Geology, geochemistry and emplacement conditions of the Vega intrusive complex: an example of large-scale crustal anatexis in north-central Norway

Abstract The c. 350 km2 Vega intrusive complex is part of the Bindal Batholith and was emplaced at c. 475 Ma into polydeformed supracrustal rocks of the Helgeland Nappe Complex. The intrusive complex is tilted towards the west, exposing asymmetrical zoning. From east to west, the complex is composed of biotite granite, garnet-biotite granite, garnet-bearing muscovite biotite granodiorite and sillimanite-bearing garnet cordierite muscovite biotite granodiorite. In addition, the complex contains small amounts of intrusive migmatite. Granodiorite and intrusive migmatite contain abundant metasedimentary, mafic and ultramafic enclaves. Granodiorite, granite and migmatite are generally peraluminous to strongly peraluminous, calcic to alkalic and magnesian, with initial 87Sr/86Sr ratios of 0.7096–0.7469 and ϵNd from −7.0 to −11.0. Emplacement of the Vega intrusive complex was coeval with the intrusion of metaluminous dioritic rocks. The intrusive mafic rocks and enclaves in the complex have MORB-like (mid-ocean ridge basalt-like) to calc-alkaline geochemical characteristics. The lack of an isotopic compositional trend between mafic and granitic rocks indicates that magmas did not mix. Instead granitic magmas formed by unmixing of residual phases from crustally derived magmas. Partial melting of supracrustal source rocks may have been related to intra- and underplating of MORB-like magmas into the lower crust during extension. Supplementary material: Detailed petrographic descriptions, photomicrographs, and field images of selected enclaves are available at http://www.geolsoc.org.uk/SUP18653.

The Earth's geothermal gradient is lower than the melting conditions of crustal rocks and, for the most part, the Earth's crust is solid. The formation of so-called 'S-type' granites from crustal source rocks (Chappell & White 1974;White & Chappell 1977) therefore requires that the temperature of crust be raised or the melting temperature lowered. Perturbation of the gradient by crustal thickening may cause partial melting and the formation of S-type leucogranites (Nabelek & Liu 2004). Large volumes of S-type granites have been formed by partial melting of predominantly greywacke protoliths. This process was accompanied by hybridization of the crustal melts with mafic, mantle-derived magmas (Collins 1996;Healy et al. 2004). Although peraluminous leucogranites may form large batholiths (e.g. Le Fort 1981;Nabelek & Liu 2004), large plutonic masses of classic S-type granite (White & Chappell 1977) that lack compositional evidence of mantle input are uncommon. One such batholith-size intrusive complex in the Bindal Batholith, the Vega intrusive complex, provides insight into the conditions necessary to form voluminous crustal magmas which are uncontaminated by hybridization with mafic magma. Our study focuses on a description of the complex, a general model for its petrogenesis and a possible tectonic setting within the context of the Scandinavian Caledonides.
The Scandinavian Caledonides are composed of a series of thrust complexes that were juxtaposed during multiple events culminating with the Laurentia-Baltica (Scandian) collision in the Early Devonian. The lowest allochthonous units originated along the margin of Baltica, whereas the highest (parts of the Upper and the Uppermost Allochthons including the Helgeland Nappe Complex) are exotic to Baltica; they were probably derived from the Laurentian margin and from arc complexes and micro-continents of uncertain origin (see review in Roberts et al. 2007). The timing of ophiolite formation, plutonism and metamorphism and the Early Ordovician evolution of the Helgeland Nappe Complex (HNC) supports a general correlation between the Uppermost Allochthon in Scandinavia with rock complexes in the British Caledonides and the Appalachians (Barnes et al. 2007). Granitoid plutons are common in many of these units, and the origin and setting of the Vega intrusive complex may therefore have important consequences for understanding the history of these rock complexes.
Bindal magmatism occurred in three or possibly four magmatic pulses: early 'S-type' magmas c. 478 -473 Ma; minor mafic magmatism at c. 465 Ma; a relatively large pulse of broadly calcalkaline magmatism during 450-442 Ma; and a final pulse of calc-alkaline magmatism between 439 and 424 Ma (Barnes et al. 2007). Most detailed geological, geochemical and structural studies have been performed on intrusions that have a relatively large mantle component, such as the diorite and related intrusions in the Hortavaer intrusive complex (c. 465 Ma; Barnes et al. 2005;McCulloch 2007;Li 2008) and gabbro and diorite intrusions in the Velfjord area (445-448 Ma; Barnes et al. 1992Barnes et al. , 2002Barnes et al. , 2004. These relatively high-T plutons are considered to be related to Taconian arc activity outboard of Laurentia (Roberts 1988(Roberts , 2003Yoshinobu et al. 2002;Barnes et al. 2002Barnes et al. , 2007Meyer et al. 2003).
The 478-473 Ma S-type plutons are relatively uncommon in the Bindal Batholith and, prior to this study, no comprehensive examination of these granites (senso lato) and their host rocks has been undertaken. The absence of such granites from later magmatism in the region suggests a secular change in the regional geotherm and a transition to more clearly arc-related magmatism in the region. The Vega intrusive complex is the largest S-type pluton in this age range. In this contribution, we present geological, geochemical and thermobarometric data that inform a model for the emplacement of the complex and the regional tectonic setting.

Geological setting
The Vega intrusive complex is situated 15 km WNW of Brønnøysund in north-central Norway and is exposed on the islands of Vega, Ylvingen, Søla and a number of smaller islands in the Vega archipelago (Fig. 2). The complex is elliptical in map view and c. 28 km long (east -west), 17 km wide and has c. 800 m of vertical relief. The complex was emplaced entirely into the HNC of the Uppermost Allochthon ( Fig. 1) (Nordgulen 1992).

Host rock structure and deformation
The Vega intrusive complex intrudes a structural boundary that separates host rocks with distinct structural features and orientations. Folded contacts between metasedimentary rocks trend c. N858W west of Gladstad, and contacts trend c. N308E east of Gladstad on the islands of Grimsøya and   Gustavson (1975Gustavson ( , 1977. Inset shows author distribution. Ylvingen (Fig. 2). Two contractional deformation events (D2-D3) are defined in host rocks west of Gladstad (Anderson et al. 2005). F2 folds are isoclinal and have formed compositional banding which has been refolded into F3 folds with wavelengths ranging from several centimetres to 1km. Amphibolite dykes have been transposed in S2 compositional banding and in some cases are disrupted. Granitic sheets intruded parallel to gneissic layers and were folded during the formation of F3 folds (Fig. 3e) (Anderson et al. 2005). Granitic dykes on NW Vega have experienced ductile (Fig. 3f) and brittle boudinage.

