Chemical Signatures of Melt–Rock Interaction in the Root of a Magmatic Arc

Chemical Signatures of Melt–Rock Interaction in the Root of a Magmatic Arc ABSTRACT Identification of melt–rock interaction during melt flux through crustal rocks is limited to field relationships and microstructural evidence, with little consideration given to characterising the geochemical signatures of this process. We examine the mineral and whole-rock geochemistry of four distinct styles of melt–rock interaction during melt flux through the Pembroke Granulite, a gabbroic gneiss from the Fiordland magmatic arc root, New Zealand. Spatial distribution, time-integrated flux of melt and stress field vary between each melt flux style. Whole-rock metasomatism is not detected in three of the four melt flux styles. The mineral assemblage and major element mineral composition in modified rocks are dictated by inferred P–T conditions, as in sub-solidus metamorphic systems, and time-integrated volumes of melt flux. Heterogeneous mineral major and trace element compositions are linked to low time-integrated volumes of melt flux, which inhibits widespread modification and equilibration. Amphibole and clinozoisite in modified rocks have igneous-like REE patterns, formed by growth and/or recrystallisation in the presence of melt and large equilibration volumes provided by the grain boundary network of melt. Heterogeneities in mineral REE compositions are linked to localisation of melt flux by deformation and resulting smaller equilibration volumes and/or variation in the composition of the fluxing melt. When combined with microstructural evidence for the former presence of melt, the presence of igneous-like mineral REE chemical signatures in a metamorphic rock are proposed as powerful indicators of melt–rock interaction during melt flux. INTRODUCTION Partial melting, melt migration and crystallisation are fundamental processes responsible for the chemical evolution of the Earth’s crust. Each process imparts a different geochemical signature on the melt and surrounding rock with which it reacts. Despite demonstrations that melt–rock interaction during melt migration produces distinctive chemical signatures in the mantle (e.g. Nicolas, 1986; Spiegelman & Elliott, 1993; Rampone et al., 1994; Lundstrom et al., 1995; Pirard & Hermann, 2015), studies of melt migration through mafic to intermediate lower crust are limited. Most studies of mafic granulites and mafic migmatites focus on the geochemical signature of sites of melt production and/or injection and crystallisation (e.g. Tait & Harley, 1988; Collins & Sawyer, 1996; Sawyer et al., 1999), or at the axes of spreading oceanic ridges (e.g. Korenaga & Kelemen, 1997; Saal & Van Orman, 2004; Lissenberg & Dick, 2008; Lissenberg et al., 2013). In contrast, little consideration is given to geochemical signatures produced during the migration of melt through the lower continental crust. Melt–rock interaction in crustal rocks is rarely documented because of the inherent complexity of crustal rock types and melt compositions, and a lack of criteria for identification of the former flux of melt. Evidence for melt–rock interaction in the field may take the form of distributed or narrow zones of modification, such as hydration (e.g. Daczko et al., 2002b; Stuart, 2017; Stuart et al., 2017), or bodies of rock with igneous features that lack intrusive or structural boundaries with adjacent metamorphic rocks (e.g. Daczko et al., 2016). The most robust microstructural evidence is observation of microstructures indicative of the former presence of melt (e.g. small dihedral angles, pseudomorphs of grain boundary melt films, pseudomorphs of melt pockets at triple junctions, see Sawyer, 1999; White et al., 2005; Holness, 2008; Vernon, 2011) in rocks where in situ partial melting reactions are not interpreted. Microstructures indicative of a replacement reaction are also common in the deep crust, especially in rocks partially modified by flux of an externally-derived melt (Stuart et al., 2016, 2017). Though physical indicators of the former presence of melt are identified via comparison with igneous rocks and migmatites, little attention has been given to establishing geochemical signatures indicative of melt–rock interaction. Geochemical signatures in exposed arc roots and plutons are widely used to infer metamorphic and igneous processes occurring at depth. For example, studies have shown that metamorphic reactions in the lower crust do not necessarily mobilise rare earth elements (REE) over large distances, due to slow diffusion rates (e.g. Chernoff & Carlson, 1999; Van Orman et al., 2001; Cherniak, 2003; Tirone et al., 2005) and weak partitioning into aqueous fluids (Cullers et al., 1973; Flynn & Burnham, 1978; Wendlandt & Harrison, 1979; Cullers & Graf, 1984). As a result, metamorphic products in sub-solidus or fluid-poor conditions (e.g. Daczko et al., 2009) may inherit their chemical signatures from reactant minerals (e.g. Schröter et al., 2004; El Korh et al., 2009; Clarke et al., 2013; Chapman et al., 2015). In contrast, the REE are mobilised and redistributed by processes involving melt (Plank & Langmuir, 1993) and REE data are commonly used to determine evolving mineral assemblages and P–T conditions in the arc root (e.g. Dalpé & Baker, 2000; Tulloch & Kimbrough, 2003; Bignold et al., 2006) and magma differentiation processes (e.g. Stevenson et al., 2005; Davidson et al., 2007; Allibone et al., 2009; Smith, 2014; Chapman et al., 2016; Cooper et al., 2016). In principle, REE ratios and patterns in rocks and minerals may be informative of processes involving melt. However, melt–rock interaction straddles the boundary between igneous and metamorphic processes, involving melt but ultimately producing a metamorphic rock. The presence of melt, as opposed to an aqueous fluid, is expected to increase chemical communication and equilibration volumes. Therefore, melt presence may produce igneous mineral chemical signatures in the modified rock, though the effects of protolith composition, melt volume, and prevailing P–T conditions during melt–rock interaction have not been extensively studied. Four styles of melt flux throughout a single, homogeneous host rock, the Pembroke Granulite in Fiordland, southern New Zealand are examined in this contribution. Throughout this work, the term ‘flux’ is used to describe the open-system passage of melt through, rather than injection into, the Pembroke Granulite. Previous studies examining melt flux have focused on field relationships and microstructures, demonstrating that melt–rock interaction involved little crystallisation of melt within the modified rocks, and that mineral assemblages and microstructures produced during melt–rock interaction are common in lower crustal rocks (Daczko et al., 2016; Stuart et al., 2016, 2017; Stuart, 2017). We analysed bulk-rock composition and mineral major and trace element compositions to investigate the link between chemical signatures, melt flux style and P–T conditions of melt–rock interaction in the lower crust. Our findings indicate that the stable assemblage and major element composition of minerals are defined by variable equilibration with the fluxing melt. In contrast, mineral REE patterns are characterised by signatures common in igneous rocks. We relate the strength of these igneous signatures and relative homogeneity to the spatial extent and connectivity of the melt network throughout the rock. GEOLOGICAL SETTING AND PREVIOUS WORK Geological setting The Pembroke Granulite, Fiordland, New Zealand (Fig. 1), is a low-strain component of the Median Batholith, a suite of Carboniferous to Early Cretaceous plutons emplaced into, and partly comprising, the lower crust of a Cordilleran magmatic arc (Blattner, 1991; Mortimer et al., 1999). Emplaced at 139–129 Ma (Hollis et al., 2003), the gabbroic protolith to the Pembroke Granulite had an igneous assemblage of enstatite, diopside, brown-green pargasite, plagioclase, and ilmenite. The whole-rock composition of the two-pyroxene–pargasite pluton varies by several weight percent for all major element oxides (Stuart et al., 2016). The Pembroke Granulite, aside from some slight grain size variations, is otherwise homogeneous. Igneous minerals were variably recrystallised during D1 to form a gneissic foliation (S1) that strikes NE and dips steeply to the north and south. Deformation was recorded by the development of undulose extinction, deformation twins, and minor subgrain formation in igneous plagioclase, amphibole, and pyroxene grains, forming a rock distinctly metamorphic in character (Stuart et al., 2016). Similar assemblages in the nearby Western Fiordland Orthogneiss formed at lower-crustal conditions of ∼850°C and <11 kbar (Daczko & Halpin, 2009). Fig. 1. View largeDownload slide Geological map of Northern Fiordland, New Zealand. Location of study area marked by black arrow. Inset: Geological provinces of South Island, New Zealand, divided by the Alpine Fault plate boundary. Note the plate boundary ranges along strike from subduction to transpressional with local extension (Daczko et al., 2003). Fig. 1. View largeDownload slide Geological map of Northern Fiordland, New Zealand. Location of study area marked by black arrow. Inset: Geological provinces of South Island, New Zealand, divided by the Alpine Fault plate boundary. Note the plate boundary ranges along strike from subduction to transpressional with local extension (Daczko et al., 2003). Emplacement and D1 preceded a major pulse of high-Sr/Y magmatism in the arc system (126–115 Ma; Hollis et al., 2003, 2004; Tulloch & Kimbrough, 2003; Allibone et al., 2009; Milan et al., 2016) representing an arc flare up (Milan et al., 2017), and as such the Pembroke Granulite represents part of the lower crust through which the high-Sr/Y melts migrated (Daczko et al., 2016; Stuart et al., 2016, 2017). Summary of styles of melt flux Melt–rock interaction, as the high-Sr/Y melts fluxed through the Pembroke Granulite, resulted in four different styles of modification of the deformed, two-pyroxene–pargasite gneiss protolith (Fig. 2). Key changes common to all melt flux styles involve hydration and an increase in the mode of amphibole (Fig. 3). The first melt flux style involved widespread growth of pargasite-bearing coronae around igneous and S1 pyroxene throughout the entire Pembroke Granulite (Style 1 [En + Di + Prg +Qtz + Pl + Czo + Rt ± Ap]; Stuart et al., 2016). Each of the later styles of melt-rock interaction (Styles 2–4) formed distinct minor rock types hosted within the Pembroke Granulite, including tschermakite–clinozoisite gneiss and migmatite (Style 2 [Ts + Grt + Czo + Pl + Ms + Qtz + Rt]; sometimes called dioritic gneiss in previous literature; Daczko et al., 2001a, 2002a; Stuart et al., 2017), melt-bearing high-grade shear zones (Style 3 [Mhb + Grt + Czo + Pl + Ms + Qtz + Rt]; called D3 and D4 shear zones in previous literature; Daczko et al., 2001b; Gardner et al., 2016; Stuart, 2017), and hornblendite (Style 4 [Prg ± Czo ± Pl ± Bt ± Grt ± Ru]; Meek, 2015; Daczko et al., 2016). Previous work suggests that each melt flux style is distinct in terms of the volume of melt in the rock at any one time, the scale of melt-fluxed rock, the degree of modification of the protolith, the interpreted time-integrated melt flux and whether the flux was deformation assisted or occurred under static conditions. The key characteristics of the four styles of melt–rock interaction are summarised below. Mineral abbreviations used throughout the manuscript follow the scheme proposed by Whitney & Evans (2010). Fig. 2. View largeDownload slide Summary of key characteristics of styles of melt–rock interaction from Daczko et al. (2001b), Daczko et al. (2016), Meek (2015), Stuart (2017), Stuart et al. (2017) and Stuart et al. (2016), including sketches of melt flux along grain boundaries and relative spatial extent of melt flux. P–T conditions for Style 1 are from Stuart et al. (2016), and for Style 3 from Daczko et al. (2001b). Fig. 2. View largeDownload slide Summary of key characteristics of styles of melt–rock interaction from Daczko et al. (2001b), Daczko et al. (2016), Meek (2015), Stuart (2017), Stuart et al. (2017) and Stuart et al. (2016), including sketches of melt flux along grain boundaries and relative spatial extent of melt flux. P–T conditions for Style 1 are from Stuart et al. (2016), and for Style 3 from Daczko et al. (2001b). Fig. 3. View largeDownload slide Styles of melt–rock interaction, showing from left to right: typical field relationships including insets of a schematic cross section approximately 100 m long of the Pembroke Granulite showing the relative scale of each melt flux style, outcrop appearance and microstructures of modified rock types (scale bars 1000 µm), including insets of BSE images of microstructures indicative of the former presence of melt (scale bars 50 µm). Mineral abbreviations follow the scheme proposed by Whitney & Evans (2010). (a–c): Style 1; (d–f): Style 2, (g–i): Style 3, (j–l): Style 4. Fig. 3. View largeDownload slide Styles of melt–rock interaction, showing from left to right: typical field relationships including insets of a schematic cross section approximately 100 m long of the Pembroke Granulite showing the relative scale of each melt flux style, outcrop appearance and microstructures of modified rock types (scale bars 1000 µm), including insets of BSE images of microstructures indicative of the former presence of melt (scale bars 50 µm). Mineral abbreviations follow the scheme proposed by Whitney & Evans (2010). (a–c): Style 1; (d–f): Style 2, (g–i): Style 3, (j–l): Style 4. Style 1: Diffuse porous melt flow The earliest style of melt–rock interaction occurred post-D1 and involved partial modification along grain boundaries throughout the bulk of the Pembroke Granulite (Figs 2, 3a;Stuart et al., 2016). Initially, small local volumes (< 5%) of in situ partial melting are inferred to have created a permeable grain boundary network, which was utilised as migration pathways by an externally derived, high-Sr/Y melt. Melt–rock interaction involved partial hydration and replacement of the two-pyroxene–pargasite gneiss (Fig. 3b), forming coronae of blue-green pargasite and quartz (Fig. 3c) that are intimately associated with clinozoisite, Sr enrichment in plagioclase and microstructures indicative of the former presence of melt (pseudomorphs of melt including small dihedral angles, grain boundary films, and aggregates of quartz, plagioclase, and K-feldspar; Fig. 3c, inset; Stuart et al., 2016). Partial hydration of the S1 assemblage is observed, to a variable extent, throughout the entirety of the Pembroke Granulite. Style 1 involves a relatively small modification of the protolith under static conditions, distributed on a large scale, with an interpreted small, time-integrated melt flux at an outcrop scale. Partially hydrated S1 in the two-pyroxene–pargasite gneiss (Style 1 melt flux) is cut by a series of sub-vertical anorthositic D2 dykes, around which the assemblage is dehydrated and transformed to garnet granulite at a scale of cm-dm, called garnet reaction zones (GRZ); these are well-documented in the literature (Blattner, 1976; Clarke et al., 2000, 2005; Daczko et al., 2001a; Schröter et al., 2004; Daczko & Halpin, 2009; Smith et al., 2015). All subsequent styles of melt flux (Styles 2, 3 and 4) overprint both the partially hydrated S1 (Style 1) as well as GRZ. Style 2: Channelised porous melt flow Style 2 melt flux is similar in nature to Style 1, except that the melt flux is concentrated into narrow (<20 m wide) channels (Figs 2, 3d;Stuart et al., 2017). Within the channels, melt–rock interaction formed microstructures overprinting S1, where intermediate modification forms tschermakite–clinozoisite gneiss, and high modification forms migmatite (Fig. 3e). Modified rocks have hydrous assemblages dominated by tschermakite, clinozoisite and plagioclase (Fig. 3f). Migmatites exhibit significant recrystallisation and contain coarse-grained garnet (Fig. 3e and f) previously interpreted as a peritectic phase, indicating in situ partial melting (Daczko et al., 2001a; Stuart et al., 2017). The limited volumes of in situ melt produced are consistent with fluid-undersaturated conditions during partial melting (Stuart et al., 2017). The modification of the protolith associated with Style 2 melt flux is more extensive relative to Style 1, indicating either prolonged melt flux and/or higher time-integrated volumes of melt moving through the channels. Style 2, therefore, involves intermediate modification of the protolith, pervasive within small channels under static conditions, with a relatively intermediate time-integrated melt flux at an outcrop scale. Style 3: Deformation assisted porous melt flow Style 3 involves flux of melt along grain boundaries, where the flux is spatially restricted within D3 shear zones (Figs 2, 3g; these include the D3 and D4 shear zones of Daczko et al. (2001b)). Two-pyroxene–pargasite gneiss, tschermakite–clinozoisite gneiss, and migmatite are all deformed into shear zones (Fig. 3g and h) where they are modified to assemblages dominated by magnesio-hornblende, clinozoisite and plagioclase. Importantly, small amounts of syn-tectonic, medium-grained garnet (Fig. 3h) are observed (Daczko et al., 2001b), interpreted as a peritectic phase suggesting melt present conditions (Stuart, 2017). Ongoing work reveals that microstructures indicative of the former presence of melt are ubiquitous throughout shear zones (pseudomorphs of melt including: small dihedral angles, grain boundary films, grains of plagioclase interconnected in 3 D, multiphase aggregates at triple-point-junctions, and felsic dykelets; Fig. 3i, inset; Stuart, 2017), consistent with published P–T conditions for deformation which lie above the solidus (Fig. 2; Daczko et al., 2001b). Studies show that melt flux is more efficient in dynamic (deformation assisted) versus static conditions (e.g. Davidson et al., 1994; Daines & Kohlstedt, 1997; Rosenberg & Handy, 2000). Therefore, volumes of melt moving through the shear zones are expected to be higher than those in Styles 1 and 2. This is consistent with the complete hydration and overprinting of the S1 assemblage within the shear zones (Fig. 3i). Style 3 is, therefore, interpreted to involve intermediate modification of the deforming protolith, concentrated within zones of active shearing, with a relatively intermediate to high time-integrated melt flux at the outcrop scale. Style 4: Reactive infiltration, followed by melt flux through an armoured zone Style 4 involves the deformation-assisted reactive infiltration of melt within a 30–40 m wide channel forming a hornblendite (Figs 2, 3j; Daczko et al., 2016). Melt-rock interaction involved dissolution of plagioclase and pyroxene, and the precipitation of pargasite with, or without, clinozoisite (Fig. 3k), resulting in complete modification of the protolith. Local bands of garnet-rich rock are present (Daczko et al., 2016). The complete dissolution of plagioclase–pyroxene and growth of pargasite ± clinozoisite resulted in the formation of an armoured channel that is unreactive with the fluxing melt. Assisted by deformation, large volumes of melt may have fluxed efficiently through the unreactive channel once it became armoured. Former melt is now pseudomorphed by minor plagioclase (Fig. 3l, inset). The Style 4 hornblendite cuts rock types formed during Styles 1, 2 and 3. Style 4 is, therefore, interpreted to involve complete modification of the deforming protolith, concentrated within an armoured channel, with a relatively high time-integrated melt flux at an outcrop scale. SAMPLE SELECTION Data presented in this study are from representative rock types of all melt flux styles. Data for Styles 1 and 2 originate from samples analysed by Stuart et al. (2016) and Stuart et al. (2017). Raw data from these studies were collated with new data for Styles 3 and 4. Where possible, samples were collected directly from outcrops, however in some cases rock types were sampled from float in creeks after careful consideration of field relationships, characteristic mineralogy and textures compared to the outcrops. METHODS Whole-rock composition Concentrations of major element oxides SiO2, TiO2, Al2O3, Fe2O3, MgO, MnO, CaO, Na2O, K2O, P2O5 and SO3 and selected trace elements, including Sr and Y, were determined for representative samples of intermediate and high modification from Styles 3 and 4 using a PANalytical PW2400 Sequential WDXRF Spectrometer using WROXI standards (Mark Wainwright Analytical Centre at the University of New South Wales, Sydney, Australia). Trace and rare earth elements were determined for selected samples. Powdered samples were digested in sealed 15 mL Savillex Teflon beakers using a 1: 1 mixture of concentrated HF and concentrated HNO3 at 120°C for 24 hours, then dried down and repeated. Samples were then digested in 6 HNO3 for 24 hours, then dried down and diluted to 10 mL using 2% HNO3 and trace HF. 1: 1000 dilutions of each sample were then individually spiked with a 15 μL aliquot of a solution of 6Li, As, Rh, In, Tm and Bi in 2% HNO3. Samples were analysed on an Agilent 7500 series inductively coupled plasma mass spectrometer (ICP-MS) using BCR-2 as a calibration standard, the spike to correct for instrument drift, and a 2% HNO3 solution to measure background. Calibration was checked using the standards