TY - JOUR
AU - Godard,, Marguerite
AB - Abstract Although the axial melt lens (AML) beneath fast-spreading mid-ocean ridges has been detected by seismic reflection for decades, its nature and role in the accretion of lower oceanic crust and the evolution and eruption of mid-ocean ridge basalts (MORB) are still poorly constrained. Plutonic rocks consisting of quartz-bearing gabbros, diorites and tonalites, which might represent the upper part of a fossilized AML, have for the first time been recovered from an intact fast-spreading oceanic crust section by Integrated Ocean Drilling Program (IODP) Hole 1256D. Whole-rock major elements show a wide and continuous compositional range (e.g. Mg# 24–70) and apparent enrichments in Ti and Fe at intermediate MgO contents (4–6 wt %). Trace element characteristics are coherent for the different lithology groups defined by petrography and mineral modes; that is, gabbro, clinopyroxene-rich diorite, amphibole-rich or oxide-rich diorite and tonalite. The gabbros and diorites are consistent with modeled products of MORB fractional crystallization, composed of mixed melt and cumulate in varying ratios. Modeled trace elements (especially with respect to Eu) support a model in which the tonalites originated from low-degree partial melting of the sheeted dikes overlying the AML, rather than extreme fractional crystallization. Enrichments in rare earth elements (REE) in clinopyroxenes from the gabbroic and dioritic intrusive rocks suggest strong assimilation of REE-rich tonalitic components by evolved MORB magmas. Hydrothermal alteration was pervasive during cooling of the plutonic system, which can be traced by petrography, mineral compositions and bulk-rock geochemistry. The upper part of AML, largely composed of low-density and high-viscosity felsic magmas, may serve as a barrier to eruptible MORB melts in the lower part of AML. INTRODUCTION Oceanic crust formed at mid-ocean ridges (MOR) covers two-thirds of the Earth’s surface, representing the largest magma generation factory on Earth. The mid-ocean ridge basalts (MORB) are the most studied lithologies, with the aim of revealing the mechanisms responsible for the generation, transport and eruption of melts involved in the accretion of oceanic crust. They are, however, generally believed to differ significantly from the primary melts generated in the source regions in the mantle (O’Hara, 1968; Klein & Langmuir, 1987; Niu & Batiza, 1991). A series of intra-mantle and/or intra-crustal processes, from melt transport in the shallow mantle to ridge volcanism, must have occurred, resulting in the evolution of the range of MOR magma compositions (e.g. Bédard, 1993; Rubin et al., 2009). Therefore, attempts to use MORB geochemical features to infer magma source affinities are not straightforward and are very challenging. Numerous efforts have been addressed at constraining mantle melting temperatures and depths (Klein & Langmuir, 1987; Niu & O’Hara, 2008), ridge thermal structures (e.g. Rubin & Sinton, 2007), crustal components recycled into the mantle (e.g. Class & le Roex, 2008), and the tectonic setting of ophiolites (Pearce, 2014). A variety of processes have been proposed to explain the discrepancy between observed MORB geochemistry and that generated solely by mantle melting. The three commonly discussed first-order models are: fractional crystallization (Grove et al., 1992; Herzberg, 2004; O’Neill & Jenner, 2012; Wanless & Shaw, 2012); melt–rock interactions in the shallow mantle (Coogan et al., 2002a; Dick et al., 2010; Sanfilippo et al., 2013) or at the mantle–crust interface (Drouin et al., 2009, 2010; Godard et al., 2009) or in the lower crust (Lissenberg & Dick, 2008; Lissenberg et al., 2013); contamination by crustal materials (Coogan et al., 2003; le Roux et al., 2006; France et al., 2009; Koepke et al., 2011). Despite numerous case studies, there is at present no consensus regarding which end-member process has exerted a major influence on MORB evolution at the global scale, which reflects our lack of knowledge on the mechanism of melt evolution and transportation from the mantle source to the eruption on the seafloor. Fractional crystallization and contamination by crustal materials can occur simultaneously in modifying magma compositions, especially in generating felsic magmas (e.g. Amri et al., 1996; Wanless et al., 2010; Freund et al., 2013). Alternatively, partial melting of previously hydrothermally altered sheeted dikes or gabbros might also be important in forming felsic melts at MORs (Gillis & Coogan, 2002; Koepke et al., 2004, 2007; France et al., 2010, 2014; Bindeman et al., 2012; Erdmann et al., 2015). Subsequent contamination of MORB by these felsic components might be common and responsible for their observed enrichments in volatile elements (e.g. Cl, B, K) and low δ18O values (Gillis et al., 2003; Wanless et al., 2011; Freund et al., 2013; Grimes et al., 2013; Sharapov et al., 2013). The volumetrically minor, but genetically important, felsic lavas, dikes or intrusions from the oceanic crust and ophiolites are usually diorites, tonalites and trondhjemites, collectively referred to as oceanic plagiogranites [see definition by Koepke et al. (2007)]. They mark the process by which silicic components can be produced independently from mafic oceanic crust, a process that might have operated in Archean times to form the earliest continents on Earth (e.g. Rollinson, 2008). Here, we present a comprehensive whole-rock and mineral major and trace elemental dataset for a variety of felsic lithologies within the plutonic part of the MOR system recovered in situ by Integrated Ocean Drilling Program (IODP) Hole 1256D, where, for the first time, the dike–gabbro transition zone of intact oceanic crust was penetrated and sampled, allowing the processes occurring within the axial melt lens (AML) to be directly studied. The purpose of this work is to better understand the petrological and geochemical complexity within the AML and its conducting boundary layer, and further to decipher the various models of petrogenesis of felsic rocks at fast-spreading MORs. THE AXIAL MELT LENS AT FAST-SPREADING RIDGES Seismic reflection profiles at fast-spreading ridges (Detrick et al., 1987; Kent et al., 1990; Toomey et al., 1990; Singh et al., 1998) have revealed a predominantly molten zone at shallow depths (1–2 km below the seafloor), with a thickness of 50–100 m and a width of less than 2 km. This shallow, small, melt-dominated zone (i.e. the AML) is generally interpreted to serve as a shallow crystallization pool and density filter, from which evolved low-density lavas erupt and where crystallized phases sink to form a lower crustal axial mush zone (e.g. Sinton & Detrick, 1992); this interpretation is known as the ‘gabbro glacier’ model (Henstock et al., 1993; Phipps Morgan & Chen, 1993; Quick & Denlinger, 1993; Coogan et al., 2002b). In contrast, an alternative ‘sheeted sill’ model considers that mantle-derived magmas crystallize pervasively along their route to the melt lens to form the whole lower crust via multiple sill intrusions, without comprehensive crystal settling (Kelemen et al., 1997; Korenaga & Kelemen, 1997). In this model, the uppermost melt lens might be mainly occupied by highly evolved magmas derived from crystallization in the deeper crust (Natland & Dick, 1996; Pan & Batiza, 2003). The ‘gabbro glacier’ model requires rapid cooling at the surface of the AML in order to extract a large amount of latent heat (Maclennan et al., 2005; Zhang et al., 2014), whereas a requirement for the ‘sheeted sill’ model is that significant heat must be removed from the lower crust (e.g. by focused fluid flow in the lower crust; Coogan et al., 2006). A hybrid model integrating the above competitive end-member models, which requires crystallization both within the uppermost melt lens and at much greater depths, might be more plausible for natural cases as suggested by abundant evidence from multiple disciplines (Grove et al., 1992; Boudier et al., 1996; Crawford et al., 1999; Coogan et al., 2002a; Maclennan et al., 2005; Rubin & Sinton, 2007; Wanless & Shaw, 2012; Marjanovic et al., 2014). Therefore, although the AML has been detected for several decades, its nature and role in the accretion and evolution of the lower oceanic crust and eruption of MORB is still controversial Although numerous geological, geophysical and geochemical studies concerning MORs evoke a shallow AML, direct sampling of it in intact oceanic crust was not achieved until the drilling of Hole 1256D by the Ocean Drilling Program (ODP) and Integrated Ocean Drilling Program (IODP) through a multi-expedition campaign in the equatorial Pacific (Teagle et al., 2006, 2012). Hole 1256D penetrated, for the first time, the whole upper oceanic crust down into the uppermost gabbros, which can be regarded as parts of the frozen AML (Wilson et al., 2006; Koepke et al., 2011). Drilling encountered a variety of felsic intrusive rocks located in the gabbros and in the roof rocks of the AML, which offer a unique opportunity to look in situ within the AML to understand the prevailing processes (e.g. magma differentiation, crustal re-melting and hydrothermal alteration) contributing to the accretion of the oceanic crust. GEOLOGICAL BACKGROUND IODP Hole 1256D (6°44·2’N, 91°56·1’W) is located in 15 Ma oceanic lithosphere of the Cocos Plate, which was formed at the East Pacific Rise during superfast spreading (Fig. 1, full spreading rate ∼220 mm a–1; Cande & Kent, 1995; Wilson et al., 2006). The drilling was initiated by ODP Leg 206 and continued by IODP Expeditions 309, 312 and 335. At the end of Expedition 335, Hole 1256D had reached a total depth of 1521·6 m below seafloor (mbsf) and had penetrated the whole upper oceanic crust (Fig. 2), composed mainly of extrusive lavas (∼800 m thick) and a sheeted dike complex (∼350 m thick), and finally into the uppermost gabbros (Teagle et al., 2006, 2012). The gabbro intrusions, believed to be part of the frozen AML, were first encountered at 1406·6 mbsf. This shallow depth is consistent with seismic surveys from MORs and the prediction that the depth of the melt lens is negatively correlated with spreading rate (Wilson et al., 2006). Fig. 1. Open in new tabDownload slide Map of the eastern Pacific Ocean showing the location of IODP Hole 1256D. Also shown is the