Magmatic rocks
The intrusive complex consists mainly of granitic and granodioritic rocks and is zoned. From east to west, the zones are biotite granite, garnet biotite granite, garnet-bearing muscovite biotite granodiorite and sillimanite-bearing garnet cordierite muscovite biotite granodiorite (Fig. 2); these units are described in the following sections. Boundaries between igneous units vary from gradational to sharp. Zircon crystallization ages for Vega granodiorite, biotite granite and migmatitic rocks are 475 + 4, 476+2 and 472 + 5 Ma, respectively Barnes et al. 2007).

Biotite and garnet-biotite granite of Ylvingen
Medium-grained biotite granite (Fig. 4a) underlies the western half of Ylvingen and skerries along the northeastern side of Vega (Fig. 2). Textures range from equigranular to porphyritic with K-feldspar phenocrysts up to c. 2 cm long. Medium-grained garnet biotite granite crops out on the northern end of Ylvingen (Fig. 2). Both units display a weak magmatic foliation defined by biotite (Oalmann et al. 2011). Leucogranite is medium-grained with textures ranging from hypidiomorphic granular to granophyric (Oalmann et al. 2011). Granitic rocks also have a subsolidus foliation defined by quartz aggregates and biotite (Oalmann et al. 2011). Enclaves of marble, pelitic schist, calc-silicate schist, gneiss, quartzite and biotite-rich schist are sparse (Oalmann et al. 2011).

Vega granodiorite
The Vega granodiorite consists of medium-grained garnet-bearing muscovite biotite granite and granodiorite and underlies most of Vega, Søla and island groups south and SW of Vega (Figs 2,. In the western two-thirds of the unit, nodular to dendritic cordierite (Fig. 4b) and fibrous to acicular sillimanite are present. Leuco-granodioritic rocks are medium-grained, contain ≤5% biotite and occur in layers or pods several metres long. The unit contains abundant enclaves that consist of tonalite, cordierite leucogranite, leucogranite, biotite diorite, hornfels, muscovite schist, muscovite biotite gneiss, calc-silicate gneiss, quartzite, metaperidotite and surmicaceous (biotite-rich) schist (see Marko 2012 for a detailed description of enclave types, sizes and shapes). Diorite and tonalite enclaves are commonly surrounded or partially rimed by leucocratic moats composed of plagioclase, K-feldspar and quartz. Some diorite enclaves are zoned and have reaction rinds of coarsened biotite.
The Vega granodiorite is locally layered; such layering occurs on Vega and the islands between Vega and Ylvingen in 1-500 m wide zones, which are adjacent to non-layered granodiorite (Fig. 2). The layering consists of millimetre -centimetrescale bands of light and dark granitic rock, with the variation due to differing proportions of biotite. Magmatic foliation is defined by the alignment of elongate enclaves and by a weak alignment of biotite grains subparallel to the enclaves. East of Fløa and Vika (Fig. 2), magmatic foliation overprints modal layering and has an opposing dip direction (Fig. 4c).

Fuglevaeret granodiorite
Skerries along the western contact of the complex at Fuglevaeret (Fig. 2) are underlain by massive, medium-grained, cordierite-bearing granite to granodiorite ( Fig. 4d), here referred to as the Fuglevaeret granodiorite. Similar rocks crop out on NW Vega (Fig. 2). Fuglevaeret granodiorite is hypidiomorphic granular and contains cordierite in three distinct habits: centimetre-scale subhedral, blocky phenocrysts, dendrites and nodules. Magmatic fabric and structures are absent and this unit contains distinctly fewer enclaves than the Vega granodiorite. The enclave types are biotite schist, biotite gneiss, quartzofeldspathic gneiss, surmicaceous enclaves and rare meta-peridotite.
Diorite, quartz diorite, tonalite and amphibolite dykes Small (,1 km 2 ), medium-to fine-grained locally porphyritic biotite + hornblende quartz diorite bodies intrude the host rocks along the northern contact of the complex and form c. 1 m thick sills NE of Gladstad between Hongset and Vegsteinen (Fig. 2). The quartz diorite is hypidiomorphic granular and contains plagioclase phenocrysts 2-7 mm long; some samples are foliated due to the planar orientation of biotite, plagioclase and amphibole. Tonalitic rocks have magmatic foliation defined by the aligned biotite. Locally, the tonalite contains dioritic enclaves with curved ( Fig. 5a) to cuspate boundaries against the tonalite, a feature suggestive of magma mingling.
Amphibolite dykes are fine-grained and intrude polydeformed metasedimentary host rocks. Dykes contain anhedral hornblende and plagioclase and are commonly folded or disrupted.

Migmatite
Migmatitic rocks crop out over ,2% of the area of the complex. They are exposed in kilometre-to sub-kilometre-scale bodies along intrusive contacts and as dykes that intrude the host rocks on NW Vega and northern Ylvingen (Fig. 2). The migmatitic rocks are divided into contact migmatite or intrusive migmatite on the basis of their contact relationships, texture and mineral assemblage.
Intrusive migmatites crop out in a discontinuous belt along the northern contact on Vega and Ylvingen and as smaller bodies within the central granodiorite on Vega (Fig. 2). These migmatites consist of layered diatexite (Fig. 5b), diatexite ( Fig. 5c) and massive migmatite that lacks melanosomes and leucosomes. Contacts between diatexite and host rocks are sharp and locally truncate gneissic banding in the host rocks. The layered diatexite forms 2-3 km long sills in the host rocks ( Fig. 2) that are parallel to the main intrusive contact and to host rock foliation (Fig. 2). The layered diatexites consist of plagioclase, quartz, K-feldspar and biotite + garnet + cordierite. Melanosomes 2-50 cm thick contain abundant surmicaceous enclaves (Fig. 5b) and fewer quartz biotite gneiss and schist enclaves. Leucocratic migmatite is medium-grained and it occurs in discontinuous zones and layers, containing plagioclase, quartz, K-feldspar and ≤5% biotite. Small 1 -3 m wide magmatic shear zones are present in the migmatites. Shear zones (Fig. 5d) are defined by deflected melanosomes, leucosomes, schlieren and enclaves. Mineral grains in melanosomes, leucosomes, schlieren and enclaves within the shear zone are not recrystallized and indicate the rocks were above their solidus when they were sheared.
Contact migmatite is exposed along the northeastern pluton-host rock contact in a 20 × 500 m strip adjacent to biotite granite (Fig. 2). The contact between this migmatite and non-migmatized host rock is gradational. The migmatite is mediumgrained and comprises quartz, plagioclase, Kfeldspar, biotite, muscovite, garnet and staurolite + kyanite. Leucosome and melanosome layers are 5-30 cm wide and are disrupted, with individual layers less than 40 cm long (Fig. 5e). Enclaves are typically surmicaceous schist.

