheses of orbicule genesis include both magmatic and metamorphic origins for these rocks; however, no single hypothesis provides a general explanation (Leveson, 1966 and references therein). Spacing of orbicule shells reflects instability of the environment during genesis. Some orbicule shells result from exchange of material between core and matrix. Conditions that cause rhythmic layering in igneous rocks may result in orbicule formation if crystallization is localized about scattered cores. Cores of such orbicules serve as crystallization centres and need not have specific or limited compositions. No correlation has been demonstrated between orbicule structure, chemical composition, or gross geologic setting.Standard orbicule terminology is suggested: orbicular rocks, orbicules, cores, shells, matrix, and country rock. Orbicular texture is primarily found in granite. Host rocks vary from fine- to coarse-grained, dykes, sills, stocks, batholiths, lavas, and gneisses. Orbicular facies are generally local, less than several hundred meters in greatest dimension. Orbicules range from <3 cm to >30 cm in diameter. Shells number from one to more than twenty, and may be spaced irregularly or in geometrical progression. Cores may be fragments of foreign rock (xenoliths), early segregations from a magma (autoliths) or, in metamorphic surroundings, recrystallized country rock.

Basalt is an extrusive maficigneous rock formed from the rapid cooling of low-viscosity lava. A wide range of textural variation can be seen in basalt. Overall texture of basalts can vary from glassy or aphanitic toequigranular and fine grained (<1 mm) orporphyritic. In porphyritic basalts, phenocrysts usually are generally of augite, olivine, pegioniteor a calcium-rich plagioclase which areembedded in a matrix consisting of fine-grained crystals (crystallites), microlites and/or glass(Fig. 1.2c, d). Olivine phenocryst, when present, may have rims of pigeonite. These phenocrysts often occur as clusters of same (glomeroporphyritic; Fig. 1.4.2e) or different minerals (cumuloporphyritic; Fig. 1.4.2d). In case of very rapid cooling, presence of dendritic crystals (i.e. olivine, plagioclase, magnetite) or ‘spiky’ (swallowtail) plagioclase crystals are common in basalts(Fig. 1.6.4b).A wide array of such textures including dendritic plagioclase (acicular, hollow, swallow-tail forms, rosettes etc; Fig. 1.2.5a), clinopyroxene (plumose, radiating to sheaf-like) and olivine (hollow, skeletal, skeletal chains and lantern-like; Gélinas& Brooks, 1974)have been reported from Archaeantholeiitic basalts.Basalts commonly show ophitic and/or subophitic texture defined by complete or partial inclusion of plagioclase laths within clinopyroxene. Intergranular (interstices of the crystals are filled by fine crystals, usually of plagioclase) and intersertal (interstices of the crystals are filled by glass; Fig. 1.4.7a) textures are also commonly found in basalts. Feldspar microlites are sometimes arranged subparallelly, reflecting flow in relatively rapidly cooled magma and defining pilotaxitic texture.Spherulitic aggregates (‘varioles’) in some basalts consist ofplagioclase and clinopyroxene.Basalts often contain xenocrysts derived from their parent magma (i.e. olivine) which can be identified from their corroded and embayed grain boundaries. Flow banding, although comparatively rare, but can be seen in basalts(Fig. 1.4.2k, l).

Vesicles are formed in basalt when dissolved gases bubble escape from the magma as it decompresses during its approach to the surface. When vesicles make up a substantial fraction of the volume of the rock, the rock is described as scoria (described in detail under pyroclastic section). Vesicles are also formedby steam bubbles enclosed in some fine-grained, moistash deposits generated by explosive eruptions. Amygdales are former vesicles thathave been partially or completely infilled with secondary minerals. Pipe vesicles are slender cylindrical cavities up toseveral millimeters across and tens of centimeters in length. They are commonly found near thebases of subaerial pahoehoe lava flows, but may also occur indykes and sills. Adjacent pipevesicles in flows occasionally coalesce upward formingan inverted Y(Fig. 1.4.10b); few subdivide upward. Pipe vesicles are formed due to the continuous growth and advancement of bubbles that are attached to the zone of solidification forming pipes perpendicular to thesolidification front.Vesicles partly filled with residual melt segregated from the surrounding magmahave been described from submarine basalts, subaerial basalts and are called as segregation vesicles.