age of the seafloor around the East Pacific Rise, which indicates that the oceanic crust at Hole 1256D was formed by superfast spreading with a full spreading rate of c. 210 mm a–1 (after Wilson et al., 2006). Fig. 1. Open in new tabDownload slide Map of the eastern Pacific Ocean showing the location of IODP Hole 1256D. Also shown is the age of the seafloor around the East Pacific Rise, which indicates that the oceanic crust at Hole 1256D was formed by superfast spreading with a full spreading rate of c. 210 mm a–1 (after Wilson et al., 2006). Fig. 2. Open in new tabDownload slide Stratigraphy of IODP Hole 1256D, showing the relation between gabbro plutonic rocks, felsic rocks (diorite and tonalite) and basaltic dikes (modified after Teagle et al., 2012). Two gabbro intrusive rocks, denoted Gabbro 1 and Gabbro 2, are sandwiched by roof hornfelsic dikes and dike screens. Fig. 2. Open in new tabDownload slide Stratigraphy of IODP Hole 1256D, showing the relation between gabbro plutonic rocks, felsic rocks (diorite and tonalite) and basaltic dikes (modified after Teagle et al., 2012). Two gabbro intrusive rocks, denoted Gabbro 1 and Gabbro 2, are sandwiched by roof hornfelsic dikes and dike screens. Two gabbro intrusions were penetrated at depths between 1406 and 1495 mbsf (Fig. 2). All samples display non-layered, isotropic textures composed of varying textural domains ranging from subophitic to coarse-grained granular (Koepke et al., 2011). It has been demonstrated that the gabbros have undergone in situ crystallization and hybridization between evolved and relatively primitive magmas (Koepke et al., 2011; Teagle et al., 2012). Both gabbro intrusions are sandwiched by, and also enclose xenoliths of, hornfelsic basalt (also named granoblastic basalt), which is the product of high-temperature recrystallization of sheeted dikes at two-pyroxene hornfels facies conditions, with heat provided by an up-moving AML (Koepke et al., 2008; Alt et al., 2010). Estimated peak equilibrium temperature for the contact metamorphism was up to c. 1050°C (Koepke et al., 2008; Zhang et al., 2014), a temperature that is high enough to trigger partial melting in basaltic rocks, leading to felsic melts and residues consisting of two pyroxenes, plagioclase and Fe–Ti oxides (Erdmann et al., 2015); the former might resemble the typical felsic rocks of the oceanic crust, whereas the latter might be analogous to the observed two-pyroxene hornfels both in modern oceanic crust and in ophiolites (Coogan et al., 2003; Gillis, 2008; Koepke et al., 2008; France et al., 2009, 2010, 2014). Extensive studies have been conducted for Hole 1256D on the geophysical, petrological, mineralogical and geochemical characteristics of lavas, dikes, gabbros and hydrothermal systems (Dziony et al., 2008, 2014; Koepke et al., 2008, 2011; Sano et al., 2008, 2011; France et al., 2009; Neo et al., 2009; Tominaga et al., 2009; Yamazaki et al., 2009; Alt et al., 2010; Alt & Shanks, 2011; Shilobreeva et al., 2011; Gao et al., 2012; Höfig et al., 2014; Veloso et al., 2014; Zhang et al., 2014; Harris et al., 2015; Koepke & Goetze, 2015), but little attention has been focused on the felsic dikelets and patchy intrusions that occur ubiquitously in the dike–gabbro transition zone. The drilled basaltic rocks display strong compositional variations from a fairly primitive high-Mg endmember to a high-Fe and low-Ca endmember (Fig. 3); they span a wide range of Mg# [100 × molar Mg/(Mg + Fe)] from 40 to 70. This overall evolutionary trend implies that pervasive magmatic fractional crystallization has occurred, an observation consistent with the widely varying whole-rock compositions of the gabbros (Wilson et al., 2006; Neo et al., 2009). Furthermore, systematic variations in texture and mineral composition observed in the gabbros suggest strong local in situ fractionation within a large temperature range (Koepke et al., 2011). The most primitive dikes (i.e. having highest Mg#) from Hole 1256D are similar to the East Pacific Rise present-day melt inclusions and glasses in terms of major elements (Fig. 3), probably indicating similar parental magmas compositions for these two magmatic systems with different ages. Fig. 3. Open in new tabDownload slide Major element compositions (Fe2O3T vs MgO and CaO vs Mg#) for the various lithologies recovered at IODP Hole 1256D. Fe2O3T is total iron expressed as Fe2O3. Data are compiled from Neo et al. (2009), Yamazaki et al. (2009), Sano et al. (2011), Expedition 335 Scientists (2012) and this study. It should be noted that the felsic lithologies discussed in this study are part of the ‘intrusive rock’ group. The average composition (±1SD) of East Pacific Rise melt inclusions and glasses (EPR MI & GL; data from Wanless & Shaw, 2012) is plotted for comparison. Fig. 3. Open in new tabDownload slide Major element compositions (Fe2O3T vs MgO and CaO vs Mg#) for the various lithologies recovered at IODP Hole 1256D. Fe2O3T is total iron expressed as Fe2O3. Data are compiled from Neo et al. (2009), Yamazaki et al. (2009), Sano et al. (2011), Expedition 335 Scientists (2012) and this study. It should be noted that the felsic lithologies discussed in this study are part of the ‘intrusive rock’ group. The average composition (±1SD) of East Pacific Rise melt inclusions and glasses (EPR MI & GL; data from Wanless & Shaw, 2012) is plotted for comparison. PETROGRAPHY The series of evolved intrusive plutonic rocks sampled at IODP 1256D includes clinopyroxene-rich diorite (cpx-rich diorite), amphibole-rich diorite (amp-rich diorite), Fe–Ti oxide-rich diorite (oxide-rich diorite) and tonalite. Their original IODP descriptions, depths below seafloor, and abbreviations used in this work are listed in Table 1. For the definitions of rock names and petrographic qualifiers we follow Teagle et al. (2012). The felsic lithologies may be associated either with one of the two gabbro screens or with intrusions within the two-pyroxene hornfelses within the depth interval between 1404 and 1522 mbsf (Fig. 2). The dioritic rocks form centimeter- to meter-sized discrete irregular intrusions or intrusive patches always with diffuse, sutured contacts to the host-rocks, implying that the hosting gabbros were in a state of a near solidus crystal mush during intrusion. Most tonalites form millimeter- to centimeter-sized veins within the hornfelses or within the gabbroic–dioritic rocks, without any chilled margins and with sutured contacts to the host-rocks, implying that the felsic melts intruded along fractured networks when the wall-rocks were still very hot, probably under near-solidus conditions. Details on contacts and intrusion relations have been given by Teagle et al. (2006, 2012). Generally, no shape-preferred orientation for any mineral is present, indicating that the intruding magma was mostly a melt (∼100% liquid) and that it underwent no, or very weak, deformation during crystallization. The petrography of the felsic lithologies varies greatly from rock to rock. Common features in mineral assemblage and texture allow grouping into lithological classes as described below. Table 1 List of samples from the IODP Hole 1256D investigated in this study Exp. . Core . Sec. . Top . Bottom . Piece . Fraction . Depth . Sample . Lithology . . . . (cm) . (cm) . . . (mbsf) . name . group . 312 212 1 26 28 6 1 1404·4 212–1-P6_1 tonalite 312 212 1 26 28 6 2 1404·4 212–1-P6_2 hornfels 312 214 1 38 41 10 — 1411·3 214–1-P10 amp-rich diorite 312 214 1 68 70 14 — 1411·6 214–1-P14 tonalite 312 214 1 138 141 27 1 1412·3 214–1-P27_1 amp-rich diorite 312 214 1 138 141 27 2 1412·3 214–1-P27_2 granular gabbro 312 216 1 69 71 13 — 1418·6 216–1-P13 granular gabbro 312 216 1 93 99 20 1 1418·9 216–1-P20_1 oxide-rich diorite 312 216 1 93 99 20 2 1418·9 216–1-P20_2 granular gabbro 312 217 1 88 94 21 — 1422·5 217–1-P21 amp-rich diorite 312 217 1 95 98 22 1 1422·6 217–1-P22_1 oxide-rich diorite 312 217 1 95 98 22 2 1422·6 217–1-P22_2 granular gabbro 312 227 1 87 81 14 — 1469·4 227–1-P14 albitite 312 227 2 12 20 2 — 1470·2 227–2-P2 cpx-rich diorite 312 227 2 78 80 12 1 1470·8 227–2–12_1 albitite 312 227 2 78 80 12 2 1470·8 227–2-P12_2 hornfels 312 230 1 16 24 5 — 1483·2 230–1-P5 cpx-rich diorite 312 230 1 122 130 11 1 1484·3 230–1-P11_1 cpx-rich diorite 312 230 1 122 130 11 2 1484·3 230–1-P11_2 granular gabbro 335 235 1 23·5 24·5 5 1 1507·3 235–1-P5_1 tonalite 335 235 1 23·5 24·5 5 2 1507·3 235–1-P5_2 oxide-rich diorite 335 235 1 30·5 33·5 7 — 1507·4 235–1-P7 oxide-rich diorite 335 — R11 A 1 — R11_1 amp-rich diorite 335 — R11 A 2 — R11_2 hornfels 335 — R12 P — — R12-P amp-rich diorite 335 — R14 A 1 — R14-A_1 amp-rich diorite 335 — R14 A 2 — R14-A_2 hornfels 335 — R14 C — — R14-C amp-rich diorite 335 — R14 D — — R14-D amp-rich diorite 335 — R17 A — — R17 oxide-rich diorite 335 — R20 E — — R20 albitite Exp. . Core . Sec. . Top . Bottom . Piece . Fraction . Depth . Sample . Lithology . . . . (cm) . (cm) . . . (mbsf) . name . group . 312 212 1 26 28 6 1 1404·4 212–1-P6_1 tonalite 312 212 1 26 28 6 2 1404·4 212–1-P6_2 hornfels 312 214 1 38 41 10 — 1411·3 214–1-P10 amp-rich diorite 312 214 1 68 70 14 — 1411·6 214–1-P14 tonalite 312 214 1 138 141 27 1 1412·3 214–1-P27_1 amp-rich diorite 312 214 1 138 141 27 2 1412·3 214–1-P27_2 granular gabbro 312 216 1 69 71 13 — 1418·6 216–1-P13 granular gabbro 312 216 1 93 99 20 1 1418·9 216–1-P20_1 oxide-rich diorite 312 216 1 93 99 20 2 1418·9 216–1-P20_2 granular gabbro 312 217 1 88 94 21 — 1422·5 217–1-P21 amp-rich diorite 312 217 1 95 98 22 1 1422·6 217–1-P22_1 oxide-rich diorite 312 217 1 95 98 22 2 1422·6 217–1-P22_2 granular gabbro 312 227 1 87 81 14 — 1469·4 227–1-P14 albitite 312 227 2 12 20 2 — 1470·2 227–2-P2 cpx-rich diorite 312 227 2 78 80 12 1 1470·8 227–2–12_1 albitite 312 227 2 78 80 12 2 1470·8 227–2-P12_2 hornfels 312 230 1 16 24 5 — 1483·2 230–1-P5 cpx-rich diorite 312 230 1 122 130 11 1 1484·3 230–1-P11_1 cpx-rich diorite 312 230 1 122 130 11 2 1484·3 230–1-P11_2 granular gabbro 335 235 1 23·5 24·5 5 1 1507·3 235–1-P5_1 tonalite 335 235 1 23·5 24·5 5 2 1507·3 235–1-P5_2 oxide-rich diorite 335 235 1 30·5 33·5 7 — 1507·4 235–1-P7 oxide-rich diorite 335 — R11 A 1 — R11_1 amp-rich diorite 335 — R11 A 2 — R11_2 hornfels 335 — R12 P — — R12-P amp-rich diorite 335 — R14 A 1 — R14-A_1 amp-rich diorite 335 — R14 A 2 — R14-A_2 hornfels 335 — R14 C — — R14-C amp-rich diorite 335 — R14 D — — R14-D amp-rich diorite 335 — R17 A — — R17 oxide-rich diorite 335 — R20 E — — R20 albitite For some samples two fractions with different lithology were separated from one rock piece by electronic saw and hand. Samples without a core number were collected from junk baskets and thus their depths were not available. Mineral abbreviations in lithology groups: amp, amphibole; cpx, clinopyroxene; oxide, Fe–Ti oxide; qtz, quartz. Table 1 List of samples from the IODP Hole 1256D investigated in this study Exp. . Core . Sec. . Top . Bottom . Piece . Fraction . Depth . Sample . Lithology . . . . (cm) . (cm) . . . (mbsf) . name . group . 