BIR-1 and BHVO-2 (Macquarie University GeoAnalytical, Sydney, Australia). Mineral major element compositions Polished thin (30 µm) and thick (100 µm) sections were made from blocks cut from representative samples; in the case of Style 3 samples blocks were cut perpendicular to foliation and parallel to lineation. A petrographic microscope was used in combination with the Virtual Petrographic Microscope (Tetley & Daczko, 2014), ImageJ (Rasband, 1997–2015), and backscattered electron (BSE) images for mineral identification. BSE images were collected using thin and thick sections coated with 10 nm of carbon in a Carl Zeiss IVO scanning electron microscope (SEM; Macquarie University GeoAnalytical, Sydney, Australia). The SEM was run with an accelerating voltage of 15–20 kV, a beam current of 10 nA and a working distance 10–12·5 mm. Major element compositions of minerals were determined using a CAMECA SX100 electron microprobe (S1, Styles 1, 2, and 4; Macquarie University GeoAnalytical, Sydney, Australia) and a JEOL JXA8100 electron microprobe (Style 3; the Institute of Geology and Geophysics, Chinese Academy of Sciences (IGGCAS), Beijing, China). Electron microprobes were run with accelerating voltages of 15 kV and beam currents of 20 nA, analysing 1–5 µm spot sizes for 3·5–5 minutes per spot. Weight percent oxide data were recalculated into cations per formula unit for individual minerals, using 23 oxygen for amphibole, 12 oxygen for garnet, 12·5 oxygen for clinozoisite and 8 oxygen for feldspar. Amphiboles were classified according to the scheme proposed by Hawthorne et al. (2012) using the spreadsheet provided by Locock (2014). The Fe3+ content in garnet was calculated by converting a portion of the FeO to Fe2O3 so that the cations per formula unit sum to 8. Garnet end-member proportions were then calculated using the following: Alm = 100 * Fe2+ / (Fe2+ + Mg + Ca + Mn); Pyr = 100 * Mg / (Fe2+ + Mg + Ca + Mn); Grs = 100 * Ca / (Fe2+ + Mg + Ca + Mn); Sps = 100 * Mn / (Fe2+ + Mg + Ca + Mn). Plagioclase end-member proportions were calculated using the following: Ab = 100 * Na / (Na + Ca + K); An = 100 * Ca / (Na + Ca + K); Or = 100 * K / (Na + Ca + K). Pistacite content in clinozoisite was calculated using: Ps = 100 * Fe3+ / (Fe3+ + Al3+). Mineral trace element compositions Trace element distribution mapping of representative samples of high modification from Styles 1, 2, and 3 was performed on the X-Ray Fluorescence Microscopy beamline, using the Maia-384 detector on the Kirkpatrick-Baez mirror microprobe at the Australian Synchrotron, Melbourne (Ryan et al., 2010b; Paterson et al., 2011). Polished thin sections mounted on glass slides were used for analysis; no additional preparation was performed. Maps were made by scanning thin sections in 4 µm step sizes in the x and y directions at a speed of 4·096 mm/s and dwell time of 0·98 ms/pixel. Spot sizes of 4 µm2 were analysed using a beam energy of 18·5 keV. The Maia detector has a minimum energy sensitivity of 3·3 keV, allowing detection of elements down to Z = 20 (Ca). Standard foils (Pt, Mn, Fe, YF3) were periodically analysed for calibration. Real time processing, using the Dynamic Analysis (DA) method (Ryan, 2000; Ryan et al., 2010a), deconvolves each individual X-ray event into element signals, allowing rapid data collection, high count rates and high sensitivity. Data reduction was performed using GeoPIXE (Ryan et al., 1990), which deconvolves the spectra using the fundamental parameter model for the layered sample, the Maia detector efficiency model, and the DA matrix method. A pure plagioclase matrix, using concentrations of values from analysed features in the samples and mineral densities, was used for matrix correction. Maps were constructed for each sample to highlight relative zoning of Sr in plagioclase. Mineral rare earth element composition Rare earth element compositions of minerals were determined using laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS; Macquarie University GeoAnalytical, Sydney, Australia). A Photon Machines Excite excimer 193 nm laser ablation microprobe system was used to ablate 25–50 µm spots from thin sections at a frequency of 5 Hz. Ablated material was transported in helium gas to the plasma at a flow rate of 0·8 Lmin-1. Material was analysed in an Agilent 7700x ICP-MS, using Ar as a carrier gas with a flow rate of 1·0 Lmin-1. Calcium (measured by EMP) was used as an internal standard for all minerals, and NIST 610 glass, basalt from the Columbia River (BCR-2), and MONGOL garnet were used as external standards. Raw signal data were reduced using GLITTER software (Griffin et al., 2008). Rare earth element concentrations obtained from GLITTER were normalised to chondrite values of McDonough & Sun (1995). WHOLE-ROCK COMPOSITION OF THE PEMBROKE GRANULITE Rocks fluxed by Styles 1, 2, and 3 are monzodioritic or monzogabbroic in composition (Table 1; Fig. 4a); some Style 3 samples fall within the gabbroic field on the (TAS) silica versus total alkalis diagram of Middlemost (1994). Style 4 samples have compositions distinct from the other styles, plotting in the foid-gabbroic field on the TAS diagram, with lower SiO2 (40·29–42·25 wt %; Table 1; Fig. 4a). All samples are peraluminous, with Aluminium saturation indices of ∼1·36 for Styles 1, 2, and 3, and ∼1·09 for Style 4, and Peacock indexes of ∼0·24 for Styles 1, 2, and 3, and ∼0·50 for Style 4. All samples have low TiO2 (0·42–1·99 wt %), high Al2O3 (16·74–20·44 wt %; Fig. 4b), and high alkali contents (Na2O + K2O = 2·55–6·18 wt %; Fig. 4a). Table 1: Representative bulk compositions of the protolith (S1) and rock types formed in each melt flux style Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1222 1320 1449 1459 1336B 1443A 1518 1514 Major elements (wt % oxide; XRF)  SiO2 52·39 53·42 51·66 51·43 51·21 51·61 40·41 41·62  TiO2 1·12 0·95 0·61 0·79 0·85 0·70 1·58 1·41  Al2O3 19·17 19·13 19·79 19·90 19·45 20·15 17·75 15·57  FeO(tot) 8·20 7·57 7·13 6·87 7·94 7·45 12·06 13·81  MnO 0·15 0·14 0·00 0·00 0·19 0·15 0·09 0·14  MgO 4·21 3·88 4·93 4·71 4·38 4·73 8·43 10·48  CaO 8·49 7·70 9·05 9·29 8·43 8·77 13·13 10·06  Na2O 5·05 5·21 4·88 4·50 4·46 5·06 2·69 3·53  K2O 0·49 0·64 0·42 0·46 0·41 0·41 0·53 0·68  P2O5 0·22 0·31 0·09 0·04 0·06 0·09 0·23 0·01  S <0·01 <0·01 <0·01 <0·01 <0·01 <0·01 <0·01  LOI 0·60 0·01 1·24 0·99 0·85 0·84 1·39 1·37  Total 101·00 99·80 100·60 99·74 99·13 100·81 99·63 100·21 Trace elements (ppm; XRF)  Sr 796·4 813·0 898·4 913·6 760·4 793·9 494·8 105·9  Y 8·00 16·80 5·80 5·50 3·80 4·20 18·20 5·30 REE data (chondrite normalised using values of McDonough & Sun (1995); solution ICP-MS)  La 44·26 14·16 12·16 64·28 0·61  Ce 37·70 12·33 10·14 53·90 0·84  Pr 34·16 10·59 9·11 49·44 0·82  Nd 31·09 9·89 8·55 44·15 1·41  Sm 22·50 7·44 6·70 36·46 2·04  Eu 21·474 12·744 13·603 45·264 2·720  Gd 16·13 5·57 5·17 27·48 2·54  Tb 13·269 4·702 4·439 23·123 1·937  Dy 10·85 4·10 3·85 17·77 2·68  Ho 10·110 3·924 3·679 13·247 2·176  Er 9·519 3·813 3·517 9·697 2·350  Yb 8·323 3·688 3·174 5·137 1·595  Lu 8·211 3·655 3·119 2·708 Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1222 1320 1449 1459 1336B 1443A 1518 1514 Major elements (wt % oxide; XRF)  SiO2 52·39 53·42 51·66 51·43 51·21 51·61 40·41 41·62  TiO2 1·12 0·95 0·61 0·79 0·85 0·70 1·58 1·41  Al2O3 19·17 19·13 19·79 19·90 19·45 20·15 17·75 15·57  FeO(tot) 8·20 7·57 7·13 6·87 7·94 7·45 12·06 13·81  MnO 0·15 0·14 0·00 0·00 0·19 0·15 0·09 0·14  MgO 4·21 3·88 4·93 4·71 4·38 4·73 8·43 10·48  CaO 8·49 7·70 9·05 9·29 8·43 8·77 13·13 10·06  Na2O 5·05 5·21 4·88 4·50 4·46 5·06 2·69 3·53  K2O 0·49 0·64 0·42 0·46 0·41 0·41 0·53 0·68  P2O5 0·22 0·31 0·09 0·04 0·06 0·09 0·23 0·01  S <0·01 <0·01 <0·01 <0·01 <0·01 <0·01 <0·01  LOI 0·60 0·01 1·24 0·99 0·85 0·84 1·39 1·37  Total 101·00 99·80 100·60 99·74 99·13 100·81 99·63 100·21 Trace elements (ppm; XRF)  Sr 796·4 813·0 898·4 913·6 760·4 793·9 494·8 105·9  Y 8·00 16·80 5·80 5·50 3·80 4·20 18·20 5·30 REE data (chondrite normalised using values of McDonough & Sun (1995); solution ICP-MS)  La 44·26 14·16 12·16 64·28 0·61  Ce 37·70 12·33 10·14 53·90 0·84  Pr 34·16 10·59 9·11 49·44 0·82  Nd 31·09 9·89 8·55 44·15 1·41  Sm 22·50 7·44 6·70 36·46 2·04  Eu 21·474 12·744 13·603 45·264 2·720  Gd 16·13 5·57 5·17 27·48 2·54  Tb 13·269 4·702 4·439 23·123 1·937  Dy 10·85 4·10 3·85 17·77 2·68  Ho 10·110 3·924 3·679 13·247 2·176  Er 9·519 3·813 3·517 9·697 2·350  Yb 8·323 3·688 3·174 5·137 1·595  Lu 8·211 3·655 3·119 2·708 The complete dataset is available in Supplementary Data Electronic Appendix 1; supplementary data are available for downloading at http://www.petrology.oxfordjournals.org. Table 1: Representative bulk compositions of the protolith (S1) and rock types formed in each melt flux style Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1222 1320 1449 1459 1336B 1443A 1518 1514 Major elements (wt % oxide; XRF)  SiO2 52·39 53·42 51·66 51·43 51·21 51·61 40·41 41·62  TiO2 1·12 0·95 0·61 0·79 0·85 0·70 1·58 1·41  Al2O3 19·17 19·13 19·79 19·90 19·45 20·15 17·75 15·57  FeO(tot) 8·20 7·57 7·13 6·87 7·94 7·45 12·06 13·81  MnO 0·15 0·14 0·00 0·00 0·19 0·15 0·09 0·14  MgO 4·21 3·88 4·93 4·71 4·38 4·73 8·43 10·48  CaO 8·49 7·70 9·05 9·29 8·43 8·77 13·13 10·06  Na2O 5·05 5·21 4·88 4·50 4·46 5·06 2·69 3·53  K2O 0·49 0·64 0·42 0·46 0·41 0·41 0·53 0·68  P2O5 0·22 0·31 0·09 0·04 0·06 0·09 0·23 0·01  S <0·01 <0·01 <0·01 <0·01 <0·01 <0·01 <0·01  LOI 0·60 0·01 1·24 0·99 0·85 0·84 1·39 1·37  Total 101·00 99·80 100·60 99·74 99·13 100·81 99·63 100·21 Trace elements (ppm; XRF)  Sr 796·4 813·0 898·4 913·6 760·4 793·9 494·8 105·9  Y 8·00 16·80 5·80 5·50 3·80 4·20 18·20 5·30 REE data (chondrite normalised using values of McDonough & Sun (1995); solution ICP-MS)  La 44·26 14·16 12·16 64·28 0·61  Ce 37·70 12·33 10·14 53·90 0·84  Pr 34·16 10·59 9·11 49·44 0·82  Nd 31·09 9·89 8·55 44·15 1·41  Sm 22·50 7·44 6·70 36·46 2·04  Eu 21·474 12·744 13·603 45·264 2·720  Gd 16·13 5·57 5·17 27·48 2·54  Tb 13·269 4·702 4·439 23·123 1·937  Dy 10·85 4·10 3·85 17·77 2·68  Ho 10·110 3·924 3·679 13·247 2·176  Er 9·519 3·813 3·517 9·697 2·350  Yb 8·323 3·688 3·174 5·137 1·595  Lu 8·211 3·655 3·119 2·708 Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1222 1320 1449 1459 1336B 1443A 1518 1514 Major elements (wt % oxide; XRF)  SiO2 52·39 53·42 51·66 51·43 51·21 51·61 40·41 41·62  TiO2 1·12 0·95 0·61 0·79 0·85 0·70 1·58 1·41  Al2O3 19·17 19·13 19·79 19·90 19·45 20·15 17·75 15·57  FeO(tot) 8·20 7·57 7·13 6·87 7·94 7·45 12·06 13·81  MnO 0·15 0·14 0·00 0·00 0·19 0·15 0·09 0·14  MgO 4·21 3·88 4·93 4·71 4·38 4·73 8·43 10·48  CaO 8·49 7·70 9·05 9·29 8·43 8·77 13·13 10·06  Na2O 5·05 5·21 4·88 4·50 4·46 5·06 2·69 3·53  K2O 0·49 0·64 0·42 0·46 0·41 0·41 0·53 0·68  P2O5 0·22 0·31 0·09 0·04 0·06 0·09 0·23 0·01  S <0·01 <0·01 <0·01 <0·01 <0·01 <0·01 <0·01  LOI 0·60 0·01 1·24 0·99 0·85 0·84 1·39 1·37  Total 101·00 99·80 100·60 99·74 99·13 100·81 99·63 100·21 Trace elements (ppm; XRF)  Sr 796·4 813·0 898·4 913·6 760·4 793·9 494·8 105·9  Y 8·00 16·80 5·80 5·50 3·80 4·20 18·20 5·30 REE data (chondrite normalised using values of McDonough & Sun (1995); solution ICP-MS)  La 44·26 14·16 12·16 64·28 0·61  Ce 37·70 12·33 10·14 53·90 0·84  Pr 34·16 10·59 9·11 49·44 0·82  Nd 31·09 9·89 8·55 44·15 1·41  Sm 22·50 7·44 6·70 36·46 2·04  Eu 21·474 12·744 13·603 45·264 2·720  Gd 16·13 5·57 5·17 27·48 2·54  Tb 13·269 4·702 4·439 23·123 1·937  Dy 10·85 4·10 3·85 17·77 2·68  Ho 10·110 3·924 3·679 13·247 2·176  Er 9·519 3·813 3·517 9·697 2·350  Yb 8·323 3·688 3·174 5·137 1·595  Lu 8·211 3·655 3·119 2·708 The complete dataset is available in Supplementary Data Electronic Appendix 1; supplementary data are available for downloading at http://www.petrology.oxfordjournals.org. Fig. 4. View largeDownload slide Whole-rock composition of the S1 protolith, and melt–rock interaction styles (int. mod., intermediate modification; high mod., high modification). (a) TAS classification diagram. (b) Aluminium Harker diagram. (c) Magnesium Harker diagram. (d) SiO2 plotted against Sr/Y. (e) Chondrite normalised REE ratios for Styles 1 and 2. (f) Chondrite normalised REE patterns for Style 4. Grey field is range of REE contents in Styles 1 and 2. Fig. 4. View largeDownload slide Whole-rock composition of the S1 protolith, and melt–rock interaction styles (int. mod., intermediate modification; high mod., high modification). (a) TAS classification diagram. (b) Aluminium Harker diagram. (c) Magnesium Harker diagram. (d) SiO2 plotted against Sr/Y. (e) Chondrite normalised REE ratios for Styles 1 and 2. (f) Chondrite normalised REE patterns for Style 4. Grey field is range of REE contents in Styles 1 and 2. Samples exhibit small variations in all major element oxide concentrations, and overlap within the range of compositional variation of Styles 1, 2, and 3 (Fig. 4b and c). Styles 1, 2, and 3, are characterised by MgO of 3·88–6·34 wt % (Fig. 4c) and Sr/Y ratios of 36·5–221·1 (Fig. 4d). In comparison, Style 4 samples have higher MgO (6·70–11·37 wt %; Fig. 4c) and lower Sr/Y ratios (4·41–26·39; Table 1; Fig. 4d). Style 4 samples exhibit small variations in all major element oxide concentrations. However, Style 4 samples containing clinozoisite tend to have higher Al2O3 and lower MgO than those without (Fig. 4b and c). All styles have whole-rock REE patterns enriched relative to chondrite (Table 1). Styles 1 and 2 have negatively sloped REE patterns (LREE > 9 times chondrite, and MREE and HREE 1–11 times chondrite) and positive Eu anomalies (Fig. 4e). The REE patterns of Style 4 clinozoisite-bearing samples have similarly shaped, negatively sloped REE patterns which are further enriched compared to Styles 1 and 2 (LREE 20–90 times chondrite, MREE and HREE 2–25 times chondrite; Fig. 4f). Style 4 samples lacking clinozoisite have convex patterns, with inflections centred on Sm and larger variations in enrichment compared to chondrite (Fig. 4f). Only one clinozoisite-bearing sample has a pronounced, positive Eu anomaly; other samples show less pronounced, negative Eu anomalies or lack Eu anomalies. MINERAL COMPOSITIONS Major elements Amphibole formed during S1 and Styles 1 and 4 is predominantly pargasite (A site (NaA+KA) > 0·60, Fig. 5a;Table 2), whereas amphibole formed during Styles 2 and 3 is tschermakite and magnesio-hornblende, respectively (Style 2 A site < 0·60, Style 3 A site < 0·50, Fig. 5a;Table 2). S1 and Style 1 pargasite have overlapping compositions and exhibit the largest compositional variation. Style 2 tschermakite, Style 3 magnesio-hornblende, and Style 4 pargasite have smaller, more distinct ranges in composition (Fig. 5a and b), where Style 2 tschermakites have low A site occupancy and low Ti (∼0·51 and ∼0·06, respectively), Style 3 magnesio-hornblendes have the lowest A site occupancy and low Ti (∼0·45 and ∼0·05, respectively) and Style 4 pargasites have high A site occupancy and high Ti (∼0·76 and ∼0·09, respectively). Within the distinct compositional clusters for each Style, tschermakites in Style 2 intermediate modification samples have consistently lower Ti compared to samples of high modification, whereas pargasite in Style 4 clinozoisite-bearing samples has consistently higher Ti compared to samples lacking clinozoisite (Fig. 5b). Table 2: Representative compositions of minerals in each metamorphic event Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Sample 1453 1453 1453 1460 1460 1460 1449 1449 1449 1449 1459 1459 1459 1459 Mineral Prg Pl En Di Prg Pl Ts Czo Pl Grt Ts Czo Pl Grt Major elements (EMP; wt % oxide)  SiO2 39·27 56·36 51·64 50·84 39·92 57·38 41·14 37·47 61·98 37·96 40·81 37·31 61·04 38·07  TiO2 1·83 0·05 0·24 0·62 0·36 0·05 0·06 0·59 0·10 0·07  Al2O3 14·82 27·18 1·94 5·59 14·53 26·66 15·06 27·12 23·64 21·94 15·55 27·27 23·99 21·84  Cr2O3 0·00 0·04  FeO(tot) 15·42 0·03 22·91 8·90 14·91 0·05 13·43 7·87 0·07 20·30 13·72 7·55 0·09  MnO 0·22 0·01 0·65 0·17 0·17 0·22 0·11 0·45 0·13 0·12 1·62  MgO 9·82 0·01 21·62 10·84 10·41 11·21 0·10 10·83 11·01 0·13 26·49  CaO 10·96 9·85 0·26 20·19 11·41 8·43 10·74 23·58 5·37 7·37 10·65 23·33 5·68 6·07  Na2O 2·16 6·00 2·18 1·73 7·12 1·98 8·95 0·03 1·97 8·94 6·68  K2O 1·53 0·15 1·33 0·19 0·38 0·04 0·52 0·06 0·02  F 0·11  Cl 0·03 0·04  NiO 0·01  Total 96·03 99·65 99·02 98·95 95·06 99·83 94·56 93·60 100·05 98·94 94·99 95·92 99·80 100·86 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 12·53 13·46 0·15 2·74 6·76 10·55 29·24 171·56  Ce 19·28 9·20 0·15 6·39 12·25 4·32 0·12 19·23 159·87  Pr 22·09 6·12 0·10 9·68 16·35 2·09 0·42 10·13 152·91  Nd 21·77 4·18 10·57 15·05 0·97 0·49 9·26 151·51  Sm 19·46 1·11 0·16 10·00 12·64 1·96 1·35 0·29 103·51  Eu 30·728 5·133 0·314 14·760 17·638 3·535 1·616 76·554 0·201 112·256  Gd 15·13 8·34 6·53 0·70 4·17 7·39 69·60  Tb 11·690 6·704 5·596 2·825 8·449 40·139 1·335  Dy 13·05 0·15 7·68 6·67 4·11 10·93 34·92 3·11  Ho 11·868 0·201 6·758 6·465 4·212 6·007 0·687 26·264 5·147  Er 14·063 0·250 6·088 6·538 5·688 11·688 0·819 22·125 8·181  Tm 12·470 0·470 6·032 8·381 4·453 18·583 1·194 14·818 12·146  Yb 13·043 0·988 5·963 8·137 0·497 5·963 17·019 1·273 12·609 18·696  Lu 16·260 1·455 4·797 11·098 5·569 23·171 3·415 2·679 9·512 25·203 Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Sample 1453 1453 1453 1460 1460 1460 1449 1449 1449 1449 1459 1459 1459 1459 Mineral Prg Pl En Di Prg Pl Ts Czo Pl Grt Ts Czo Pl Grt Major elements (EMP; wt % oxide)  SiO2 39·27 56·36 51·64 50·84 39·92 57·38 41·14 37·47 61·98 37·96 40·81 37·31 61·04 38·07  TiO2 1·83 0·05 0·24 0·62 0·36 0·05 0·06 0·59 0·10 0·07  Al2O3 14·82 27·18 1·94 5·59 14·53 26·66 15·06 27·12 23·64 21·94 15·55 27·27 23·99 21·84  Cr2O3 0·00 0·04  FeO(tot) 15·42 0·03 22·91 8·90 14·91 0·05 13·43 7·87 0·07 20·30 13·72 7·55 0·09  MnO 0·22 0·01 0·65 0·17 0·17 0·22 0·11 0·45 0·13 0·12 1·62  MgO 9·82 0·01 21·62 10·84 10·41 11·21 0·10 10·83 11·01 0·13 26·49  CaO 10·96 9·85 0·26 20·19 11·41 8·43 10·74 23·58 5·37 7·37 10·65 23·33 5·68 6·07  Na2O 2·16 6·00 2·18 1·73 7·12 1·98 8·95 0·03 1·97 8·94 6·68  K2O 1·53 0·15 1·33 0·19 0·38 0·04 0·52 0·06 0·02  F 0·11  Cl 0·03 0·04  NiO 0·01  Total 96·03 99·65 99·02 98·95 95·06 99·83 94·56 93·60 100·05 98·94 94·99 95·92 99·80 100·86 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 12·53 13·46 0·15 2·74 6·76 10·55 29·24 171·56  Ce 19·28 9·20 0·15 6·39 12·25 4·32 0·12 19·23 159·87  Pr 22·09 6·12 0·10 9·68 16·35 2·09 0·42 10·13 152·91  Nd 21·77 4·18 10·57 15·05 0·97 0·49 9·26 151·51  Sm 19·46 1·11 0·16 10·00 12·64 1·96 1·35 0·29 103·51  Eu 30·728 5·133 0·314 14·760 17·638 3·535 1·616 76·554 0·201 112·256  Gd 15·13 8·34 6·53 0·70 4·17 7·39 69·60  Tb 11·690 6·704 5·596 2·825 8·449 40·139 1·335  Dy 13·05 0·15 7·68 6·67 4·11 10·93 34·92 3·11  Ho 11·868 0·201 6·758 6·465 4·212 6·007 0·687 26·264 5·147  Er 14·063 0·250 6·088 6·538 5·688 11·688 0·819 22·125 8·181  Tm 12·470 0·470 6·032 8·381 4·453 18·583 1·194 14·818 12·146  Yb 13·043 0·988 5·963 8·137 0·497 5·963 17·019 1·273 12·609 18·696  Lu 16·260 1·455 4·797 11·098 5·569 23·171 3·415 2·679 9·512 25·203 Timing Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1336b 1336b 1336b 1336b 1443a 1443a 1443a 1443a 1509 1509 1509 1509 1515 1515 1515 Mineral Mhb Czo Pl Grt Mhb Czo Pl Grt Prg Czo Pl Grt Prg Pl Grt Major elements (EMP; wt % oxide)  SiO2 42·01 38·68 61·52 38·95 42·08 38·79 61·54 38·50 42·46 38·18 63·12 38·53 40·30 63·16 37·94  TiO2 0·56 0·12 0·01 0·07 0·46 0·11 0·03 0·94 0·26 0·05 0·74 0·13 0·15  Al2O3 16·06 27·43 24·14 21·74 15·94 26·97 23·90 21·60 14·86 27·54 23·49 22·10 16·58 23·94 21·23  Cr2O3 0·01 0·01 0·02 0·02 0·00  FeO(tot) 14·30 7·36 0·15 27·63 15·83 7·40 0·08 26·03 13·03 8·03 0·18 21·96 14·30 0·16 24·46  MnO 0·21 0·05 0·03 1·43 0·51 0·05 4·51 0·08 0·08 1·12 0·13 1·95  MgO 9·68 0·11 6·33 8·87 0·12 0·00 3·93 11·79 0·18 5·61 10·61 4·44  CaO 10·67 23·47 5·47 5·32 10·89 23·31 5·67 6·09 10·69 23·84 4·54 10·27 10·60 4·25 8·88  Na2O 1·98 0·04 8·54 0·04 1·65 0·02 8·34 0·06 2·70 9·65 0·02 3·14 9·87  K2O 0·72 0·02 0·06 0·03 0·47 0·08 0·01 0·69 0·08 0·50 0·05  F 0·03  Cl 0·06 0·06  NiO  Total 96·19 97·29 99·92 101·54 96·73 96·77 99·63 100·79 97·30 97·31 101·06 101·65 96·96 101·56 100·72 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 0·06 54·94 0·35 0·09 13·63 1·54 0·16 26·08 0·02 0·16 0·07  Ce 0·11 51·99 0·22 0·21 0·10 11·21 1·66 0·39 19·04 0·05 0·40 0·08  Pr 50·54 0·16 0·83 0·21 10·38 0·95 0·07 0·58 14·11 0·07 0·06 0·56  Nd 0·15 54·18 0·30 4·03 0·30 9·34 0·64 1·02 12·19 0·18 0·05 1·06  Sm 64·59 15·20 0·53 8·65 0·65 0·39 1·30 8·85 0·53 2·31 0·49  Eu 1·439 183·304 0·995 20·071 2·416 55·773 1·510 0·716 3·535 23·108 0·760 1·687 3·481 0·263 1·794  Gd 51·31 26·63 0·67 5·22 0·31 0·64 1·86 4·92 0·43 1·44 1·54 2·48  Tb 54·155 0·288 38·144 1·587 6·233 0·211 1·781 2·715 5·873 0·706 6·427 2·327 0·072 5·152  Dy 1·15 45·33 40·45 2·48 7·64 0·17 4·20 2·19 6·18 0·76 13·21 2·85 10·98  Ho 1·170 35·714 37·033 2·033 5·458 0·176 9·176 2·088 6·520 0·489 29·121 2·564 14·689  Er 1·244 23·250 36·063 2·394 3·750 0·073 11·900 1·594 5·806 0·625 39·438 2·350 18·750  Tm 1·000 14·858 30·040 2·826 2·215 15·223 1·603 5·263 0·413 44·656 1·344 0·097 20·972  Yb 1·354 8·590 33·540 3·509 1·820 0·317 19·689 1·093 5·497 0·416 57·888 1·826 30·000  Lu 1·163 6·057 35·122 3·293 1·447 18·049 1·065 3·455 0·276 56·504 1·163 0·207 32·114 Timing Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1336b 1336b 1336b 1336b 1443a 1443a 1443a 1443a 1509 1509 1509 1509 1515 1515 1515 Mineral Mhb Czo Pl Grt Mhb Czo Pl Grt Prg Czo Pl Grt Prg Pl Grt Major elements (EMP; wt % oxide)  SiO2 42·01 38·68 61·52 38·95 42·08 38·79 61·54 38·50 42·46 38·18 63·12 38·53 40·30 63·16 37·94  TiO2 0·56 0·12 0·01 0·07 