UNA CONEXIÓN TCP COMPLETA

From the two previous sub-sections it is evident that the AMS ellipsoid size (K_mean), shape (m), strength (H) and the orientations of its principal axes, referred to as the magnetic fabric, may parallel the mineral petrofabric (i.e. K₁ = A₁; K₂ = A₂; K₃ = A₃) (Knight and Walker, 1988). Several studies have supported this assumption through combined AMS and petrographical analysis (e.g. Khan, 1962; Launeau and Cruden, 1998; Abelson et al., 2001;

Callot et al., 2001; Féménias et al., 2004; Horsman et al., 2005; O’Driscoll et al., 2008). Even in weakly anisotropic igneous rocks, it is now widely accepted (e.g. Bouchez, 1997) that the magnetic lineations and foliations potentially reflect the magmatic fabric and can provide information on magma migration, flow geometries, and regional strain (King 1966; Owens and Bamford 1976; Hrouda, 1982; Borradaile 1987, 1988; Rochette 1987; Rochette et al., 1992; Tarling and Hrouda 1993; Petronis et al., 2004; Horsman et al., 2005; O’Driscoll et al., 2006; Stevenson et al., 2007a;). However, some magnetic fabrics, particularly the magnetic lineation, have been measured that significantly deviate from the mineral petrofabric (e.g.

Geoffroy et al., 2002; Callot and Guichet, 2003; Aubourg et al., 2008; Creixell et al., 2009).

Although processes responsible for sub-fabric formation have been suggested to account for these disparities (discussed further in section 2.2.4), there are several potential explanations related to the magnetic mineralogy that first need to be contemplated. An awareness of these possible caveats is further supported following a consideration of the broad range of mineralogical compositions and crystallisation histories that igneous rocks may have (section 2.1.3).

2.2.3:1 Magnetic fraction composition

The magnetic fraction of igneous rocks is typically dominated by either Fe-bearing silicate phases or by magnetite. In Fe-bearing silicate dominated rocks, the magnetocrystalline anisotropy of paramagnetic silicates such as olivine, biotite or amphibole will control the AMS fabric when magnetite is absent or of very low volume percent. Comparatively, for magnetite-bearing rocks (>0.1 vol. % magnetite) the contribution of Fe-Mg silicates is negligible due to the high intrinsic magnetic susceptibility of magnetite (Tarling and Hrouda, 1993). Stacey (1960) postulated the magnetic fabric of magnetite-bearing rocks (e.g. most mafic igneous rocks), therefore results primarily from the shape anisotropy of the grains of magnetite. This effect has been examined and corroborated by microscopic observations on various rock types (Khan, 1962; Uyeda et al., 1963; Ellwood and Whitney, 1980; Grégoire et al., 1995). However, some mineral phases (e.g. haematite, Tarling and Hrouda, 1993; tourmaline, Borradaile and Jackson, 2004) produce a magnetic fabric that is geometrically inverse to its petrofabric (i.e. K₁ = A₃). For haematite, this occurs because its basal plane (parallel to its magnetocrystalline ‘easy’ axis) commonly defines an oblate crystal shape (Tarling and Hrouda, 1993) and haematite displays canted antiferromagnetic behaviour (Dunlop and Özdemir, 1997). In contrast, paramagnetic tourmaline has a minimum susceptibility axis parallel to its longest crystallographic axis (Borradaile and Jackson, 2004).

Identification of potential exsolution or alteration products is also important. Kissel et al., (2010) report ilmenite laths exsolved parallel to the short axis of coarse titanomagnetite crystals in dykes in E Iceland have divided the host crystal into small MD titanomagnetite laths with long axes orthogonal to the magma flow fabrics.

2.2.3:2 Controls on magnetite petrofabrics

Magnetic fabrics measured in magnetite-bearing rocks are at times difficult to interpret because of the unknown relationship between the magnetite fabric and the mineral fabrics of the volumetrically dominant silicate phases (e.g. feldspars, amphibole and biotite).

When attempting to interpret magma flow fabrics this issue is compounded as magnetite frequently crystallises late. The shape of magnetite crystallised in association with K-feldspar, hornblende, and biotite is often observed to be controlled by the crystalline faces of the dominant silicate phases (Hrouda et al., 1971; Borradaile, 1994). Stacey (1960) originally suggested the primary silicate framework may control the growth and orientation of later crystallising magnetic mineral phases. Therefore, the magnetic fabric likely reflects the petrofabric of the silicate phases. This “template” behaviour has been demonstrated with quantitative image analysis of thin sections (e.g. Cruden and Launeau, 1994; Archanjo et al., 1995; Launeau and Cruden, 1998). Hargraves et al., (1991), suggest that the primary silicate framework control on magnetite distribution may also create a network of closely spaced magnetites that are able to magnetically interact to an extent where shape anisotropy may be replaced or accentuated by a distribution anisotropy. Cañón-Tapia (2001) concluded shape anisotropy should dominate, unless the magnetite grains were equant and/or clustered.

2.2.3:3 Grainsize influence on AMS

The magnetic response of titanomagnetites is controlled by its grainsize as well as its shape anisotropy (Tarling and Hrouda, 1993). MD titanomagnetites (>100 µm) have a strong shape preferred anisotropy, thus the maximum susceptibility axis will parallel the long axis of the grain. In contrast, SD magnetites (<1 µm) are more susceptible to magnetization along the magnetocrystalline ‘easy’ axis, orthogonal to the shape long axis (Hrouda, 1982; O’Reilly, 1984; Potter and Stephenson, 1988; Dunlop and Özdemir, 1997). From the dependence of principal susceptibility axis orientation on grainsize, titanomagnetite populations consisting purely of MD or SD grainsizes are interpreted as producing normal and inverse magnetic fabrics respectively (Rochette et al., 1999; Ferré, 2002). A normal magnetic fabric implies that the magnetic fabric reflects the mineral fabric and that the magnetic foliation in sheet intrusions is parallel to the dyke plane. Distinguishing the presence and proportion of magnetic domain states is an important distinction as inverse magnetic fabrics orthogonal to the actual petrofabric and likely perpendicular to the dyke plane (Potter and Stephenson,

1988; Hargraves et al., 1991; Rochette et al, 1992; Dragoni et al., 1997; Rochette et al., 1999;

Aubourg et al., 2002, Ferré, 2002; Cañón-Tapia and Chávez-Álvarez, 2004). A mixture of SD and MD titanomagnetites may yield intermediate fabrics, where two of the principal susceptibility axes are switched (Rochette et al., 1999; Ferré, 2002).