312 212 1 26 28 6 1 1404·4 212–1-P6_1 tonalite 312 212 1 26 28 6 2 1404·4 212–1-P6_2 hornfels 312 214 1 38 41 10 — 1411·3 214–1-P10 amp-rich diorite 312 214 1 68 70 14 — 1411·6 214–1-P14 tonalite 312 214 1 138 141 27 1 1412·3 214–1-P27_1 amp-rich diorite 312 214 1 138 141 27 2 1412·3 214–1-P27_2 granular gabbro 312 216 1 69 71 13 — 1418·6 216–1-P13 granular gabbro 312 216 1 93 99 20 1 1418·9 216–1-P20_1 oxide-rich diorite 312 216 1 93 99 20 2 1418·9 216–1-P20_2 granular gabbro 312 217 1 88 94 21 — 1422·5 217–1-P21 amp-rich diorite 312 217 1 95 98 22 1 1422·6 217–1-P22_1 oxide-rich diorite 312 217 1 95 98 22 2 1422·6 217–1-P22_2 granular gabbro 312 227 1 87 81 14 — 1469·4 227–1-P14 albitite 312 227 2 12 20 2 — 1470·2 227–2-P2 cpx-rich diorite 312 227 2 78 80 12 1 1470·8 227–2–12_1 albitite 312 227 2 78 80 12 2 1470·8 227–2-P12_2 hornfels 312 230 1 16 24 5 — 1483·2 230–1-P5 cpx-rich diorite 312 230 1 122 130 11 1 1484·3 230–1-P11_1 cpx-rich diorite 312 230 1 122 130 11 2 1484·3 230–1-P11_2 granular gabbro 335 235 1 23·5 24·5 5 1 1507·3 235–1-P5_1 tonalite 335 235 1 23·5 24·5 5 2 1507·3 235–1-P5_2 oxide-rich diorite 335 235 1 30·5 33·5 7 — 1507·4 235–1-P7 oxide-rich diorite 335 — R11 A 1 — R11_1 amp-rich diorite 335 — R11 A 2 — R11_2 hornfels 335 — R12 P — — R12-P amp-rich diorite 335 — R14 A 1 — R14-A_1 amp-rich diorite 335 — R14 A 2 — R14-A_2 hornfels 335 — R14 C — — R14-C amp-rich diorite 335 — R14 D — — R14-D amp-rich diorite 335 — R17 A — — R17 oxide-rich diorite 335 — R20 E — — R20 albitite Exp. . Core . Sec. . Top . Bottom . Piece . Fraction . Depth . Sample . Lithology . . . . (cm) . (cm) . . . (mbsf) . name . group . 312 212 1 26 28 6 1 1404·4 212–1-P6_1 tonalite 312 212 1 26 28 6 2 1404·4 212–1-P6_2 hornfels 312 214 1 38 41 10 — 1411·3 214–1-P10 amp-rich diorite 312 214 1 68 70 14 — 1411·6 214–1-P14 tonalite 312 214 1 138 141 27 1 1412·3 214–1-P27_1 amp-rich diorite 312 214 1 138 141 27 2 1412·3 214–1-P27_2 granular gabbro 312 216 1 69 71 13 — 1418·6 216–1-P13 granular gabbro 312 216 1 93 99 20 1 1418·9 216–1-P20_1 oxide-rich diorite 312 216 1 93 99 20 2 1418·9 216–1-P20_2 granular gabbro 312 217 1 88 94 21 — 1422·5 217–1-P21 amp-rich diorite 312 217 1 95 98 22 1 1422·6 217–1-P22_1 oxide-rich diorite 312 217 1 95 98 22 2 1422·6 217–1-P22_2 granular gabbro 312 227 1 87 81 14 — 1469·4 227–1-P14 albitite 312 227 2 12 20 2 — 1470·2 227–2-P2 cpx-rich diorite 312 227 2 78 80 12 1 1470·8 227–2–12_1 albitite 312 227 2 78 80 12 2 1470·8 227–2-P12_2 hornfels 312 230 1 16 24 5 — 1483·2 230–1-P5 cpx-rich diorite 312 230 1 122 130 11 1 1484·3 230–1-P11_1 cpx-rich diorite 312 230 1 122 130 11 2 1484·3 230–1-P11_2 granular gabbro 335 235 1 23·5 24·5 5 1 1507·3 235–1-P5_1 tonalite 335 235 1 23·5 24·5 5 2 1507·3 235–1-P5_2 oxide-rich diorite 335 235 1 30·5 33·5 7 — 1507·4 235–1-P7 oxide-rich diorite 335 — R11 A 1 — R11_1 amp-rich diorite 335 — R11 A 2 — R11_2 hornfels 335 — R12 P — — R12-P amp-rich diorite 335 — R14 A 1 — R14-A_1 amp-rich diorite 335 — R14 A 2 — R14-A_2 hornfels 335 — R14 C — — R14-C amp-rich diorite 335 — R14 D — — R14-D amp-rich diorite 335 — R17 A — — R17 oxide-rich diorite 335 — R20 E — — R20 albitite For some samples two fractions with different lithology were separated from one rock piece by electronic saw and hand. Samples without a core number were collected from junk baskets and thus their depths were not available. Mineral abbreviations in lithology groups: amp, amphibole; cpx, clinopyroxene; oxide, Fe–Ti oxide; qtz, quartz. Gabbros Gabbros have been described in detail by Teagle et al. (2006, 2012) and Koepke et al. (2011). A striking feature of the gabbro sections is the occurrence of various domains with different mineralogy and textures, which, at the thin-section scale, are mainly represented by intensely interfingered sub-ophitic and granular domains. The sub-ophitic gabbros contain millimeter-sized poikilitic clinopyroxenes that enclose skeletal plagioclase, indicating fast crystal growth. The granular (medium-grained) gabbros are composed mainly of prismatic plagioclase and granular clinopyroxene and orthopyroxene, as well as occasional Fe–Ti oxides. Interstitial late-stage minerals are common, including albite-rich plagioclase, apatite, and quartz. Diorites Clinopyroxene-rich diorites have seriate inequigranular textures, composed of medium-grained granular pyroxene, mainly clinopyroxene and minor orthopyroxene, euhedral plagioclase laths and interstitial plagioclase and quartz (Fig. 4a and b). They can alternatively be termed oxide gabbronorite. A characteristic feature of the clinopyroxenes is that they contain abundant micro-inclusions of Fe–Ti oxides. Plagioclase shows strong normal zoning with a range in composition from An65 to An40. It is likely that both pyroxene and plagioclase (cores) are part of an early stage cumulate and crystallized simultaneously. Later, interstitial albite-rich plagioclase, apatite and quartz formed among the framework of clinopyroxene and plagioclase. The intrusive contact between a cpx-rich diorite and the host-rock, a former basalt, which has been recrystallized to hornfels and pervasively altered, is shown in Fig. 4a. Hydrothermal overprinting of the hornfels is more pronounced close to the contact with the cpx-rich diorite. This implies that fluids migrated preferentially along the lithological boundary at low temperature or, more probably, that some fluids were released during the crystallization of the cpx-rich diorite, triggering alteration of the wall-rocks. Fig. 4. Open in new tabDownload slide Representative photomicrographs of felsic lithologies. (a) Cpx-rich diorite (sample 230–1-P5; lower part) in contact with a heavily altered fine-grained, pyroxene hornfels (upper part). (b) Close-up view of the cpx-rich diorite shown in (a). In the hornfels, clinopyroxene has been totally altered to a mix of fibrous actinolite, chlorite and oxide (dark green grains). In the cpx-rich diorite, medium-grained granular plagioclase and clinopyroxene both show near euhedral crystal shapes, with albite-rich plagioclase as interstitial growth. Numerous tiny oxides occur as inclusions in clinopyroxene. (c) Oxide-rich diorite with about 15% Fe–Ti oxides in the mode (sample 235–1-P7). Fe–Ti oxides are pervasively distributed, but no preferential orientation is observable. (d) Close-up view of the oxide-rich diorite shown in (c). Pervasively altered plagioclases laths form a framework with granular Fe–Ti oxides and granular to prismatic clinopyroxene. Acicular apatite is visible. (e) Amp-rich diorite (sample 214–1-P10). Pseudomorphs of actinolite aggregates after subhedral magmatic hornblende surrounded by a framework of plagioclase and quartz. (f) Close-up view of the amp-rich diorite shown in (e). Tabular, moderately to strongly altered plagioclase and former subhedral, magmatic amphibole associated with granular Fe–Ti oxides form the framework of the rock. The amphibole is now composed of aggregates of fibrous actinolite intergrown with tiny oxide grains of secondary origin. (g) Cpx-bearing tonalite (sample 235–1-P5_1), which occurs as a vein intruding into the host oxide-rich diorite (sample 235–1-P5_2). (h) Close-up view of the tonalite in (g). (i) Altered rock (pseudo felsic rock) (sample 227–1-P14), which is characterized by a framework of granular albite intergrown with greenish chlorite, yellow–green poikiloblastic epidote and quartz. (j) Close-up view of the altered rock shown in (i), showing poikiloblastic epidote with a radial structure. Mineral abbreviations: ab, albite; act, actinolite; chl, chlorite; cpx, clinopyroxene; oxi, Fe–Ti oxide; plg, plagioclase; qtz, quartz; ru, rutile; ttn, titanite. Fig. 4. Open in new tabDownload slide Representative photomicrographs of felsic lithologies. (a) Cpx-rich diorite (sample 230–1-P5; lower part) in contact with a heavily altered fine-grained, pyroxene hornfels (upper part). (b) Close-up view of the cpx-rich diorite shown in (a). In the hornfels, clinopyroxene has been totally altered to a mix of fibrous actinolite, chlorite and oxide (dark green grains). In the cpx-rich diorite, medium-grained granular plagioclase and clinopyroxene both show near euhedral crystal shapes, with albite-rich plagioclase as interstitial growth. Numerous tiny oxides occur as inclusions in clinopyroxene. (c) Oxide-rich diorite with about 15% Fe–Ti oxides in the mode (sample 235–1-P7). Fe–Ti oxides are pervasively distributed, but no preferential orientation is observable. (d) Close-up view of the oxide-rich diorite shown in (c). Pervasively altered plagioclases laths form a framework with granular Fe–Ti oxides and granular to prismatic clinopyroxene. Acicular apatite is visible. (e) Amp-rich diorite (sample 214–1-P10). Pseudomorphs of actinolite aggregates after subhedral magmatic hornblende surrounded by a framework of plagioclase and quartz. (f) Close-up view of the amp-rich diorite shown in (e). Tabular, moderately to strongly altered plagioclase and former subhedral, magmatic amphibole associated with granular Fe–Ti oxides form the framework of the rock. The amphibole is now composed of aggregates of fibrous actinolite intergrown with tiny oxide grains of secondary origin. (g) Cpx-bearing tonalite (sample 235–1-P5_1), which occurs as a vein intruding into the host oxide-rich diorite (sample 235–1-P5_2). (h) Close-up view of the tonalite in (g). (i) Altered rock (pseudo felsic rock) (sample 227–1-P14), which is characterized by a framework of granular albite intergrown with greenish chlorite, yellow–green poikiloblastic epidote and quartz. (j) Close-up view of the altered rock shown in (i), showing poikiloblastic epidote with a radial structure. Mineral abbreviations: ab, albite; act, actinolite; chl, chlorite; cpx, clinopyroxene; oxi, Fe–Ti