0·46 0·11 0·03 0·94 0·26 0·05 0·74 0·13 0·15  Al2O3 16·06 27·43 24·14 21·74 15·94 26·97 23·90 21·60 14·86 27·54 23·49 22·10 16·58 23·94 21·23  Cr2O3 0·01 0·01 0·02 0·02 0·00  FeO(tot) 14·30 7·36 0·15 27·63 15·83 7·40 0·08 26·03 13·03 8·03 0·18 21·96 14·30 0·16 24·46  MnO 0·21 0·05 0·03 1·43 0·51 0·05 4·51 0·08 0·08 1·12 0·13 1·95  MgO 9·68 0·11 6·33 8·87 0·12 0·00 3·93 11·79 0·18 5·61 10·61 4·44  CaO 10·67 23·47 5·47 5·32 10·89 23·31 5·67 6·09 10·69 23·84 4·54 10·27 10·60 4·25 8·88  Na2O 1·98 0·04 8·54 0·04 1·65 0·02 8·34 0·06 2·70 9·65 0·02 3·14 9·87  K2O 0·72 0·02 0·06 0·03 0·47 0·08 0·01 0·69 0·08 0·50 0·05  F 0·03  Cl 0·06 0·06  NiO  Total 96·19 97·29 99·92 101·54 96·73 96·77 99·63 100·79 97·30 97·31 101·06 101·65 96·96 101·56 100·72 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 0·06 54·94 0·35 0·09 13·63 1·54 0·16 26·08 0·02 0·16 0·07  Ce 0·11 51·99 0·22 0·21 0·10 11·21 1·66 0·39 19·04 0·05 0·40 0·08  Pr 50·54 0·16 0·83 0·21 10·38 0·95 0·07 0·58 14·11 0·07 0·06 0·56  Nd 0·15 54·18 0·30 4·03 0·30 9·34 0·64 1·02 12·19 0·18 0·05 1·06  Sm 64·59 15·20 0·53 8·65 0·65 0·39 1·30 8·85 0·53 2·31 0·49  Eu 1·439 183·304 0·995 20·071 2·416 55·773 1·510 0·716 3·535 23·108 0·760 1·687 3·481 0·263 1·794  Gd 51·31 26·63 0·67 5·22 0·31 0·64 1·86 4·92 0·43 1·44 1·54 2·48  Tb 54·155 0·288 38·144 1·587 6·233 0·211 1·781 2·715 5·873 0·706 6·427 2·327 0·072 5·152  Dy 1·15 45·33 40·45 2·48 7·64 0·17 4·20 2·19 6·18 0·76 13·21 2·85 10·98  Ho 1·170 35·714 37·033 2·033 5·458 0·176 9·176 2·088 6·520 0·489 29·121 2·564 14·689  Er 1·244 23·250 36·063 2·394 3·750 0·073 11·900 1·594 5·806 0·625 39·438 2·350 18·750  Tm 1·000 14·858 30·040 2·826 2·215 15·223 1·603 5·263 0·413 44·656 1·344 0·097 20·972  Yb 1·354 8·590 33·540 3·509 1·820 0·317 19·689 1·093 5·497 0·416 57·888 1·826 30·000  Lu 1·163 6·057 35·122 3·293 1·447 18·049 1·065 3·455 0·276 56·504 1·163 0·207 32·114 The complete dataset is available in Supplementary Data Electronic Appendix 2. Table 2: Representative compositions of minerals in each metamorphic event Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Sample 1453 1453 1453 1460 1460 1460 1449 1449 1449 1449 1459 1459 1459 1459 Mineral Prg Pl En Di Prg Pl Ts Czo Pl Grt Ts Czo Pl Grt Major elements (EMP; wt % oxide)  SiO2 39·27 56·36 51·64 50·84 39·92 57·38 41·14 37·47 61·98 37·96 40·81 37·31 61·04 38·07  TiO2 1·83 0·05 0·24 0·62 0·36 0·05 0·06 0·59 0·10 0·07  Al2O3 14·82 27·18 1·94 5·59 14·53 26·66 15·06 27·12 23·64 21·94 15·55 27·27 23·99 21·84  Cr2O3 0·00 0·04  FeO(tot) 15·42 0·03 22·91 8·90 14·91 0·05 13·43 7·87 0·07 20·30 13·72 7·55 0·09  MnO 0·22 0·01 0·65 0·17 0·17 0·22 0·11 0·45 0·13 0·12 1·62  MgO 9·82 0·01 21·62 10·84 10·41 11·21 0·10 10·83 11·01 0·13 26·49  CaO 10·96 9·85 0·26 20·19 11·41 8·43 10·74 23·58 5·37 7·37 10·65 23·33 5·68 6·07  Na2O 2·16 6·00 2·18 1·73 7·12 1·98 8·95 0·03 1·97 8·94 6·68  K2O 1·53 0·15 1·33 0·19 0·38 0·04 0·52 0·06 0·02  F 0·11  Cl 0·03 0·04  NiO 0·01  Total 96·03 99·65 99·02 98·95 95·06 99·83 94·56 93·60 100·05 98·94 94·99 95·92 99·80 100·86 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 12·53 13·46 0·15 2·74 6·76 10·55 29·24 171·56  Ce 19·28 9·20 0·15 6·39 12·25 4·32 0·12 19·23 159·87  Pr 22·09 6·12 0·10 9·68 16·35 2·09 0·42 10·13 152·91  Nd 21·77 4·18 10·57 15·05 0·97 0·49 9·26 151·51  Sm 19·46 1·11 0·16 10·00 12·64 1·96 1·35 0·29 103·51  Eu 30·728 5·133 0·314 14·760 17·638 3·535 1·616 76·554 0·201 112·256  Gd 15·13 8·34 6·53 0·70 4·17 7·39 69·60  Tb 11·690 6·704 5·596 2·825 8·449 40·139 1·335  Dy 13·05 0·15 7·68 6·67 4·11 10·93 34·92 3·11  Ho 11·868 0·201 6·758 6·465 4·212 6·007 0·687 26·264 5·147  Er 14·063 0·250 6·088 6·538 5·688 11·688 0·819 22·125 8·181  Tm 12·470 0·470 6·032 8·381 4·453 18·583 1·194 14·818 12·146  Yb 13·043 0·988 5·963 8·137 0·497 5·963 17·019 1·273 12·609 18·696  Lu 16·260 1·455 4·797 11·098 5·569 23·171 3·415 2·679 9·512 25·203 Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Sample 1453 1453 1453 1460 1460 1460 1449 1449 1449 1449 1459 1459 1459 1459 Mineral Prg Pl En Di Prg Pl Ts Czo Pl Grt Ts Czo Pl Grt Major elements (EMP; wt % oxide)  SiO2 39·27 56·36 51·64 50·84 39·92 57·38 41·14 37·47 61·98 37·96 40·81 37·31 61·04 38·07  TiO2 1·83 0·05 0·24 0·62 0·36 0·05 0·06 0·59 0·10 0·07  Al2O3 14·82 27·18 1·94 5·59 14·53 26·66 15·06 27·12 23·64 21·94 15·55 27·27 23·99 21·84  Cr2O3 0·00 0·04  FeO(tot) 15·42 0·03 22·91 8·90 14·91 0·05 13·43 7·87 0·07 20·30 13·72 7·55 0·09  MnO 0·22 0·01 0·65 0·17 0·17 0·22 0·11 0·45 0·13 0·12 1·62  MgO 9·82 0·01 21·62 10·84 10·41 11·21 0·10 10·83 11·01 0·13 26·49  CaO 10·96 9·85 0·26 20·19 11·41 8·43 10·74 23·58 5·37 7·37 10·65 23·33 5·68 6·07  Na2O 2·16 6·00 2·18 1·73 7·12 1·98 8·95 0·03 1·97 8·94 6·68  K2O 1·53 0·15 1·33 0·19 0·38 0·04 0·52 0·06 0·02  F 0·11  Cl 0·03 0·04  NiO 0·01  Total 96·03 99·65 99·02 98·95 95·06 99·83 94·56 93·60 100·05 98·94 94·99 95·92 99·80 100·86 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 12·53 13·46 0·15 2·74 6·76 10·55 29·24 171·56  Ce 19·28 9·20 0·15 6·39 12·25 4·32 0·12 19·23 159·87  Pr 22·09 6·12 0·10 9·68 16·35 2·09 0·42 10·13 152·91  Nd 21·77 4·18 10·57 15·05 0·97 0·49 9·26 151·51  Sm 19·46 1·11 0·16 10·00 12·64 1·96 1·35 0·29 103·51  Eu 30·728 5·133 0·314 14·760 17·638 3·535 1·616 76·554 0·201 112·256  Gd 15·13 8·34 6·53 0·70 4·17 7·39 69·60  Tb 11·690 6·704 5·596 2·825 8·449 40·139 1·335  Dy 13·05 0·15 7·68 6·67 4·11 10·93 34·92 3·11  Ho 11·868 0·201 6·758 6·465 4·212 6·007 0·687 26·264 5·147  Er 14·063 0·250 6·088 6·538 5·688 11·688 0·819 22·125 8·181  Tm 12·470 0·470 6·032 8·381 4·453 18·583 1·194 14·818 12·146  Yb 13·043 0·988 5·963 8·137 0·497 5·963 17·019 1·273 12·609 18·696  Lu 16·260 1·455 4·797 11·098 5·569 23·171 3·415 2·679 9·512 25·203 Timing Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1336b 1336b 1336b 1336b 1443a 1443a 1443a 1443a 1509 1509 1509 1509 1515 1515 1515 Mineral Mhb Czo Pl Grt Mhb Czo Pl Grt Prg Czo Pl Grt Prg Pl Grt Major elements (EMP; wt % oxide)  SiO2 42·01 38·68 61·52 38·95 42·08 38·79 61·54 38·50 42·46 38·18 63·12 38·53 40·30 63·16 37·94  TiO2 0·56 0·12 0·01 0·07 0·46 0·11 0·03 0·94 0·26 0·05 0·74 0·13 0·15  Al2O3 16·06 27·43 24·14 21·74 15·94 26·97 23·90 21·60 14·86 27·54 23·49 22·10 16·58 23·94 21·23  Cr2O3 0·01 0·01 0·02 0·02 0·00  FeO(tot) 14·30 7·36 0·15 27·63 15·83 7·40 0·08 26·03 13·03 8·03 0·18 21·96 14·30 0·16 24·46  MnO 0·21 0·05 0·03 1·43 0·51 0·05 4·51 0·08 0·08 1·12 0·13 1·95  MgO 9·68 0·11 6·33 8·87 0·12 0·00 3·93 11·79 0·18 5·61 10·61 4·44  CaO 10·67 23·47 5·47 5·32 10·89 23·31 5·67 6·09 10·69 23·84 4·54 10·27 10·60 4·25 8·88  Na2O 1·98 0·04 8·54 0·04 1·65 0·02 8·34 0·06 2·70 9·65 0·02 3·14 9·87  K2O 0·72 0·02 0·06 0·03 0·47 0·08 0·01 0·69 0·08 0·50 0·05  F 0·03  Cl 0·06 0·06  NiO  Total 96·19 97·29 99·92 101·54 96·73 96·77 99·63 100·79 97·30 97·31 101·06 101·65 96·96 101·56 100·72 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 0·06 54·94 0·35 0·09 13·63 1·54 0·16 26·08 0·02 0·16 0·07  Ce 0·11 51·99 0·22 0·21 0·10 11·21 1·66 0·39 19·04 0·05 0·40 0·08  Pr 50·54 0·16 0·83 0·21 10·38 0·95 0·07 0·58 14·11 0·07 0·06 0·56  Nd 0·15 54·18 0·30 4·03 0·30 9·34 0·64 1·02 12·19 0·18 0·05 1·06  Sm 64·59 15·20 0·53 8·65 0·65 0·39 1·30 8·85 0·53 2·31 0·49  Eu 1·439 183·304 0·995 20·071 2·416 55·773 1·510 0·716 3·535 23·108 0·760 1·687 3·481 0·263 1·794  Gd 51·31 26·63 0·67 5·22 0·31 0·64 1·86 4·92 0·43 1·44 1·54 2·48  Tb 54·155 0·288 38·144 1·587 6·233 0·211 1·781 2·715 5·873 0·706 6·427 2·327 0·072 5·152  Dy 1·15 45·33 40·45 2·48 7·64 0·17 4·20 2·19 6·18 0·76 13·21 2·85 10·98  Ho 1·170 35·714 37·033 2·033 5·458 0·176 9·176 2·088 6·520 0·489 29·121 2·564 14·689  Er 1·244 23·250 36·063 2·394 3·750 0·073 11·900 1·594 5·806 0·625 39·438 2·350 18·750  Tm 1·000 14·858 30·040 2·826 2·215 15·223 1·603 5·263 0·413 44·656 1·344 0·097 20·972  Yb 1·354 8·590 33·540 3·509 1·820 0·317 19·689 1·093 5·497 0·416 57·888 1·826 30·000  Lu 1·163 6·057 35·122 3·293 1·447 18·049 1·065 3·455 0·276 56·504 1·163 0·207 32·114 Timing Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1336b 1336b 1336b 1336b 1443a 1443a 1443a 1443a 1509 1509 1509 1509 1515 1515 1515 Mineral Mhb Czo Pl Grt Mhb Czo Pl Grt Prg Czo Pl Grt Prg Pl Grt Major elements (EMP; wt % oxide)  SiO2 42·01 38·68 61·52 38·95 42·08 38·79 61·54 38·50 42·46 38·18 63·12 38·53 40·30 63·16 37·94  TiO2 0·56 0·12 0·01 0·07 0·46 0·11 0·03 0·94 0·26 0·05 0·74 0·13 0·15  Al2O3 16·06 27·43 24·14 21·74 15·94 26·97 23·90 21·60 14·86 27·54 23·49 22·10 16·58 23·94 21·23  Cr2O3 0·01 0·01 0·02 0·02 0·00  FeO(tot) 14·30 7·36 0·15 27·63 15·83 7·40 0·08 26·03 13·03 8·03 0·18 21·96 14·30 0·16 24·46  MnO 0·21 0·05 0·03 1·43 0·51 0·05 4·51 0·08 0·08 1·12 0·13 1·95  MgO 9·68 0·11 6·33 8·87 0·12 0·00 3·93 11·79 0·18 5·61 10·61 4·44  CaO 10·67 23·47 5·47 5·32 10·89 23·31 5·67 6·09 10·69 23·84 4·54 10·27 10·60 4·25 8·88  Na2O 1·98 0·04 8·54 0·04 1·65 0·02 8·34 0·06 2·70 9·65 0·02 3·14 9·87  K2O 0·72 0·02 0·06 0·03 0·47 0·08 0·01 0·69 0·08 0·50 0·05  F 0·03  Cl 0·06 0·06  NiO  Total 96·19 97·29 99·92 101·54 96·73 96·77 99·63 100·79 97·30 97·31 101·06 101·65 96·96 101·56 100·72 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 0·06 54·94 0·35 0·09 13·63 1·54 0·16 26·08 0·02 0·16 0·07  Ce 0·11 51·99 0·22 0·21 0·10 11·21 1·66 0·39 19·04 0·05 0·40 0·08  Pr 50·54 0·16 0·83 0·21 10·38 0·95 0·07 0·58 14·11 0·07 0·06 0·56  Nd 0·15 54·18 0·30 4·03 0·30 9·34 0·64 1·02 12·19 0·18 0·05 1·06  Sm 64·59 15·20 0·53 8·65 0·65 0·39 1·30 8·85 0·53 2·31 0·49  Eu 1·439 183·304 0·995 20·071 2·416 55·773 1·510 0·716 3·535 23·108 0·760 1·687 3·481 0·263 1·794  Gd 51·31 26·63 0·67 5·22 0·31 0·64 1·86 4·92 0·43 1·44 1·54 2·48  Tb 54·155 0·288 38·144 1·587 6·233 0·211 1·781 2·715 5·873 0·706 6·427 2·327 0·072 5·152  Dy 1·15 45·33 40·45 2·48 7·64 0·17 4·20 2·19 6·18 0·76 13·21 2·85 10·98  Ho 1·170 35·714 37·033 2·033 5·458 0·176 9·176 2·088 6·520 0·489 29·121 2·564 14·689  Er 1·244 23·250 36·063 2·394 3·750 0·073 11·900 1·594 5·806 0·625 39·438 2·350 18·750  Tm 1·000 14·858 30·040 2·826 2·215 15·223 1·603 5·263 0·413 44·656 1·344 0·097 20·972  Yb 1·354 8·590 33·540 3·509 1·820 0·317 19·689 1·093 5·497 0·416 57·888 1·826 30·000  Lu 1·163 6·057 35·122 3·293 1·447 18·049 1·065 3·455 0·276 56·504 1·163 0·207 32·114 The complete dataset is available in Supplementary Data Electronic Appendix 2. Fig. 5. View largeDownload slide Mineral major element composition for the S1 protolith, and each of the melt flux style styles (int. mod., intermediate modification; high mod., high modification). (a) A-site occupancy in amphiboles. Error bars shown to the right, in the same scale as the graph. (b) Ti (c.p.f.u.) in amphiboles. Error bars shown to the right, in the same scale as the graph. (c) Garnet ternary classification diagram. (d) Pistacite content in clinozoisite. (e) Albite content in plagioclase. Fig. 5. View largeDownload slide Mineral major element composition for the S1 protolith, and each of the melt flux style styles (int. mod., intermediate modification; high mod., high modification). (a) A-site occupancy in amphiboles. Error bars shown to the right, in the same scale as the graph. (b) Ti (c.p.f.u.) in amphiboles. Error bars shown to the right, in the same scale as the graph. (c) Garnet ternary classification diagram. (d) Pistacite content in clinozoisite. (e) Albite content in plagioclase. Garnet formed during Style 2 is almandine with composition Alm50–57Pyr18–26Grs16–23Sps2–6 (Fig. 5c;Table 2), with rims that are enriched in pyrope and depleted in grossular by 1–3% respectively. Garnet in partially modified Style 2 samples is almandine–pyrope with Alm36–45Pyr32–42Grs20–21Sps1–3. Garnet formed during Style 3 is almandine with Alm51–60Pyr15–30Grs14–21Sps1–10, with minor zoning of Ca rich rims observed. Garnet formed during Style 4 is almandine–grossular with Alm44–53Pyr17–25Grs25–32Sps2–4, and lacks chemical zoning. The Fe3+ content in Style 2 garnet is 0·06–0·20 cpfu, whereas Style 3 and 4 garnets are within error of no Fe3+ content. Clinozoisite shows minor variations in Fe3+ substitution by Al3+ (Table 2), with little appreciable compositional difference between those formed in the different melt flux styles. Pistacite contents range between 12 and 18 (Fig. 5d), for which Style 1 clinozoisites have higher values (16–18), and Styles 2, 3, and 4 have wider ranges in composition. Plagioclase deformed in S1 has XAb = 58–66 (Table 2; Fig. 5e), and classifies as andesine. In contrast, plagioclase formed in the different melt flux styles is typically more albitic; Styles 1, 2, and 3 plagioclase analyses are classified as andesine or oligoclase (XAb = 53–80, 66–76, and 67–75 respectively), and Style 4 plagioclase has high XAb (77–87) and is either oligoclase or albite. Within Style 2, samples of intermediate modification have a wider compositional range (XAb = 53–76) compared to samples of high modification (XAb = 73–76). Minor K-feldspar occurring in melt pseudomorphs in Style 4 hornblendites are orthoclase (XOr = 98). Sr trace element mapping Synchrotron element mapping shows significant zoning of Sr both at the map scale and within individual grains of plagioclase associated with the different melt flux styles. Zonation of Sr shows no correlation with Ca or Ba, and Na was not detected as it is outside the energy sensitivity of the detector. Within a Style 1 sample, plagioclase (approx. 70 vol.%) can be divided into 1–2 mm wide bands of high- and low-Sr traversing the thin section (Fig. 6a). High-Sr bands closely follow pyroxene–plagioclase boundaries and are spatially associated with replacement microstructures of pargasite and quartz, Style 1 melt–rock interaction products. Coarse-grained plagioclase (0·9–2·0 mm) is zoned in Sr, where high-Sr grain boundaries are adjacent to replacement microstructures forming the high-Sr bands, and low-Sr grain boundaries are next to other plagioclase grains that are also low in Sr. Fig. 6. View largeDownload slide Range of Sr zoning in plagioclase. Note different Sr scales for each map. Mineral abbreviations follow the scheme proposed by Whitney & Evans (2010). (a) Style 1, field of view (FOV) = 18 mm across. (b) Style 2 (high modification), FOV = 21·5 mm across. (c) Style 3 (high modification), FOV = 20 mm across. Fig. 6. View largeDownload slide Range of Sr zoning in plagioclase. Note different Sr scales for each map. Mineral abbreviations follow the scheme proposed by Whitney & Evans (2010). (a) Style 1, field of view (FOV) = 18 mm across. (b) Style 2 (high modification), FOV = 21·5 mm across. (c) Style 3 (high modification), FOV = 20 mm across. Similar map-scale and grain-scale zoning are observed in a Style 2 sample (Fig. 6b). Sr zoning in plagioclase (approx. 40 vol. %) is present throughout the map area and is not spatially associated with microstructures or minerals. The map can be divided into bands of higher and lower Sr enrichment, up to 5 mm wide that are oriented vertically across the map and are sub-parallel with a composite S1 / S3. The high-Sr bands comprise many plagioclase grains (100–800 µm) that are enriched in Sr. Similarly, low-Sr bands comprise many grains that have relatively low Sr. Individual grains in both the high- and low-Sr bands are themselves asymmetrically zoned in Sr, where grains have high- and low-Sr grain boundaries that bear no relation to the grain boundaries of adjacent grains (i.e. a high-Sr boundary of one grain may be adjacent to a low-Sr boundary of a neighbouring grain). Style 3 samples have less pronounced map-scale and grain-scale zoning compared to Styles 1 and 2 (Fig. 6c). Map-scale zoning involves only slight relative changes to the overall Sr content in plagioclase (approx. 40 vol. %). One high-Sr band, approximately 5 mm wide, is associated with coarse-grained garnet on the far right of the map. Many individual grains exhibit homogeneous Sr contents; however, some rare grains have minor asymmetric Sr zoning. Rare earth element mineral compositions Amphiboles from S1 and Style 1 have enriched REE patterns with gentle convex shapes from La to Sm (1–27 times chondrite; Table 2; Fig. 7), small, positive Eu anomalies and flat MREE and HREE (4–19 times chondrite). In comparison, Style 2 amphiboles become progressively depleted in REE, with less pronounced Eu anomalies. Partially modified Style 2 samples have a wide range of depleted LREE values (0·01–1 times chondrite) which increase towards flat, enriched MREE to HREE (1–9 times chondrite). Style 2 samples of high modification have LREE values below detection limits, and increasing patterns from MREE to HREE (0·4–5 times chondrite). Style 3 amphiboles from samples of intermediate and high modification have similar patterns, with wide ranges in LREE from enriched, flat patterns (1–5 times chondrite) to depleted, sloped patterns (0·01–1 times chondrite). MREE and HREE are flat, approximately 1 times chondrite in partially modified samples and slightly more enriched in samples of high modification (1–4 times chondrite). All patterns have weak, positive Eu anomalies. Style 4 amphiboles have increasing trends from La to Eu (0·04–2·6 times chondrite) to flat, enriched MREE and HREE (up to 6 times chondrite). Fig. 7. View largeDownload slide Rare earth element profiles for amphibole, garnet and clinozoisite in the protolith, S1, and each melt–rock interaction style (int. mod., intermediate modification; high mod., high modification). Graphs are formatted consistent with the example shown in the top right. The range in 1σ error for each mineral is shown at the top right. Fig. 7. View largeDownload slide Rare earth element profiles for amphibole, garnet and clinozoisite in the protolith, S1, and each melt–rock interaction style (int. mod., intermediate modification; high mod., high modification). Graphs are formatted consistent with the example shown in the top right. The range in 1σ error for each mineral is shown at the top right. Garnets from Style 2 samples have LREE below detection limits, and the MREE define an increasing trend to HREE (1–56 times chondrite; Table 2; Fig. 7). Style 3 garnets have three different REE patterns. One is defined by increasing REE patterns from depleted LREE to enriched MREE, with an inflection point at Sm (∼25 times chondrite), a weak positive or neutral Eu anomaly and flat, enriched HREE patterns (∼30 times chondrite). A second pattern is comparatively depleted in HREE compared to the first (1–10 times chondrite), with decreasing patterns from Dy to Ho. A third pattern is comparatively depleted in LREE and MREE (below detection limits and 0·6–5 times chondrite, respectively), and has increasing trends from Dy to Lu (up to ∼40 times chondrite). In clinozoisite-bearing Style 4 samples, garnets have two patterns. One pattern is defined by an overall increasing trend from depleted LREE (below detection limits to 6 times chondrite) to enriched HREE (up to 23 times chondrite), with a decrease between Eu and Gd defining a strong, positive Eu anomaly. The second pattern also has an increasing trend from depleted LREE to enriched HREE, however it has either a neutral or weak, positive Eu anomaly, depleted LREE and a slight inflection between La and Ce, forming an inverted spoon-shaped pattern. This second, inverted spoon shaped pattern is typical of garnets in Style 4 samples lacking clinozoisite, which may be further enriched up to 250 times chondrite. Eu anomalies may be either weakly positive, weakly negative or neutral. In all samples, clinozoisite has enriched REE patterns (1–1300 times chondrite; Table 2; Fig. 7) defined by decreasing trends from La to Lu, with positive Eu anomalies of varying magnitude. Clinozoisite grains in Style 2 samples of intermediate modification have strong Eu anomalies, and the HREE tend to be flat rather than sloped. Clinozoisite grains from Style 2 samples of high modification are more enriched than those in samples of intermediate modification and have steeper HREE patterns. Style 3 clinozoisites have a mix of flat and steep HREE trends in samples of both intermediate and high modification and range between 1 and 225 times chondrite. Style 4 clinozoisites also show flat and steep HREE trends, where flat HREE profiles come from cores of grains and steep profiles from rims. TIME-INTEGRATED MELT FLUX ESTIMATION In this section, we estimate the time-integrated flux of melt during each melt flux style. The open system nature of melt–rock interaction and the passage of melt through, rather than injection into, the Pembroke Granulite precludes calculation of the composition of the melt and absolute volume of melt flux. However, the observed increase in the proportion of hydrous phases during each melt flux style can be utilised to interpret a melt-driven H2O metasomatism and then to calculate a minimum time-integrated melt flux. To calculate a minimum time-integrated melt flux we consider a theoretical cube of gabbroic (pyroxene-plagioclase) gneiss, which has essentially no hydrous phases and an H2O content of close to 0 wt %, as the initial composition of the Pembroke Granulite for each melt flux style. The final H2O composition of the cube resulting from melt flux is determined using the proportions of hydrous phases formed during melt–rock interaction (amphibole, clinozoisite, and muscovite), and their average H2O content. Proportions of hydrous phases have been published previously (Daczko et al., 2016; Stuart et al., 2017). The average H2O content of each phase is calculated by averaging the totals of EMP analyses and assuming that any missing mass represents H2O. Calculated H2O contents for amphibole, clinozoisite, and muscovite