Andesite lavas usually have porphyritic or vitrophyric textures. Here, various phenocrysts, such as plagioclase, clinopyroxene (mostly augite), orthopyroxene, hornblende, biotite and olivine occur within a fine grained matrix mostly consisting of plagioclase and clinopyroxenemicrolites and glass. However, in altered or low-grade metamorphosed andesites, the glass may be devitrified into fine-grained aggregate of muscovite and/or clay minerals where the plagioclase phenocrysts may be saussuritized(Fig. 1.4.11a).Plagioclase crystals are usually complexly zoned.Thecore of the plagioclase phenocrysts may be homogenous, patchyor oscillatory-zoned or may be inclusion-rich. Boundary of the plagioclase phenocrysts may often be resorbed(Fig. 1.4.11c).At several instances, the core is surrounded by a clear normally zoned mantle and thin rim, usually similar in composition to groundmass microlites. The inner feldspar core often shows resorbed boundary. Abrupt changes in zoning pattern, resorption of the boundary, reverse zoning are typical characters of plagioclase phenocrysts and groundmass crystals. These features are attributed to the physico-chemical processes related to the magma mixing or assimilation which is the prime process responsible for producing andesitic magma.Augite is the second most abundant type of phenocryst in andesite, andis also the most abundant crystals in the groundmass.Augite crystals in the groundmass and phenocrysts are generally compositionally similar.Orthopyroxene, commonlybronzite and hypersthene are common, especially in basaltic andesite.Green to brown, magmatichornblende is the typical amphibole found as phenocrysts in andesite which normally does not appear in the groundmass. Coarse, subhedral biotite flakes may also occur as phenocryst (Fig. 1.4.11b)while fine, more or less euhedral biotite may occur in the groundmass. Andesitic magma with relatively higher vapour and alkali components may stabilize hornblende and biotite.Experimental studies indicate the outer rim ofhornblende crystals show evidence of resorption and may be replaced by a corona of plagioclase, pyroxene and magnetite (Vernon, 2004).Olivine occurs in minor (less than 1%) amounts normally as phenocrysts. Both ilmenite and titanomagnetite occur as primary minerals in andesite and they occur both in phenocryst and in the groundmass. Garnet is a rare mineral found in some andesite and the distribution of such andesite is normally restricted to continental margins that contain pelitic sediments. Some have plagioclase-magnetite and chlorite corona that indicate reaction between magma and crystals, followed by the precipitation of a rind of minerals that are stable at lower pressures. Feldspar microlites in groundmass are sometimes arranged subparallelly, reflecting flow in relatively rapidly cooled magma and defining pilotaxitic texture (Fig. 1.4.8a; 1.4.11d).

Felsic volcanic rocks collectively represent both solidified lava (i.e. rhyolite) and pyroclastic rock (i.e. tuff). Felsic tuff is composed of volcanic ash, glass shards and lithic fragments(discussed under the section on Pyroclastics in this Atlas) while texture of felsic lava can vary from equigranular (i.e. fine grained or aphanitic) to more commonly porphyritic.

Porphyritic texture consists of relatively large, euhedral(Fig. 1.4.13d)to subhedralphenocrysts dispersed in much finergrained (microcrystalline) or glassy groundmass. Among phenocrysts, bipyramidal quartz (Fig. 1.4.13c)is a diagnostic feature. Glomeroporphyritictexture may also be present consisting of a numerous phenocrysts clustered together. Phenocrystare sometimes cracked and broken apartas a result of shear duringflowage, rapid vesiculation of the enclosing melt, quenching and hydration of the host lava orpressure release during magma rise and eruption. Resorption of phenocryst is a common feature.

Embayment or resorption: Theoriginal shapes of phenocrysts can be modified if thechemical or physical environment changes. Shapes are modified by (i) partial resorption, whichresults in embayed and rounded outlines, and (ii) reactionwith the melt, which generates rims of newly-crystallized, fine-grainedminerals around the phenocrysts. Quartz phenocrysts insilicic lavas and syn-volcanic intrusions commonlyshow the effects of resorption. They typically have abi-pyramidal habit but are embayed and partly rounded. Quartz phenocrysts are often resorbed as silica solubility in the melt increases asthe pressure decreasesduring rise and eruption of the magma.

Xenocrysts:These are crystals which didnot crystallize from the host magma but wereaccidentally incorporated from a foreign source, such asdisintegrating wall rocks. They are in disequilibrium with the melt and are identified by resorption,embayments and reaction rims. Xenocrysts can comprisemineral phases incompatible with thehost magma composition.

Flow bands: These are present in massive, uniform felsic lava flow units. When the viscous lava flow encounters a surface, frictional drag produces internal banding within the mobile lava. Flow linesappear most clearly in glassy rocks which are often contorted and folded during the flow and may be delineated by concentrationsof crystallites or microlites.

Acid volcanic are often associated with pyroclastic rocks. For a detailed description of pyroclastics, interested readers may go through Chapter 3 : Pyroclastic Rocks in this Atlas.