oxide; plg, plagioclase; qtz, quartz; ru, rutile; ttn, titanite. Oxide-rich diorites have equigranular medium-grained textures, with granular Fe–Ti oxides closely associated with subhedral to anhedral clinopyroxene (Fig. 4c and d). They can alternatively be termed amphibole-bearing oxide gabbronorites. Clinopyroxenes, oxides and moderately altered euhedral plagioclase form a framework with albite, quartz, and prismatic apatite. Intergranular amphibole is locally present. Compared with those in cpx-rich diorites, the plagioclase here is smaller and strongly zoned, and displays more evolved compositions (An55–An20). The clinopyroxenes have more anhedral shapes and the Fe–Ti oxide inclusions are less common than in the cpx-rich diorites. Interstitial late-stage minerals such as apatite and quartz are common and heterogeneously distributed. Amphibole-rich diorites are characterized by a high amount of initially primary magmatic, subhedral, medium-grained amphibole (30% on average) forming a granular framework together with tabular plagioclase, granular oxides and quartz (Fig. 4e and f). These rocks can alternatively be termed hornblende-rich oxide quartz diorites. The amphiboles are now composed of aggregates of secondary fibrous actinolite intergrown with tiny Fe–Ti oxides, often arranged along the cleavage planes of the amphiboles, which might be the result of the destabilization of Ti-rich magmatic amphibole to lower temperature actinolite and associated secondary Fe–Ti oxides. Minor amounts of quartz are present as an interstitial phase. Compared with those in cpx-rich diorites and in oxide-rich diorites, the plagioclase here is more evolved and is strongly and continuously zoned with compositions of An40 to An20. Tonalites Tonalites may occur as dikelets within other lithologies (Fig. 4g); their grain size is smaller than that of the other felsic rocks studied here. They usually contain a granular plagioclase–quartz framework, associated with minor amounts of magmatic amphibole (now pervasively altered to secondary amphibole), Fe–Ti oxide, and locally anhedral to subhedral clinopyroxene (Fig. 4h). Plagioclase usually shows compositional zoning from An56 to An17, similar to that in the oxide-rich diorites. Accessory minerals include apatite and zircon as primary magmatic phases. The tonalites are pervasively altered and contain a high amount of secondary minerals (chlorite, secondary amphibole, titanite, albite). Therefore, the influence of hydrothermal alteration on the whole-rock major and trace element compositions cannot be neglected. Albitite Another lithological group, which might be regarded as pseudo felsic rocks in appearance (i.e. leucocratic in color), are extremely altered rocks, probably representing former granoblastic hornfelses that have been extensively leached by hydrothermal fluids (Teagle et al., 2012). They are characterized by pervasive alteration, which has resulted in a dusty, cloudy appearance. They are composed exclusively of typical alteration minerals for greenschist-facies conditions, including mainly albite intergrown with poikiloblastic epidote, greenish actinolite, rutile–titanite–oxide intergrowth, chlorite, and quartz (Fig. 4i and j). In the following discussion we will refer to these as ‘albitite’. Mineral proportions To estimate mineral proportions, we first examined thin sections under a transmitted light microscope to determine the phases, and then scanned the thin sections and processed the photographs with the program ImageJ (http://imagej.nih.gov/ij/). The photographs were converted to grayscale and different phases were then chosen according to their different grayscale ranges so that their area fractions could be calculated using the program ImageJ. The results are presented in Fig. 5. It is worth noting that the cpx-rich diorites, oxide-rich diorites and gabbros (not shown) contain a similar mineral assemblage composed of clinopyroxene, plagioclase and Fe–Ti oxides as major phases. In contrast, amp-rich diorites and tonalites are similar to each other in containing amphibole and quartz as major phases, although in varying proportions. Fig. 5. Open in new tabDownload slide Mineral proportions in the various lithologies estimated by image analysis on thin sections. The proportions are average values for each lithology group. Minor quartz in cpx-rich diorite and oxide-rich diorite is neglected. The Fe–Ti oxide consists mainly of ilmenite and lesser amounts of magnetite for all lithologies. The pyroxene consists predominantly of clinopyroxene and minor orthopyroxene for both cpx-rich diorite and oxide-rich diorite. (See Table 1 for lithology groups). Fig. 5. Open in new tabDownload slide Mineral proportions in the various lithologies estimated by image analysis on thin sections. The proportions are average values for each lithology group. Minor quartz in cpx-rich diorite and oxide-rich diorite is neglected. The Fe–Ti oxide consists mainly of ilmenite and lesser amounts of magnetite for all lithologies. The pyroxene consists predominantly of clinopyroxene and minor orthopyroxene for both cpx-rich diorite and oxide-rich diorite. (See Table 1 for lithology groups). ANALYTICAL METHODS Whole-rock major and trace elements All investigated samples were examined for lithological homogeneity (see Table 1) before crushing. If more than one lithological domain was observed in a hand-specimen, the felsic rock was separated carefully by sawing and then ground into a powder using an agate mill. The powders were analyzed at SARM (CRPG-CNRS, Nancy) for major and trace element concentrations. For each sample, about 200 mg powder was fused together with LiBO2 and subsequently dissolved in a diluted HNO3 solution. The major element oxide contents were measured by inductively coupled plasma atomic emission spectrometry (ICP-AES) following the method of Govindaraju & Mevelle (1987), and the uncertainties are better than 5% for values >0·5 wt % and better than 15% for values >0·01wt %. The trace element concentrations were measured by inductively coupled plasma mass spectrometry (ICP-MS) following the method of Carignan et al. (2001); the uncertainties were usually better than 10% for most trace elements whose concentrations are higher than the detection limits by more than 10 times, but were probably >25% for those with concentrations similar to the detection limits. Analytical accuracy and reproducibility were monitored by repeat analyses of five geochemical reference standards [see details given by Carignan et al. (2001)]. Mineral major elements Electron microprobe analysis (EMPA) of minerals was conducted using a Cameca SX100 electron microprobe equipped with five spectrometers and the operating system ‘Peak sight’ at the Institute of Mineralogy, Leibniz University of Hannover. The standard materials include wollastonite (Si and Ca), albite (Na), orthoclase (K), apatite (P), and synthetic oxides including Al2O3, TiO2, Fe2O3, MgO, Mn3O4, and Cr2O3. Raw data were corrected using standard PAP procedures (Pouchou & Pichoir, 1991). A 15 kV acceleration voltage was used for all analyses. Pyroxene, plagioclase and Fe–Ti oxide were measured with a focused beam, whereas amphibole was measured with a defocused beam of 2 μm diameter. The beam current and acquisition time were set at 15 nA and 10 s at peak, respectively. Mineral trace elements All trace element measurements were carried out in situ on polished thin sections, by laser ablation (LA)-ICP-MS at the University of Montpellier, France. The LA system is a Geolas (Microlas) automated platform containing a 193 nm Compex 102 laser from LambdaPhysik. Thin sections were ablated in a small cell of c. 30 cm3, and then the particles from the samples were transported by helium gas and mixed with argon prior to entering the plasma. A laser pulse with an energy density of 10 J cm–2 at a sensitivity of 8 Hz was used, and the laser spot size was 26 μm for most mineral analyses and 15 μm for smaller minerals. The CaO contents of clinopyroxene and plagioclase determined by electron microprobe were used as an internal standard (43Ca) for quantification; the TiO2 content of ilmenite was used as an internal standard (49Ti) for quantification; NIST 612 (Pearce et al., 1997) was used as internal standard for trace element concentration calibration; BIR-1G basalt glass was used as an external reference material for checking signal stability (see Supplementary Data Table 1; supplementary data are available for downloading at http://www.petrology.oxfordjournals.org). Data were processed using GLITTER software (Macquarie Research Ltd), applying a linear fit across all standard analyses (Van Achterbergh et al., 2001). WHOLE-ROCK GEOCHEMISTRY Major elements The whole-rock major and trace element compositions are listed in Table 2. Gabbros, cpx-rich diorites and oxide-rich diorites contain 1–7 wt % normative quartz. Most amp-rich diorites and all tonalites have normative quartz higher than 10 wt %, and plot near the trondhjemite–tonalite boundary in the normative An–Ab–Or diagram of Barker (1979). Table 2 Whole-rock element compositions of selected samples from the IODP Hole 1256D Lithology: . albitite . tonalite . amp-rich diorite . Depth (mbsf): . 1469·4 . 1470·8 . . 1404·4 . 1411·6 . 1507·3 . 1411·3 . 1412·3 . 1422·5 . . . . . . Sample: . 227–1- . 227–2- . R20 . 212–1- . 214–1- . 235–1- . 214–1- . 214–1- . 