are 4·09 wt % (n = 210), 3·65 wt % (n = 100), and 4·21 wt % (n = 3), respectively. Combining the proportions of hydrous minerals in each melt flux style produces the final H2O content that ranges from 1·6 wt % for Style 1, to 4·1 wt % for Style 4 (Table 3). Table 3: Calculated estimates of melt flux Style 1 Style 2 Style 3 Style 4 (−czo) Proportion of minerals (%) Amphibole 30 38 43 100 Clinozoisite 9·0 13 3·0 0 Muscovite 0 4·0 0 0 Total H2O (wt %) 1·6 2·2 1·9 4·1 Melt H2O (wt %) m3 of melt flux per m3 rock 2·0 0·77 1·1 0·93 2·0 4·0 0·39 0·55 0·47 1·0 6·0 0·26 0·37 0·31 0·68 Style 1 Style 2 Style 3 Style 4 (−czo) Proportion of minerals (%) Amphibole 30 38 43 100 Clinozoisite 9·0 13 3·0 0 Muscovite 0 4·0 0 0 Total H2O (wt %) 1·6 2·2 1·9 4·1 Melt H2O (wt %) m3 of melt flux per m3 rock 2·0 0·77 1·1 0·93 2·0 4·0 0·39 0·55 0·47 1·0 6·0 0·26 0·37 0·31 0·68 Table 3: Calculated estimates of melt flux Style 1 Style 2 Style 3 Style 4 (−czo) Proportion of minerals (%) Amphibole 30 38 43 100 Clinozoisite 9·0 13 3·0 0 Muscovite 0 4·0 0 0 Total H2O (wt %) 1·6 2·2 1·9 4·1 Melt H2O (wt %) m3 of melt flux per m3 rock 2·0 0·77 1·1 0·93 2·0 4·0 0·39 0·55 0·47 1·0 6·0 0·26 0·37 0·31 0·68 Style 1 Style 2 Style 3 Style 4 (−czo) Proportion of minerals (%) Amphibole 30 38 43 100 Clinozoisite 9·0 13 3·0 0 Muscovite 0 4·0 0 0 Total H2O (wt %) 1·6 2·2 1·9 4·1 Melt H2O (wt %) m3 of melt flux per m3 rock 2·0 0·77 1·1 0·93 2·0 4·0 0·39 0·55 0·47 1·0 6·0 0·26 0·37 0·31 0·68 The time-integrated melt flux is estimated by calculating the volume of melt required to drive the increase in H2O content. Three melts of variable H2O content are considered (2, 4, and 6 wt %), which represent the range of typical values for subduction-related arc melts (Wallace, 2005; Plank et al., 2013 and references therein). The model assumes that all H2O in the melt is partitioned into the new hydrous phases, and that melt flux is homogeneous throughout the cube. In nature, only small amounts of the H2O would be partitioned into the solid phases, and melt flux is likely to be heterogeneously distributed. The model is limited by not considering anhydrous phases, such as plagioclase, that were stable during melt flux Styles 1–3. Therefore, calculated volumes of melt flux are significantly underestimated, especially in the case of Style 4 where unreactive melt flux is interpreted (Daczko et al., 2016). However, they provide a baseline for comparison between the four melt flux styles. Calculated volumes of melt flux range between 0·26 and 2·0 m3 of melt per m3 of rock (Table 3), depending on the initial H2O content of the fluxing melt and the melt flux style. Style 1 involves the smallest volume of melt flux, consistent with the observed partial hydration and replacement of pyroxene grains. Melt flux Styles 2 and 3 involve higher volumes of melt flux compared to Style 1. In contrast, Style 4 involves volumes of melt that are approximately double that calculated for Styles 1, 2, and 3. DISCUSSION Time-integrated volumes of melt flux Calculated time-integrated volumes of melt flux significantly underestimate the true volumes of melt flux, due to assumptions made in the model set up. However, the calculations can be indicative of relative differences between the individual melt flux styles. Styles 1, 2, and 3 have similar time-integrated volumes of melt flux, which stems from the similar assemblages in the modified rock types. Style 1 has the lowest time-integrated melt flux, consistent with the wide spread in mineral compositions (Fig. 7) and variable degree of modification (Stuart et al., 2016). Styles 2 and 3 both have higher time-integrated melt flux volumes, consistent with the homogeneous mineral compositions (Fig. 7), evidence for in situ partial melting (Daczko et al., 2001a, 2001b; Stuart et al., 2017), and associated higher degree of modification by melt–rock interaction. Calculated time-integrated volumes of melt flux for Styles 2 and 3 are likely to be further underestimated due to the in situ partial melting and breakdown of hydrous phases. The time-integrated volume of melt flux during Style 4 is calculated to be approximately double the volumes for Styles 1, 2, and 3 (Table 3). This is the minimum volume required to drive mineralogical changes within the host-rock, forming the armoured channel of hornblendite (Daczko et al., 2016). Continuation of melt flux after a rock has achieved a high degree of modification and reactant minerals are largely exhausted may obscure the ‘true’ time-integrated melt flux. For example, once the hornblendite has formed, melt flux is interpreted to occur without chemical exchange with the host-rock, and so realistic volumes of melt involved in Style 4 may be significantly higher than twice the amount of melt involved in Styles 1, 2, and 3. Equally, given that amphibole and plagioclase are the dominant stable minerals during Styles 2 and 3 melt flux, once these assemblages have formed, significant continued melt flux could also be unreactive, further obscuring the ‘true’ time-integrated melt flux in Styles 2 and 3. Different melt flux styles—different P–T conditions? Garnet-bearing assemblages in mafic to intermediate metaigneous rocks are indicative of high-P metamorphism and, in some cases, partial melting (O'Brien & Rötzler, 2003; Pattison, 2003; De Paoli et al., 2012; Chapman et al., 2017). Previous studies have shown that large volumes of the Fiordland lower crust, including the Pembroke Granulite, experienced burial and recrystallisation at high-P during an arc flare up, a period of voluminous (>100 km3/Ma per km of arc) high-Sr/Y pluton emplacement from 125 to 114 Ma (Bradshaw, 1989; Clarke et al., 2000; Daczko et al., 2002a, 2002b, 2009; Hollis et al., 2003, 2004; Daczko & Halpin, 2009; De Paoli et al., 2009; Stowell et al., 2010, 2014; Chapman et al., 2015; Milan et al., 2016, 2017). The four styles of melt flux observed in the Pembroke Granulite are interpreted as records of migration of these high-Sr/Y magmas (Stuart et al., 2016), albeit with different chemical signatures and variable time-integrated volumes of melt flux. P–T conditions of melt flux could not be calculated as the assemblages required for amphibole thermobarometry (e.g. Otten, 1984; Hammarstrom & Zen, 1986; Hollister et al., 1987; Anderson & Smith, 1995) were not present. However, interpretation of in situ melting (or a lack thereof) can be used to infer temperature differences between melt flux styles. Style 1 samples do not show any evidence for in situ partial melting (Stuart et al., 2016), whereas samples from Styles 2 and 3 with high modification include peritectic garnet evidencing minor (<5 vol. %) amounts of in situ partial melting (Fig. 3e–h; Daczko et al., 2001b; Stuart et al., 2017), suggesting that Styles 2 and 3 occurred at higher temperatures than Style 1. Plagioclase-absent assemblages from Style 4 are distinct from the plagioclase-bearing assemblages from Styles 2 and 3. Crystallisation experiments in mafic to intermediate magmas show that plagioclase is stable at lower temperatures and pressures close to the solidus and that the plagioclase-out line is at higher temperature for intermediate compared to mafic compositions (Green, 1982). Therefore, subtly higher temperatures and/or a more mafic melt accompanied Style 4 melt flux and dramatically changed the character of melt–rock interaction by destabilising plagioclase in the presence of the externally-derived melt. This switch may have been caused by advection of heat during extended or higher time-integrated melt flux of the externally derived melt and/or a shift to a more mafic character of the fluxing melt as the channel became armoured and less reactive. A localised temperature increase is supported by evidence for partial melting in a metre-scale transition zone that separates the Style 4 hornblendite body from the unmelted, precursor two-pyroxene–pargasite gneiss (Daczko et al., 2016). As such, Style 4 also describes a history of melt flux at conditions above the solidus within the Fiordland lower crust, which we infer was likely higher-temperature compared to Styles 1–3. Homogeneous versus heterogeneous whole rock compositions Preservation of S1 assemblages and partial modification during each style of melt–rock interaction provides the opportunity to evaluate the chemical evolution of both the bulk-rock and the mineralogy during different styles of melt–rock interaction. Each style of melt–rock interaction is subtly different; however the presence of a grain boundary network of melt, a bulk hydration effect and the growth of amphibole are common to all four styles. During flux, melt is interpreted to pass through the Pembroke Granulite with only small volumes of melt crystallising in place, implying that any variation in bulk composition is a result of melt-driven metasomatism. Rock types formed by melt–rock interaction during Styles 1, 2, and 3 have bulk-rock compositions indistinct from the precursor two-pyroxene–pargasite gneiss (Fig. 4; excluding volatile content). This includes a few samples of Style 3 that plot in the gabbroic field of Fig. 4a and are interpreted as sampling of slightly more mafic primary components of the Pembroke Granulite. Homogeneous bulk compositions in Styles 1, 2, and 3 overlapping with the precursor two-pyroxene–pargasite gneiss are consistent with nearly isochemical melt–rock interaction. In contrast, major and REE element compositions of Style 4 samples are different to that of the precursor rock (Fig. 4), consistent with melt-driven metasomatism during flux. In the case of Style 4, the melt–rock interaction is likely to have occurred at higher temperatures and involved a larger time-integrated melt flux, facilitating significant reaction and mass exchange during flux. Major element mineral chemistry and partitioning of rare earth elements in the presence of melt Despite significant melt-driven metasomatism during melt–rock interaction, the compositions of minerals from Style 4 are similar to, or overlap, the compositions of minerals from the precursor rock and Styles 1, 2, and 3 (Fig. 5). Variations in mineral compositions between Styles 1, 2, and 3 are difficult to link to melt-driven metasomatism when bulk-rock compositions are homogeneous. However, each melt flux style involved different time-integrated volumes of fluxed melt, and equilibration with variable volumes of melt may have imparted distinct mineral assemblages and compositions for each style. Style 1 involved low time-integrated volumes of melt flux (Table 3). The composition of Style 1 amphiboles (Fig. 5a and b) and distribution of Sr in Style 1 plagioclase (Fig. 6a) varies considerably, indicating a lack of widespread equilibration during melt-rock interaction. On the other hand, small ranges in mineral compositions (Fig. 5), less pronounced Sr zoning in plagioclase (Fig. 6b and c), and higher time-integrated volumes of melt flux (Table 3) for in Styles 2, 3, and 4 are indicative of widespread modification and better equilibration during melt-rock interaction. Homogeneity in major element mineral composition is most likely related to higher time-integrated volumes of melt flux (Table 3) and a higher degree of modification and recrystallisation. Rare earth element (REE) mineral chemistry (Fig. 7) suggests there was significant redistribution of REE between minerals during melt-rock interaction, but little enrichment or depletion occurred at a bulk-rock scale, except for Style 4 (Fig. 4f). This is likely due to the fact that the protolith for the Pembroke Granulite was an igneous rock, with a lack of accessory minerals in significant abundances that strongly partition trace elements and REE. In most cases, minerals have homogeneous REE patterns, despite forming in reactions where they replace precursor minerals with heterogeneous REE contents (e.g. Stuart et al., 2016). Homogenisation of REE patterns in metamorphic products is facilitated by the presence of an intergranular network of melt, providing more efficient diffusion pathways compared to solid-state reactions (Mann, 1980; Lesher, 1994; Acosta-Vigil et al., 2012) and enabling connectivity at an outcrop scale. In contrast to minerals from Styles 1, 2, and 4, the amphibole and garnet in high modification Style 3 samples have a significant variation in their REE patterns; a mix of positive and negative slopes are observed in both garnet HREE and amphibole LREE (Fig. 7). Given that REE are strongly partitioned into clinozoisite (Frei et al., 2003, 2004; Mulrooney & Rivers, 2005; Beard et al., 2006), heterogeneous REE patterns are likely a result of equilibration with varying proportions of clinozoisite. Style 3 is distinguished from the other styles by significant deformation during melt flux, which may localise melt flux (Bauer et al., 2000; Rosenberg & Handy, 2000; Holtzman et al., 2003; Baltzell et al., 2015), limiting connectivity and effectively reducing the volume of rock in chemical communication. Therefore, the spatial extent or connectivity of the melt network may dictate the equilibration volume and resulting homogeneity of REE compositions for each mineral. In this dynamic case, Style 3 may also have experienced multiple episodes of structural reactivation and repeated periods of melt flux, possibly by melts of variable composition. This may also contribute to the heterogeneous nature of mineral REE compositions. The broadly homogeneous REE patterns in melt–rock interaction products, such as garnet, amphibole, and clinozoisite, highlight the extent to which the REE are partitioned between the solid minerals and melt, in this P–T space straddling the boundary between igneous and sub-solidus metamorphic processes. We have evaluated published REE patterns of amphibole and clinozoisite formed under igneous versus sub-solidus metamorphic conditions (Gromet & Silver, 1983; Dalpé & Baker, 2000; El Korh et al., 2009) and compared them to those obtained in this contribution involving melt–rock interaction. Published values for sub-solidus metamorphic clinozoisite/epidote (Fig. 8a) are all enriched relative to chondrite with overlapping, flat REE patterns. The pattern for igneous clinozoisite/epidote is different to the metamorphic patterns, with enriched LREE relative to HREE forming a sloped pattern. This closely matches the average REE patterns of clinozoisite grains formed during melt-rock interaction in the Pembroke Granulite (Fig. 8b). Published amphibole REE patterns can be clearly divided into igneous, which are enriched with humps from La to Dy and flat HREE, versus sub-solidus metamorphic, which are flat and depleted relative to chondrite (Fig. 8c). Amphiboles formed during melt-rock interaction in the Pembroke Granulite are slightly more ambiguous, with characteristics of both types of published pattern. In general, the amphibole grains formed during Styles 1–4 melt flux have igneous-like, flat, enriched patterns from Gd to Lu (Fig. 8d). S1 and Style 1 amphiboles have a LREE hump, like igneous patterns from the literature. It is important to note that the Style 1 amphiboles are forming in an assemblage where only minor amounts of clinozoisite are stable, and garnet is not stable. Thus, amphiboles are not in competition for the REE available. Amphibole in Styles 2, 3, and 4 have more depleted, flat to sloped LREE. Depletion is more characteristic of a metamorphic signature, and as discussed above may be a result of partitioning with varying amounts of clinozoisite. However, unlike the metamorphic patterns, the LREE have a slope from La to Eu, which is closer to the shape of igneous-like patterns. Overall, amphibole and clinozoisite REE patterns share more similarities with published igneous REE patterns. Recrystallisation in the presence of melt, and the large equilibration volume provided by the melt network are two factors which have likely contributed to the formation of these igneous-like REE signatures. Fig. 8. View largeDownload slide Igneous and metamorphic signatures of mineral REE patterns. (a) Published values for igneous and metamorphic clinozoisite/epidote. Data from Gromet & Silver (1983) and El Korh et al. (2009). (b) Average REE patterns for clinozoisite from each melt–rock interaction style. (c) Published values for igneous and metamorphic amphibole. Data from Dalpé & Baker (2000) and El Korh et al. (2009). (d) Average REE patterns for amphibole from S1 and each melt–rock interaction style. Fig. 8. View largeDownload slide Igneous and metamorphic signatures of mineral REE patterns. (a) Published values for igneous and metamorphic clinozoisite/epidote. Data from Gromet & Silver (1983) and El Korh et al. (2009). (b) Average REE patterns for clinozoisite from each melt–rock interaction style. (c) Published values for igneous and metamorphic amphibole. Data from Dalpé & Baker (2000) and El Korh et al. (2009). (d) Average REE patterns for amphibole from S1 and each melt–rock interaction style. Generating igneous-like mineral chemical signatures in a metamorphic rock Differences in melt flux between each style highlight a role for physical processes in the formation of an igneous-like mineral chemical signature during melt–rock interaction. P–T conditions, time scales, and time-integrated volumes of melt flux may control both the stable assemblage and the major element composition of the minerals. In the case of the Pembroke Granulite, temperature is also inferred to play a role in the resulting assemblage of modified rock types, where the major element compositions of minerals formed during melt–rock interaction are determined by the P–T–X conditions of melt and host-rock, as in sub-solidus metamorphic systems. The homogeneity of mineral major element compositions, or the closeness to equilibrium, is here inferred to relate to the time-integrated volume of melt flux, where smaller time-integrated melt fluxes inhibit extensive equilibration, resulting in heterogeneous mineral major element compositions as in Style 1. On the other hand, the REE compositions of minerals are more homogeneous and have igneous-like signatures, where networks of melt provide large equilibration volumes for each mineral. In this case, the strength of the igneous-like signature and degree of homogeneity relies on the spatial distribution and connectivity of the melt network, which may be limited by deformation, such as in Style 3. CONCLUSIONS Melt–rock interaction during melt flux through the root of a magmatic arc has produced new assemblages and broadly homogeneous mineral compositions at an outcrop scale. Time-integrated volume of melt flux and variable equilibration between the host-rock and fluxing melt generates new mineral assemblages; temperature is inferred to control major element compositions. Igneous-like REE patterns in minerals formed in the presence of a grain boundary network of melt, and are identified as geochemical signatures recording the former flux of melt. The network of melt enhanced equilibration volumes and the mobility of REE at an outcrop scale. The time-integrated melt flux and spatial distribution of melt networks influence the degree of homogeneity and strength of the igneous-like mineral REE signature. Comparison of the four different melt flux styles examined in this study highlights the role that the physical characteristics of melt flux plays in generating the geochemical signatures. ACKNOWLEDGEMENTS We thank the Department of Conservation, New Zealand for permission to visit and sample localities in the Fiordland National Park. Editorial handling by G. Zellmer and reviews by S. Harley, M Pistone, and O. Jagoutz helped to improve this paper. This is contribution 1155 from the ARC Centre of Excellence for Core to Crust Fluid Systems (www.ccfs.mq.edu.au) and 1221 from the GEMOC Key Centre (www.gemoc.mq.edu.au). FUNDING This work was supported by an Australian Research Council Future Fellowship to S.P. (grant number FT110100070); an Australian Research Council Discovery Project to S.P. and N.R.D. (grant number DP120102060); and Australian Government Research Training Program Scholarships to C.A.S. and U.M. Part of this research was undertaken on the X-Ray Fluorescence Microscopy beamline at the Australian Synchrotron, Victoria, Australia. This work was supported by the Multi-modal Australian ScienceS Imaging and Visualisation Environment (MASSIVE) (www.massive.org.au). This study used instrumentation funded by ARC LIEF and DEST Systemic Infrastructure Grants, Macquarie University and Industry. SUPPLEMENTARY DATA Supplementary data for this paper, including whole rock composition, and mineral major and REE compositions are available at Journal of Petrology online. REFERENCES Acosta-Vigil A. , London D. , Morgan G. B. VI . ( 2012 ). Chemical diffusion of major components in granitic liquids: implications for the rates of homogenization of crustal melts . Lithos 153 , 308 – 323 . Google Scholar CrossRef Search ADS Allibone A. H. , Jongens R. , Turnbull I. M. , Milan L. 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For permissions, please e-mail: journals.permissions@oup.com This article is published and distributed under the terms of the Oxford University Press, Standard Journals Publication Model (https://academic.oup.com/journals/pages/about_us/legal/notices) http://www.deepdyve.com/assets/images/DeepDyve-Logo-lg.png Journal of Petrology Oxford University Press

Chemical Signatures of Melt–Rock Interaction in the Root of a Magmatic Arc