Granitic dykes
Granitic dykes from 20 cm to 10 m wide are widely distributed in the intrusive complex, where they cross-cut igneous layering and enclaves. Leucogranitic dykes are medium-grained and hypidiomorphic granular, and some contain cordierite. In rare cases, centimetre-scale tourmaline dendrites are present, as are surmicaceous enclaves.
Medium-grained leucocratic quartz syenite dykes are sparse and range in width from a few centimetres to 1 m. They intrude calc-silicate enclaves and calc-silicate host rocks. In some calcsilicate enclaves dykes are folded into open upright folds, either forming centimetre-scale anastomosing bodies or occurring as 30 -40 cm round to elliptical pods in calc-silicate hosts.

Intrusive contacts with host rocks
Intrusive contacts range from discordant to concordant and are commonly sharp. On the northwestern side of Vega, layered diatexite intrudes calc-silicate host rocks along a sharp contact that dips 408NE, subparallel to host rock foliation (Fig. 3d). In this area, a series of granitic dykes intruded subparallel to the pluton-host rock contact (Fig. 2). Approximately 5 km to the east near Gladstad the contact strikes NE, perpendicular to host rock foliation and banding (Fig. 2). The sinuous shape of the contact suggests that granodiorite interfingers with the host rocks, although poor exposure makes it unclear whether the contact is folded or if the map pattern was produced by dyking.
One kilometre east of Gladstad, metre-scale sheets of biotite granite intrude pelitic schist, strike c. N758E and dip moderately to the SE. The resultant sheeted zone is c. 750 m wide. Between Grimsøya and Bø (Fig. 2), the pluton host rock contact strikes east -west and truncates host rock foliation and gneissic banding. On Ylvingen, the contact between metasedimentary rocks and granite is a sinistral shear zone, but was interpreted by Oalmann (2010) to be a deformed intrusive contact subparallel to the steeply east-dipping host rock foliation.

Contacts between intrusive units
Contacts between biotite granite of Ylvingen and enclave-bearing Vega granodiorite are sharp (Fig. 4a) and easily recognized by the decrease in colour index from granodiorite to granite, as well as by the decrease in enclave abundance. On eastern Vega, near station VGWM-67 ( Fig. 2), garnet biotite granite is in contact with Vega granodiorite along a curved to cuspate boundary. A similar curved contact is present near VGWM-113 (Fig. 2). Magmatic foliation and enclave orientation in the granodiorite are parallel to the contact and the enclaves are not truncated at the contact. At station VGWM-99 (Fig. 2), biotite granite and Vega granodiorite are in sharp contact and an enclave of Vega granodiorite c. 20 cm long is incorporated into biotite granite (Fig. 4a).
Contacts between Fuglevaeret granodiorite and Vega granodiorite are subtle and are interpreted to be gradational. Other contact relations among plutonic and migmatitic rocks are not exposed.

Thermobarometry
Contact migmatite (sample VGWM-20-1) and massive diatexite (sample VGWM-34) contain phases appropriate for application of the garnetaluminosilicate-silica-plagioclase (GASP) barometer (Wu & Cheng 2006). The equilibrium reaction is defined: Calculations were made using the software winTWQ v2.3 (Berman 1991(Berman , 2007, which used the thermodynamic database of Berman (1988). Core and rim compositions of garnet were tested and the highest measured anorthite contents of plagioclase were used in the calculations. Activity models for biotite used data from Berman et al. (2007). Models for garnet and cordierite were those from the database of Berman & Aranovich (1996). The activity model for plagioclase was that of Fuhrman & Lindsley (1988). All other phases were assumed to have unit activities unless otherwise stated. Compositions used in calculations are presented in Tables 1 and 2 and analytical  methods are presented with Table 1.
In sample VGWM-20-1, anhedral plagioclase is in contact with kyanite, euhedral to subhedral garnet and quartz, although kyanite, garnet, plagioclase and quartz do not occur in mutual contact in the thin section. The equilibrium boundaries calculated on the basis of a range of garnet core compositions intersect the pelite wet solidus (Spear 1993) between c. 575 and c. 625 MPa (Fig. 6a). Assuming all phases were in equilibrium with melt, and because sillimanite did not crystallize in these rocks, pressures in the kyanite field range from 750 to 875 MPa (Fig. 6a). Calculations in which garnet rim compositions were used resulted in equilibrium boundaries in the sillimanite stability field (Fig. 6a) and suggest disequilibrium between garnet rims and plagioclase cores.
From the massive migmatite sample (VGWM-34), GASP calculations that use garnet and plagioclase core compositions yield equilibrium boundaries that intersect the pelite wet solidus between 450 and 550 MPa in the sillimanite stability field and intersect the first biotite dehydration melting reaction between c. 600 and c. 750 MPa in the sillimanite stability field (Fig. 6b). When garnet rim compositions are used, the calculated equilibrium boundaries intersect the pelite solidus curve between 650 and 750 MPa in the kyanite stability field and intersect the biotite dehydration melting reaction at pressures greater than 925 MPa in the kyanite stability field (Fig. 6b).
Diatexite sample VGWM-19-1 lacks aluminosilicate and biotite but GASP equilibrium boundaries were calculated using the highest measured An content of plagioclase, garnet core compositions and aluminosilicate activities of 0.5 and 0.9. Pressure -temperature (P -T ) conditions of diatexite formation were thus constrained by the GASP equilibrium boundaries and the wet pelite solidus within the range 600-725 MPa, or by the muscovite dehydration melting reaction in the range 725-950 MPa (Fig. 6c).

Cordierite stability
Cordierite thermometers were applied to dendritic cordierite in sample VGWM-4a-5 and prismatic cordierite in sample F5. The Na-in-cordierite thermometer (Mirwald 1986) was used to estimate Na equilibrium between bulk rock compositions and cordierite channel sites in nodular and prismatic cordierite. Average Na abundances from dendritic cordierite in sample VGWM-4a-5 (Table 3) gave temperatures between 675 8C and 713 + 30 8C, whereas prismatic crystals give temperature estimates from 739 8C to 755 + 30 8C. Cordieritegarnet Mg -Fe exchange thermometry (Bhattacharya et al. 1988) for sample VGWM-4a-5 resulted in temperature estimates from c. 634 8C to 715 + 65 8C. This range of temperatures is the result of variable compositions of garnet grains included in cordierite dendrites.  Kretz (1983). Quantitative compositional analyses of plagioclase were carried out by wavelength-dispersive analysis on a JOEL JXA 8900 electron microprobe at the Department of Geology and Geophysics at the University of Wyoming using a nominal accelerating voltage of 15 kV, a beam current of 10 nA and count time of 10 s. Data were reduced using the ZAF correction.
Textural equilibrium between cordierite and garnet allows use of the reaction: to estimate the pressure of equilibrium. Although sillimanite was not observed in the analysed samples, it does occur as a secondary phase replacing cordierite and as an accessory phase in Vega granodiorite samples VGWM-47 and VGWM-59. It is possible that sillimanite was the reactionlimiting phase or that the activity of Al 2 SiO 5 was ,1. Barbey et al. (1999) demonstrated that magmatic cordierite may be stable when Al 2 SiO 5 activity is less than unity. Pressures calculated for equilibrium (2) assuming aluminosilicate activities of 1.0 and 0.5 indicate that dendritic cordierite and garnet from sample VGWM-4a-5 equilibrated at pressures no greater than 500 MPa (Fig. 6d).