217–1- . R11- . R12-P . R14- . R14-C . R14-D . . P14 . P12 . . P6_1 . P14 . P5_1 . P10 . P27_1 . P21 . A_1 . . A_1 . . . Major elements (oxide wt %) SiO2 62·61 63·89 62·91 70·27 61·83 69·70 50·59 53·34 47·78 52·55 55·77 59·75 55·94 56·46 TiO2 0·60 1·54 0·50 0·71 1·44 0·98 3·74 2·24 2·70 2·90 2·27 1·95 2·25 2·18 Al2O3 18·04 16·87 20·62 10·78 10·74 10·86 11·91 12·43 12·57 10·70 11·99 10·89 11·28 11·05 Fe2O3 5·57 5·72 2·20 6·65 13·27 6·92 16·26 15·13 17·02 9·08 11·92 9·86 14·31 7·75 MnO 0·07 0·06 0·04 0·04 0·08 0·04 0·14 0·11 0·18 0·17 0·09 0·09 0·10 0·14 MgO 1·20 1·36 0·77 1·90 2·04 1·08 3·77 2·54 5·91 6·03 2·75 3·00 2·25 6·21 CaO 4·28 3·19 2·81 4·60 4·34 3·71 7·00 6·62 9·42 11·18 7·15 6·52 6·57 10·30 Na2O 7·63 7·34 9·78 3·56 4·06 4·86 3·55 3·91 3·15 4·68 5·20 4·48 4·52 4·64 K2O 0·00 0·03 0·38 0·09 0·09 0·12 0·18 0·18 0·17 0·06 0·20 0·10 0·26 0·10 P2O5 0·00 0·00 0·00 0·19 0·59 0·31 0·38 1·00 0·20 0·88 0·83 0·62 0·79 0·54 LOI 1·72 1·89 1·67 1·01 0·92 0·98 1·68 2·29 0·92 0·98 0·91 1·90 1·74 0·80 Total 100·46 99·50 99·60 99·79 99·40 99·55 99·19 99·79 100·01 99·20 99·08 99·16 100·01 100·17 Mg# 29·84 31·97 40·92 36·1 23·3 23·7 31·5 24·9 40·8 56·8 31·4 37·6 23·8 61·3 Trace elements (ppm) Be 0·598 0·658 0·707 1·19 1·17 1·25 0·89 1·23 0·58 0·96 0·90 0·82 0·67 0·60 Sc 1·81 1·82 4·76 15·4 21·7 13·1 38·4 26·7 51·3 50·0 30·6 29·5 30·8 1·81 V 107·2 84·33 38·87 120 150 251 298 154 586 179 561 398 552 107 Cr 19·55 43·33 42·37 b.d.l. b.d.l. 5·0 5·2 b.d.l. 25·6 70·7 22·5 9·2 b.d.l. 19·6 Co 21·78 23·58 5·203 15·4 23·2 8·9 33·9 26·0 57·2 26·9 16·7 14·8 14·7 21·8 Ni 15·25 8·362 b.d.l. 15·0 28·3 19·9 34·3 33·6 58·4 47·1 19·3 17·9 20·9 15·3 Cu b.d.l. b.d.l. 7·588 5·5 b.d.l. b.d.l. 9·1 5·6 134·6 199·4 b.d.l. 5·7 5·3 b.d.l. Zn 19·44 19·2 15·03 17·2 32·6 14·0 41·6 32·5 51·7 28·5 22·6 20·8 25·8 19·4 Ga 20·14 19·3 19·22 16·7 18·2 19·7 19·0 19·3 21·4 19·0 25·0 19·9 23·0 20·1 Ge 0·738 0·93 0·637 1·15 1·11 1·15 1·50 1·23 1·81 1·85 1·30 1·45 1·33 0·74 Rb b.d.l. b.d.l. 4·57 0·58 0·73 1·10 1·74 1·73 1·51 b.d.l. 2·52 0·94 3·24 b.d.l. Sr 37·38 92·51 128·5 93·9 89·5 78·3 86·9 84·4 97·4 79·6 96·9 79·1 92·5 37·4 Y 11·86 36·66 2·337 63·7 85·7 49·0 86·9 81·2 52·9 108 74·6 67·8 61·8 11·9 Zr 44·55 71·06 29·88 675 501 1424 291 444 139 357 174 545 117 44·6 Nb 2·149 6·214 1·452 7·46 10·7 3·76 10·7 14·4 5·48 12·1 6·81 8·14 5·76 2·15 Sn 1·317 9·586 0·866 3·72 2·91 0·94 3·58 2·42 1·51 1·33 1·51 1·92 8·88 1·32 Ba 1·629 4·11 19·56 16·3 18·2 9·06 24·5 15·0 42·9 8·92 15·6 9·56 20·6 1·63 La 2·71 11·04 0·92 8·62 12·8 6·15 10·1 14·3 5·58 12·7 12·4 9·54 11·2 2·71 Ce 11·08 49·23 1·60 23·5 38·1 16·1 30·4 38·3 16·0 35·9 38·7 27·2 31·9 11·1 Pr 1·72 8·07 0·18 3·46 5·67 2·40 4·65 5·45 2·59 5·63 5·74 4·12 4·85 1·72 Nd 8·90 39·34 0·84 19·7 32·1 13·8 26·9 30·3 15·7 33·7 33·4 23·8 27·9 8·90 Sm 2·213 7·839 0·219 6·37 9·83 4·19 9·21 9·09 5·41 11·18 9·81 7·26 8·17 2·21 Eu 1·031 5·416 0·628 2·39 3·44 1·40 2·77 2·74 1·58 1·88 2·02 1·72 1·76 1·03 Gd 2·28 7·471 0·267 8·60 12·9 5·85 12·40 12·08 7·61 15·3 12·3 9·80 10·6 2·28 Tb 0·361 1·152 0·054 1·49 2·15 1·01 2·16 1·98 1·32 2·64 1·96 1·64 1·65 0·36 Dy 2·168 6·975 0·352 9·87 13·8 6·94 14·2 12·7 8·70 17·3 12·2 10·5 10·2 2·17 Ho 0·431 1·421 0·079 2·16 2·90 1·63 3·03 2·69 1·87 3·71 2·49 2·27 2·09 0·43 Er 1·241 4·018 0·256 6·66 8·45 5·17 8·76 7·84 5·38 10·8 7·05 6·62 5·82 1·24 Tm 0·178 0·576 0·042 1·06 1·28 0·86 1·36 1·16 0·82 1·64 1·02 1·00 0·84 0·18 Yb 1·176 3·763 0·29 7·44 8·33 6·42 9·13 7·78 5·57 11·2 6·58 6·79 5·31 1·18 Lu 0·165 0·536 0·049 1·22 1·25 1·16 1·37 1·25 0·85 1·77 1·00 1·09 0·81 0·17 Hf 1·32 1·93 0·84 17·66 11·81 34·59 7·61 10·20 3·76 8·79 5·71 11·51 3·34 1·32 Ta 0·25 0·36 0·09 0·81 1·00 0·33 0·86 1·01 0·41 1·03 0·68 0·64 0·47 0·25 Pb b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. 0·80 b.d.l. Th 0·123 0·287 b.d.l. 0·86 1·07 0·84 0·68 1·00 0·37 1·02 1·01 0·71 0·74 0·12 U 0·141 0·616 b.d.l. 0·27 0·38 0·18 0·25 0·31 0·11 0·13 0·12 0·13 0·09 0·14 EuN/Eu* 1·45 2·24 8·20 0·98 0·93 0·86 0·79 0·80 0·75 0·56 0·62 0·58 0·52 0·44 LaN/SmN 1·02 1·17 3·48 1·12 1·08 1·22 0·91 1·30 0·86 1·05 1·09 1·13 0·78 0·94 GdN/LuN 1·46 1·47 0·57 0·75 1·09 0·53 0·95 1·02 0·94 1·30 0·95 1·38 0·95 0·91 Lithology: . albitite . tonalite . amp-rich diorite . Depth (mbsf): . 1469·4 . 1470·8 . . 1404·4 . 1411·6 . 1507·3 . 1411·3 . 1412·3 . 1422·5 . . . . . . Sample: . 227–1- . 227–2- . R20 . 212–1- . 214–1- . 235–1- . 214–1- . 214–1- . 217–1- . R11- . R12-P . R14- . R14-C . R14-D . . P14 . P12 . . P6_1 . P14 . P5_1 . P10 . P27_1 . P21 . A_1 . . A_1 . . . Major elements (oxide wt %) SiO2 62·61 63·89 62·91 70·27 61·83 69·70 50·59 53·34 47·78 52·55 55·77 59·75 55·94 56·46 TiO2 0·60 1·54 0·50 0·71 1·44 0·98 3·74 2·24 2·70 2·90 2·27 1·95 2·25 2·18 Al2O3 18·04 16·87 20·62 10·78 10·74 10·86 11·91 12·43 12·57 10·70 11·99 10·89 11·28 11·05 Fe2O3 5·57 5·72 2·20 6·65 13·27 6·92 16·26 15·13 17·02 9·08 11·92 9·86 14·31 7·75 MnO 0·07 0·06 0·04 0·04 0·08 0·04 0·14 0·11 0·18 0·17 0·09 0·09 0·10 0·14 MgO 1·20 1·36 0·77 1·90 2·04 1·08 3·77 2·54 5·91 6·03 2·75 3·00 2·25 6·21 CaO 4·28 3·19 2·81 4·60 4·34 3·71 7·00 6·62 9·42 11·18 7·15 6·52 6·57 10·30 Na2O 7·63 7·34 9·78 3·56 4·06 4·86 3·55 3·91 3·15 4·68 5·20 4·48 4·52 4·64 K2O 0·00 0·03 0·38 0·09 0·09 0·12 0·18 0·18 0·17 0·06 0·20 0·10 0·26 0·10 P2O5 0·00 0·00 0·00 0·19 0·59 0·31 0·38 1·00 0·20 0·88 0·83 0·62 0·79 0·54 LOI 1·72 1·89 1·67 1·01 0·92 0·98 1·68 2·29 0·92 0·98 0·91 1·90 1·74 0·80 Total 100·46 99·50 99·60 99·79 99·40 99·55 99·19 99·79 100·01 99·20 99·08 99·16 100·01 100·17 Mg# 29·84 31·97 40·92 36·1 23·3 23·7 31·5 24·9 40·8 56·8 31·4 37·6 23·8 61·3 Trace elements (ppm) Be 0·598 0·658 0·707 1·19 1·17 1·25 0·89 1·23 0·58 0·96 0·90 0·82 0·67 0·60 Sc 1·81 1·82 4·76 15·4 21·7 13·1 38·4 26·7 51·3 50·0 30·6 29·5 30·8 1·81 V 107·2 84·33 38·87 120 150 251 298 154 586 179 561 398 552 107 Cr 19·55 43·33 42·37 b.d.l. b.d.l. 5·0 5·2 b.d.l. 25·6 70·7 22·5 9·2 b.d.l. 19·6 Co 21·78 23·58 5·203 15·4 23·2 8·9 33·9 26·0 57·2 26·9 16·7 14·8 14·7 21·8 Ni 15·25 8·362 b.d.l. 15·0 28·3 19·9 34·3 33·6 58·4 47·1 19·3 17·9 20·9 15·3 Cu b.d.l. b.d.l. 7·588 5·5 b.d.l. b.d.l. 9·1 5·6 134·6 199·4 b.d.l. 5·7 5·3 b.d.l. Zn 19·44 19·2 15·03 17·2 32·6 14·0 41·6 32·5 51·7 28·5 22·6 20·8 25·8 19·4 Ga 20·14 19·3 19·22 16·7 18·2 19·7 19·0 19·3 21·4 19·0 25·0 19·9 23·0 20·1 Ge 0·738 0·93 0·637 1·15 1·11 1·15 1·50 1·23 1·81 1·85 1·30 1·45 1·33 0·74 Rb b.d.l. b.d.l. 4·57 0·58 0·73 1·10 1·74 1·73 1·51 b.d.l. 2·52 0·94 3·24 b.d.l. Sr 37·38 92·51 128·5 93·9 89·5 78·3 86·9 84·4 97·4 79·6 96·9 79·1 92·5 37·4 Y 11·86 36·66 2·337 63·7 85·7 49·0 86·9 81·2 52·9 108 74·6 67·8 61·8 11·9 Zr 44·55 71·06 29·88 675 501 1424 291 444 139 357 174 545 117 44·6 Nb 2·149 6·214 1·452 7·46 10·7 3·76 10·7 14·4 5·48 12·1 6·81 8·14 5·76 2·15 Sn 1·317 9·586 0·866 3·72 2·91 0·94 3·58 2·42 1·51 1·33 1·51 1·92 8·88 1·32 Ba 1·629 4·11 19·56 16·3 18·2 9·06 24·5 15·0 42·9 8·92 15·6 9·56 20·6 1·63 La 2·71 11·04 0·92 8·62 12·8 6·15 10·1 14·3 5·58 12·7 12·4 9·54 11·2 2·71 Ce 11·08 49·23 1·60 23·5 38·1 16·1 30·4 38·3 16·0 35·9 38·7 27·2 31·9 11·1 Pr 1·72 8·07 0·18 3·46 5·67 2·40 4·65 5·45 2·59 5·63 5·74 4·12 4·85 1·72 Nd 8·90 39·34 0·84 19·7 32·1 13·8 26·9 30·3 15·7 33·7 33·4 23·8 27·9 8·90 Sm 2·213 7·839 0·219 6·37 9·83 4·19 9·21 9·09 5·41 11·18 9·81 7·26 8·17 2·21 Eu 1·031 5·416 0·628 2·39 3·44 1·40 2·77 2·74 1·58 1·88 2·02 1·72 1·76 1·03 Gd 2·28 7·471 0·267 8·60 12·9 5·85 12·40 12·08 7·61 15·3 12·3 9·80 10·6 2·28 Tb 0·361 1·152 0·054 1·49 2·15 1·01 2·16 1·98 1·32 2·64 1·96 1·64 1·65 0·36 Dy 2·168 6·975 0·352 9·87 13·8 6·94 14·2 12·7 8·70 17·3 12·2 10·5 10·2 2·17 Ho 0·431 1·421 0·079 2·16 2·90 1·63 3·03 2·69 1·87 3·71 2·49 2·27 2·09 0·43 Er 1·241 4·018 0·256 6·66 8·45 5·17 8·76 7·84 5·38 10·8 7·05 6·62 5·82 1·24 Tm 0·178 0·576 0·042 1·06 1·28 0·86 1·36 1·16 0·82 1·64 1·02 1·00 0·84 0·18 Yb 1·176 3·763 0·29 7·44 8·33 6·42 9·13 7·78 5·57 11·2 6·58 6·79 5·31 1·18 Lu 0·165 0·536 0·049 1·22 1·25 1·16 1·37 1·25 0·85 1·77 1·00 1·09 0·81 0·17 Hf 1·32 1·93 0·84 17·66 11·81 34·59 7·61 10·20 3·76 8·79 5·71 11·51 3·34 1·32 Ta 0·25 0·36 0·09 0·81 1·00 0·33 0·86 1·01 0·41 1·03 0·68 0·64 0·47 0·25 Pb b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. 0·80 b.d.l. Th 0·123 0·287 b.d.l. 0·86 1·07 0·84 0·68 1·00 0·37 1·02 1·01 0·71 0·74 0·12 U 0·141 0·616 b.d.l. 0·27 0·38 0·18 0·25 0·31 0·11 0·13 0·12 0·13 0·09 0·14 EuN/Eu* 1·45 2·24 8·20 0·98 0·93 0·86 0·79 0·80 0·75 0·56 0·62 0·58 0·52 0·44 LaN/SmN 1·02 1·17 3·48 1·12 1·08 1·22 0·91 1·30 0·86 1·05 1·09 1·13 0·78 0·94 GdN/LuN 1·46 1·47 0·57 0·75 1·09 0·53 0·95 1·02 0·94 1·30 0·95 1·38 0·95 0·91 Lithology: . oxide-rich diorite . cpx-rich diorite . granular gabbro . d.l. . Depth (mbsf): . 1418·9 . 1422·6 . 1507·3 . 1507·4 . . 1470·2 . 1483·2 . 1484·3 . 1412·3 . 1418·6 . 1418·9 . 1422·6 . 1484·3 . . Sample: . 216–1- . 217–1- . 235–1- . 235–1- . R17-A . 227–2- . 230–1- . 230–1- . 214–1- . 216–1- . 216–1- . 217–1- . 230–1- . . . P20_1 . P22_1 . P5_2 . P7 . . P2 . P5 . P11_1 . P27_2 . P13 . P20_2 . P22_2 . P11_2 . . Major elements (oxide wt %) SiO2 41·59 44·71 49·05 43·91 51·05 55·83 52·23 55·52 49·25 50·34 49·74 50·58 49·06 0·02 TiO2 4·96 4·99 2·67 3·68 2·74 1·71 3·02 1·63 1·44 0·88 0·87 0·61 1·07 0·02 Al2O3 8·49 12·33 12·88 11·74 11·91 12·33 12·29 12·72 16·25 16·05 16·21 15·33 15·18 0·02 Fe2O3 26·51 21·31 17·22 20·94 15·94 8·69 8·94 7·13 10·85 8·68 8·67 8·72 11·15 0·02 MnO 0·27 0·19 0·13 0·15 0·11 0·11 0·23 0·11 0·14 0·14 0·13 0·13 0·15 0·0002 MgO 5·45 3·95 3·63 3·68 2·86 5·83 5·29 6·04 6·65 7·60 7·55 8·13 7·41 0·015 CaO 6·40 7·43 8·56 9·53 7·98 8·69 10·19 10·12 11·12 13·46 13·00 12·67 11·39 0·03 Na2O 3·58 3·52 4·33 4·54 5·13 5·49 4·83 5·60 3·09 2·57 2·62 2·55 2·84 0·01 K2O 0·05 0·17 0·16 0·23 0·21 0·07 0·07 0·09 0·25 0·07 0·14 0·10 0·04 0·01 P2O5 0·51 0·34 0·84 1·12 0·96 b.d.l. b.d.l. b.d.l. 0·10 0·06 0·04 0·04 b.d.l. 0·04 LOI 1·53 0·52 0·17 0·60 0·83 0·94 1·92 0·59 1·13 0·67 1·00 0·64 1·59 Total 99·34 99·44 99·64 100·13 99·71 99·68 99·02 99·55 100·28 100·50 99·97 99·49 99·87 Mg# 28·9 26·8 29·5 25·9 26·2 57·1 54·0 62·7 54·8 63·4 63·3 64·9 56·8 Trace elements (ppm) Be 0·46 0·91 0·88 0·87 0·84 0·86 0·85 1·01 0·44 b.d.l. b.d.l. b.d.l. 0·45 0·40 Sc 40·3 36·8 35·7 42·0 36·7 24·2 47·5 27·9 38·2 48·0 44·7 44·3 42·0 1·00 V 788 778 275 441 392 233 200 177 273 235 221 180 305 0·70 Cr 8·4 24·5 18·5 8·9 b.d.l. 11·4 17·2 43·1 227·4 348·6 423·7 414·1 227·5 4·00 Co 56·0 39·5 34·9 29·2 17·6 32·9 22·9 24·5 31·1 33·8 30·4 38·3 44·8 0·40 Ni 87·9 46·5 47·7 37·9 26·9 61·1 67·3 32·4 70·9 76·0 73·4 103·3 52·0 5·00 Cu 15·6 54·7 91·6 20·7 5·8 b.d.l. b.d.l. b.d.l. 10·1 70·1 27·7 72·3 60·2 5·00 Zn 60·8 49·2 31·3 33·9 24·7 22·6 33·6 21·3 33·1 40·4 32·1 36·0 50·1 11·00 Ga 14·1 22·4 28·1 28·3 24·7 18·4 18·6 17·8 17·5 14·9 15·0 15·2 17·7 0·20 Ge 1·37 1·39 1·57 1·67 1·42 1·69 1·58 1·69 1·38 1·41 1·39 1·46 1·34 0·15 Rb 0·55 1·46 1·77 2·92 2·72 0·57 0·53 0·66 2·68 b.d.l. 1·25 0·60 b.d.l. 0·40 Sr 42·8 94·7 103 97·3 95·7 89·8 94·9 94·0 106 92·3 94·6 89·7 86·8 2·00 Y 64·5 66·9 79·1 80·9 76·8 48·1 57·0 71·7 28·9 19·3 18·5 28·6 29·5 0·20 Zr 184 205 225 129 214 551 270 561 69·3 36·2 35·6 34·9 43·6 1·00 Nb 12·2 8·44 7·70 8·12 7·17 8·22 11·4 8·43 2·79 1·18 1·16 1·04 1·68 0·09 Sn 2·03 1·52 1·41 4·64 1·68 36·84 5·49 4·91 1·32 b.d.l. 0·65 0·75 0·85 0·45 Ba 4·76 31·4 13·7 12·1 13·6 6·21 6·15 11·0 32·1 7·76 34·9 30·5 6·64 1·60 La 5·41 7·30 11·2 15·4 13·9 2·05 1·19 2·45 2·84 1·50 1·71 2·08 1·40 0·09 Ce 18·0 20·0 30·7 42·9 39·6 7·69 4·85 7·85 8·42 4·56 4·93 6·43 4·57 0·14 Pr 3·08 3·17 4·71 6·35 5·88 1·34 1·03 1·59 1·35 0·71 0·80 1·12 0·87 0·015 Nd 19·6 19·3 27·9 35·3 33·6 8·28 7·91 11·8 8·12 4·70 4·86 7·29 6·10 0·06 Sm 7·05 6·58 9·02 10·56 9·82 3·30 3·87 5·40 2·81 1·75 1·78 2·80 2·63 0·015 Eu 1·39 1·71 2·26 2·10 1·93 1·11 1·04 1·26 1·27 0·82 0·89 0·99 1·01 0·005 Gd 9·63 9·36 12·2 13·5 12·7 5·02 6·31 8·38 4·00 2·63 2·62 4·03 3·96 0·013 Tb 1·62 1·64 2·03 2·15 2·00 0·96 1·22 1·59 0·70 0·47 0·46 0·72 0·72 0·003 Dy 10·4 10·8 13·1 13·6 12·6 6·89 8·74 10·8 4·70 3·13 3·10 4·86 4·90 0·01 Ho 2·19 2·29 2·75 2·81 2·62 1·54 1·89 2·40 1·01 0·67 0·66 1·03 1·07 0·002 Er 6·36 6·71 7·79 7·91 7·31 4·88 5·82 7·21 2·95 1·98 1·90 3·02 3·17 0·01 Tm 0·95 1·02 1·15 1·16 1·06 0·81 0·95 1·14 0·45 0·30 0·29 0·46 0·49 0·001 Yb 6·53 6·91 7·60 7·39 6·82 6·01 6·83 8·03 3·11 1·99 1·95 3·04 3·30 0·007 Lu 1·11 1·09 1·18 1·11 1·05 1·07 1·17 1·34 0·48 0·31 0·29 0·44 0·51 0·003 Hf 5·01 5·44 6·09 3·77 5·75 13·87 6·84 12·25 1·93 1·12 1·01 1·01 1·29 0·03 Ta 1·13 0·66 0·62 0·51 0·51 0·69 0·68 0·74 0·20 0·12 0·10 0·08 0·14 0·01 Pb b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. 