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Oxford University Press
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© The Author(s) 2018. Published by Oxford University Press. All rights reserved. For permissions, please e-mail: journals.permissions@oup.com
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0022-3530
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10.1093/petrology/egy029
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Abstract

ABSTRACT Identification of melt–rock interaction during melt flux through crustal rocks is limited to field relationships and microstructural evidence, with little consideration given to characterising the geochemical signatures of this process. We examine the mineral and whole-rock geochemistry of four distinct styles of melt–rock interaction during melt flux through the Pembroke Granulite, a gabbroic gneiss from the Fiordland magmatic arc root, New Zealand. Spatial distribution, time-integrated flux of melt and stress field vary between each melt flux style. Whole-rock metasomatism is not detected in three of the four melt flux styles. The mineral assemblage and major element mineral composition in modified rocks are dictated by inferred P–T conditions, as in sub-solidus metamorphic systems, and time-integrated volumes of melt flux. Heterogeneous mineral major and trace element compositions are linked to low time-integrated volumes of melt flux, which inhibits widespread modification and equilibration. Amphibole and clinozoisite in modified rocks have igneous-like REE patterns, formed by growth and/or recrystallisation in the presence of melt and large equilibration volumes provided by the grain boundary network of melt. Heterogeneities in mineral REE compositions are linked to localisation of melt flux by deformation and resulting smaller equilibration volumes and/or variation in the composition of the fluxing melt. When combined with microstructural evidence for the former presence of melt, the presence of igneous-like mineral REE chemical signatures in a metamorphic rock are proposed as powerful indicators of melt–rock interaction during melt flux. INTRODUCTION Partial melting, melt migration and crystallisation are fundamental processes responsible for the chemical evolution of the Earth’s crust. Each process imparts a different geochemical signature on the melt and surrounding rock with which it reacts. Despite demonstrations that melt–rock interaction during melt migration produces distinctive chemical signatures in the mantle (e.g. Nicolas, 1986; Spiegelman & Elliott, 1993; Rampone et al., 1994; Lundstrom et al., 1995; Pirard & Hermann, 2015), studies of melt migration through mafic to intermediate lower crust are limited. Most studies of mafic granulites and mafic migmatites focus on the geochemical signature of sites of melt production and/or injection and crystallisation (e.g. Tait & Harley, 1988; Collins & Sawyer, 1996; Sawyer et al., 1999), or at the axes of spreading oceanic ridges (e.g. Korenaga & Kelemen, 1997; Saal & Van Orman, 2004; Lissenberg & Dick, 2008; Lissenberg et al., 2013). In contrast, little consideration is given to geochemical signatures produced during the migration of melt through the lower continental crust. Melt–rock interaction in crustal rocks is rarely documented because of the inherent complexity of crustal rock types and melt compositions, and a lack of criteria for identification of the former flux of melt. Evidence for melt–rock interaction in the field may take the form of distributed or narrow zones of modification, such as hydration (e.g. Daczko et al., 2002b; Stuart, 2017; Stuart et al., 2017), or bodies of rock with igneous features that lack intrusive or structural boundaries with adjacent metamorphic rocks (e.g. Daczko et al., 2016). The most robust microstructural evidence is observation of microstructures indicative of the former presence of melt (e.g. small dihedral angles, pseudomorphs of grain boundary melt films, pseudomorphs of melt pockets at triple junctions, see Sawyer, 1999; White et al., 2005; Holness, 2008; Vernon, 2011) in rocks where in situ partial melting reactions are not interpreted. Microstructures indicative of a replacement reaction are also common in the deep crust, especially in rocks partially modified by flux of an externally-derived melt (Stuart et al., 2016, 2017). Though physical indicators of the former presence of melt are identified via comparison with igneous rocks and migmatites, little attention has been given to establishing geochemical signatures indicative of melt–rock interaction. Geochemical signatures in exposed arc roots and plutons are widely used to infer metamorphic and igneous processes occurring at depth. For example, studies have shown that metamorphic reactions in the lower crust do not necessarily mobilise rare earth elements (REE) over large distances, due to slow diffusion rates (e.g. Chernoff & Carlson, 1999; Van Orman et al., 2001; Cherniak, 2003; Tirone et al., 2005) and weak partitioning into aqueous fluids (Cullers et al., 1973; Flynn & Burnham, 1978; Wendlandt & Harrison, 1979; Cullers & Graf, 1984). As a result, metamorphic products in sub-solidus or fluid-poor conditions (e.g. Daczko et al., 2009) may inherit their chemical signatures from reactant minerals (e.g. Schröter et al., 2004; El Korh et al., 2009; Clarke et al., 2013; Chapman et al., 2015). In contrast, the REE are mobilised and redistributed by processes involving melt (Plank & Langmuir, 1993) and REE data are commonly used to determine evolving mineral assemblages and P–T conditions in the arc root (e.g. Dalpé & Baker, 2000; Tulloch & Kimbrough, 2003; Bignold et al., 2006) and magma differentiation processes (e.g. Stevenson et al., 2005; Davidson et al., 2007; Allibone et al., 2009; Smith, 2014; Chapman et al., 2016; Cooper et al., 2016). In principle, REE ratios and patterns in rocks and minerals may be informative of processes involving melt. However, melt–rock interaction straddles the boundary between igneous and metamorphic processes, involving melt but ultimately producing a metamorphic rock. The presence of melt, as opposed to an aqueous fluid, is expected to increase chemical communication and equilibration volumes. Therefore, melt presence may produce igneous mineral chemical signatures in the modified rock, though the effects of protolith composition, melt volume, and prevailing P–T conditions during melt–rock interaction have not been extensively studied. Four styles of melt flux throughout a single, homogeneous host rock, the Pembroke Granulite in Fiordland, southern New Zealand are examined in this contribution. Throughout this work, the term ‘flux’ is used to describe the open-system passage of melt through, rather than injection into, the Pembroke Granulite. Previous studies examining melt flux have focused on field relationships and microstructures, demonstrating that melt–rock interaction involved little crystallisation of melt within the modified rocks, and that mineral assemblages and microstructures produced during melt–rock interaction are common in lower crustal rocks (Daczko et al., 2016; Stuart et al., 2016, 2017; Stuart, 2017). We analysed bulk-rock composition and mineral major and trace element compositions to investigate the link between chemical signatures, melt flux style and P–T conditions of melt–rock interaction in the lower crust. Our findings indicate that the stable assemblage and major element composition of minerals are defined by variable equilibration with the fluxing melt. In contrast, mineral REE patterns are characterised by signatures common in igneous rocks. We relate the strength of these igneous signatures and relative homogeneity to the spatial extent and connectivity of the melt network throughout the rock. GEOLOGICAL SETTING AND PREVIOUS WORK Geological setting The Pembroke Granulite, Fiordland, New Zealand (Fig. 1), is a low-strain component of the Median Batholith, a suite of Carboniferous to Early Cretaceous plutons emplaced into, and partly comprising, the lower crust of a Cordilleran magmatic arc (Blattner, 1991; Mortimer et al., 1999). Emplaced at 139–129 Ma (Hollis et al., 2003), the gabbroic protolith to the Pembroke Granulite had an igneous assemblage of enstatite, diopside, brown-green pargasite, plagioclase, and ilmenite. The whole-rock composition of the two-pyroxene–pargasite pluton varies by several weight percent for all major element oxides (Stuart et al., 2016). The Pembroke Granulite, aside from some slight grain size variations, is otherwise homogeneous. Igneous minerals were variably recrystallised during D1 to form a gneissic foliation (S1) that strikes NE and dips steeply to the north and south. Deformation was recorded by the development of undulose extinction, deformation twins, and minor subgrain formation in igneous plagioclase, amphibole, and pyroxene grains, forming a rock distinctly metamorphic in character (Stuart et al., 2016). Similar assemblages in the nearby Western Fiordland Orthogneiss formed at lower-crustal conditions of ∼850°C and <11 kbar (Daczko & Halpin, 2009). Fig. 1. View largeDownload slide Geological map of Northern Fiordland, New Zealand. Location of study area marked by black arrow. Inset: Geological provinces of South Island, New Zealand, divided by the Alpine Fault plate boundary. Note the plate boundary ranges along strike from subduction to transpressional with local extension (Daczko et al., 2003). Fig. 1. View largeDownload slide Geological map of Northern Fiordland, New Zealand. Location of study area marked by black arrow. Inset: Geological provinces of South Island, New Zealand, divided by the Alpine Fault plate boundary. Note the plate boundary ranges along strike from subduction to transpressional with local extension (Daczko et al., 2003). Emplacement and D1 preceded a major pulse of high-Sr/Y magmatism in the arc system (126–115 Ma; Hollis et al., 2003, 2004; Tulloch & Kimbrough, 2003; Allibone et al., 2009; Milan et al., 2016) representing an arc flare up (Milan et al., 2017), and as such the Pembroke Granulite represents part of the lower crust through which the high-Sr/Y melts migrated (Daczko et al., 2016; Stuart et al., 2016, 2017). Summary of styles of melt flux Melt–rock interaction, as the high-Sr/Y melts fluxed through the Pembroke Granulite, resulted in four different styles of modification of the deformed, two-pyroxene–pargasite gneiss protolith (Fig. 2). Key changes common to all melt flux styles involve hydration and an increase in the mode of amphibole (Fig. 3). The first melt flux style involved widespread growth of pargasite-bearing coronae around igneous and S1 pyroxene throughout the entire Pembroke Granulite (Style 1 [En + Di + Prg +Qtz + Pl + Czo + Rt ± Ap]; Stuart et al., 2016). Each of the later styles of melt-rock interaction (Styles 2–4) formed distinct minor rock types hosted within the Pembroke Granulite, including tschermakite–clinozoisite gneiss and migmatite (Style 2 [Ts + Grt + Czo + Pl + Ms + Qtz + Rt]; sometimes called dioritic gneiss in previous literature; Daczko et al., 2001a, 2002a; Stuart et al., 2017), melt-bearing high-grade shear zones (Style 3 [Mhb + Grt + Czo + Pl + Ms + Qtz + Rt]; called D3 and D4 shear zones in previous literature; Daczko et al., 2001b; Gardner et al., 2016; Stuart, 2017), and hornblendite (Style 4 [Prg ± Czo ± Pl ± Bt ± Grt ± Ru]; Meek, 2015; Daczko et al., 2016). Previous work suggests that each melt flux style is distinct in terms of the volume of melt in the rock at any one time, the scale of melt-fluxed rock, the degree of modification of the protolith, the interpreted time-integrated melt flux and whether the flux was deformation assisted or occurred under static conditions. The key characteristics of the four styles of melt–rock interaction are summarised below. Mineral abbreviations used throughout the manuscript follow the scheme proposed by Whitney & Evans (2010). Fig. 2. View largeDownload slide Summary of key characteristics of styles of melt–rock interaction from Daczko et al. (2001b), Daczko et al. (2016), Meek (2015), Stuart (2017), Stuart et al. (2017) and Stuart et al. (2016), including sketches of melt flux along grain boundaries and relative spatial extent of melt flux. P–T conditions for Style 1 are from Stuart et al. (2016), and for Style 3 from Daczko et al. (2001b). Fig. 2. View largeDownload slide Summary of key characteristics of styles of melt–rock interaction from Daczko et al. (2001b), Daczko et al. (2016), Meek (2015), Stuart (2017), Stuart et al. (2017) and Stuart et al. (2016), including sketches of melt flux along grain boundaries and relative spatial extent of melt flux. P–T conditions for Style 1 are from Stuart et al. (2016), and for Style 3 from Daczko et al. (2001b). Fig. 3. View largeDownload slide Styles of melt–rock interaction, showing from left to right: typical field relationships including insets of a schematic cross section approximately 100 m long of the Pembroke Granulite showing the relative scale of each melt flux style, outcrop appearance and microstructures of modified rock types (scale bars 1000 µm), including insets of BSE images of microstructures indicative of the former presence of melt (scale bars 50 µm). Mineral abbreviations follow the scheme proposed by Whitney & Evans (2010). (a–c): Style 1; (d–f): Style 2, (g–i): Style 3, (j–l): Style 4. Fig. 3. View largeDownload slide Styles of melt–rock interaction, showing from left to right: typical field relationships including insets of a schematic cross section approximately 100 m long of the Pembroke Granulite showing the relative scale of each melt flux style, outcrop appearance and microstructures of modified rock types (scale bars 1000 µm), including insets of BSE images of microstructures indicative of the former presence of melt (scale bars 50 µm). Mineral abbreviations follow the scheme proposed by Whitney & Evans (2010). (a–c): Style 1; (d–f): Style 2, (g–i): Style 3, (j–l): Style 4. Style 1: Diffuse porous melt flow The earliest style of melt–rock interaction occurred post-D1 and involved partial modification along grain boundaries throughout the bulk of the Pembroke Granulite (Figs 2, 3a;Stuart et al., 2016). Initially, small local volumes (< 5%) of in situ partial melting are inferred to have created a permeable grain boundary network, which was utilised as migration pathways by an externally derived, high-Sr/Y melt. Melt–rock interaction involved partial hydration and replacement of the two-pyroxene–pargasite gneiss (Fig. 3b), forming coronae of blue-green pargasite and quartz (Fig. 3c) that are intimately associated with clinozoisite, Sr enrichment in plagioclase and microstructures indicative of the former presence of melt (pseudomorphs of melt including small dihedral angles, grain boundary films, and aggregates of quartz, plagioclase, and K-feldspar; Fig. 3c, inset; Stuart et al., 2016). Partial hydration of the S1 assemblage is observed, to a variable extent, throughout the entirety of the Pembroke Granulite. Style 1 involves a relatively small modification of the protolith under static conditions, distributed on a large scale, with an interpreted small, time-integrated melt flux at an outcrop scale. Partially hydrated S1 in the two-pyroxene–pargasite gneiss (Style 1 melt flux) is cut by a series of sub-vertical anorthositic D2 dykes, around which the assemblage is dehydrated and transformed to garnet granulite at a scale of cm-dm, called garnet reaction zones (GRZ); these are well-documented in the literature (Blattner, 1976; Clarke et al., 2000, 2005; Daczko et al., 2001a; Schröter et al., 2004; Daczko & Halpin, 2009; Smith et al., 2015). All subsequent styles of melt flux (Styles 2, 3 and 4) overprint both the partially hydrated S1 (Style 1) as well as GRZ. Style 2: Channelised porous melt flow Style 2 melt flux is similar in nature to Style 1, except that the melt flux is concentrated into narrow (<20 m wide) channels (Figs 2, 3d;Stuart et al., 2017). Within the channels, melt–rock interaction formed microstructures overprinting S1, where intermediate modification forms tschermakite–clinozoisite gneiss, and high modification forms migmatite (Fig. 3e). Modified rocks have hydrous assemblages dominated by tschermakite, clinozoisite and plagioclase (Fig. 3f). Migmatites exhibit significant recrystallisation and contain coarse-grained garnet (Fig. 3e and f) previously interpreted as a peritectic phase, indicating in situ partial melting (Daczko et al., 2001a; Stuart et al., 2017). The limited volumes of in situ melt produced are consistent with fluid-undersaturated conditions during partial melting (Stuart et al., 2017). The modification of the protolith associated with Style 2 melt flux is more extensive relative to Style 1, indicating either prolonged melt flux and/or higher time-integrated volumes of melt moving through the channels. Style 2, therefore, involves intermediate modification of the protolith, pervasive within small channels under static conditions, with a relatively intermediate time-integrated melt flux at an outcrop scale. Style 3: Deformation assisted porous melt flow Style 3 involves flux of melt along grain boundaries, where the flux is spatially restricted within D3 shear zones (Figs 2, 3g; these include the D3 and D4 shear zones of Daczko et al. (2001b)). Two-pyroxene–pargasite gneiss, tschermakite–clinozoisite gneiss, and migmatite are all deformed into shear zones (Fig. 3g and h) where they are modified to assemblages dominated by magnesio-hornblende, clinozoisite and plagioclase. Importantly, small amounts of syn-tectonic, medium-grained garnet (Fig. 3h) are observed (Daczko et al., 2001b), interpreted as a peritectic phase suggesting melt present conditions (Stuart, 2017). Ongoing work reveals that microstructures indicative of the former presence of melt are ubiquitous throughout shear zones (pseudomorphs of melt including: small dihedral angles, grain boundary films, grains of plagioclase interconnected in 3 D, multiphase aggregates at triple-point-junctions, and felsic dykelets; Fig. 3i, inset; Stuart, 2017), consistent with published P–T conditions for deformation which lie above the solidus (Fig. 2; Daczko et al., 2001b). Studies show that melt flux is more efficient in dynamic (deformation assisted) versus static conditions (e.g. Davidson et al., 1994; Daines & Kohlstedt, 1997; Rosenberg & Handy, 2000). Therefore, volumes of melt moving through the shear zones are expected to be higher than those in Styles 1 and 2. This is consistent with the complete hydration and overprinting of the S1 assemblage within the shear zones (Fig. 3i). Style 3 is, therefore, interpreted to involve intermediate modification of the deforming protolith, concentrated within zones of active shearing, with a relatively intermediate to high time-integrated melt flux at the outcrop scale. Style 4: Reactive infiltration, followed by melt flux through an armoured zone Style 4 involves the deformation-assisted reactive infiltration of melt within a 30–40 m wide channel forming a hornblendite (Figs 2, 3j; Daczko et al., 2016). Melt-rock interaction involved dissolution of plagioclase and pyroxene, and the precipitation of pargasite with, or without, clinozoisite (Fig. 3k), resulting in complete modification of the protolith. Local bands of garnet-rich rock are present (Daczko et al., 2016). The complete dissolution of plagioclase–pyroxene and growth of pargasite ± clinozoisite resulted in the formation of an armoured channel that is unreactive with the fluxing melt. Assisted by deformation, large volumes of melt may have fluxed efficiently through the unreactive channel once it became armoured. Former melt is now pseudomorphed by minor plagioclase (Fig. 3l, inset). The Style 4 hornblendite cuts rock types formed during Styles 1, 2 and 3. Style 4 is, therefore, interpreted to involve complete modification of the deforming protolith, concentrated within an armoured channel, with a relatively high time-integrated melt flux at an outcrop scale. SAMPLE SELECTION Data presented in this study are from representative rock types of all melt flux styles. Data for Styles 1 and 2 originate from samples analysed by Stuart et al. (2016) and Stuart et al. (2017). Raw data from these studies were collated with new data for Styles 3 and 4. Where possible, samples were collected directly from outcrops, however in some cases rock types were sampled from float in creeks after careful consideration of field relationships, characteristic mineralogy and textures compared to the outcrops. METHODS Whole-rock composition Concentrations of major element oxides SiO2, TiO2, Al2O3, Fe2O3, MgO, MnO, CaO, Na2O, K2O, P2O5 and SO3 and selected trace elements, including Sr and Y, were determined for representative samples of intermediate and high modification from Styles 3 and 4 using a PANalytical PW2400 Sequential WDXRF Spectrometer using WROXI standards (Mark Wainwright Analytical Centre at the University of New South Wales, Sydney, Australia). Trace and rare earth elements were determined for selected samples. Powdered samples were digested in sealed 15 mL Savillex Teflon beakers using a 1: 1 mixture of concentrated HF and concentrated HNO3 at 120°C for 24 hours, then dried down and repeated. Samples were then digested in 6 HNO3 for 24 hours, then dried down and diluted to 10 mL using 2% HNO3 and trace HF. 1: 1000 dilutions of each sample were then individually spiked with a 15 μL aliquot of a solution of 6Li, As, Rh, In, Tm and