Muscovite stability
Muscovite is a common secondary phase in many of the igneous rocks on Vega and replaces Kfeldspar. Euhedral muscovite is present in the samples F2 and F5 from the Fuglevaeret granodiorite. Miller et al. (1981) concluded that euhedral to subhedral muscovite that is not included in other

Timing, emplacement conditions and tilting
Zircon crystallization ages for Vega granodiorite, biotite granite and migmatitic rocks are within error of each other, making it difficult to interpret the relative emplacement timing of these units. Near station VGWM-99 (Fig. 2), an enclave of Vega granodiorite is mingled in biotite granite. Other contacts between Vega granodiorite and biotite granite are curved and suggest that biotite granite and Vega granodiorite co-mingled at the time of their emplacement. Contact relationships between Vega and Fuglevaeret granodiorites are not exposed; however, emplacement of a small intrusion of Fuglevaeret granodiorite within the Vega granodiorite and of a dyke containing cordierite phenocrysts into Vega granodiorite suggest that Fuglevaeret magmas were synmagmatic with or post-dated Vega granodiorite. Contacts between intrusive migmatite and Vega granodiorite appear to be gradational and a body of migmatite was emplaced between Vega granodiorite and biotite granite (Fig. 2), consistent with geochronologic results suggesting that emplacement of some migmatite rocks could have postdated Vega granodiorite and biotite granite rocks. The presence of deformed and undeformed mafic rocks in the aureole suggests that emplacement of mafic magmas both pre-and post-dated emplacement of Vega magmas. The estimates of emplacement pressure for the eastern and northeastern part of the complex reported above, yield values of 450-950 MPa, with most estimates in the range 650-950 MPa. Oalmann et al. (2011) determined peak pressures in Ylvingen metamorphic rocks to be between 700 and 800 MPa. The fact that some migmatites are intrusive and clearly coeval with the remainder of the complex indicates that emplacement of the eastern part of the complex was between c. 700 and 800 MPa. In contrast, cordierite is stable in the western part of the intrusion (Fig. 2), and pressure estimates of cordierite crystallization are c. 500 MPa. These pressure estimates represent distinct conditions of magma emplacement within the complex. The increase in calculated pressure from west to east is therefore interpreted to indicate that the intrusive complex and its host rocks were tilted to the west. Although it is possible that a westwards decrease in emplacement pressure resulted from thrusting of eastern structurally deeper parts of the complex over western shallower parts, no evidence of a fault separating cordierite-bearing and non-cordierite-bearing plutonic rocks was found.
If tilting is assumed and if the change in pressure due to burial was 100 MPa per 3.7 km, then the range of post-emplacement tilting is therefore 14-358, implying structural relief across the complex of 7-16 km.

Geochemical characterization
Major and trace element abundances were measured by inductively coupled plasma atomic emission spectrometry, x-ray fluorescence and inductively coupled plasma mass spectrometry. Analytical methods and representative analyses are presented with Table 4. The range of SiO 2 concentrations in the Vega intrusive complex is 48-77 wt%. Vega and Fuglevaeret granodiorites, Ylvingen granites and migmatitic rocks are mostly alkali-calcic to calcic (Fig. 7a). The majority of the rocks are magnesian in the terminology of Frost et al. (2001) (Fig. 7b). Aluminium saturation indices (ASI ¼ molecular ratio Al/(Ca-1.67P + Na + K) ;Shand 1943;Frost et al. 2001) range from 0.95 to 2.5 and show that rock compositions vary from metaluminous to strongly peraluminous (Fig. 7c). All but two migmatitic rocks are peraluminous. The granitic dykes range from 64 to 77 wt% SiO 2 and are alkali-calcic to calcic. Their aluminium saturation indices range from metaluminous (0.99) to strongly peraluminous (1.49). Diorite, quartz diorite and tonalite have SiO 2 contents of 48 -60 wt%. Some of these rocks lie outside the modified alkali -lime index (MALI) classification fields (,50 wt% SiO 2 ) (Frost et al. 2001) but mafic rocks with SiO 2 .50 wt% are calc-alkalic to calcic (Fig. 7a). Dioritic samples are magnesian and metaluminous. Amphibolitic dykes, diorite rocks and mafic magmatic enclaves that contain ,52 wt% SiO 2 are plotted on the basalt discrimination diagrams of Pearce & Cann (1973). Amphibolitic dykes and mafic enclaves plot in the fields for mid-ocean ridge basalts (MORB), island-arc tholeiites (IAT), calc-alkaline basalts and within plate basalts (Fig. 8a). Diorite, quartz diorite and tonalite samples from the north side of the complex (Fig. 2) plot in the fields of MORB, IAT and calc-alkaline basalts. Figure 8b shows that amphibolite dykes plot outside of the defined fields, but that some mafic rocks and enclaves plot within the MORB and calc-alkaline fields. Figure 8c shows that amphibolite dykes plot within the MORB field and that mafic rocks and enclaves plot in the MORB and calc-alkali fields. Figure 8d shows amphibolite dykes in the normal mid-ocean ridge basalt (N-MORB) field, but two mafic enclaves plot in the IAT field and calc-alkaline basalt field.