0·70 Th 0·68 0·61 0·72 0·78 1·08 0·59 0·38 0·58 0·18 0·09 0·08 0·12 0·14 0·06 U 0·22 0·15 0·13 0·14 0·20 0·17 0·10 0·15 0·06 0·03 0·03 0·04 b.d.l. 0·03 EuN/Eu* 0·51 0·66 0·66 0·54 0·53 0·83 0·64 0·57 1·16 1·17 1·26 0·90 0·96 LaN/SmN 0·64 0·92 1·03 1·21 1·18 0·52 0·26 0·38 0·84 0·71 0·80 0·62 0·44 GdN/LuN 0·92 0·90 1·09 1·29 1·27 0·50 0·57 0·66 0·88 0·90 0·94 0·96 0·82 Lithology: . oxide-rich diorite . cpx-rich diorite . granular gabbro . d.l. . Depth (mbsf): . 1418·9 . 1422·6 . 1507·3 . 1507·4 . . 1470·2 . 1483·2 . 1484·3 . 1412·3 . 1418·6 . 1418·9 . 1422·6 . 1484·3 . . Sample: . 216–1- . 217–1- . 235–1- . 235–1- . R17-A . 227–2- . 230–1- . 230–1- . 214–1- . 216–1- . 216–1- . 217–1- . 230–1- . . . P20_1 . P22_1 . P5_2 . P7 . . P2 . P5 . P11_1 . P27_2 . P13 . P20_2 . P22_2 . P11_2 . . Major elements (oxide wt %) SiO2 41·59 44·71 49·05 43·91 51·05 55·83 52·23 55·52 49·25 50·34 49·74 50·58 49·06 0·02 TiO2 4·96 4·99 2·67 3·68 2·74 1·71 3·02 1·63 1·44 0·88 0·87 0·61 1·07 0·02 Al2O3 8·49 12·33 12·88 11·74 11·91 12·33 12·29 12·72 16·25 16·05 16·21 15·33 15·18 0·02 Fe2O3 26·51 21·31 17·22 20·94 15·94 8·69 8·94 7·13 10·85 8·68 8·67 8·72 11·15 0·02 MnO 0·27 0·19 0·13 0·15 0·11 0·11 0·23 0·11 0·14 0·14 0·13 0·13 0·15 0·0002 MgO 5·45 3·95 3·63 3·68 2·86 5·83 5·29 6·04 6·65 7·60 7·55 8·13 7·41 0·015 CaO 6·40 7·43 8·56 9·53 7·98 8·69 10·19 10·12 11·12 13·46 13·00 12·67 11·39 0·03 Na2O 3·58 3·52 4·33 4·54 5·13 5·49 4·83 5·60 3·09 2·57 2·62 2·55 2·84 0·01 K2O 0·05 0·17 0·16 0·23 0·21 0·07 0·07 0·09 0·25 0·07 0·14 0·10 0·04 0·01 P2O5 0·51 0·34 0·84 1·12 0·96 b.d.l. b.d.l. b.d.l. 0·10 0·06 0·04 0·04 b.d.l. 0·04 LOI 1·53 0·52 0·17 0·60 0·83 0·94 1·92 0·59 1·13 0·67 1·00 0·64 1·59 Total 99·34 99·44 99·64 100·13 99·71 99·68 99·02 99·55 100·28 100·50 99·97 99·49 99·87 Mg# 28·9 26·8 29·5 25·9 26·2 57·1 54·0 62·7 54·8 63·4 63·3 64·9 56·8 Trace elements (ppm) Be 0·46 0·91 0·88 0·87 0·84 0·86 0·85 1·01 0·44 b.d.l. b.d.l. b.d.l. 0·45 0·40 Sc 40·3 36·8 35·7 42·0 36·7 24·2 47·5 27·9 38·2 48·0 44·7 44·3 42·0 1·00 V 788 778 275 441 392 233 200 177 273 235 221 180 305 0·70 Cr 8·4 24·5 18·5 8·9 b.d.l. 11·4 17·2 43·1 227·4 348·6 423·7 414·1 227·5 4·00 Co 56·0 39·5 34·9 29·2 17·6 32·9 22·9 24·5 31·1 33·8 30·4 38·3 44·8 0·40 Ni 87·9 46·5 47·7 37·9 26·9 61·1 67·3 32·4 70·9 76·0 73·4 103·3 52·0 5·00 Cu 15·6 54·7 91·6 20·7 5·8 b.d.l. b.d.l. b.d.l. 10·1 70·1 27·7 72·3 60·2 5·00 Zn 60·8 49·2 31·3 33·9 24·7 22·6 33·6 21·3 33·1 40·4 32·1 36·0 50·1 11·00 Ga 14·1 22·4 28·1 28·3 24·7 18·4 18·6 17·8 17·5 14·9 15·0 15·2 17·7 0·20 Ge 1·37 1·39 1·57 1·67 1·42 1·69 1·58 1·69 1·38 1·41 1·39 1·46 1·34 0·15 Rb 0·55 1·46 1·77 2·92 2·72 0·57 0·53 0·66 2·68 b.d.l. 1·25 0·60 b.d.l. 0·40 Sr 42·8 94·7 103 97·3 95·7 89·8 94·9 94·0 106 92·3 94·6 89·7 86·8 2·00 Y 64·5 66·9 79·1 80·9 76·8 48·1 57·0 71·7 28·9 19·3 18·5 28·6 29·5 0·20 Zr 184 205 225 129 214 551 270 561 69·3 36·2 35·6 34·9 43·6 1·00 Nb 12·2 8·44 7·70 8·12 7·17 8·22 11·4 8·43 2·79 1·18 1·16 1·04 1·68 0·09 Sn 2·03 1·52 1·41 4·64 1·68 36·84 5·49 4·91 1·32 b.d.l. 0·65 0·75 0·85 0·45 Ba 4·76 31·4 13·7 12·1 13·6 6·21 6·15 11·0 32·1 7·76 34·9 30·5 6·64 1·60 La 5·41 7·30 11·2 15·4 13·9 2·05 1·19 2·45 2·84 1·50 1·71 2·08 1·40 0·09 Ce 18·0 20·0 30·7 42·9 39·6 7·69 4·85 7·85 8·42 4·56 4·93 6·43 4·57 0·14 Pr 3·08 3·17 4·71 6·35 5·88 1·34 1·03 1·59 1·35 0·71 0·80 1·12 0·87 0·015 Nd 19·6 19·3 27·9 35·3 33·6 8·28 7·91 11·8 8·12 4·70 4·86 7·29 6·10 0·06 Sm 7·05 6·58 9·02 10·56 9·82 3·30 3·87 5·40 2·81 1·75 1·78 2·80 2·63 0·015 Eu 1·39 1·71 2·26 2·10 1·93 1·11 1·04 1·26 1·27 0·82 0·89 0·99 1·01 0·005 Gd 9·63 9·36 12·2 13·5 12·7 5·02 6·31 8·38 4·00 2·63 2·62 4·03 3·96 0·013 Tb 1·62 1·64 2·03 2·15 2·00 0·96 1·22 1·59 0·70 0·47 0·46 0·72 0·72 0·003 Dy 10·4 10·8 13·1 13·6 12·6 6·89 8·74 10·8 4·70 3·13 3·10 4·86 4·90 0·01 Ho 2·19 2·29 2·75 2·81 2·62 1·54 1·89 2·40 1·01 0·67 0·66 1·03 1·07 0·002 Er 6·36 6·71 7·79 7·91 7·31 4·88 5·82 7·21 2·95 1·98 1·90 3·02 3·17 0·01 Tm 0·95 1·02 1·15 1·16 1·06 0·81 0·95 1·14 0·45 0·30 0·29 0·46 0·49 0·001 Yb 6·53 6·91 7·60 7·39 6·82 6·01 6·83 8·03 3·11 1·99 1·95 3·04 3·30 0·007 Lu 1·11 1·09 1·18 1·11 1·05 1·07 1·17 1·34 0·48 0·31 0·29 0·44 0·51 0·003 Hf 5·01 5·44 6·09 3·77 5·75 13·87 6·84 12·25 1·93 1·12 1·01 1·01 1·29 0·03 Ta 1·13 0·66 0·62 0·51 0·51 0·69 0·68 0·74 0·20 0·12 0·10 0·08 0·14 0·01 Pb b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. 0·70 Th 0·68 0·61 0·72 0·78 1·08 0·59 0·38 0·58 0·18 0·09 0·08 0·12 0·14 0·06 U 0·22 0·15 0·13 0·14 0·20 0·17 0·10 0·15 0·06 0·03 0·03 0·04 b.d.l. 0·03 EuN/Eu* 0·51 0·66 0·66 0·54 0·53 0·83 0·64 0·57 1·16 1·17 1·26 0·90 0·96 LaN/SmN 0·64 0·92 1·03 1·21 1·18 0·52 0·26 0·38 0·84 0·71 0·80 0·62 0·44 GdN/LuN 0·92 0·90 1·09 1·29 1·27 0·50 0·57 0·66 0·88 0·90 0·94 0·96 0·82 Total iron is expressed as Fe2O3; Mg# is calculated as molar Mg/(Mg + Fe) × 100. b.d.l., below detection limit; the data for detection limit are listed in the last column. (For more information of samples, see Table 1.) Table 2 Whole-rock element compositions of selected samples from the IODP Hole 1256D Lithology: . albitite . tonalite . amp-rich diorite . Depth (mbsf): . 1469·4 . 1470·8 . . 1404·4 . 1411·6 . 1507·3 . 1411·3 . 1412·3 . 1422·5 . . . . . . Sample: . 227–1- . 227–2- . R20 . 212–1- . 214–1- . 235–1- . 214–1- . 214–1- . 217–1- . R11- . R12-P . R14- . R14-C . R14-D . . P14 . P12 . . P6_1 . P14 . P5_1 . P10 . P27_1 . P21 . A_1 . . A_1 . . . Major elements (oxide wt %) SiO2 62·61 63·89 62·91 70·27 61·83 69·70 50·59 53·34 47·78 52·55 55·77 59·75 55·94 56·46 TiO2 0·60 1·54 0·50 0·71 1·44 0·98 3·74 2·24 2·70 2·90 2·27 1·95 2·25 2·18 Al2O3 18·04 16·87 20·62 10·78 10·74 10·86 11·91 12·43 12·57 10·70 11·99 10·89 11·28 11·05 Fe2O3 5·57 5·72 2·20 6·65 13·27 6·92 16·26 15·13 17·02 9·08 11·92 9·86 14·31 7·75 MnO 0·07 0·06 0·04 0·04 0·08 0·04 0·14 0·11 0·18 0·17 0·09 0·09 0·10 0·14 MgO 1·20 1·36 0·77 1·90 2·04 1·08 3·77 2·54 5·91 6·03 2·75 3·00 2·25 6·21 CaO 4·28 3·19 2·81 4·60 4·34 3·71 7·00 6·62 9·42 11·18 7·15 6·52 6·57 10·30 Na2O 7·63 7·34 9·78 3·56 4·06 4·86 3·55 3·91 3·15 4·68 5·20 4·48 4·52 4·64 K2O 0·00 0·03 0·38 0·09 0·09 0·12 0·18 0·18 0·17 0·06 0·20 0·10 0·26 0·10 P2O5 0·00 0·00 0·00 0·19 0·59 0·31 0·38 1·00 0·20 0·88 0·83 0·62 0·79 0·54 LOI 1·72 1·89 1·67 1·01 0·92 0·98 1·68 2·29 0·92 0·98 0·91 1·90 1·74 0·80 Total 100·46 99·50 99·60 99·79 99·40 99·55 99·19 99·79 100·01 99·20 99·08 99·16 100·01 100·17 Mg# 29·84 31·97 40·92 36·1 23·3 23·7 31·5 24·9 40·8 56·8 31·4 37·6 23·8 61·3 Trace elements (ppm) Be 0·598 0·658 0·707 1·19 1·17 1·25 0·89 1·23 0·58 0·96 0·90 0·82 0·67 0·60 Sc 1·81 1·82 4·76 15·4 21·7 13·1 38·4 26·7 51·3 50·0 30·6 29·5 30·8 1·81 V 107·2 84·33 38·87 120 150 251 298 154 586 179 561 398 552 107 Cr 19·55 43·33 42·37 b.d.l. b.d.l. 5·0 5·2 b.d.l. 25·6 70·7 22·5 9·2 b.d.l. 19·6 Co 21·78 23·58 5·203 15·4 23·2 8·9 33·9 26·0 57·2 26·9 16·7 14·8 14·7 21·8 Ni 15·25 8·362 b.d.l. 15·0 28·3 19·9 34·3 33·6 58·4 47·1 19·3 17·9 20·9 15·3 Cu b.d.l. b.d.l. 7·588 5·5 b.d.l. b.d.l. 9·1 5·6 134·6 199·4 b.d.l. 5·7 5·3 b.d.l. Zn 19·44 19·2 15·03 17·2 32·6 14·0 41·6 32·5 51·7 28·5 22·6 20·8 25·8 19·4 Ga 20·14 19·3 19·22 16·7 18·2 19·7 19·0 19·3 21·4 19·0 25·0 19·9 23·0 20·1 Ge 0·738 0·93 0·637 1·15 1·11 1·15 1·50 1·23 1·81 1·85 1·30 1·45 1·33 0·74 Rb b.d.l. b.d.l. 4·57 0·58 0·73 1·10 1·74 1·73 1·51 b.d.l. 2·52 0·94 3·24 b.d.l. Sr 37·38 92·51 128·5 93·9 89·5 78·3 86·9 84·4 97·4 79·6 96·9 79·1 92·5 37·4 Y 11·86 36·66 2·337 63·7 85·7 49·0 86·9 81·2 52·9 108 74·6 67·8 61·8 11·9 Zr 44·55 71·06 29·88 675 501 1424 291 444 139 357 174 545 117 44·6 Nb 2·149 6·214 1·452 7·46 10·7 3·76 10·7 14·4 5·48 12·1 6·81 8·14 5·76 2·15 Sn 1·317 9·586 0·866 3·72 2·91 0·94 3·58 2·42 1·51 1·33 1·51 1·92 8·88 1·32 Ba 1·629 4·11 19·56 16·3 18·2 9·06 24·5 15·0 42·9 8·92 15·6 9·56 20·6 1·63 La 2·71 11·04 0·92 8·62 12·8 6·15 10·1 14·3 5·58 12·7 12·4 9·54 11·2 2·71 Ce 11·08 49·23 1·60 23·5 38·1 16·1 30·4 38·3 16·0 35·9 38·7 27·2 31·9 11·1 Pr 1·72 8·07 0·18 3·46 5·67 2·40 4·65 5·45 2·59 5·63 5·74 4·12 4·85 1·72 Nd 8·90 39·34 0·84 19·7 32·1 13·8 26·9 30·3 15·7 33·7 33·4 23·8 27·9 8·90 Sm 2·213 7·839 0·219 6·37 9·83 4·19 9·21 9·09 5·41 11·18 9·81 7·26 8·17 2·21 Eu 1·031 5·416 0·628 2·39 3·44 1·40 2·77 2·74 1·58 1·88 2·02 1·72 1·76 1·03 Gd 2·28 7·471 0·267 8·60 12·9 5·85 12·40 12·08 7·61 15·3 12·3 9·80 10·6 2·28 Tb 0·361 1·152 0·054 1·49 2·15 1·01 2·16 1·98 1·32 2·64 1·96 1·64 1·65 0·36 Dy 2·168 6·975 0·352 9·87 13·8 6·94 14·2 12·7 8·70 17·3 12·2 10·5 10·2 2·17 Ho 0·431 1·421 0·079 2·16 2·90 1·63 3·03 2·69 1·87 3·71 2·49 2·27 2·09 0·43 Er 1·241 4·018 0·256 6·66 8·45 5·17 8·76 7·84 5·38 10·8 7·05 6·62 5·82 1·24 Tm 0·178 0·576 0·042 1·06 1·28 0·86 1·36 1·16 0·82 1·64 1·02 1·00 0·84 0·18 Yb 1·176 3·763 0·29 7·44 8·33 6·42 9·13 7·78 5·57 11·2 6·58 6·79 5·31 1·18 Lu 0·165 0·536 0·049 1·22 1·25 1·16 1·37 1·25 0·85 1·77 1·00 1·09 0·81 0·17 Hf 1·32 1·93 0·84 17·66 11·81 34·59 7·61 10·20 3·76 8·79 5·71 11·51 3·34 1·32 Ta 0·25 0·36 0·09 0·81 1·00 0·33 0·86 1·01 0·41 1·03 0·68 0·64 0·47 0·25 Pb b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. 0·80 b.d.l. Th 0·123 0·287 b.d.l. 0·86 1·07 0·84 0·68 1·00 0·37 1·02 1·01 0·71 0·74 0·12 U 0·141 0·616 b.d.l. 0·27 0·38 0·18 0·25 0·31 0·11 0·13 0·12 0·13 0·09 0·14 EuN/Eu* 1·45 2·24 8·20 0·98 0·93 0·86 0·79 0·80 0·75 0·56 0·62 0·58 0·52 0·44 LaN/SmN 1·02 1·17 3·48 1·12 1·08 1·22 0·91 1·30 0·86 1·05 1·09 1·13 0·78 0·94 GdN/LuN 1·46 1·47 0·57 0·75 1·09 0·53 0·95 1·02 0·94 1·30 0·95 1·38 0·95 0·91 Lithology: . albitite . tonalite . amp-rich diorite . Depth (mbsf): . 