Bi in 2% HNO3. Samples were analysed on an Agilent 7500 series inductively coupled plasma mass spectrometer (ICP-MS) using BCR-2 as a calibration standard, the spike to correct for instrument drift, and a 2% HNO3 solution to measure background. Calibration was checked using the standards BIR-1 and BHVO-2 (Macquarie University GeoAnalytical, Sydney, Australia). Mineral major element compositions Polished thin (30 µm) and thick (100 µm) sections were made from blocks cut from representative samples; in the case of Style 3 samples blocks were cut perpendicular to foliation and parallel to lineation. A petrographic microscope was used in combination with the Virtual Petrographic Microscope (Tetley & Daczko, 2014), ImageJ (Rasband, 1997–2015), and backscattered electron (BSE) images for mineral identification. BSE images were collected using thin and thick sections coated with 10 nm of carbon in a Carl Zeiss IVO scanning electron microscope (SEM; Macquarie University GeoAnalytical, Sydney, Australia). The SEM was run with an accelerating voltage of 15–20 kV, a beam current of 10 nA and a working distance 10–12·5 mm. Major element compositions of minerals were determined using a CAMECA SX100 electron microprobe (S1, Styles 1, 2, and 4; Macquarie University GeoAnalytical, Sydney, Australia) and a JEOL JXA8100 electron microprobe (Style 3; the Institute of Geology and Geophysics, Chinese Academy of Sciences (IGGCAS), Beijing, China). Electron microprobes were run with accelerating voltages of 15 kV and beam currents of 20 nA, analysing 1–5 µm spot sizes for 3·5–5 minutes per spot. Weight percent oxide data were recalculated into cations per formula unit for individual minerals, using 23 oxygen for amphibole, 12 oxygen for garnet, 12·5 oxygen for clinozoisite and 8 oxygen for feldspar. Amphiboles were classified according to the scheme proposed by Hawthorne et al. (2012) using the spreadsheet provided by Locock (2014). The Fe3+ content in garnet was calculated by converting a portion of the FeO to Fe2O3 so that the cations per formula unit sum to 8. Garnet end-member proportions were then calculated using the following: Alm = 100 * Fe2+ / (Fe2+ + Mg + Ca + Mn); Pyr = 100 * Mg / (Fe2+ + Mg + Ca + Mn); Grs = 100 * Ca / (Fe2+ + Mg + Ca + Mn); Sps = 100 * Mn / (Fe2+ + Mg + Ca + Mn). Plagioclase end-member proportions were calculated using the following: Ab = 100 * Na / (Na + Ca + K); An = 100 * Ca / (Na + Ca + K); Or = 100 * K / (Na + Ca + K). Pistacite content in clinozoisite was calculated using: Ps = 100 * Fe3+ / (Fe3+ + Al3+). Mineral trace element compositions Trace element distribution mapping of representative samples of high modification from Styles 1, 2, and 3 was performed on the X-Ray Fluorescence Microscopy beamline, using the Maia-384 detector on the Kirkpatrick-Baez mirror microprobe at the Australian Synchrotron, Melbourne (Ryan et al., 2010b; Paterson et al., 2011). Polished thin sections mounted on glass slides were used for analysis; no additional preparation was performed. Maps were made by scanning thin sections in 4 µm step sizes in the x and y directions at a speed of 4·096 mm/s and dwell time of 0·98 ms/pixel. Spot sizes of 4 µm2 were analysed using a beam energy of 18·5 keV. The Maia detector has a minimum energy sensitivity of 3·3 keV, allowing detection of elements down to Z = 20 (Ca). Standard foils (Pt, Mn, Fe, YF3) were periodically analysed for calibration. Real time processing, using the Dynamic Analysis (DA) method (Ryan, 2000; Ryan et al., 2010a), deconvolves each individual X-ray event into element signals, allowing rapid data collection, high count rates and high sensitivity. Data reduction was performed using GeoPIXE (Ryan et al., 1990), which deconvolves the spectra using the fundamental parameter model for the layered sample, the Maia detector efficiency model, and the DA matrix method. A pure plagioclase matrix, using concentrations of values from analysed features in the samples and mineral densities, was used for matrix correction. Maps were constructed for each sample to highlight relative zoning of Sr in plagioclase. Mineral rare earth element composition Rare earth element compositions of minerals were determined using laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS; Macquarie University GeoAnalytical, Sydney, Australia). A Photon Machines Excite excimer 193 nm laser ablation microprobe system was used to ablate 25–50 µm spots from thin sections at a frequency of 5 Hz. Ablated material was transported in helium gas to the plasma at a flow rate of 0·8 Lmin-1. Material was analysed in an Agilent 7700x ICP-MS, using Ar as a carrier gas with a flow rate of 1·0 Lmin-1. Calcium (measured by EMP) was used as an internal standard for all minerals, and NIST 610 glass, basalt from the Columbia River (BCR-2), and MONGOL garnet were used as external standards. Raw signal data were reduced using GLITTER software (Griffin et al., 2008). Rare earth element concentrations obtained from GLITTER were normalised to chondrite values of McDonough & Sun (1995). WHOLE-ROCK COMPOSITION OF THE PEMBROKE GRANULITE Rocks fluxed by Styles 1, 2, and 3 are monzodioritic or monzogabbroic in composition (Table 1; Fig. 4a); some Style 3 samples fall within the gabbroic field on the (TAS) silica versus total alkalis diagram of Middlemost (1994). Style 4 samples have compositions distinct from the other styles, plotting in the foid-gabbroic field on the TAS diagram, with lower SiO2 (40·29–42·25 wt %; Table 1; Fig. 4a). All samples are peraluminous, with Aluminium saturation indices of ∼1·36 for Styles 1, 2, and 3, and ∼1·09 for Style 4, and Peacock indexes of ∼0·24 for Styles 1, 2, and 3, and ∼0·50 for Style 4. All samples have low TiO2 (0·42–1·99 wt %), high Al2O3 (16·74–20·44 wt %; Fig. 4b), and high alkali contents (Na2O + K2O = 2·55–6·18 wt %; Fig. 4a). Table 1: Representative bulk compositions of the protolith (S1) and rock types formed in each melt flux style Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1222 1320 1449 1459 1336B 1443A 1518 1514 Major elements (wt % oxide; XRF)  SiO2 52·39 53·42 51·66 51·43 51·21 51·61 40·41 41·62  TiO2 1·12 0·95 0·61 0·79 0·85 0·70 1·58 1·41  Al2O3 19·17 19·13 19·79 19·90 19·45 20·15 17·75 15·57  FeO(tot) 8·20 7·57 7·13 6·87 7·94 7·45 12·06 13·81  MnO 0·15 0·14 0·00 0·00 0·19 0·15 0·09 0·14  MgO 4·21 3·88 4·93 4·71 4·38 4·73 8·43 10·48  CaO 8·49 7·70 9·05 9·29 8·43 8·77 13·13 10·06  Na2O 5·05 5·21 4·88 4·50 4·46 5·06 2·69 3·53  K2O 0·49 0·64 0·42 0·46 0·41 0·41 0·53 0·68  P2O5 0·22 0·31 0·09 0·04 0·06 0·09 0·23 0·01  S <0·01 <0·01 <0·01 <0·01 <0·01 <0·01 <0·01  LOI 0·60 0·01 1·24 0·99 0·85 0·84 1·39 1·37  Total 101·00 99·80 100·60 99·74 99·13 100·81 99·63 100·21 Trace elements (ppm; XRF)  Sr 796·4 813·0 898·4 913·6 760·4 793·9 494·8 105·9  Y 8·00 16·80 5·80 5·50 3·80 4·20 18·20 5·30 REE data (chondrite normalised using values of McDonough & Sun (1995); solution ICP-MS)  La 44·26 14·16 12·16 64·28 0·61  Ce 37·70 12·33 10·14 53·90 0·84  Pr 34·16 10·59 9·11 49·44 0·82  Nd 31·09 9·89 8·55 44·15 1·41  Sm 22·50 7·44 6·70 36·46 2·04  Eu 21·474 12·744 13·603 45·264 2·720  Gd 16·13 5·57 5·17 27·48 2·54  Tb 13·269 4·702 4·439 23·123 1·937  Dy 10·85 4·10 3·85 17·77 2·68  Ho 10·110 3·924 3·679 13·247 2·176  Er 9·519 3·813 3·517 9·697 2·350  Yb 8·323 3·688 3·174 5·137 1·595  Lu 8·211 3·655 3·119 2·708 Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1222 1320 1449 1459 1336B 1443A 1518 1514 Major elements (wt % oxide; XRF)  SiO2 52·39 53·42 51·66 51·43 51·21 51·61 40·41 41·62  TiO2 1·12 0·95 0·61 0·79 0·85 0·70 1·58 1·41  Al2O3 19·17 19·13 19·79 19·90 19·45 20·15 17·75 15·57  FeO(tot) 8·20 7·57 7·13 6·87 7·94 7·45 12·06 13·81  MnO 0·15 0·14 0·00 0·00 0·19 0·15 0·09 0·14  MgO 4·21 3·88 4·93 4·71 4·38 4·73 8·43 10·48  CaO 8·49 7·70 9·05 9·29 8·43 8·77 13·13 10·06  Na2O 5·05 5·21 4·88 4·50 4·46 5·06 2·69 3·53  K2O 0·49 0·64 0·42 0·46 0·41 0·41 0·53 0·68  P2O5 0·22 0·31 0·09 0·04 0·06 0·09 0·23 0·01  S <0·01 <0·01 <0·01 <0·01 <0·01 <0·01 <0·01  LOI 0·60 0·01 1·24 0·99 0·85 0·84 1·39 1·37  Total 101·00 99·80 100·60 99·74 99·13 100·81 99·63 100·21 Trace elements (ppm; XRF)  Sr 796·4 813·0 898·4 913·6 760·4 793·9 494·8 105·9  Y 8·00 16·80 5·80 5·50 3·80 4·20 18·20 5·30 REE data (chondrite normalised using values of McDonough & Sun (1995); solution ICP-MS)  La 44·26 14·16 12·16 64·28 0·61  Ce 37·70 12·33 10·14 53·90 0·84  Pr 34·16 10·59 9·11 49·44 0·82  Nd 31·09 9·89 8·55 44·15 1·41  Sm 22·50 7·44 6·70 36·46 2·04  Eu 21·474 12·744 13·603 45·264 2·720  Gd 16·13 5·57 5·17 27·48 2·54  Tb 13·269 4·702 4·439 23·123 1·937  Dy 10·85 4·10 3·85 17·77 2·68  Ho 10·110 3·924 3·679 13·247 2·176  Er 9·519 3·813 3·517 9·697 2·350  Yb 8·323 3·688 3·174 5·137 1·595  Lu 8·211 3·655 3·119 2·708 The complete dataset is available in Supplementary Data Electronic Appendix 1; supplementary data are available for downloading at http://www.petrology.oxfordjournals.org. Table 1: Representative bulk compositions of the protolith (S1) and rock types formed in each melt flux style Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1222 1320 1449 1459 1336B 1443A 1518 1514 Major elements (wt % oxide; XRF)  SiO2 52·39 53·42 51·66 51·43 51·21 51·61 40·41 41·62  TiO2 1·12 0·95 0·61 0·79 0·85 0·70 1·58 1·41  Al2O3 19·17 19·13 19·79 19·90 19·45 20·15 17·75 15·57  FeO(tot) 8·20 7·57 7·13 6·87 7·94 7·45 12·06 13·81  MnO 0·15 0·14 0·00 0·00 0·19 0·15 0·09 0·14  MgO 4·21 3·88 4·93 4·71 4·38 4·73 8·43 10·48  CaO 8·49 7·70 9·05 9·29 8·43 8·77 13·13 10·06  Na2O 5·05 5·21 4·88 4·50 4·46 5·06 2·69 3·53  K2O 0·49 0·64 0·42 0·46 0·41 0·41 0·53 0·68  P2O5 0·22 0·31 0·09 0·04 0·06 0·09 0·23 0·01  S <0·01 <0·01 <0·01 <0·01 <0·01 <0·01 <0·01  LOI 0·60 0·01 1·24 0·99 0·85 0·84 1·39 1·37  Total 101·00 99·80 100·60 99·74 99·13 100·81 99·63 100·21 Trace elements (ppm; XRF)  Sr 796·4 813·0 898·4 913·6 760·4 793·9 494·8 105·9  Y 8·00 16·80 5·80 5·50 3·80 4·20 18·20 5·30 REE data (chondrite normalised using values of McDonough & Sun (1995); solution ICP-MS)  La 44·26 14·16 12·16 64·28 0·61  Ce 37·70 12·33 10·14 53·90 0·84  Pr 34·16 10·59 9·11 49·44 0·82  Nd 31·09 9·89 8·55 44·15 1·41  Sm 22·50 7·44 6·70 36·46 2·04  Eu 21·474 12·744 13·603 45·264 2·720  Gd 16·13 5·57 5·17 27·48 2·54  Tb 13·269 4·702 4·439 23·123 1·937  Dy 10·85 4·10 3·85 17·77 2·68  Ho 10·110 3·924 3·679 13·247 2·176  Er 9·519 3·813 3·517 9·697 2·350  Yb 8·323 3·688 3·174 5·137 1·595  Lu 8·211 3·655 3·119 2·708 Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1222 1320 1449 1459 1336B 1443A 1518 1514 Major elements (wt % oxide; XRF)  SiO2 52·39 53·42 51·66 51·43 51·21 51·61 40·41 41·62  TiO2 1·12 0·95 0·61 0·79 0·85 0·70 1·58 1·41  Al2O3 19·17 19·13 19·79 19·90 19·45 20·15 17·75 15·57  FeO(tot) 8·20 7·57 7·13 6·87 7·94 7·45 12·06 13·81  MnO 0·15 0·14 0·00 0·00 0·19 0·15 0·09 0·14  MgO 4·21 3·88 4·93 4·71 4·38 4·73 8·43 10·48  CaO 8·49 7·70 9·05 9·29 8·43 8·77 13·13 10·06  Na2O 5·05 5·21 4·88 4·50 4·46 5·06 2·69 3·53  K2O 0·49 0·64 0·42 0·46 0·41 0·41 0·53 0·68  P2O5 0·22 0·31 0·09 0·04 0·06 0·09 0·23 0·01  S <0·01 <0·01 <0·01 <0·01 <0·01 <0·01 <0·01  LOI 0·60 0·01 1·24 0·99 0·85 0·84 1·39 1·37  Total 101·00 99·80 100·60 99·74 99·13 100·81 99·63 100·21 Trace elements (ppm; XRF)  Sr 796·4 813·0 898·4 913·6 760·4 793·9 494·8 105·9  Y 8·00 16·80 5·80 5·50 3·80 4·20 18·20 5·30 REE data (chondrite normalised using values of McDonough & Sun (1995); solution ICP-MS)  La 44·26 14·16 12·16 64·28 0·61  Ce 37·70 12·33 10·14 53·90 0·84  Pr 34·16 10·59 9·11 49·44 0·82  Nd 31·09 9·89 8·55 44·15 1·41  Sm 22·50 7·44 6·70 36·46 2·04  Eu 21·474 12·744 13·603 45·264 2·720  Gd 16·13 5·57 5·17 27·48 2·54  Tb 13·269 4·702 4·439 23·123 1·937  Dy 10·85 4·10 3·85 17·77 2·68  Ho 10·110 3·924 3·679 13·247 2·176  Er 9·519 3·813 3·517 9·697 2·350  Yb 8·323 3·688 3·174 5·137 1·595  Lu 8·211 3·655 3·119 2·708 The complete dataset is available in Supplementary Data Electronic Appendix 1; supplementary data are available for downloading at http://www.petrology.oxfordjournals.org. Fig. 4. View largeDownload slide Whole-rock composition of the S1 protolith, and melt–rock interaction styles (int. mod., intermediate modification; high mod., high modification). (a) TAS classification diagram. (b) Aluminium Harker diagram. (c) Magnesium Harker diagram. (d) SiO2 plotted against Sr/Y. (e) Chondrite normalised REE ratios for Styles 1 and 2. (f) Chondrite normalised REE patterns for Style 4. Grey field is range of REE contents in Styles 1 and 2. Fig. 4. View largeDownload slide Whole-rock composition of the S1 protolith, and melt–rock interaction styles (int. mod., intermediate modification; high mod., high modification). (a) TAS classification diagram. (b) Aluminium Harker diagram. (c) Magnesium Harker diagram. (d) SiO2 plotted against Sr/Y. (e) Chondrite normalised REE ratios for Styles 1 and 2. (f) Chondrite normalised REE patterns for Style 4. Grey field is range of REE contents in Styles 1 and 2. Samples exhibit small variations in all major element oxide concentrations, and overlap within the range of compositional variation of Styles 1, 2, and 3 (Fig. 4b and c). Styles 1, 2, and 3, are characterised by MgO of 3·88–6·34 wt % (Fig. 4c) and Sr/Y ratios of 36·5–221·1 (Fig. 4d). In comparison, Style 4 samples have higher MgO (6·70–11·37 wt %; Fig. 4c) and lower Sr/Y ratios (4·41–26·39; Table 1; Fig. 4d). Style 4 samples exhibit small variations in all major element oxide concentrations. However, Style 4 samples containing clinozoisite tend to have higher Al2O3 and lower MgO than those without (Fig. 4b and c). All styles have whole-rock REE patterns enriched relative to chondrite (Table 1). Styles 1 and 2 have negatively sloped REE patterns (LREE > 9 times chondrite, and MREE and HREE 1–11 times chondrite) and positive Eu anomalies (Fig. 4e). The REE patterns of Style 4 clinozoisite-bearing samples have similarly shaped, negatively sloped REE patterns which are further enriched compared to Styles 1 and 2 (LREE 20–90 times chondrite, MREE and HREE 2–25 times chondrite; Fig. 4f). Style 4 samples lacking clinozoisite have convex patterns, with inflections centred on Sm and larger variations in enrichment compared to chondrite (Fig. 4f). Only one clinozoisite-bearing sample has a pronounced, positive Eu anomaly; other samples show less pronounced, negative Eu anomalies or lack Eu anomalies. MINERAL COMPOSITIONS Major elements Amphibole formed during S1 and Styles 1 and 4 is predominantly pargasite (A site (NaA+KA) > 0·60, Fig. 5a;Table 2), whereas amphibole formed during Styles 2 and 3 is tschermakite and magnesio-hornblende, respectively (Style 2 A site < 0·60, Style 3 A site < 0·50, Fig. 5a;Table 2). S1 and Style 1 pargasite have overlapping compositions and exhibit the largest compositional variation. Style 2 tschermakite, Style 3 magnesio-hornblende, and Style 4 pargasite have smaller, more distinct ranges in composition (Fig. 5a and b), where Style 2 tschermakites have low A site occupancy and low Ti (∼0·51 and ∼0·06, respectively), Style 3 magnesio-hornblendes have the lowest A site occupancy and low Ti (∼0·45 and ∼0·05, respectively) and Style 4 pargasites have high A site occupancy and high Ti (∼0·76 and ∼0·09, respectively). Within the distinct compositional clusters for each Style, tschermakites in Style 2 intermediate modification samples have consistently lower Ti compared to samples of high modification, whereas pargasite in Style 4 clinozoisite-bearing samples has consistently higher Ti compared to samples lacking clinozoisite (Fig. 5b). Table 2: Representative compositions of minerals in each metamorphic event Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Sample 1453 1453 1453 1460 1460 1460 1449 1449 1449 1449 1459 1459 1459 1459 Mineral Prg Pl En Di Prg Pl Ts Czo Pl Grt Ts Czo Pl Grt Major elements (EMP; wt % oxide)  SiO2 39·27 56·36 51·64 50·84 39·92 57·38 41·14 37·47 61·98 37·96 40·81 37·31 61·04 38·07  TiO2 1·83 0·05 0·24 0·62 0·36 0·05 0·06 0·59 0·10 0·07  Al2O3 14·82 27·18 1·94 5·59 14·53 26·66 15·06 27·12 23·64 21·94 15·55 27·27 23·99 21·84  Cr2O3 0·00 0·04  FeO(tot) 15·42 0·03 22·91 8·90 14·91 0·05 13·43 7·87 0·07 20·30 13·72 7·55 0·09  MnO 0·22 0·01 0·65 0·17 0·17 0·22 0·11 0·45 0·13 0·12 1·62  MgO 9·82 0·01 21·62 10·84 10·41 11·21 0·10 10·83 11·01 0·13 26·49  CaO 10·96 9·85 0·26 20·19 11·41 8·43 10·74 23·58 5·37 7·37 10·65 23·33 5·68 6·07  Na2O 2·16 6·00 2·18 1·73 7·12 1·98 8·95 0·03 1·97 8·94 6·68  K2O 1·53 0·15 1·33 0·19 0·38 0·04 0·52 0·06 0·02  F 0·11  Cl 0·03 0·04  NiO 0·01  Total 96·03 99·65 99·02 98·95 95·06 99·83 94·56 93·60 100·05 98·94 94·99 95·92 99·80 100·86 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 12·53 13·46 0·15 2·74 6·76 10·55 29·24 171·56  Ce 19·28 9·20 0·15 6·39 12·25 4·32 0·12 19·23 159·87  Pr 22·09 6·12 0·10 9·68 16·35 2·09 0·42 10·13 152·91  Nd 21·77 4·18 10·57 15·05 0·97 0·49 9·26 151·51  Sm 19·46 1·11 0·16 10·00 12·64 1·96 1·35 0·29 103·51  Eu 30·728 5·133 0·314 14·760 17·638 3·535 1·616 76·554 0·201 112·256  Gd 15·13 8·34 6·53 0·70 4·17 7·39 69·60  Tb 11·690 6·704 5·596 2·825 8·449 40·139 1·335  Dy 13·05 0·15 7·68 6·67 4·11 10·93 34·92 3·11  Ho 11·868 0·201 6·758 6·465 4·212 6·007 0·687 26·264 5·147  Er 14·063 0·250 6·088 6·538 5·688 11·688 0·819 22·125 8·181  Tm 12·470 0·470 6·032 8·381 4·453 18·583 1·194 14·818 12·146  Yb 13·043 0·988 5·963 8·137 0·497 5·963 17·019 1·273 12·609 18·696  Lu 16·260 1·455 4·797 11·098 5·569 23·171 3·415 2·679 9·512 25·203 Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Sample 1453 1453 1453 1460 1460 1460 1449 1449 1449 1449 1459 1459 1459 1459 Mineral Prg Pl En Di Prg Pl Ts Czo Pl Grt Ts Czo Pl Grt Major elements (EMP; wt % oxide)  SiO2 39·27 56·36 51·64 50·84 39·92 57·38 41·14 37·47 61·98 37·96 40·81 37·31 61·04 38·07  TiO2 1·83 0·05 0·24 0·62 0·36 0·05 0·06 0·59 0·10 0·07  Al2O3 14·82 27·18 1·94 5·59 14·53 26·66 15·06 27·12 23·64 21·94 15·55 27·27 23·99 21·84  Cr2O3 0·00 0·04  FeO(tot) 15·42 0·03 22·91 8·90 14·91 0·05 13·43 7·87 0·07 20·30 13·72 7·55 0·09  MnO 0·22 0·01 0·65 0·17 0·17 0·22 0·11 0·45 0·13 0·12 1·62  MgO 9·82 0·01 21·62 10·84 10·41 11·21 0·10 10·83 11·01 0·13 26·49  CaO 10·96 9·85 0·26 20·19 11·41 8·43 10·74 23·58 5·37 7·37 10·65 23·33 5·68 6·07  Na2O 2·16 6·00 2·18 1·73 7·12 1·98 8·95 0·03 1·97 8·94 6·68  K2O 1·53 0·15 1·33 0·19 0·38 0·04 0·52 0·06 0·02  F 0·11  Cl 0·03 0·04  NiO 0·01  Total 96·03 99·65 99·02 98·95 95·06 99·83 94·56 93·60 100·05 98·94 94·99 95·92 99·80 100·86 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 12·53 13·46 0·15 2·74 6·76 10·55 29·24 171·56  Ce 19·28 9·20 0·15 6·39 12·25 4·32 0·12 19·23 159·87  Pr 22·09 6·12 0·10 9·68 16·35 2·09 0·42 10·13 152·91  Nd 21·77 4·18 10·57 15·05 0·97 0·49 9·26 151·51  Sm 19·46 1·11 0·16 10·00 12·64 1·96 1·35 0·29 103·51  Eu 30·728 5·133 0·314 14·760 17·638 3·535 1·616 76·554 0·201 112·256  Gd 15·13 8·34 6·53 0·70 4·17 7·39 69·60  Tb 11·690 6·704 5·596 2·825 8·449 40·139 1·335  Dy 13·05 0·15 7·68 6·67 4·11 10·93 34·92 3·11  Ho 11·868 0·201 6·758 6·465 4·212 6·007 0·687 26·264 5·147  Er 14·063 0·250 6·088 6·538 5·688 11·688 0·819 22·125 8·181  Tm 12·470 0·470 6·032 8·381 4·453 18·583 1·194 14·818 12·146  Yb 13·043 0·988 5·963 8·137 0·497 5·963 17·019 1·273 12·609 18·696  Lu 16·260 1·455 4·797 11·098 5·569 23·171 3·415 2·679 9·512 25·203 Timing Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1336b 1336b 1336b 1336b 1443a 1443a 1443a 1443a 1509 1509 1509 1509 1515 1515 1515 Mineral Mhb Czo Pl Grt Mhb Czo Pl Grt Prg Czo Pl Grt Prg Pl Grt Major elements (EMP; wt % oxide)  SiO2 42·01 38·68 61·52 38·95 42·08 38·79 61·54 38·50 42·46 38·18 63·12 38·53 40·30 63·16 37·94  TiO2 0·56 0·12 0·01 0·07 0·46 0·11 0·03 0·94 0·26 0·05 0·74 0·13 0·15  Al2O3 16·06 27·43 24·14 21·74 15·94 26·97 23·90 21·60 14·86 27·54 23·49 22·10 16·58 23·94 21·23  Cr2O3 0·01 0·01 0·02 0·02 0·00  FeO(tot) 14·30 7·36 0·15 27·63 15·83 7·40 0·08 26·03 13·03 8·03 0·18 21·96 14·30 0·16 24·46  MnO 0·21 0·05 0·03 1·43 0·51 0·05 4·51 0·08 0·08 1·12 0·13 1·95  MgO 9·68 0·11 6·33 8·87 0·12 0·00 3·93 11·79 0·18 5·61 10·61 4·44  CaO 10·67 23·47 5·47 5·32 10·89 23·31 5·67 6·09 10·69 23·84 4·54 10·27 10·60 4·25 8·88  Na2O 1·98 0·04 8·54 0·04 1·65 0·02 8·34 0·06 2·70 9·65 0·02 3·14 9·87  K2O 0·72 0·02 0·06 0·03 0·47 0·08 0·01 0·69 0·08 0·50 0·05  F 0·03  Cl 0·06 0·06  NiO  Total 96·19 97·29 99·92 101·54 96·73 96·77 99·63 100·79 97·30 97·31 101·06 101·65 96·96 101·56 100·72 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 0·06 54·94 0·35 0·09 13·63 1·54 0·16 26·08 0·02 0·16 0·07  Ce 0·11 51·99 0·22 0·21 0·10 11·21 1·66 0·39 19·04 0·05 0·40 0·08  Pr 50·54 0·16 0·83 0·21 10·38 0·95 0·07 0·58 14·11 0·07 0·06 0·56  Nd 0·15 54·18 0·30 4·03 0·30 9·34 0·64 1·02 12·19 0·18 0·05 1·06  Sm 64·59 15·20 0·53 8·65 0·65 0·39 1·30 8·85 0·53 2·31 0·49  Eu 1·439 183·304 0·995 20·071 2·416 55·773 1·510 0·716 3·535 23·108 0·760 1·687 3·481 0·263 1·794  Gd 51·31 26·63 0·67 5·22 0·31 0·64 1·86 4·92 0·43 1·44 1·54 2·48  Tb 54·155 0·288 38·144 1·587 6·233 0·211 1·781 2·715 5·873 0·706 6·427 2·327 0·072 5·152  Dy 1·15 45·33 40·45 2·48 7·64 0·17 4·20 2·19 6·18 0·76 13·21 2·85 10·98  Ho 1·170 35·714 37·033 2·033 5·458 0·176 9·176 2·088 6·520 0·489 29·121 2·564 14·689  Er 1·244 23·250 36·063 2·394 3·750 0·073 11·900 1·594 5·806 0·625 39·438 2·350 18·750  Tm 1·000 14·858 30·040 2·826 2·215 15·223 1·603 5·263 0·413 44·656 1·344 0·097 20·972  Yb 1·354 8·590 33·540 3·509 1·820 0·317 19·689 1·093 5·497 0·416 57·888 1·826 30·000  Lu 1·163 6·057 35·122 3·293 1·447 18·049 1·065 3·455 0·276 56·504 1·163 0·207 32·114 Timing Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1336b 1336b 