Major and trace elements
Amphibolite dyke compositions cluster with respect to Al 2 O 3 , FeO T , CaO and Na 2 O (Fig. 9). Dykes have higher FeO T , CaO and TiO 2 and lower Al 2 O 3 contents than most other rock compositions. Concentrations of Rb, Zr and Ba are clustered but Sr has a relatively wide degree of variation (Fig. 9). Rare earth element (REE) patterns are concave down and have very small negative Eu anomalies (Fig. 10). Concave patterns are similar to those of MORB (Rollinson 1993).
Quartz diorite and tonalite sample compositions plot in broad arrays. The major oxides CaO, TiO 2 , Al 2 O 3 and FeO T are negatively correlated with SiO 2 and Na 2 O is positively correlated with increasing SiO 2 (Fig. 9). Quartz diorite and tonalite trace element concentrations define a broad, scattered array that extends toward amphibolitic dyke compositions in plots of Zr and Ba (Fig. 9), whereas Sr contents cluster between compositions of mafic dykes and Vega, Fuglevaeret and migmatite. The quartz diorite/tonalite sample has a REE pattern with a negative slope and prominent negative europium anomaly (Fig. 10).
The compositions of Vega and Fuglevaeret granodiorites and migmatitic rocks are similar with respect to CaO, TiO 2 , Al 2 O 3 , FeO T and Na 2 O (Fig. 9). Migmatitic rocks from Ylvingen have the lowest silica contents (≤60 wt%). Vega and Fuglevaeret granodiorites and migmatitic rock CaO compositions are almost constant, whereas Al 2 O 3 , FeO T and TiO 2 are negatively correlated with SiO 2 . Na 2 O and K 2 O compositions are clustered within ranges of 1.2-5 wt%. Leucocratic rock compositions cluster, with SiO 2 abundances greater than 70 wt% (Fig. 9).
The abundances of Rb, Zr, Ba and Sr in the Vega and Fuglevaeret granodiorites and the migmatites are generally similar (Fig. 9). Sr decreases with increasing SiO 2 as does Zr in the Vega granodiorite and intrusive migmatites. REE patterns for Vega and Fuglevaeret granodiorite have negative slopes and negative Eu anomalies (Fig. 10). There is little distinction between Fuglevaeret or Vega granodioritic rocks or between cordierite and noncordierite bearing rocks, but Fuglevaeret rocks have the highest REE abundances (Fig. 10). A Vega HClO4, evaporated and re-dissolved in HClO4, evaporated and brought to solution in a mixture of water, HNO3, H2O2 and HF for analyses. Long-term reported precision for REE is better than +5% (RSD) and +10% (RSD) for all other trace elements.
granodiorite and a leuco-granodiorite sample have concave-up patterns and the lowest REE contents.
The leuco-granodiorite sample has a positive Eu anomaly (Fig. 10); the granodiorite has a negative anomaly. Garnet biotite granite and biotite granite compositions are clustered with respect to Al 2 O 3 and CaO, and these clusters overlap compositions of Vega and Fuglevaeret granodiorites and the migmatitic rocks (Fig. 9). The TiO 2 and FeO T concentrations show a linear decrease with increasing SiO 2 (Fig. 9) and, in general, these oxides have lower abundances in the biotite granite compared to the granodioritic units and the migmatites. Zr contents are approximately constant among these rocks (Fig. 9). The only analysed sample of biotite granite from NE Vega (VGWM-35) has as negative REE slope with a negative Eu anomaly and REE abundances slightly lower than those of Vega and Fuglevaeret granodiorites (Fig. 10).
Granitic dykes are depleted, relative to the granodiorites, in Al 2 O 3 , FeO T and TiO 2 and plot in a cluster between Ylvingen leucogranite and migmatitic rocks (Fig. 9). These dykes have variable abundances of the alkalis, with 2-5 wt% Na 2 O and show large variations of Sr and Rb (Fig. 9).

Summary of Sr and Nd isotopes
Bulk rock isotopic ratios, analytical error and analytical methods are presented with Table 5. Initial 87 Sr/ 86 Sr and epsilon Nd (1 Nd ) values are calculated to the approximate age of the complex at 475 Ma. Initial 87 Sr/ 86 Sr ratios analysed by Nordgulen & Sundvoll (1992) are included in Figure 11. Corresponding 1 Nd analyses are not available for these samples, so data are plotted outside of the 1 Nd range presented on Figure 11.
In general, Vega and Fuglevaeret granodiorites have initial 87 Sr/ 86 Sr values of 0.7183-0.7469 and a small range of 1 Nd ratios from 27.5 to 29.0 (Fig. 11). Migmatitic rocks have initial 87 Sr/ 86 Sr ratios that range from 0.7096 to 0.7226 (Fig. 11) and 1 Nd from 27.0 to 211.0 (Fig. 11). Two biotite granites from Vega (samples VG20 and VG24; Fig. 2  from 0.7191 to 0.7401 and 1 Nd of 28.0 to 212.7 (Fig. 11). A disrupted basaltic dyke has an 1 Nd of +9.0 and a basaltic (now biotite hornfels) enclave hosted by Vega granodiorite has an 1 Nd of +8.2 with an initial 87 Sr/ 86 Sr ratio of 0.7227. A diorite enclave hosted in Vega granodiorite has an initial 87 Sr/ 86 Sr value of 0.7220 and an 1 Nd value of 22.9.

Coeval mafic and felsic magmas
The presence of tourmaline, sillimanite, kyanite, cordierite, muscovite and garnet is common in plutonic rocks that originated by partial melting of greywacke and/or pelitic rocks (Zeck 1970;Chappell & White 1974, 1992Wyborn et al. 1981;White & Chappell 1983;Munksgaard 1984;Chappell et al. 1987;Pichavant et al. 1988;Scheepers & Poujol 2002). The mineral assemblage and the diverse suite of enclaves and xenoliths (mafic magmatic enclaves, metasedimentary xenoliths and surmicaceous schist enclaves) are indicative of so-called 'S-type' granite (Chen et al. 1989;Elburg 1996). The abundant inherited zircon Barnes et al. 2007) and strongly peraluminous character of the granitic rocks (Nordgulen 1992;Oalmann et al. 2011) are also characteristic of crust-derived rocks. In contrast, the hornblendebearing diorite, quartz diorite and tonalite intrusive rocks and enclaves have SiO 2 contents, REE patterns, isotopic ratios and trace element compositions that suggest ultimate mantle sources (see below). The largest volume of the complex (.95%) therefore consists of crust-derived granitic and migmatitic rocks, whereas a much smaller volume consists of mafic/intermediate rocks with mantle origins.