1469·4 . 1470·8 . . 1404·4 . 1411·6 . 1507·3 . 1411·3 . 1412·3 . 1422·5 . . . . . . Sample: . 227–1- . 227–2- . R20 . 212–1- . 214–1- . 235–1- . 214–1- . 214–1- . 217–1- . R11- . R12-P . R14- . R14-C . R14-D . . P14 . P12 . . P6_1 . P14 . P5_1 . P10 . P27_1 . P21 . A_1 . . A_1 . . . Major elements (oxide wt %) SiO2 62·61 63·89 62·91 70·27 61·83 69·70 50·59 53·34 47·78 52·55 55·77 59·75 55·94 56·46 TiO2 0·60 1·54 0·50 0·71 1·44 0·98 3·74 2·24 2·70 2·90 2·27 1·95 2·25 2·18 Al2O3 18·04 16·87 20·62 10·78 10·74 10·86 11·91 12·43 12·57 10·70 11·99 10·89 11·28 11·05 Fe2O3 5·57 5·72 2·20 6·65 13·27 6·92 16·26 15·13 17·02 9·08 11·92 9·86 14·31 7·75 MnO 0·07 0·06 0·04 0·04 0·08 0·04 0·14 0·11 0·18 0·17 0·09 0·09 0·10 0·14 MgO 1·20 1·36 0·77 1·90 2·04 1·08 3·77 2·54 5·91 6·03 2·75 3·00 2·25 6·21 CaO 4·28 3·19 2·81 4·60 4·34 3·71 7·00 6·62 9·42 11·18 7·15 6·52 6·57 10·30 Na2O 7·63 7·34 9·78 3·56 4·06 4·86 3·55 3·91 3·15 4·68 5·20 4·48 4·52 4·64 K2O 0·00 0·03 0·38 0·09 0·09 0·12 0·18 0·18 0·17 0·06 0·20 0·10 0·26 0·10 P2O5 0·00 0·00 0·00 0·19 0·59 0·31 0·38 1·00 0·20 0·88 0·83 0·62 0·79 0·54 LOI 1·72 1·89 1·67 1·01 0·92 0·98 1·68 2·29 0·92 0·98 0·91 1·90 1·74 0·80 Total 100·46 99·50 99·60 99·79 99·40 99·55 99·19 99·79 100·01 99·20 99·08 99·16 100·01 100·17 Mg# 29·84 31·97 40·92 36·1 23·3 23·7 31·5 24·9 40·8 56·8 31·4 37·6 23·8 61·3 Trace elements (ppm) Be 0·598 0·658 0·707 1·19 1·17 1·25 0·89 1·23 0·58 0·96 0·90 0·82 0·67 0·60 Sc 1·81 1·82 4·76 15·4 21·7 13·1 38·4 26·7 51·3 50·0 30·6 29·5 30·8 1·81 V 107·2 84·33 38·87 120 150 251 298 154 586 179 561 398 552 107 Cr 19·55 43·33 42·37 b.d.l. b.d.l. 5·0 5·2 b.d.l. 25·6 70·7 22·5 9·2 b.d.l. 19·6 Co 21·78 23·58 5·203 15·4 23·2 8·9 33·9 26·0 57·2 26·9 16·7 14·8 14·7 21·8 Ni 15·25 8·362 b.d.l. 15·0 28·3 19·9 34·3 33·6 58·4 47·1 19·3 17·9 20·9 15·3 Cu b.d.l. b.d.l. 7·588 5·5 b.d.l. b.d.l. 9·1 5·6 134·6 199·4 b.d.l. 5·7 5·3 b.d.l. Zn 19·44 19·2 15·03 17·2 32·6 14·0 41·6 32·5 51·7 28·5 22·6 20·8 25·8 19·4 Ga 20·14 19·3 19·22 16·7 18·2 19·7 19·0 19·3 21·4 19·0 25·0 19·9 23·0 20·1 Ge 0·738 0·93 0·637 1·15 1·11 1·15 1·50 1·23 1·81 1·85 1·30 1·45 1·33 0·74 Rb b.d.l. b.d.l. 4·57 0·58 0·73 1·10 1·74 1·73 1·51 b.d.l. 2·52 0·94 3·24 b.d.l. Sr 37·38 92·51 128·5 93·9 89·5 78·3 86·9 84·4 97·4 79·6 96·9 79·1 92·5 37·4 Y 11·86 36·66 2·337 63·7 85·7 49·0 86·9 81·2 52·9 108 74·6 67·8 61·8 11·9 Zr 44·55 71·06 29·88 675 501 1424 291 444 139 357 174 545 117 44·6 Nb 2·149 6·214 1·452 7·46 10·7 3·76 10·7 14·4 5·48 12·1 6·81 8·14 5·76 2·15 Sn 1·317 9·586 0·866 3·72 2·91 0·94 3·58 2·42 1·51 1·33 1·51 1·92 8·88 1·32 Ba 1·629 4·11 19·56 16·3 18·2 9·06 24·5 15·0 42·9 8·92 15·6 9·56 20·6 1·63 La 2·71 11·04 0·92 8·62 12·8 6·15 10·1 14·3 5·58 12·7 12·4 9·54 11·2 2·71 Ce 11·08 49·23 1·60 23·5 38·1 16·1 30·4 38·3 16·0 35·9 38·7 27·2 31·9 11·1 Pr 1·72 8·07 0·18 3·46 5·67 2·40 4·65 5·45 2·59 5·63 5·74 4·12 4·85 1·72 Nd 8·90 39·34 0·84 19·7 32·1 13·8 26·9 30·3 15·7 33·7 33·4 23·8 27·9 8·90 Sm 2·213 7·839 0·219 6·37 9·83 4·19 9·21 9·09 5·41 11·18 9·81 7·26 8·17 2·21 Eu 1·031 5·416 0·628 2·39 3·44 1·40 2·77 2·74 1·58 1·88 2·02 1·72 1·76 1·03 Gd 2·28 7·471 0·267 8·60 12·9 5·85 12·40 12·08 7·61 15·3 12·3 9·80 10·6 2·28 Tb 0·361 1·152 0·054 1·49 2·15 1·01 2·16 1·98 1·32 2·64 1·96 1·64 1·65 0·36 Dy 2·168 6·975 0·352 9·87 13·8 6·94 14·2 12·7 8·70 17·3 12·2 10·5 10·2 2·17 Ho 0·431 1·421 0·079 2·16 2·90 1·63 3·03 2·69 1·87 3·71 2·49 2·27 2·09 0·43 Er 1·241 4·018 0·256 6·66 8·45 5·17 8·76 7·84 5·38 10·8 7·05 6·62 5·82 1·24 Tm 0·178 0·576 0·042 1·06 1·28 0·86 1·36 1·16 0·82 1·64 1·02 1·00 0·84 0·18 Yb 1·176 3·763 0·29 7·44 8·33 6·42 9·13 7·78 5·57 11·2 6·58 6·79 5·31 1·18 Lu 0·165 0·536 0·049 1·22 1·25 1·16 1·37 1·25 0·85 1·77 1·00 1·09 0·81 0·17 Hf 1·32 1·93 0·84 17·66 11·81 34·59 7·61 10·20 3·76 8·79 5·71 11·51 3·34 1·32 Ta 0·25 0·36 0·09 0·81 1·00 0·33 0·86 1·01 0·41 1·03 0·68 0·64 0·47 0·25 Pb b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. 0·80 b.d.l. Th 0·123 0·287 b.d.l. 0·86 1·07 0·84 0·68 1·00 0·37 1·02 1·01 0·71 0·74 0·12 U 0·141 0·616 b.d.l. 0·27 0·38 0·18 0·25 0·31 0·11 0·13 0·12 0·13 0·09 0·14 EuN/Eu* 1·45 2·24 8·20 0·98 0·93 0·86 0·79 0·80 0·75 0·56 0·62 0·58 0·52 0·44 LaN/SmN 1·02 1·17 3·48 1·12 1·08 1·22 0·91 1·30 0·86 1·05 1·09 1·13 0·78 0·94 GdN/LuN 1·46 1·47 0·57 0·75 1·09 0·53 0·95 1·02 0·94 1·30 0·95 1·38 0·95 0·91 Lithology: . oxide-rich diorite . cpx-rich diorite . granular gabbro . d.l. . Depth (mbsf): . 1418·9 . 1422·6 . 1507·3 . 1507·4 . . 1470·2 . 1483·2 . 1484·3 . 1412·3 . 1418·6 . 1418·9 . 1422·6 . 1484·3 . . Sample: . 216–1- . 217–1- . 235–1- . 235–1- . R17-A . 227–2- . 230–1- . 230–1- . 214–1- . 216–1- . 216–1- . 217–1- . 230–1- . . . P20_1 . P22_1 . P5_2 . P7 . . P2 . P5 . P11_1 . P27_2 . P13 . P20_2 . P22_2 . P11_2 . . Major elements (oxide wt %) SiO2 41·59 44·71 49·05 43·91 51·05 55·83 52·23 55·52 49·25 50·34 49·74 50·58 49·06 0·02 TiO2 4·96 4·99 2·67 3·68 2·74 1·71 3·02 1·63 1·44 0·88 0·87 0·61 1·07 0·02 Al2O3 8·49 12·33 12·88 11·74 11·91 12·33 12·29 12·72 16·25 16·05 16·21 15·33 15·18 0·02 Fe2O3 26·51 21·31 17·22 20·94 15·94 8·69 8·94 7·13 10·85 8·68 8·67 8·72 11·15 0·02 MnO 0·27 0·19 0·13 0·15 0·11 0·11 0·23 0·11 0·14 0·14 0·13 0·13 0·15 0·0002 MgO 5·45 3·95 3·63 3·68 2·86 5·83 5·29 6·04 6·65 7·60 7·55 8·13 7·41 0·015 CaO 6·40 7·43 8·56 9·53 7·98 8·69 10·19 10·12 11·12 13·46 13·00 12·67 11·39 0·03 Na2O 3·58 3·52 4·33 4·54 5·13 5·49 4·83 5·60 3·09 2·57 2·62 2·55 2·84 0·01 K2O 0·05 0·17 0·16 0·23 0·21 0·07 0·07 0·09 0·25 0·07 0·14 0·10 0·04 0·01 P2O5 0·51 0·34 0·84 1·12 0·96 b.d.l. b.d.l. b.d.l. 0·10 0·06 0·04 0·04 b.d.l. 0·04 LOI 1·53 0·52 0·17 0·60 0·83 0·94 1·92 0·59 1·13 0·67 1·00 0·64 1·59 Total 99·34 99·44 99·64 100·13 99·71 99·68 99·02 99·55 100·28 100·50 99·97 99·49 99·87 Mg# 28·9 26·8 29·5 25·9 26·2 57·1 54·0 62·7 54·8 63·4 63·3 64·9 56·8 Trace elements (ppm) Be 0·46 0·91 0·88 0·87 0·84 0·86 0·85 1·01 0·44 b.d.l. b.d.l. b.d.l. 0·45 0·40 Sc 40·3 36·8 35·7 42·0 36·7 24·2 47·5 27·9 38·2 48·0 44·7 44·3 42·0 1·00 V 788 778 275 441 392 233 200 177 273 235 221 180 305 0·70 Cr 8·4 24·5 18·5 8·9 b.d.l. 11·4 17·2 43·1 227·4 348·6 423·7 414·1 227·5 4·00 Co 56·0 39·5 34·9 29·2 17·6 32·9 22·9 24·5 31·1 33·8 30·4 38·3 44·8 0·40 Ni 87·9 46·5 47·7 37·9 26·9 61·1 67·3 32·4 70·9 76·0 73·4 103·3 52·0 5·00 Cu 15·6 54·7 91·6 20·7 5·8 b.d.l. b.d.l. b.d.l. 10·1 70·1 27·7 72·3 60·2 5·00 Zn 60·8 49·2 31·3 33·9 24·7 22·6 33·6 21·3 33·1 40·4 32·1 36·0 50·1 11·00 Ga 14·1 22·4 28·1 28·3 24·7 18·4 18·6 17·8 17·5 14·9 15·0 15·2 17·7 0·20 Ge 1·37 1·39 1·57 1·67 1·42 1·69 1·58 1·69 1·38 1·41 1·39 1·46 1·34 0·15 Rb 0·55 1·46 1·77 2·92 2·72 0·57 0·53 0·66 2·68 b.d.l. 1·25 0·60 b.d.l. 0·40 Sr 42·8 94·7 103 97·3 95·7 89·8 94·9 94·0 106 92·3 94·6 89·7 86·8 2·00 Y 64·5 66·9 79·1 80·9 76·8 48·1 57·0 71·7 28·9 19·3 18·5 28·6 29·5 0·20 Zr 184 205 225 129 214 551 270 561 69·3 36·2 35·6 34·9 43·6 1·00 Nb 12·2 8·44 7·70 8·12 7·17 8·22 11·4 8·43 2·79 1·18 1·16 1·04 1·68 0·09 Sn 2·03 1·52 1·41 4·64 1·68 36·84 5·49 4·91 1·32 b.d.l. 0·65 0·75 0·85 0·45 Ba 4·76 31·4 13·7 12·1 13·6 6·21 6·15 11·0 32·1 7·76 34·9 30·5 6·64 1·60 La 5·41 7·30 11·2 15·4 13·9 2·05 1·19 2·45 2·84 1·50 1·71 2·08 1·40 0·09 Ce 18·0 20·0 30·7 42·9 39·6 7·69 4·85 7·85 8·42 4·56 4·93 6·43 4·57 0·14 Pr 3·08 3·17 4·71 6·35 5·88 1·34 1·03 1·59 1·35 0·71 0·80 1·12 0·87 0·015 Nd 19·6 19·3 27·9 35·3 33·6 8·28 7·91 11·8 8·12 4·70 4·86 7·29 6·10 0·06 Sm 7·05 6·58 9·02 10·56 9·82 3·30 3·87 5·40 2·81 1·75 1·78 2·80 2·63 0·015 Eu 1·39 1·71 2·26 2·10 1·93 1·11 1·04 1·26 1·27 0·82 0·89 0·99 1·01 0·005 Gd 9·63 9·36 12·2 13·5 12·7 5·02 6·31 8·38 4·00 2·63 2·62 4·03 3·96 0·013 Tb 1·62 1·64 2·03 2·15 2·00 0·96 1·22 1·59 0·70 0·47 0·46 0·72 0·72 0·003 Dy 10·4 10·8 13·1 13·6 12·6 6·89 8·74 10·8 4·70 3·13 3·10 4·86 4·90 0·01 Ho 2·19 2·29 2·75 2·81 2·62 1·54 1·89 2·40 1·01 0·67 0·66 1·03 1·07 0·002 Er 6·36 6·71 7·79 7·91 7·31 4·88 5·82 7·21 2·95 1·98 1·90 3·02 3·17 0·01 Tm 0·95 1·02 1·15 1·16 1·06 0·81 0·95 1·14 0·45 0·30 0·29 0·46 0·49 0·001 Yb 6·53 6·91 7·60 7·39 6·82 6·01 6·83 8·03 3·11 1·99 1·95 3·04 3·30 0·007 Lu 1·11 1·09 1·18 1·11 1·05 1·07 1·17 1·34 0·48 0·31 0·29 0·44 0·51 0·003 Hf 5·01 5·44 6·09 3·77 5·75 13·87 6·84 12·25 1·93 1·12 1·01 1·01 1·29 0·03 Ta 1·13 0·66 0·62 0·51 0·51 0·69 0·68 0·74 0·20 0·12 0·10 0·08 0·14 0·01 Pb b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. 0·70 Th 0·68 0·61 0·72 0·78 1·08 0·59 0·38 0·58 0·18 0·09 0·08 0·12 0·14 0·06 U 0·22 0·15 0·13 0·14 0·20 0·17 0·10 0·15 0·06 0·03 0·03 0·04 b.d.l. 0·03 EuN/Eu* 0·51 0·66 0·66 0·54 0·53 0·83 0·64 0·57 1·16 1·17 1·26 0·90 0·96 LaN/SmN 0·64 0·92 1·03 1·21 1·18 0·52 0·26 0·38 0·84 0·71 0·80 0·62 0·44 GdN/LuN 0·92 0·90 1·09 1·29 1·27 0·50 0·57 0·66 0·88 0·90 0·94 0·96 0·82 Lithology: . oxide-rich diorite . cpx-rich diorite . granular gabbro . d.l. . Depth (mbsf): . 1418·9 . 1422·6 . 1507·3 . 1507·4 . . 1470·2 . 1483·2 . 1484·3 . 1412·3 . 1418·6 . 1418·9 . 1422·6 . 1484·3 . . Sample: . 216–1- . 217–1- . 235–1- . 235–1- . R17-A . 227–2- . 230–1- . 230–1- . 214–1- . 216–1- . 216–1- . 217–1- . 