1336b 1336b 1443a 1443a 1443a 1443a 1509 1509 1509 1509 1515 1515 1515 Mineral Mhb Czo Pl Grt Mhb Czo Pl Grt Prg Czo Pl Grt Prg Pl Grt Major elements (EMP; wt % oxide)  SiO2 42·01 38·68 61·52 38·95 42·08 38·79 61·54 38·50 42·46 38·18 63·12 38·53 40·30 63·16 37·94  TiO2 0·56 0·12 0·01 0·07 0·46 0·11 0·03 0·94 0·26 0·05 0·74 0·13 0·15  Al2O3 16·06 27·43 24·14 21·74 15·94 26·97 23·90 21·60 14·86 27·54 23·49 22·10 16·58 23·94 21·23  Cr2O3 0·01 0·01 0·02 0·02 0·00  FeO(tot) 14·30 7·36 0·15 27·63 15·83 7·40 0·08 26·03 13·03 8·03 0·18 21·96 14·30 0·16 24·46  MnO 0·21 0·05 0·03 1·43 0·51 0·05 4·51 0·08 0·08 1·12 0·13 1·95  MgO 9·68 0·11 6·33 8·87 0·12 0·00 3·93 11·79 0·18 5·61 10·61 4·44  CaO 10·67 23·47 5·47 5·32 10·89 23·31 5·67 6·09 10·69 23·84 4·54 10·27 10·60 4·25 8·88  Na2O 1·98 0·04 8·54 0·04 1·65 0·02 8·34 0·06 2·70 9·65 0·02 3·14 9·87  K2O 0·72 0·02 0·06 0·03 0·47 0·08 0·01 0·69 0·08 0·50 0·05  F 0·03  Cl 0·06 0·06  NiO  Total 96·19 97·29 99·92 101·54 96·73 96·77 99·63 100·79 97·30 97·31 101·06 101·65 96·96 101·56 100·72 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 0·06 54·94 0·35 0·09 13·63 1·54 0·16 26·08 0·02 0·16 0·07  Ce 0·11 51·99 0·22 0·21 0·10 11·21 1·66 0·39 19·04 0·05 0·40 0·08  Pr 50·54 0·16 0·83 0·21 10·38 0·95 0·07 0·58 14·11 0·07 0·06 0·56  Nd 0·15 54·18 0·30 4·03 0·30 9·34 0·64 1·02 12·19 0·18 0·05 1·06  Sm 64·59 15·20 0·53 8·65 0·65 0·39 1·30 8·85 0·53 2·31 0·49  Eu 1·439 183·304 0·995 20·071 2·416 55·773 1·510 0·716 3·535 23·108 0·760 1·687 3·481 0·263 1·794  Gd 51·31 26·63 0·67 5·22 0·31 0·64 1·86 4·92 0·43 1·44 1·54 2·48  Tb 54·155 0·288 38·144 1·587 6·233 0·211 1·781 2·715 5·873 0·706 6·427 2·327 0·072 5·152  Dy 1·15 45·33 40·45 2·48 7·64 0·17 4·20 2·19 6·18 0·76 13·21 2·85 10·98  Ho 1·170 35·714 37·033 2·033 5·458 0·176 9·176 2·088 6·520 0·489 29·121 2·564 14·689  Er 1·244 23·250 36·063 2·394 3·750 0·073 11·900 1·594 5·806 0·625 39·438 2·350 18·750  Tm 1·000 14·858 30·040 2·826 2·215 15·223 1·603 5·263 0·413 44·656 1·344 0·097 20·972  Yb 1·354 8·590 33·540 3·509 1·820 0·317 19·689 1·093 5·497 0·416 57·888 1·826 30·000  Lu 1·163 6·057 35·122 3·293 1·447 18·049 1·065 3·455 0·276 56·504 1·163 0·207 32·114 The complete dataset is available in Supplementary Data Electronic Appendix 2. Table 2: Representative compositions of minerals in each metamorphic event Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Sample 1453 1453 1453 1460 1460 1460 1449 1449 1449 1449 1459 1459 1459 1459 Mineral Prg Pl En Di Prg Pl Ts Czo Pl Grt Ts Czo Pl Grt Major elements (EMP; wt % oxide)  SiO2 39·27 56·36 51·64 50·84 39·92 57·38 41·14 37·47 61·98 37·96 40·81 37·31 61·04 38·07  TiO2 1·83 0·05 0·24 0·62 0·36 0·05 0·06 0·59 0·10 0·07  Al2O3 14·82 27·18 1·94 5·59 14·53 26·66 15·06 27·12 23·64 21·94 15·55 27·27 23·99 21·84  Cr2O3 0·00 0·04  FeO(tot) 15·42 0·03 22·91 8·90 14·91 0·05 13·43 7·87 0·07 20·30 13·72 7·55 0·09  MnO 0·22 0·01 0·65 0·17 0·17 0·22 0·11 0·45 0·13 0·12 1·62  MgO 9·82 0·01 21·62 10·84 10·41 11·21 0·10 10·83 11·01 0·13 26·49  CaO 10·96 9·85 0·26 20·19 11·41 8·43 10·74 23·58 5·37 7·37 10·65 23·33 5·68 6·07  Na2O 2·16 6·00 2·18 1·73 7·12 1·98 8·95 0·03 1·97 8·94 6·68  K2O 1·53 0·15 1·33 0·19 0·38 0·04 0·52 0·06 0·02  F 0·11  Cl 0·03 0·04  NiO 0·01  Total 96·03 99·65 99·02 98·95 95·06 99·83 94·56 93·60 100·05 98·94 94·99 95·92 99·80 100·86 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 12·53 13·46 0·15 2·74 6·76 10·55 29·24 171·56  Ce 19·28 9·20 0·15 6·39 12·25 4·32 0·12 19·23 159·87  Pr 22·09 6·12 0·10 9·68 16·35 2·09 0·42 10·13 152·91  Nd 21·77 4·18 10·57 15·05 0·97 0·49 9·26 151·51  Sm 19·46 1·11 0·16 10·00 12·64 1·96 1·35 0·29 103·51  Eu 30·728 5·133 0·314 14·760 17·638 3·535 1·616 76·554 0·201 112·256  Gd 15·13 8·34 6·53 0·70 4·17 7·39 69·60  Tb 11·690 6·704 5·596 2·825 8·449 40·139 1·335  Dy 13·05 0·15 7·68 6·67 4·11 10·93 34·92 3·11  Ho 11·868 0·201 6·758 6·465 4·212 6·007 0·687 26·264 5·147  Er 14·063 0·250 6·088 6·538 5·688 11·688 0·819 22·125 8·181  Tm 12·470 0·470 6·032 8·381 4·453 18·583 1·194 14·818 12·146  Yb 13·043 0·988 5·963 8·137 0·497 5·963 17·019 1·273 12·609 18·696  Lu 16·260 1·455 4·797 11·098 5·569 23·171 3·415 2·679 9·512 25·203 Timing S1 Style 1 Style 2 int. mod. Style 2 high mod. Sample 1453 1453 1453 1460 1460 1460 1449 1449 1449 1449 1459 1459 1459 1459 Mineral Prg Pl En Di Prg Pl Ts Czo Pl Grt Ts Czo Pl Grt Major elements (EMP; wt % oxide)  SiO2 39·27 56·36 51·64 50·84 39·92 57·38 41·14 37·47 61·98 37·96 40·81 37·31 61·04 38·07  TiO2 1·83 0·05 0·24 0·62 0·36 0·05 0·06 0·59 0·10 0·07  Al2O3 14·82 27·18 1·94 5·59 14·53 26·66 15·06 27·12 23·64 21·94 15·55 27·27 23·99 21·84  Cr2O3 0·00 0·04  FeO(tot) 15·42 0·03 22·91 8·90 14·91 0·05 13·43 7·87 0·07 20·30 13·72 7·55 0·09  MnO 0·22 0·01 0·65 0·17 0·17 0·22 0·11 0·45 0·13 0·12 1·62  MgO 9·82 0·01 21·62 10·84 10·41 11·21 0·10 10·83 11·01 0·13 26·49  CaO 10·96 9·85 0·26 20·19 11·41 8·43 10·74 23·58 5·37 7·37 10·65 23·33 5·68 6·07  Na2O 2·16 6·00 2·18 1·73 7·12 1·98 8·95 0·03 1·97 8·94 6·68  K2O 1·53 0·15 1·33 0·19 0·38 0·04 0·52 0·06 0·02  F 0·11  Cl 0·03 0·04  NiO 0·01  Total 96·03 99·65 99·02 98·95 95·06 99·83 94·56 93·60 100·05 98·94 94·99 95·92 99·80 100·86 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 12·53 13·46 0·15 2·74 6·76 10·55 29·24 171·56  Ce 19·28 9·20 0·15 6·39 12·25 4·32 0·12 19·23 159·87  Pr 22·09 6·12 0·10 9·68 16·35 2·09 0·42 10·13 152·91  Nd 21·77 4·18 10·57 15·05 0·97 0·49 9·26 151·51  Sm 19·46 1·11 0·16 10·00 12·64 1·96 1·35 0·29 103·51  Eu 30·728 5·133 0·314 14·760 17·638 3·535 1·616 76·554 0·201 112·256  Gd 15·13 8·34 6·53 0·70 4·17 7·39 69·60  Tb 11·690 6·704 5·596 2·825 8·449 40·139 1·335  Dy 13·05 0·15 7·68 6·67 4·11 10·93 34·92 3·11  Ho 11·868 0·201 6·758 6·465 4·212 6·007 0·687 26·264 5·147  Er 14·063 0·250 6·088 6·538 5·688 11·688 0·819 22·125 8·181  Tm 12·470 0·470 6·032 8·381 4·453 18·583 1·194 14·818 12·146  Yb 13·043 0·988 5·963 8·137 0·497 5·963 17·019 1·273 12·609 18·696  Lu 16·260 1·455 4·797 11·098 5·569 23·171 3·415 2·679 9·512 25·203 Timing Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1336b 1336b 1336b 1336b 1443a 1443a 1443a 1443a 1509 1509 1509 1509 1515 1515 1515 Mineral Mhb Czo Pl Grt Mhb Czo Pl Grt Prg Czo Pl Grt Prg Pl Grt Major elements (EMP; wt % oxide)  SiO2 42·01 38·68 61·52 38·95 42·08 38·79 61·54 38·50 42·46 38·18 63·12 38·53 40·30 63·16 37·94  TiO2 0·56 0·12 0·01 0·07 0·46 0·11 0·03 0·94 0·26 0·05 0·74 0·13 0·15  Al2O3 16·06 27·43 24·14 21·74 15·94 26·97 23·90 21·60 14·86 27·54 23·49 22·10 16·58 23·94 21·23  Cr2O3 0·01 0·01 0·02 0·02 0·00  FeO(tot) 14·30 7·36 0·15 27·63 15·83 7·40 0·08 26·03 13·03 8·03 0·18 21·96 14·30 0·16 24·46  MnO 0·21 0·05 0·03 1·43 0·51 0·05 4·51 0·08 0·08 1·12 0·13 1·95  MgO 9·68 0·11 6·33 8·87 0·12 0·00 3·93 11·79 0·18 5·61 10·61 4·44  CaO 10·67 23·47 5·47 5·32 10·89 23·31 5·67 6·09 10·69 23·84 4·54 10·27 10·60 4·25 8·88  Na2O 1·98 0·04 8·54 0·04 1·65 0·02 8·34 0·06 2·70 9·65 0·02 3·14 9·87  K2O 0·72 0·02 0·06 0·03 0·47 0·08 0·01 0·69 0·08 0·50 0·05  F 0·03  Cl 0·06 0·06  NiO  Total 96·19 97·29 99·92 101·54 96·73 96·77 99·63 100·79 97·30 97·31 101·06 101·65 96·96 101·56 100·72 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 0·06 54·94 0·35 0·09 13·63 1·54 0·16 26·08 0·02 0·16 0·07  Ce 0·11 51·99 0·22 0·21 0·10 11·21 1·66 0·39 19·04 0·05 0·40 0·08  Pr 50·54 0·16 0·83 0·21 10·38 0·95 0·07 0·58 14·11 0·07 0·06 0·56  Nd 0·15 54·18 0·30 4·03 0·30 9·34 0·64 1·02 12·19 0·18 0·05 1·06  Sm 64·59 15·20 0·53 8·65 0·65 0·39 1·30 8·85 0·53 2·31 0·49  Eu 1·439 183·304 0·995 20·071 2·416 55·773 1·510 0·716 3·535 23·108 0·760 1·687 3·481 0·263 1·794  Gd 51·31 26·63 0·67 5·22 0·31 0·64 1·86 4·92 0·43 1·44 1·54 2·48  Tb 54·155 0·288 38·144 1·587 6·233 0·211 1·781 2·715 5·873 0·706 6·427 2·327 0·072 5·152  Dy 1·15 45·33 40·45 2·48 7·64 0·17 4·20 2·19 6·18 0·76 13·21 2·85 10·98  Ho 1·170 35·714 37·033 2·033 5·458 0·176 9·176 2·088 6·520 0·489 29·121 2·564 14·689  Er 1·244 23·250 36·063 2·394 3·750 0·073 11·900 1·594 5·806 0·625 39·438 2·350 18·750  Tm 1·000 14·858 30·040 2·826 2·215 15·223 1·603 5·263 0·413 44·656 1·344 0·097 20·972  Yb 1·354 8·590 33·540 3·509 1·820 0·317 19·689 1·093 5·497 0·416 57·888 1·826 30·000  Lu 1·163 6·057 35·122 3·293 1·447 18·049 1·065 3·455 0·276 56·504 1·163 0·207 32·114 Timing Style 3 int. mod. Style 3 high mod. Style 4 (+ czo) Style 4 (− czo) Sample 1336b 1336b 1336b 1336b 1443a 1443a 1443a 1443a 1509 1509 1509 1509 1515 1515 1515 Mineral Mhb Czo Pl Grt Mhb Czo Pl Grt Prg Czo Pl Grt Prg Pl Grt Major elements (EMP; wt % oxide)  SiO2 42·01 38·68 61·52 38·95 42·08 38·79 61·54 38·50 42·46 38·18 63·12 38·53 40·30 63·16 37·94  TiO2 0·56 0·12 0·01 0·07 0·46 0·11 0·03 0·94 0·26 0·05 0·74 0·13 0·15  Al2O3 16·06 27·43 24·14 21·74 15·94 26·97 23·90 21·60 14·86 27·54 23·49 22·10 16·58 23·94 21·23  Cr2O3 0·01 0·01 0·02 0·02 0·00  FeO(tot) 14·30 7·36 0·15 27·63 15·83 7·40 0·08 26·03 13·03 8·03 0·18 21·96 14·30 0·16 24·46  MnO 0·21 0·05 0·03 1·43 0·51 0·05 4·51 0·08 0·08 1·12 0·13 1·95  MgO 9·68 0·11 6·33 8·87 0·12 0·00 3·93 11·79 0·18 5·61 10·61 4·44  CaO 10·67 23·47 5·47 5·32 10·89 23·31 5·67 6·09 10·69 23·84 4·54 10·27 10·60 4·25 8·88  Na2O 1·98 0·04 8·54 0·04 1·65 0·02 8·34 0·06 2·70 9·65 0·02 3·14 9·87  K2O 0·72 0·02 0·06 0·03 0·47 0·08 0·01 0·69 0·08 0·50 0·05  F 0·03  Cl 0·06 0·06  NiO  Total 96·19 97·29 99·92 101·54 96·73 96·77 99·63 100·79 97·30 97·31 101·06 101·65 96·96 101·56 100·72 REE data (LA-ICP-MS; chondrite normalised using values of McDonough & Sun (1995))  La 0·06 54·94 0·35 0·09 13·63 1·54 0·16 26·08 0·02 0·16 0·07  Ce 0·11 51·99 0·22 0·21 0·10 11·21 1·66 0·39 19·04 0·05 0·40 0·08  Pr 50·54 0·16 0·83 0·21 10·38 0·95 0·07 0·58 14·11 0·07 0·06 0·56  Nd 0·15 54·18 0·30 4·03 0·30 9·34 0·64 1·02 12·19 0·18 0·05 1·06  Sm 64·59 15·20 0·53 8·65 0·65 0·39 1·30 8·85 0·53 2·31 0·49  Eu 1·439 183·304 0·995 20·071 2·416 55·773 1·510 0·716 3·535 23·108 0·760 1·687 3·481 0·263 1·794  Gd 51·31 26·63 0·67 5·22 0·31 0·64 1·86 4·92 0·43 1·44 1·54 2·48  Tb 54·155 0·288 38·144 1·587 6·233 0·211 1·781 2·715 5·873 0·706 6·427 2·327 0·072 5·152  Dy 1·15 45·33 40·45 2·48 7·64 0·17 4·20 2·19 6·18 0·76 13·21 2·85 10·98  Ho 1·170 35·714 37·033 2·033 5·458 0·176 9·176 2·088 6·520 0·489 29·121 2·564 14·689  Er 1·244 23·250 36·063 2·394 3·750 0·073 11·900 1·594 5·806 0·625 39·438 2·350 18·750  Tm 1·000 14·858 30·040 2·826 2·215 15·223 1·603 5·263 0·413 44·656 1·344 0·097 20·972  Yb 1·354 8·590 33·540 3·509 1·820 0·317 19·689 1·093 5·497 0·416 57·888 1·826 30·000  Lu 1·163 6·057 35·122 3·293 1·447 18·049 1·065 3·455 0·276 56·504 1·163 0·207 32·114 The complete dataset is available in Supplementary Data Electronic Appendix 2. Fig. 5. View largeDownload slide Mineral major element composition for the S1 protolith, and each of the melt flux style styles (int. mod., intermediate modification; high mod., high modification). (a) A-site occupancy in amphiboles. Error bars shown to the right, in the same scale as the graph. (b) Ti (c.p.f.u.) in amphiboles. Error bars shown to the right, in the same scale as the graph. (c) Garnet ternary classification diagram. (d) Pistacite content in clinozoisite. (e) Albite content in plagioclase. Fig. 5. View largeDownload slide Mineral major element composition for the S1 protolith, and each of the melt flux style styles (int. mod., intermediate modification; high mod., high modification). (a) A-site occupancy in amphiboles. Error bars shown to the right, in the same scale as the graph. (b) Ti (c.p.f.u.) in amphiboles. Error bars shown to the right, in the same scale as the graph. (c) Garnet ternary classification diagram. (d) Pistacite content in clinozoisite. (e) Albite content in plagioclase. Garnet formed during Style 2 is almandine with composition Alm50–57Pyr18–26Grs16–23Sps2–6 (Fig. 5c;Table 2), with rims that are enriched in pyrope and depleted in grossular by 1–3% respectively. Garnet in partially modified Style 2 samples is almandine–pyrope with Alm36–45Pyr32–42Grs20–21Sps1–3. Garnet formed during Style 3 is almandine with Alm51–60Pyr15–30Grs14–21Sps1–10, with minor zoning of Ca rich rims observed. Garnet formed during Style 4 is almandine–grossular with Alm44–53Pyr17–25Grs25–32Sps2–4, and lacks chemical zoning. The Fe3+ content in Style 2 garnet is 0·06–0·20 cpfu, whereas Style 3 and 4 garnets are within error of no Fe3+ content. Clinozoisite shows minor variations in Fe3+ substitution by Al3+ (Table 2), with little appreciable compositional difference between those formed in the different melt flux styles. Pistacite contents range between 12 and 18 (Fig. 5d), for which Style 1 clinozoisites have higher values (16–18), and Styles 2, 3, and 4 have wider ranges in composition. Plagioclase deformed in S1 has XAb = 58–66 (Table 2; Fig. 5e), and classifies as andesine. In contrast, plagioclase formed in the different melt flux styles is typically more albitic; Styles 1, 2, and 3 plagioclase analyses are classified as andesine or oligoclase (XAb = 53–80, 66–76, and 67–75 respectively), and Style 4 plagioclase has high XAb (77–87) and is either oligoclase or albite. Within Style 2, samples of intermediate modification have a wider compositional range (XAb = 53–76) compared to samples of high modification (XAb = 73–76). Minor K-feldspar occurring in melt pseudomorphs in Style 4 hornblendites are orthoclase (XOr = 98). Sr trace element mapping Synchrotron element mapping shows significant zoning of Sr both at the map scale and within individual grains of plagioclase associated with the different melt flux styles. Zonation of Sr shows no correlation with Ca or Ba, and Na was not detected as it is outside the energy sensitivity of the detector. Within a Style 1 sample, plagioclase (approx. 70 vol.%) can be divided into 1–2 mm wide bands of high- and low-Sr traversing the thin section (Fig. 6a). High-Sr bands closely follow pyroxene–plagioclase boundaries and are spatially associated with replacement microstructures of pargasite and quartz, Style 1 melt–rock interaction products. Coarse-grained plagioclase (0·9–2·0 mm) is zoned in Sr, where high-Sr grain boundaries are adjacent to replacement microstructures forming the high-Sr bands, and low-Sr grain boundaries are next to other plagioclase grains that are also low in Sr. Fig. 6. View largeDownload slide Range of Sr zoning in plagioclase. Note different Sr scales for each map. Mineral abbreviations follow the scheme proposed by Whitney & Evans (2010). (a) Style 1, field of view (FOV) = 18 mm across. (b) Style 2 (high modification), FOV = 21·5 mm across. (c) Style 3 (high modification), FOV = 20 mm across. Fig. 6. View largeDownload slide Range of Sr zoning in plagioclase. Note different Sr scales for each map. Mineral abbreviations follow the scheme proposed by Whitney & Evans (2010). (a) Style 1, field of view (FOV) = 18 mm across. (b) Style 2 (high modification), FOV = 21·5 mm across. (c) Style 3 (high modification), FOV = 20 mm across. Similar map-scale and grain-scale zoning are observed in a Style 2 sample (Fig. 6b). Sr zoning in plagioclase (approx. 40 vol. %) is present throughout the map area and is not spatially associated with microstructures or minerals. The map can be divided into bands of higher and lower Sr enrichment, up to 5 mm wide that are oriented vertically across the map and are sub-parallel with a composite S1 / S3. The high-Sr bands comprise many plagioclase grains (100–800 µm) that are enriched in Sr. Similarly, low-Sr bands comprise many grains that have relatively low Sr. Individual grains in both the high- and low-Sr bands are themselves asymmetrically zoned in Sr, where grains have high- and low-Sr grain boundaries that bear no relation to the grain boundaries of adjacent grains (i.e. a high-Sr boundary of one grain may be adjacent to a low-Sr boundary of a neighbouring grain). Style 3 samples have less pronounced map-scale and grain-scale zoning compared to Styles 1 and 2 (Fig. 6c). Map-scale zoning involves only slight relative changes to the overall Sr content in plagioclase (approx. 40 vol. %). One high-Sr band, approximately 5 mm wide, is associated with coarse-grained garnet on the far right of the map. Many individual grains exhibit homogeneous Sr contents; however, some rare grains have minor asymmetric Sr zoning. Rare earth element mineral compositions Amphiboles from S1 and Style 1 have enriched REE patterns with gentle convex shapes from La to Sm (1–27 times chondrite; Table 2; Fig. 7), small, positive Eu anomalies and flat MREE and HREE (4–19 times chondrite). In comparison, Style 2 amphiboles become progressively depleted in REE, with less pronounced Eu anomalies. Partially modified Style 2 samples have a wide range of depleted LREE values (0·01–1 times chondrite) which increase towards flat, enriched MREE to HREE (1–9 times chondrite). Style 2 samples of high modification have LREE values below detection limits, and increasing patterns from MREE to HREE (0·4–5 times chondrite). Style 3 amphiboles from samples of intermediate and high modification have similar patterns, with wide ranges in LREE from enriched, flat patterns (1–5 times chondrite) to depleted, sloped patterns (0·01–1 times chondrite). MREE and HREE are flat, approximately 1 times chondrite in partially modified samples and slightly more enriched in samples of high modification (1–4 times chondrite). All patterns have weak, positive Eu anomalies. Style 4 amphiboles have increasing trends from La to Eu (0·04–2·6 times chondrite) to flat, enriched MREE and HREE (up to 6 times chondrite). Fig. 7. View largeDownload slide Rare earth element profiles for amphibole, garnet and clinozoisite in the protolith, S1, and each melt–rock interaction style (int. mod., intermediate modification; high mod., high modification). Graphs are formatted consistent with the example shown in the top right. The range in 1σ error for each mineral is shown at the top right. Fig. 7. View largeDownload slide Rare earth element profiles for amphibole, garnet and clinozoisite in the protolith, S1, and each melt–rock interaction style (int. mod., intermediate modification; high mod., high modification). Graphs are formatted consistent with the example shown in the top right. The range in 1σ error for each mineral is shown at the top right. Garnets from Style 2 samples have LREE below detection limits, and the MREE define an increasing trend to HREE (1–56 times chondrite; Table 2; Fig. 7). Style 3 garnets have three different REE patterns. One is defined by increasing REE patterns from depleted LREE to enriched MREE, with an inflection point at Sm (∼25 times chondrite), a weak positive or neutral Eu anomaly and flat, enriched HREE patterns (∼30 times chondrite). A second pattern is comparatively depleted in HREE compared to the first (1–10 times chondrite), with decreasing patterns from Dy to Ho. A third pattern is comparatively depleted in LREE and MREE (below detection limits and 0·6–5 times chondrite, respectively), and has increasing trends from Dy to Lu (up to ∼40 times chondrite). In clinozoisite-bearing Style 4 samples, garnets have two patterns. One pattern is defined by an overall increasing trend from depleted LREE (below detection limits to 6 times chondrite) to enriched HREE (up to 23 times chondrite), with a decrease between Eu and Gd defining a strong, positive Eu anomaly. The second pattern also has an increasing trend from depleted LREE to enriched HREE, however it has either a neutral or weak, positive Eu anomaly, depleted LREE and a slight inflection between La and Ce, forming an inverted spoon-shaped pattern. This second, inverted spoon shaped pattern is typical of garnets in Style 4 samples lacking clinozoisite, which may be further enriched up to 250 times chondrite. Eu anomalies may be either weakly positive, weakly negative or neutral. In all samples, clinozoisite has enriched REE patterns (1–1300 times chondrite; Table 2; Fig. 7) defined by decreasing trends from La to Lu, with positive Eu anomalies of varying magnitude. Clinozoisite grains in Style 2 samples of intermediate modification have strong Eu anomalies, and the HREE tend to be flat rather than sloped. Clinozoisite grains from Style 2 samples of high modification are more enriched than those in samples of intermediate modification and have steeper HREE patterns. Style 3 clinozoisites have a mix of flat and steep HREE trends in samples of both intermediate and high modification and range between 1 and 225 times chondrite. Style 4 clinozoisites also show flat and steep HREE trends, where flat HREE profiles come from cores of grains and steep profiles from rims. TIME-INTEGRATED MELT FLUX ESTIMATION In this section, we estimate the time-integrated flux of melt during each melt flux style. The open system nature of melt–rock interaction and the passage of melt through, rather than injection into, the Pembroke Granulite precludes calculation of the composition of the melt and absolute volume of melt flux. However, the observed increase in the proportion of hydrous phases during each melt flux style can be utilised to interpret a melt-driven H2O metasomatism and then to calculate a minimum time-integrated melt flux. To calculate a minimum time-integrated melt flux we consider a theoretical cube of gabbroic (pyroxene-plagioclase) gneiss, which has essentially no hydrous phases and an H2O content of close to 0 wt %, as the initial composition of the Pembroke Granulite for each melt flux style. The final H2O composition of the cube resulting from melt flux is determined using the proportions of hydrous phases formed during melt–rock interaction (amphibole, clinozoisite, and muscovite), and