Magma differentiation and mafic magma input
Compositional variation of peraluminous magmas may be controlled by crystal fractionation (Dahlquist et al. 2007), assimilation (Clarke et al. 2004), hybridization (Keay et al. 1997), unmixing (Stevens et al. 2007) or some combination of these processes. Variation in major and trace element compositions for Vega and Fuglevaeret granodiorites are scattered and lie along broadly linear  (Wood 1980). trends (Fig. 9). Negative Eu anomalies are consistent with fractional crystallization but could also result from some combination of restite unmixing and partial melting.
The effects of fractional crystallization were evaluated with Ba-Rb and Sr -Rb Rayleigh fraction models using the equation (Rollinson 1993): where C L is the concentration of the element in the liquid, C 0 is the initial concentration of the element in the parent magma, F is the fraction of melt remaining and D is the bulk partition coefficient for the fractionating phases. Low-silica granodiorite samples F5 and VGWM-11 were selected as possible parental magma compositions because they contain less residuum than intrusive migmatite but are compositionally similar to these rocks. Modelled fractionation trends (Fig. 12) for both end-members intersect several granodiorite, intrusive migmatite and granitic dyke rock compositions and these few samples could have differentiated by fractional crystallization. However, the majority of peraluminous to strongly peraluminous rocks do not lie along a clear fractionation trend. Instead, granodiorite and biotite granite rocks exhibit broad trends towards lower Ba concentrations and higher Sr. Such trends may reflect plagioclase accumulation in biotite granite (see discussion of biotite granite below). However, CaO is invariant with respect to SiO 2 (Fig. 9) and does not support plagioclase accumulation in granodiorite and most intrusive migmatite. Samples VCF03.05 and VGWM-8a (Fig. 10) are depleted in REEs relative to most granodioritic rocks but have distinct REE patterns. The positive Eu anomaly in sample VCF03.05 suggests that plagioclase or orthoclase accumulated in this rock and the negative anomaly in sample VGWM-8a is consistent with removal of one or more feldspars. The distinct pattern shapes of these two samples and their low Zr contents are suggestive of control by zircon along with garnet and/or REE-rich accessory phases, such as monazite and apatite. All of these phases can affect REE distribution in peraluminous and strongly peraluminous rocks (e.g. Bea 1996). Detailed modelling of REE behaviour in these samples awaits trace element analysis of garnet and the accessory phases in the samples.
Assimilation of host rocks during magma migration or at the level of emplacement may occur by partial melting or by dissolution (Clarke 2007). The presence of massive quartz enclaves in granodiorite indicates that dissolution of quartz was limited and therefore quartzite enclaves would be preserved. Assimilation of carbonate is limited by low-P saturation of CO 2 , which causes decarbonation reactions to cease (Spera & Bergman 1980). In cases where decarbonation reactions between carbonate rocks and silicate magmas have occurred, the resultant magmas are alkaline and are associated with magmatic skarns (e.g. Barnes et al. 2005). In this regard, it is possible that the quartz syenite rinds adjacent to carbonate screens formed by local carbonate assimilation, as such rocks are lacking elsewhere in the complex. If carbonate assimilation occurred, it therefore affected small volumes of magma in close contact with calcareous xenoliths.
Assimilation of pelitic and semipelitic rocks should leave residual minerals in the magma (Didier 1973) or form new minerals as a result of assimilation reactions (Beard et al. 2005). The lack of partially digested host rock enclaves indicates that host rock assimilation was not significant at the level of emplacement.  Rb, 84 Sr, 149 Sm and 146 Nd. Rb, Sr and the REEs were separated by conventional cation-exchange procedures. Samarium and Nd were further separated in di-ethyl-hexyl orthophosphoric acid columns. All isotopic measurements were made on a VG sector multi-collector mass spectrometer at the University of Wyoming. Uncertainties in Sr isotopic ratio measurements are +0.00002 and Nd isotopic measurements are +0.00001. Blanks contained ,50 pg of Rb, Sr, Nd and Sm and no blank correction was made. Uncertainties in Rb, Sr, Nd and Sm concentrations are +2% of the measured value. Uncertainties on 1 are +0.3.
A migmatitic host rock screen (VGWM-19-1) has a strongly negative 1 Nd (212.7) and an initial 87 Sr/ 86 Sr ratio of 0.7198, below the initial 87 Sr/ 86 Sr ratio for most granodioritic rocks. If low 1 Nd values are common for pelitic and semi-pelitic host rocks and xenoliths derived from them, then assimilation of these rocks should result in isotopic values among the granodioritic rocks that trend towards a 1 Nd of 213 and initial 87 Sr/ 86 Sr values near 0.720 (Fig. 11). The lack of such trends supports field observations that in situ assimilation of pelitic and semi-pelitic host rocks and xenoliths did not strongly modify magma compositions.
It could be argued that the range of initial 87 Sr/ 86 Sr ratios among the granitic and intrusive migmatite resulted from mixing magmas with high 87 Sr/ 86 Sr ratios and low 1 Nd values with magmas with MORB-like 87 Sr/ 86 Sr and 1 Nd values, similar to the amphibolitic dykes and mafic enclaves. However, Figure 11 shows that no rock compositions lie along a mixing trend between two such end-members. This lack of hybrid rock compositions argues against magma mixing between basaltic and peraluminous magmas in the Vega intrusive complex, as does the fact that on bivariant plots of CaO and Al 2 O 3 basaltic and peraluminous rocks do not lie along a linear trend (Fig. 9).  The presence of biotite and garnet with textures indicative of resorption in the Vega and Fuglevaeret granodiorites suggests that some biotite and garnet grains were not stable in the magma and could, along with sillimanite, be residual phases. These minerals and the biotite-rich schist and finegrained gneiss enclaves that lack correlative host rocks may be related to the source of the magma. If so, then these minerals and enclaves were entrained from the source. In ratio-ratio diagrams (not shown), mixing or unmixing will result in trends that plot in hyperbolas such that rock compositions should plot in a fixed order relative to each other on every such diagram. Rock compositions for the Vega and Fuglevaeret granodiorites, garnet biotite granite and migmatite plot in hyperbolic arrays; however, the order of samples in such arrays is not the same from one plot to another, which means that neither simple mixing nor unmixing can explain the compositional arrays. It is possible that magmas in the complex arose from heterogeneous source rocks. Umixing of magmas from heterogeneous source rocks can account for the broadly linear compositional trends (e.g. White & Chappell 1977) for Al 2 O 3 , CaO, TiO 2 and FeO T on bivariant plots (Fig. 9) in the absence of substantial evidence for fractional crystallization and no evidence for hybridization. If so, then multiple source rock compositions would produce multiple simple unmixing relationships that generally follow hyperbolic trends and form complex compositional data arrays.
The restite unmixing model of White & Chappell (1977) assumes that unmelted source rock phases are separated from a fixed melt composition. If unmixing was the primary differentiation mechanism for granodiorite and intrusive migmatite in the intrusive complex, then separation of residuum may have influenced textural development of the complex. For example, the biotite-rich nature of compositional layers in granodiorite (Fig. 4c) and intrusive migmatite (Fig. 5c) may be related to disaggregation of residual surmicaceous enclaves.
Elemental concentrations (Fig. 9) and isotopic ratios (Fig. 11) for granodioritic rocks generally occur over similar ranges. As such, neither the Fuglevaeret nor the Vega granodiorite represent the parent magma to the other. Similarly, it is reasonable that intrusive migmatite magmas were not parental to granodiorite, given that these rocks have initial 87 Sr/ 86 Sr ratios less than 0.725 (Fig. 11). However, granodiorite samples VG17 and VG21 (Fig. 11) have initial 87 Sr/ 86 Sr ratios that are similar to those of intrusive migmatite rocks. These granodioritic samples are proximal to an intensely layered zone in granodiorite and to intrusive migmatite (Fig. 2), and could have been derived from magmas emplaced into these areas.
Field relationships coupled with additional isotopic analyses could be used to identify isotopically distinct zones of granodiorite and intrusive migmatite, if such zones are prominent across the intrusive complex. Possible explanations for the range in initial Sr ratios are discussed below.
Biotite granite compositions lie along a linear trend between diorite and leucogranite (Fig. 9). This linear trend could be interpreted to represent mixing between leucogranite and dioritic magmas, but a fixed compositional order is not maintained on bivariant plots. Instead, Oalmann et al. (2011) suggested that the biotite granite and garnet biotite granite crystallized from residual melts expelled from a peraluminous magma similar to the Vega granodiorite, accompanied by local accumulation of plagioclase and biotite. The separation of residual melt from the Vega granodiorite to form biotite granite is a possible alternative explanation for the variable source model, and could be evaluated with measurement of the initial 87 Sr/ 86 Sr isotope ratios for the biotite granite on Ylvingen.
In summary, geochemical and isotopic data exclude the possibility of hybridization of mafic and felsic magmas and initial Sr isotopic trends and field observations argue against extensive host rock assimilation as a mechanism to control compositional variation in the complex. Trace element Rayleigh fractionation models do not exclude the possibility that some granitic dykes as well as leucogranodioritic and leucocratic migmatitic rocks were differentiated by crystal fractionation, but most granodioritic and migmatitic rocks did not form by this process. Furthermore, leuco-granodioritic and leucocratic migmatitic magmas were differentiated from more than one chemically distinct parental magma. Linear trends on bivariant plots and unmixing models generally support differentiation of granodioritic and migmatitic magmas by the separation of residual phases from a melt, whereas the order of samples supports unmixing from a crustal source rock with a variable composition. The biotite granite and garnet biotite granite may have crystallized from residual melt extacted from the Vega granodiorite, subsequently being modified by plagioclase accumulation.