An ophiolite is a section of the oceanic lithosphere emplaced upon continental crust or within the accretionary prismsediments of a subduction zone. From bottom to top, a typicaland complete ophiolite sequence comprises depleted mantleperidotite with tectonite at the base, layered ultramafic–mafic cumulates, massive (isotropic) gabbro, sheeted dikesand extrusive volcanic rocks represented by pillow basalts(Coleman 1977).The sequence is typicallyoverlain by deep-sea pelagic sediments and chert.Petrographic textures and mineral assemblages of ophioliticrocks provide important clues to their environment offormation. There is a large array of microstructural and petrographic characters associated with each above mentioned rock types of any ophiolite suite. In the following sections, petrographic characters of the basal tectoniteshave been described:

The peridotite tectonite, also designated as meta-peridotites, are dominated by dunite followed by peridotite (dunite, harzburgite(Fig. 1.4.16a)lherzolite andminor wehrlite), pyroxenite (olivine websterite) websterite, olivine clinopyroxenite and clinopyroxenite), gabbroids (olivine gabbro, olivine gabbronorite, norite, gabbro and hornblende gabbro), plagiogranite and anorthosite.These are emplaced as tectonic slices. Lateplagiogranite intrusions cut through the entire magmaticsequence.

The cumulate peridotites show a protogranular texturewith medium- to fine-grained, granular, tabular or smoothlycurved olivine crystals(Fig. 1.4.16g). These rocks are classified on the basis of abundance of olivine (~40 volume %), ortho- and clinopyroxenes,which are variably altered to serpentine(Fig. 1.4.16i). Olivineshows preferred orientation and together with serpentinedevelops a strong foliation. Olivine occurs in two distinctmodes: (a) large porphyroclasts (sheared type; Fig. 1.4.16h, l) and protogranular (~2 mm) olivine and (b) small neoblasts and fine-grained (~1 mm) olivine(Fig. 1.4.16b, k). Theolivine porphyroclasts are flattened or curved, and often brecciated. Flattening of olivine grains, strain shadow in olivineporphyroclasts and kink bands in pyroxenes are common features observed as evidence of severestress. Olivine of dunitetectonitecommonly has amorphous greyish core defined by streaks of opaque mineralstraversing diagonally across the core and confined within the crystal outline while their outer border is commonly altered to red-brown Fe-oxide due to post-emplacementoxidation. Peridotite or dunitetectonitesin ophiolites are essentially spinel-bearing and devoidof plagioclase. The spinel grains are usually euhedral in shape, sometimes lensoid, and equidimensionalor skeletal. The cores of the serpentine veins oftenentrap trails of fine magnetite granules. Spinel may contain olivine inclusions as they crystallise after olivine. The neoblasts show uniform extinction.

Serpentinite: The serpentinites display a wide variety of textures, viz., massive, mesh or reticulate, ribbon, fibrous, bladed, sheared and brecciated (Ghose et al., 2014). Besides serpentineand relict olivine and pyroxene, other minor phasesare opaque/chrome-spinel, talc, tremolite, chrysotile, calcite,biotite, muscovite and chlorite. Mesh (honeycomb)texture is common, consisting of primary, secondaryand tertiary chords (veins) of serpentine enclosingfractured grains of olivine and pyroxene(Fig. 1.4.16l). In some cases,serpentine exhibits a ribbon texture composed of parallel andundulating bands. The bladed/mat texture shows blended andfelted mass of fine-to-medium size serpentine bladesbelonging to more than one generation. Thin chrysotile veinsoccur as fracture fillings. Both magnetite and chromite occuras disseminations, fracture fillings and also along cleavagetraces.

Gabbroids:The mafic cumulates occurring both as massive and layeredgabbroid complexes are represented by olivine gabbro,hornblende-olivine (serpentinised) gabbro, norite, gabbronorite,gabbro, hornblende gabbro and pegmatitic gabbro withveins and small bodies of plagiogranite and anorthosite (Ghose et al., 2014). The gabbros are medium- to coarse-grained,commonly showing hypidiomorphic texture, andoccasionally porphyritic, poikilitic, equigranular, sub-ophiticand brecciated.

Plagiogranite:The plagiogranites range in mineralogical composition fromdiorite, tonalite, trondhjemite and albite granite (Agrawal and Ghose 1986). They are coarse-grained, showing panidiomorphicto hypidiomorphic, granular and porphyritic or even gneissic texture.

Raiolarian chert: Radiolarian chert is a siliceous, comparatively hard, fine-grained, chemogenic sedimentary rock that is composed predominantly of the microscopic remains of radiolarians. It is well-bedded, microcrystalline radiolarite that has a well-developed siliceous cement or groundmass.Radiolarian chert may be deposited in relatively shallow depths butmore commonly are pelagic, deep water sediments thatforms from indurated radiolarian oozes.They are a frequent component in ophiolite sequences.

The compiler sincerely thanks Dr. Basab Chakraborty, Director, GSI and Dr. Biswajit Ghosh, Associate Professor, Calcutta University are thanked for their thoughtful and detailed review that significantly improved the quality of the compilation. Finally, all the field geologists of Geological Survey of India, who have contributed valuable field photographs, photomicrographs and their description are thanked as their contributions materialized this compilation. The work is dedicated to all the sincere geologists who spent hours working in field and studying rocks under microscope whose effort made this work successful.

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