their average H2O content. Proportions of hydrous phases have been published previously (Daczko et al., 2016; Stuart et al., 2017). The average H2O content of each phase is calculated by averaging the totals of EMP analyses and assuming that any missing mass represents H2O. Calculated H2O contents for amphibole, clinozoisite, and muscovite are 4·09 wt % (n = 210), 3·65 wt % (n = 100), and 4·21 wt % (n = 3), respectively. Combining the proportions of hydrous minerals in each melt flux style produces the final H2O content that ranges from 1·6 wt % for Style 1, to 4·1 wt % for Style 4 (Table 3). Table 3: Calculated estimates of melt flux Style 1 Style 2 Style 3 Style 4 (−czo) Proportion of minerals (%) Amphibole 30 38 43 100 Clinozoisite 9·0 13 3·0 0 Muscovite 0 4·0 0 0 Total H2O (wt %) 1·6 2·2 1·9 4·1 Melt H2O (wt %) m3 of melt flux per m3 rock 2·0 0·77 1·1 0·93 2·0 4·0 0·39 0·55 0·47 1·0 6·0 0·26 0·37 0·31 0·68 Style 1 Style 2 Style 3 Style 4 (−czo) Proportion of minerals (%) Amphibole 30 38 43 100 Clinozoisite 9·0 13 3·0 0 Muscovite 0 4·0 0 0 Total H2O (wt %) 1·6 2·2 1·9 4·1 Melt H2O (wt %) m3 of melt flux per m3 rock 2·0 0·77 1·1 0·93 2·0 4·0 0·39 0·55 0·47 1·0 6·0 0·26 0·37 0·31 0·68 Table 3: Calculated estimates of melt flux Style 1 Style 2 Style 3 Style 4 (−czo) Proportion of minerals (%) Amphibole 30 38 43 100 Clinozoisite 9·0 13 3·0 0 Muscovite 0 4·0 0 0 Total H2O (wt %) 1·6 2·2 1·9 4·1 Melt H2O (wt %) m3 of melt flux per m3 rock 2·0 0·77 1·1 0·93 2·0 4·0 0·39 0·55 0·47 1·0 6·0 0·26 0·37 0·31 0·68 Style 1 Style 2 Style 3 Style 4 (−czo) Proportion of minerals (%) Amphibole 30 38 43 100 Clinozoisite 9·0 13 3·0 0 Muscovite 0 4·0 0 0 Total H2O (wt %) 1·6 2·2 1·9 4·1 Melt H2O (wt %) m3 of melt flux per m3 rock 2·0 0·77 1·1 0·93 2·0 4·0 0·39 0·55 0·47 1·0 6·0 0·26 0·37 0·31 0·68 The time-integrated melt flux is estimated by calculating the volume of melt required to drive the increase in H2O content. Three melts of variable H2O content are considered (2, 4, and 6 wt %), which represent the range of typical values for subduction-related arc melts (Wallace, 2005; Plank et al., 2013 and references therein). The model assumes that all H2O in the melt is partitioned into the new hydrous phases, and that melt flux is homogeneous throughout the cube. In nature, only small amounts of the H2O would be partitioned into the solid phases, and melt flux is likely to be heterogeneously distributed. The model is limited by not considering anhydrous phases, such as plagioclase, that were stable during melt flux Styles 1–3. Therefore, calculated volumes of melt flux are significantly underestimated, especially in the case of Style 4 where unreactive melt flux is interpreted (Daczko et al., 2016). However, they provide a baseline for comparison between the four melt flux styles. Calculated volumes of melt flux range between 0·26 and 2·0 m3 of melt per m3 of rock (Table 3), depending on the initial H2O content of the fluxing melt and the melt flux style. Style 1 involves the smallest volume of melt flux, consistent with the observed partial hydration and replacement of pyroxene grains. Melt flux Styles 2 and 3 involve higher volumes of melt flux compared to Style 1. In contrast, Style 4 involves volumes of melt that are approximately double that calculated for Styles 1, 2, and 3. DISCUSSION Time-integrated volumes of melt flux Calculated time-integrated volumes of melt flux significantly underestimate the true volumes of melt flux, due to assumptions made in the model set up. However, the calculations can be indicative of relative differences between the individual melt flux styles. Styles 1, 2, and 3 have similar time-integrated volumes of melt flux, which stems from the similar assemblages in the modified rock types. Style 1 has the lowest time-integrated melt flux, consistent with the wide spread in mineral compositions (Fig. 7) and variable degree of modification (Stuart et al., 2016). Styles 2 and 3 both have higher time-integrated melt flux volumes, consistent with the homogeneous mineral compositions (Fig. 7), evidence for in situ partial melting (Daczko et al., 2001a, 2001b; Stuart et al., 2017), and associated higher degree of modification by melt–rock interaction. Calculated time-integrated volumes of melt flux for Styles 2 and 3 are likely to be further underestimated due to the in situ partial melting and breakdown of hydrous phases. The time-integrated volume of melt flux during Style 4 is calculated to be approximately double the volumes for Styles 1, 2, and 3 (Table 3). This is the minimum volume required to drive mineralogical changes within the host-rock, forming the armoured channel of hornblendite (Daczko et al., 2016). Continuation of melt flux after a rock has achieved a high degree of modification and reactant minerals are largely exhausted may obscure the ‘true’ time-integrated melt flux. For example, once the hornblendite has formed, melt flux is interpreted to occur without chemical exchange with the host-rock, and so realistic volumes of melt involved in Style 4 may be significantly higher than twice the amount of melt involved in Styles 1, 2, and 3. Equally, given that amphibole and plagioclase are the dominant stable minerals during Styles 2 and 3 melt flux, once these assemblages have formed, significant continued melt flux could also be unreactive, further obscuring the ‘true’ time-integrated melt flux in Styles 2 and 3. Different melt flux styles—different P–T conditions? Garnet-bearing assemblages in mafic to intermediate metaigneous rocks are indicative of high-P metamorphism and, in some cases, partial melting (O'Brien & Rötzler, 2003; Pattison, 2003; De Paoli et al., 2012; Chapman et al., 2017). Previous studies have shown that large volumes of the Fiordland lower crust, including the Pembroke Granulite, experienced burial and recrystallisation at high-P during an arc flare up, a period of voluminous (>100 km3/Ma per km of arc) high-Sr/Y pluton emplacement from 125 to 114 Ma (Bradshaw, 1989; Clarke et al., 2000; Daczko et al., 2002a, 2002b, 2009; Hollis et al., 2003, 2004; Daczko & Halpin, 2009; De Paoli et al., 2009; Stowell et al., 2010, 2014; Chapman et al., 2015; Milan et al., 2016, 2017). The four styles of melt flux observed in the Pembroke Granulite are interpreted as records of migration of these high-Sr/Y magmas (Stuart et al., 2016), albeit with different chemical signatures and variable time-integrated volumes of melt flux. P–T conditions of melt flux could not be calculated as the assemblages required for amphibole thermobarometry (e.g. Otten, 1984; Hammarstrom & Zen, 1986; Hollister et al., 1987; Anderson & Smith, 1995) were not present. However, interpretation of in situ melting (or a lack thereof) can be used to infer temperature differences between melt flux styles. Style 1 samples do not show any evidence for in situ partial melting (Stuart et al., 2016), whereas samples from Styles 2 and 3 with high modification include peritectic garnet evidencing minor (<5 vol. %) amounts of in situ partial melting (Fig. 3e–h; Daczko et al., 2001b; Stuart et al., 2017), suggesting that Styles 2 and 3 occurred at higher temperatures than Style 1. Plagioclase-absent assemblages from Style 4 are distinct from the plagioclase-bearing assemblages from Styles 2 and 3. Crystallisation experiments in mafic to intermediate magmas show that plagioclase is stable at lower temperatures and pressures close to the solidus and that the plagioclase-out line is at higher temperature for intermediate compared to mafic compositions (Green, 1982). Therefore, subtly higher temperatures and/or a more mafic melt accompanied Style 4 melt flux and dramatically changed the character of melt–rock interaction by destabilising plagioclase in the presence of the externally-derived melt. This switch may have been caused by advection of heat during extended or higher time-integrated melt flux of the externally derived melt and/or a shift to a more mafic character of the fluxing melt as the channel became armoured and less reactive. A localised temperature increase is supported by evidence for partial melting in a metre-scale transition zone that separates the Style 4 hornblendite body from the unmelted, precursor two-pyroxene–pargasite gneiss (Daczko et al., 2016). As such, Style 4 also describes a history of melt flux at conditions above the solidus within the Fiordland lower crust, which we infer was likely higher-temperature compared to Styles 1–3. Homogeneous versus heterogeneous whole rock compositions Preservation of S1 assemblages and partial modification during each style of melt–rock interaction provides the opportunity to evaluate the chemical evolution of both the bulk-rock and the mineralogy during different styles of melt–rock interaction. Each style of melt–rock interaction is subtly different; however the presence of a grain boundary network of melt, a bulk hydration effect and the growth of amphibole are common to all four styles. During flux, melt is interpreted to pass through the Pembroke Granulite with only small volumes of melt crystallising in place, implying that any variation in bulk composition is a result of melt-driven metasomatism. Rock types formed by melt–rock interaction during Styles 1, 2, and 3 have bulk-rock compositions indistinct from the precursor two-pyroxene–pargasite gneiss (Fig. 4; excluding volatile content). This includes a few samples of Style 3 that plot in the gabbroic field of Fig. 4a and are interpreted as sampling of slightly more mafic primary components of the Pembroke Granulite. Homogeneous bulk compositions in Styles 1, 2, and 3 overlapping with the precursor two-pyroxene–pargasite gneiss are consistent with nearly isochemical melt–rock interaction. In contrast, major and REE element compositions of Style 4 samples are different to that of the precursor rock (Fig. 4), consistent with melt-driven metasomatism during flux. In the case of Style 4, the melt–rock interaction is likely to have occurred at higher temperatures and involved a larger time-integrated melt flux, facilitating significant reaction and mass exchange during flux. Major element mineral chemistry and partitioning of rare earth elements in the presence of melt Despite significant melt-driven metasomatism during melt–rock interaction, the compositions of minerals from Style 4 are similar to, or overlap, the compositions of minerals from the precursor rock and Styles 1, 2, and 3 (Fig. 5). Variations in mineral compositions between Styles 1, 2, and 3 are difficult to link to melt-driven metasomatism when bulk-rock compositions are homogeneous. However, each melt flux style involved different time-integrated volumes of fluxed melt, and equilibration with variable volumes of melt may have imparted distinct mineral assemblages and compositions for each style. Style 1 involved low time-integrated volumes of melt flux (Table 3). The composition of Style 1 amphiboles (Fig. 5a and b) and distribution of Sr in Style 1 plagioclase (Fig. 6a) varies considerably, indicating a lack of widespread equilibration during melt-rock interaction. On the other hand, small ranges in mineral compositions (Fig. 5), less pronounced Sr zoning in plagioclase (Fig. 6b and c), and higher time-integrated volumes of melt flux (Table 3) for in Styles 2, 3, and 4 are indicative of widespread modification and better equilibration during melt-rock interaction. Homogeneity in major element mineral composition is most likely related to higher time-integrated volumes of melt flux (Table 3) and a higher degree of modification and recrystallisation. Rare earth element (REE) mineral chemistry (Fig. 7) suggests there was significant redistribution of REE between minerals during melt-rock interaction, but little enrichment or depletion occurred at a bulk-rock scale, except for Style 4 (Fig. 4f). This is likely due to the fact that the protolith for the Pembroke Granulite was an igneous rock, with a lack of accessory minerals in significant abundances that strongly partition trace elements and REE. In most cases, minerals have homogeneous REE patterns, despite forming in reactions where they replace precursor minerals with heterogeneous REE contents (e.g. Stuart et al., 2016). Homogenisation of REE patterns in metamorphic products is facilitated by the presence of an intergranular network of melt, providing more efficient diffusion pathways compared to solid-state reactions (Mann, 1980; Lesher, 1994; Acosta-Vigil et al., 2012) and enabling connectivity at an outcrop scale. In contrast to minerals from Styles 1, 2, and 4, the amphibole and garnet in high modification Style 3 samples have a significant variation in their REE patterns; a mix of positive and negative slopes are observed in both garnet HREE and amphibole LREE (Fig. 7). Given that REE are strongly partitioned into clinozoisite (Frei et al., 2003, 2004; Mulrooney & Rivers, 2005; Beard et al., 2006), heterogeneous REE patterns are likely a result of equilibration with varying proportions of clinozoisite. Style 3 is distinguished from the other styles by significant deformation during melt flux, which may localise melt flux (Bauer et al., 2000; Rosenberg & Handy, 2000; Holtzman et al., 2003; Baltzell et al., 2015), limiting connectivity and effectively reducing the volume of rock in chemical communication. Therefore, the spatial extent or connectivity of the melt network may dictate the equilibration volume and resulting homogeneity of REE compositions for each mineral. In this dynamic case, Style 3 may also have experienced multiple episodes of structural reactivation and repeated periods of melt flux, possibly by melts of variable composition. This may also contribute to the heterogeneous nature of mineral REE compositions. The broadly homogeneous REE patterns in melt–rock interaction products, such as garnet, amphibole, and clinozoisite, highlight the extent to which the REE are partitioned between the solid minerals and melt, in this P–T space straddling the boundary between igneous and sub-solidus metamorphic processes. We have evaluated published REE patterns of amphibole and clinozoisite formed under igneous versus sub-solidus metamorphic conditions (Gromet & Silver, 1983; Dalpé & Baker, 2000; El Korh et al., 2009) and compared them to those obtained in this contribution involving melt–rock interaction. Published values for sub-solidus metamorphic clinozoisite/epidote (Fig. 8a) are all enriched relative to chondrite with overlapping, flat REE patterns. The pattern for igneous clinozoisite/epidote is different to the metamorphic patterns, with enriched LREE relative to HREE forming a sloped pattern. This closely matches the average REE patterns of clinozoisite grains formed during melt-rock interaction in the Pembroke Granulite (Fig. 8b). Published amphibole REE patterns can be clearly divided into igneous, which are enriched with humps from La to Dy and flat HREE, versus sub-solidus metamorphic, which are flat and depleted relative to chondrite (Fig. 8c). Amphiboles formed during melt-rock interaction in the Pembroke Granulite are slightly more ambiguous, with characteristics of both types of published pattern. In general, the amphibole grains formed during Styles 1–4 melt flux have igneous-like, flat, enriched patterns from Gd to Lu (Fig. 8d). S1 and Style 1 amphiboles have a LREE hump, like igneous patterns from the literature. It is important to note that the Style 1 amphiboles are forming in an assemblage where only minor amounts of clinozoisite are stable, and garnet is not stable. Thus, amphiboles are not in competition for the REE available. Amphibole in Styles 2, 3, and 4 have more depleted, flat to sloped LREE. Depletion is more characteristic of a metamorphic signature, and as discussed above may be a result of partitioning with varying amounts of clinozoisite. However, unlike the metamorphic patterns, the LREE have a slope from La to Eu, which is closer to the shape of igneous-like patterns. Overall, amphibole and clinozoisite REE patterns share more similarities with published igneous REE patterns. Recrystallisation in the presence of melt, and the large equilibration volume provided by the melt network are two factors which have likely contributed to the formation of these igneous-like REE signatures. Fig. 8. View largeDownload slide Igneous and metamorphic signatures of mineral REE patterns. (a) Published values for igneous and metamorphic clinozoisite/epidote. Data from Gromet & Silver (1983) and El Korh et al. (2009). (b) Average REE patterns for clinozoisite from each melt–rock interaction style. (c) Published values for igneous and metamorphic amphibole. Data from Dalpé & Baker (2000) and El Korh et al. (2009). (d) Average REE patterns for amphibole from S1 and each melt–rock interaction style. Fig. 8. View largeDownload slide Igneous and metamorphic signatures of mineral REE patterns. (a) Published values for igneous and metamorphic clinozoisite/epidote. Data from Gromet & Silver (1983) and El Korh et al. (2009). (b) Average REE patterns for clinozoisite from each melt–rock interaction style. (c) Published values for igneous and metamorphic amphibole. Data from Dalpé & Baker (2000) and El Korh et al. (2009). (d) Average REE patterns for amphibole from S1 and each melt–rock interaction style. Generating igneous-like mineral chemical signatures in a metamorphic rock Differences in melt flux between each style highlight a role for physical processes in the formation of an igneous-like mineral chemical signature during melt–rock interaction. P–T conditions, time scales, and time-integrated volumes of melt flux may control both the stable assemblage and the major element composition of the minerals. In the case of the Pembroke Granulite, temperature is also inferred to play a role in the resulting assemblage of modified rock types, where the major element compositions of minerals formed during melt–rock interaction are determined by the P–T–X conditions of melt and host-rock, as in sub-solidus metamorphic systems. The homogeneity of mineral major element compositions, or the closeness to equilibrium, is here inferred to relate to the time-integrated volume of melt flux, where smaller time-integrated melt fluxes inhibit extensive equilibration, resulting in heterogeneous mineral major element compositions as in Style 1. On the other hand, the REE compositions of minerals are more homogeneous and have igneous-like signatures, where networks of melt provide large equilibration volumes for each mineral. In this case, the strength of the igneous-like signature and degree of homogeneity relies on the spatial distribution and connectivity of the melt network, which may be limited by deformation, such as in Style 3. CONCLUSIONS Melt–rock interaction during melt flux through the root of a magmatic arc has produced new assemblages and broadly homogeneous mineral compositions at an outcrop scale. Time-integrated volume of melt flux and variable equilibration between the host-rock and fluxing melt generates new mineral assemblages; temperature is inferred to control major element compositions. Igneous-like REE patterns in minerals formed in the presence of a grain boundary network of melt, and are identified as geochemical signatures recording the former flux of melt. The network of melt enhanced equilibration volumes and the mobility of REE at an outcrop scale. The time-integrated melt flux and spatial distribution of melt networks influence the degree of homogeneity and strength of the igneous-like mineral REE signature. Comparison of the four different melt flux styles examined in this study highlights the role that the physical characteristics of melt flux plays in generating the geochemical signatures. ACKNOWLEDGEMENTS We thank the Department of Conservation, New Zealand for permission to visit and sample localities in the Fiordland National Park. Editorial handling by G. Zellmer and reviews by S. Harley, M Pistone, and O. Jagoutz helped to improve this paper. This is contribution 1155 from the ARC Centre of Excellence for Core to Crust Fluid Systems (www.ccfs.mq.edu.au) and 1221 from the GEMOC Key Centre (www.gemoc.mq.edu.au). FUNDING This work was supported by an Australian Research Council Future Fellowship to S.P. (grant number FT110100070); an Australian Research Council Discovery Project to S.P. and N.R.D. (grant number DP120102060); and Australian Government Research Training Program Scholarships to C.A.S. and U.M. Part of this research was undertaken on the X-Ray Fluorescence Microscopy beamline at the Australian Synchrotron, Victoria, Australia. This work was supported by the Multi-modal Australian ScienceS Imaging and Visualisation Environment (MASSIVE) (www.massive.org.au). This study used instrumentation funded by ARC LIEF and DEST Systemic Infrastructure Grants, Macquarie University and Industry. SUPPLEMENTARY DATA Supplementary data for this paper, including whole rock composition, and mineral major and REE compositions are available at Journal of Petrology online. REFERENCES Acosta-Vigil A. , London D. , Morgan G. B. VI . ( 2012 ). Chemical diffusion of major components in granitic liquids: implications for the rates of homogenization of crustal melts . Lithos 153 , 308 – 323 . Google Scholar CrossRef Search ADS Allibone A. H. , Jongens R. , Turnbull I. M. , Milan L. 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Journal of PetrologyOxford University Press

Published: Mar 10, 2018

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