Explanation of isotopic variation
Mafic and felsic rocks exhibit a wide range of isotopic values (Fig. 11). In this section, explanations that are considered for the isotopic variation of rocks in the Vega intrusive complex include: hydrothermal alteration (e.g. Fourcade et al. 2001); isotopic exchange between mafic enclaves and their surrounding rocks (Barbarin & Didier 1992); mixing and assimilation fractional crystallization (AFC) (e.g. DePaolo 1981); isotopic disequilibrium during partial melting (Barbero et al. 1995); and an isotopically heterogeneous source rock.
Secondary calcite and sericitic alteration of feldspars in felsic rocks and the replacement of plagioclase by zoisite and calcite in mafic rocks are evidence of some degree of hydrothermal alteration. Shaw (1968) suggested that independence of Rb with respect to K 2 O is evidence of the importance of fluid transport within rocks. Rb and K 2 O are strongly correlated in mafic and felsic rocks in the Vega intrusive complex. In addition, the lack of water-rich magmatic features such as aplitic or pegmatitic dykes suggests that deuteric alteration of intrusive migmatite and granodioritic rocks was not intense.
A biotite-rich mafic enclave (VGWM-18-2) has a large positive 1 Nd value and a large initial 87 Sr/ 86 Sr ratio (Fig. 11) that is not easily explained by magma mixing or AFC. Instead, the reaction rind suggests that a chemical exchange occurred between the enclave and granodiorite (e.g. Barbarin & Didier 1992). Stephens et al. (1991) suggested that Sr equilibration between a mafic enclave and host magma was faster than that of Nd and could raise the initial 87 Sr/ 86 Sr in an enclave before strongly lowering 1 Nd .
Diorite enclave VQ17 and mafic samples analysed by Nordgulen & Sundvoll (1992) have initial 87 Sr/ 86 Sr isotopic ratios (Fig. 11) that are larger than those typical of rocks with a mantle source. Mixing models used to evaluate the relationship between amphibolite dykes and peraluminous rocks do not support simple mixing to form the enclave, although isotopic exchange with a crustally derived host magma cannot be excluded as a mechanism that modified the enclave's initial 87 Sr/ 86 Sr ratios. In addition, AFC of crustal rocks by a basaltic parent cannot be ruled out with the current data. Additional isotopic analyses of mafic enclaves and dioritic rocks would provide more conclusive insight into the origin and evolution of mafic rocks and enclaves related to the intrusive complex.
If unmixing is the primary mechanism of differentiation for granodiorite and migmatitic rocks, then disequilibrium melting of an isotopically homogeneous metapelite or metagreywacke source involving various proportions of biotite and plagioclase could produce a trend in which intrusive migmatite and granodiorite share a positive colinear relationship with respect to initial 87 Sr/ 86 Sr ratios and Rb/Sr (e.g. Barbero et al. 1995). However, in the Vega intrusive complex, the intrusive migmatites have a poor negative correlation and granodioritic rocks have a poor positive correlation with respect to initial 87 Sr/ 86 Sr ratios and Rb/Sr (Fig. 13a). We conclude that dehydration melting and melt migration at rates fast enough to cause isotopic disequilibrium between the melt and residual phases and to create migmatite magmas with lower initial 87 Sr/ 86 Sr ratios than granodiorite did not occur. It is possible that among the granodioritic magmas entrainment or mixing of residual biotite or biotite-rich enclaves in disequilibrium with the melt could have caused the positive correlation between initial 87 Sr/ 86 Sr ratios and Rb/Sr. However, initial Sr ratios have a negative correlation with MgO + FeO T (Fig. 13b), which indicates that biotite was in isotopic equilibrium with the melt; entrainment or unmixing of residual biotite therefore cannot explain the isotopic variations.
In contrast, intrusive migmatite samples show a range of initial 87 Sr/ 86 Sr ratios that is positively correlated with MgO + FeO T (Fig. 13b). Residual biotite enriched in radiogenic Sr relative to the melt could have caused a variation in the initial 87 Sr/ 86 Sr ratios in intrusive migmatite either by variable entrainment of residual phases or by unmixing. However, if a source rock with a homogeneous composition was melted to create both granodioritic and migmatitic rocks, then the distinctive low range in initial Sr for migmatitic rocks was not caused by dehydration melting. To create a melt with lower initial 87 Sr/ 86 Sr than the crustal source, feldspars are expected to contribute to the melt (Davies & Tommasini 2000). Melting under water-saturated conditions that involve the breakdown of plagioclase and quartz (Conrad et al. 1988) could form a melt in isotopic disequilibrium with residual phases. Intrusive migmatite on NW Vega lacks residual sillimanite, consistent with partial melting under 'wet' conditions, but intrusive migmatite on Ylvingen (VGWM-34) contains sillimanite after kyanite, consistent with the breakdown of muscovite during dehydration melting (Oalmann et al. 2011). This observation indicates that the melting conditions were not solely responsible for the lower range of initial 87 Sr/ 86 Sr ratios in intrusive migmatite.
As a final alternative, the range in isotopic ratios may be related to isotopic and elemental heterogeneity of the crustal source rock. Isotopic diversity in Vega granodiorite and intrusive migmatite is consistent with the range in initial 87 Sr/ 86 Sr ratios for Horta nappe migmatite (see following section) (Fig. 11) as well as with other plutonic rocks in the HNC. For example, tourmaline granite on the Holm peninsula in the Lower Nappe has initial 87 Sr/ 86 Sr ratios within the range 0.7150-0.7312 (Nordgulen & Sundvoll 1992). In the case of intrusive migmatite rocks with initial 87 Sr/ 86 Sr ratios less than c. 0.725, source rock isotopic heterogeneity cannot be excluded as a cause of the isotopic range and is testable with a petrologic study of the source rocks.
Likely crustal source rocks and crustal melting Barnes et al. (2007) suggested that the Vega intrusive complex was emplaced in the Horta nappe. The Horta nappe is overturned and consists of an upper section of Neoproterozoic-age rocks and a lower section of Palaeozoic-age rocks (Barnes et al. 2007) (Fig. 14). U -Pb dating of zircon in the Palaeozoic-age supracrustal rocks showed that they experienced regional migmatization at c. 477 Ma (Barnes et al. 2007), within error of the emplacement age of the Vega intrusive complex. Moreover, the Horta migmatites and the Vega intrusive complex have similar Nd and Sr isotopic ratios (Fig. 11) and elemental concentrations (Fig. 9), and the inherited zircon populations in Vega plutonic rocks are similar to the detrital zircon population in the Palaeozoic-age rocks of the Horta nappe (Barnes et al. 2007). In addition, the Horta migmatites contain gneissic enclaves identical to those in the Vega granodiorite, but such gneissic rocks are absent from the Vega host rocks (Marko 2012). The migmatitic Ordovician rocks of the Horta nappe are therefore the most likely source rocks to the Vega intrusive complex (Fig. 14). The rocks of the Vega intrusive complex are largely undeformed despite the polyphase deformation of the host rocks. The folded and boudined dykes in the host rocks indicate the Vega intrusive complex was emplaced toward the end of the D3 deformation (Anderson et al. 2005). The presence of kyanite, staurolite and garnet in host rocks outside the aureole to the Vega intrusive complex indicate that these rocks followed a clockwise Barrovian P-T path during crustal thickening (see Oalmann et al. 2011 for details).
Crustal thickening such as that which occurred in the HNC is commonly associated with crustal melting. In general, the rates of crustal thickening outpace conductive heating and result in the depression of the regional geotherm. Melting can occur after the geotherm relaxes, the deep-seated rocks are heated and exhumation raises these rocks through their solidi. In the most efficient case, thermal relaxation allows the geotherm to intersect the wet solidus for crustal rocks in c. 30 myr and muscovite dehydration melting begins after c. 60 myr (England & Thompson 1986). Given the geochronological uncertainties, the maximum time span between the end of sediment deposition in the Middle Nappe at c. 481 Ma (Barnes et al. 2007) and crystallization of zircon in the Vega intrusive complex at c. 476 Ma is c. 14 myr, suggesting that anatexis was unlikely to have occurred in response to crustal thickening alone.
Mafic magmatism must have continued during formation of granodioritic magmas because MORB-like basaltic enclaves are mingled with peraluminous granodiorite. Intra-plating of such mafic magmas into a thickened package of back-arc metasedimentary rocks was suggested as a mechanism of forming 'S-type' magmas in the Lachlan Fold Belt (Collins & Richards 2008). Intra-/underplating of hot basaltic magmas in the lower crust or mantle lithosphere (Annen & Sparks 2002;Dufek & Bergantz 2005;Annen et al. 2006) provides a significantly more efficient means of crustal melting in time frames of 10 5 to 10 6 years, compatible with the sequence of crustal thickening and partial melting in the HNC. In the Lachlan Fold Belt example, the S-type magmas were interpreted as hybrids between basaltic magmas and crustal melts/mushes (Collins 1996(Collins , 1998. Hybridization in the Lachlan Fold Belt required multiple steps (Healy et al. 2004), partly due to the unlikely scenario of small amounts of high-T basalt and large amounts of low-T granitic magmas successfully mixing. The Vega and Fuglevaeret granodiorites and associated migmatite rocks provide an example in which the chemical evidence for hybridization with mafic magmas is lacking (Fig. 14), despite the presence of MORB-like mafic enclaves and synplutonic dykes.
intermediate magmas with broadly arc-like characteristics (Nordgulen & Sundvoll 1992;Birkeland et al. 1993;Barnes et al. 2002) have isotopic signatures that indicate a mantle contribution, but lack the MORB-like signature observed in the Vega complex. Such a trend could be accomplished during slab roll-back with intra-/under-plating of MORB-like magmas into the crust during westdirected subduction (e.g. Yoshinobu et al. 2002;McArthur 2007).

Conclusions
The Vega intrusive complex was emplaced into polydeformed metasedimentary rocks of the Horta nappe at pressures of 350-800 MPa. The intrusive complex and its host rocks were tilted 14 -358 to the west and subsequent erosion has exposed a 7-17 km thick crustal section, providing a view of the compositional variation, textural differences and relative enclave abundances in the intrusive complex as a function of depth. The Vega granodiorite, Fuglevaeret granodiorite and intrusive migmatite units formed separately by partial melting of a geochemically and isotopically heterogeneous crustal source. Differentiation of individual batches of granodioritic and migmatitic magmas occurred by unmixing of melt from residual phases and enclaves and created magmatic layering in some granodioritic and migmatitic rocks.
Magmatism in the Vega intrusive complex was preceded by regional contractional deformation that involved juxtaposition of multiple nappes, some of which were floored by ophiolitic crust. This tectonic burial resulted in metasedimentary host rocks following a clockwise, Barrovian metamorphic path in the garnet-staurolite-kyanite stability field. In the deeper regions of the HNC (at pressures .800 MPa), metasedimentary rocks were heated not only by burial but also by intra-/ under-plating of MORB-like basaltic magmas which resulted in partial melting to form S-type enclave-rich granodioritic magmas and biotite granite magmas. Our present interpretation is that the MORB-like magmas formed due to extension related to slab roll-back in a broadly back-arc setting. Future research on the Vega intrusive complex will investigate relationships among magmatic structures, enclave origins and magma transfer.