230–1- . . . P20_1 . P22_1 . P5_2 . P7 . . P2 . P5 . P11_1 . P27_2 . P13 . P20_2 . P22_2 . P11_2 . . Major elements (oxide wt %) SiO2 41·59 44·71 49·05 43·91 51·05 55·83 52·23 55·52 49·25 50·34 49·74 50·58 49·06 0·02 TiO2 4·96 4·99 2·67 3·68 2·74 1·71 3·02 1·63 1·44 0·88 0·87 0·61 1·07 0·02 Al2O3 8·49 12·33 12·88 11·74 11·91 12·33 12·29 12·72 16·25 16·05 16·21 15·33 15·18 0·02 Fe2O3 26·51 21·31 17·22 20·94 15·94 8·69 8·94 7·13 10·85 8·68 8·67 8·72 11·15 0·02 MnO 0·27 0·19 0·13 0·15 0·11 0·11 0·23 0·11 0·14 0·14 0·13 0·13 0·15 0·0002 MgO 5·45 3·95 3·63 3·68 2·86 5·83 5·29 6·04 6·65 7·60 7·55 8·13 7·41 0·015 CaO 6·40 7·43 8·56 9·53 7·98 8·69 10·19 10·12 11·12 13·46 13·00 12·67 11·39 0·03 Na2O 3·58 3·52 4·33 4·54 5·13 5·49 4·83 5·60 3·09 2·57 2·62 2·55 2·84 0·01 K2O 0·05 0·17 0·16 0·23 0·21 0·07 0·07 0·09 0·25 0·07 0·14 0·10 0·04 0·01 P2O5 0·51 0·34 0·84 1·12 0·96 b.d.l. b.d.l. b.d.l. 0·10 0·06 0·04 0·04 b.d.l. 0·04 LOI 1·53 0·52 0·17 0·60 0·83 0·94 1·92 0·59 1·13 0·67 1·00 0·64 1·59 Total 99·34 99·44 99·64 100·13 99·71 99·68 99·02 99·55 100·28 100·50 99·97 99·49 99·87 Mg# 28·9 26·8 29·5 25·9 26·2 57·1 54·0 62·7 54·8 63·4 63·3 64·9 56·8 Trace elements (ppm) Be 0·46 0·91 0·88 0·87 0·84 0·86 0·85 1·01 0·44 b.d.l. b.d.l. b.d.l. 0·45 0·40 Sc 40·3 36·8 35·7 42·0 36·7 24·2 47·5 27·9 38·2 48·0 44·7 44·3 42·0 1·00 V 788 778 275 441 392 233 200 177 273 235 221 180 305 0·70 Cr 8·4 24·5 18·5 8·9 b.d.l. 11·4 17·2 43·1 227·4 348·6 423·7 414·1 227·5 4·00 Co 56·0 39·5 34·9 29·2 17·6 32·9 22·9 24·5 31·1 33·8 30·4 38·3 44·8 0·40 Ni 87·9 46·5 47·7 37·9 26·9 61·1 67·3 32·4 70·9 76·0 73·4 103·3 52·0 5·00 Cu 15·6 54·7 91·6 20·7 5·8 b.d.l. b.d.l. b.d.l. 10·1 70·1 27·7 72·3 60·2 5·00 Zn 60·8 49·2 31·3 33·9 24·7 22·6 33·6 21·3 33·1 40·4 32·1 36·0 50·1 11·00 Ga 14·1 22·4 28·1 28·3 24·7 18·4 18·6 17·8 17·5 14·9 15·0 15·2 17·7 0·20 Ge 1·37 1·39 1·57 1·67 1·42 1·69 1·58 1·69 1·38 1·41 1·39 1·46 1·34 0·15 Rb 0·55 1·46 1·77 2·92 2·72 0·57 0·53 0·66 2·68 b.d.l. 1·25 0·60 b.d.l. 0·40 Sr 42·8 94·7 103 97·3 95·7 89·8 94·9 94·0 106 92·3 94·6 89·7 86·8 2·00 Y 64·5 66·9 79·1 80·9 76·8 48·1 57·0 71·7 28·9 19·3 18·5 28·6 29·5 0·20 Zr 184 205 225 129 214 551 270 561 69·3 36·2 35·6 34·9 43·6 1·00 Nb 12·2 8·44 7·70 8·12 7·17 8·22 11·4 8·43 2·79 1·18 1·16 1·04 1·68 0·09 Sn 2·03 1·52 1·41 4·64 1·68 36·84 5·49 4·91 1·32 b.d.l. 0·65 0·75 0·85 0·45 Ba 4·76 31·4 13·7 12·1 13·6 6·21 6·15 11·0 32·1 7·76 34·9 30·5 6·64 1·60 La 5·41 7·30 11·2 15·4 13·9 2·05 1·19 2·45 2·84 1·50 1·71 2·08 1·40 0·09 Ce 18·0 20·0 30·7 42·9 39·6 7·69 4·85 7·85 8·42 4·56 4·93 6·43 4·57 0·14 Pr 3·08 3·17 4·71 6·35 5·88 1·34 1·03 1·59 1·35 0·71 0·80 1·12 0·87 0·015 Nd 19·6 19·3 27·9 35·3 33·6 8·28 7·91 11·8 8·12 4·70 4·86 7·29 6·10 0·06 Sm 7·05 6·58 9·02 10·56 9·82 3·30 3·87 5·40 2·81 1·75 1·78 2·80 2·63 0·015 Eu 1·39 1·71 2·26 2·10 1·93 1·11 1·04 1·26 1·27 0·82 0·89 0·99 1·01 0·005 Gd 9·63 9·36 12·2 13·5 12·7 5·02 6·31 8·38 4·00 2·63 2·62 4·03 3·96 0·013 Tb 1·62 1·64 2·03 2·15 2·00 0·96 1·22 1·59 0·70 0·47 0·46 0·72 0·72 0·003 Dy 10·4 10·8 13·1 13·6 12·6 6·89 8·74 10·8 4·70 3·13 3·10 4·86 4·90 0·01 Ho 2·19 2·29 2·75 2·81 2·62 1·54 1·89 2·40 1·01 0·67 0·66 1·03 1·07 0·002 Er 6·36 6·71 7·79 7·91 7·31 4·88 5·82 7·21 2·95 1·98 1·90 3·02 3·17 0·01 Tm 0·95 1·02 1·15 1·16 1·06 0·81 0·95 1·14 0·45 0·30 0·29 0·46 0·49 0·001 Yb 6·53 6·91 7·60 7·39 6·82 6·01 6·83 8·03 3·11 1·99 1·95 3·04 3·30 0·007 Lu 1·11 1·09 1·18 1·11 1·05 1·07 1·17 1·34 0·48 0·31 0·29 0·44 0·51 0·003 Hf 5·01 5·44 6·09 3·77 5·75 13·87 6·84 12·25 1·93 1·12 1·01 1·01 1·29 0·03 Ta 1·13 0·66 0·62 0·51 0·51 0·69 0·68 0·74 0·20 0·12 0·10 0·08 0·14 0·01 Pb b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. b.d.l. 0·70 Th 0·68 0·61 0·72 0·78 1·08 0·59 0·38 0·58 0·18 0·09 0·08 0·12 0·14 0·06 U 0·22 0·15 0·13 0·14 0·20 0·17 0·10 0·15 0·06 0·03 0·03 0·04 b.d.l. 0·03 EuN/Eu* 0·51 0·66 0·66 0·54 0·53 0·83 0·64 0·57 1·16 1·17 1·26 0·90 0·96 LaN/SmN 0·64 0·92 1·03 1·21 1·18 0·52 0·26 0·38 0·84 0·71 0·80 0·62 0·44 GdN/LuN 0·92 0·90 1·09 1·29 1·27 0·50 0·57 0·66 0·88 0·90 0·94 0·96 0·82 Total iron is expressed as Fe2O3; Mg# is calculated as molar Mg/(Mg + Fe) × 100. b.d.l., below detection limit; the data for detection limit are listed in the last column. (For more information of samples, see Table 1.) The granular gabbros analyzed in this study have a basaltic composition in the total alkalis vs SiO2 (TAS) diagram of Le Bas & Streckeisen (1991) (Fig. 6a), consistent with that indicated by the Nb/Y vs Zr/Ti diagram (Pearce, 1996). These data form a low-Mg part (MgO = 6·5–8·5 wt %) of the whole dataset of mafic plutonic rocks reported by Neo et al. (2009), Yamazaki et al. (2009) and Sano et al. (2011), which have MgO up to 12·8 wt %, with a mean value of 8·7 wt % (Fig. 7), except for two samples with MgO of <4 wt % and a Mg# of ∼30, which should be classified as felsic rocks but were included in the group of gabbroic plutonic rocks by Neo et al. (2009). Fig. 6. Open in new tabDownload slide Geochemical classification of the recovered magmatic rocks from IODP Hole 1256D at depths of 1404–1507 mbsf. (a) Total alkalis (Na2O + K2O) vs silica diagram (after Le Bas & Streckeisen, 1991). The reference data for ‘intrusive rock’ are from Neo et al. (2009), Yamazaki et al. (2009) and Sano et al. (2011). (b) Nb/Y vs Zr/Ti (after Pearce, 1996). Fig. 6. Open in new tabDownload slide Geochemical classification of the recovered magmatic rocks from IODP Hole 1256D at depths of 1404–1507 mbsf. (a) Total alkalis (Na2O + K2O) vs silica diagram (after Le Bas & Streckeisen, 1991). The reference data for ‘intrusive rock’ are from Neo et al. (2009), Yamazaki et al. (2009) and Sano et al. (2011). (b) Nb/Y vs Zr/Ti (after Pearce, 1996). Compared with gabbros, cpx-rich diorites have similar Mg# (55–65), with higher SiO2, TiO2 and Na2O contents, but lower MgO, Al2O3 and CaO contents (Fig. 7). They are classified in the TAS diagram (Fig. 6a) as basaltic andesitic, in accordance with the Nb/Y vs Zr/Ti diagram (Fig. 6b). Fig. 7. Open in new tabDownload slide Major element variation vs MgO content for the various lithologies recovered from IODP Hole 1256D. It should be noted that the reference data for ‘intrusive rock’ from Neo et al. (2009), Yamazaki et al. (2009) and Sano et al. (2011) contain felsic lithologies associated with gabbro. Fig. 7. Open in new tabDownload slide Major element variation vs MgO content for the various lithologies recovered from IODP Hole 1256D. It should be noted that the reference data for ‘intrusive rock’ from Neo et al. (2009), Yamazaki et al. (2009) and Sano et al. (2011) contain felsic lithologies associated with gabbro. The oxide-rich diorites are relatively less enriched in SiO2 than the amp-rich diorites (Fig. 6a), but large overlaps in major elements can be observed for these two lithologies, with similar ranges of Al2O3 (c. 10–13 wt %), Na2O (c. 3–5 wt %), CaO (c. 6–11 wt %) and TiO2 (c. 2–5 wt %) (Fig. 7). In the Nb/Y vs Zr/Ti diagram, the oxide-rich diorites are all classified as basalt, whereas the amp-rich diorites are basaltic andesite and basalt (Fig. 6b), which is consistent with their major element compositions. The oxide-rich diorites have much higher total iron contents than the amp-rich diorites (Fe2O3T, total iron expressed as Fe2O3, being 16–27 wt % and 9–17 wt %, respectively) but similar MgO contents (2–6 wt % for both). The oxide-rich diorites show restricted Mg# of 26–30, whereas the amp-rich diorites have a large range of Mg# of 24–61 (Table 2). The tonalites have high SiO2 of 60–70 wt %, but moderate Na2O contents of 3–5 wt % and very low K2O (<0·15 wt %); they are classified as dacitic in the SiO2 vs total alkalis diagram of Le Bas & Streckeisen (1991) (Fig. 6a) and confirmed by the immobile element Nb/Y vs Zr/Ti classification diagram of Pearce (1996) (Fig. 6b). The tonalites are low in MgO (≤2 wt %), TiO2 (≤1·5 wt %) and CaO (<5 wt %), whereas Na2O and Al2O3 are distinctly lower than in the Al-type felsites. In comparison with other lithological groups (Fig. 7), the tonalites show a coherent trend with the oxide-rich diorites, especially apparent in the MgO vs SiO2, TiO2 and Fe2O3T diagrams. Both of them have Mg# values of c. 30. The albitites have intermediate SiO2 contents around 62 wt %, but much higher Na2O contents of 7–10 wt %. These rocks are classified as trachytic in the SiO2 vs total alkalis diagram (Fig. 6a) but are revealed as basaltic in the Nb/Y vs Zr/Ti diagram (Fig. 6b). High fluid flux during an alteration event might have increased the contents of SiO2 and Na2O, as well as reducing the contents of MgO, Fe2O3T and CaO (Fig. 7). Trace elements The normal (N)-MORB normalized rare earth element (REE) and multi-element patterns of all lithological groups are shown in Fig. 8, from which characteristic features for the different lithologies can be observed. Patterns for REE and other trace elements in gabbros and cpx-rich diorites are plotted together in Fig. 8 for comparison. The average composition of lavas and dikes (named AvLD in the following) at IODP Hole 1256D (data from Neo et al., 2009) is also plotted for comparison, which shows consistency in heavy REE (HREE, Gd to Lu) compared with the global N-MORB mean value (Gale et al., 2013) but also slight depletion in light REE (LREE, La to Sm). The granular gabbros analyzed in this study have REE and trace element patterns consistent with those of the mafic plutonic rocks reported by Neo et al. (2009) and Yamazaki et al. (2009), whereas the latter dataset involves both sub-ophitic and granular gabbros with varying mineralogies. From this comparison, we note that there is no systematical difference in bulk trace elements between gabbros with different textures. The gabbros show parallel patterns to the AvLD, but extend to lower abundance and display slight positive Eu anomalies [expressed by EuN/Eu*, Eu* = (SmN × GdN)0·5, where subscript N indicates normalization to N-MORB]. In contrast, the cpx-rich diorites exhibit strong, gradually enhanced, depletion in LREE and enrichment in HREE (LaN/SmN = 0·3–0·5, GdN/LuN = 0·5–0·7) as well as clear negative Eu anomalies (EuN/Eu* = 0·6–0·9) (Fig. 8). For other trace elements, the gabbros are depleted in Th, U and high field strength elements (HFSE) such as Nb, Ta, Zr, Hf and Ti, relative to the AvLD, whereas the cpx-rich diorites are highly enriched in these incompatible elements and show positive anomalies relative to neighboring elements (Fig. 8). The gabbros show positive Sr anomalies relative to Pr and Nd; the Sr concentrations of all gabbro samples are only slightly lower than those for the AvLD. Fig. 8. Open in new tabDownload slide Whole-rock REE and trace element compositions. Left column: REE patterns. Right column: trace element patterns. Normalization is based on N-MORB (using data of Gale et al., 2013). Also plotted for comparison are the average compositions of lavas and dikes (AvLD) including data for sheeted lava flows, massive lavas and sheeted dikes from Hole 1256D (sample total 130) recovered by Expeditions 309 and 312 (data from Neo et al., 2009). The reference data for gabbro are from Neo et al. (2009) and Yamazaki et al. (2009), excluding low-MgO and high-SiO2 intrusive rocks, which have elevated REE abundances and strong negative Eu anomalies (as in Fig. 9). Fig. 8. Open in new tabDownload slide Whole-rock REE and trace element compositions. Left column: REE patterns. Right column: trace element patterns. Normalization is based on N-MORB (using data of Gale et al., 2013). Also plotted for comparison are the average compositions of lavas and dikes (AvLD) including data for sheeted lava flows, massive lavas and sheeted dikes from Hole 1256D (sample total 130) recovered by Expeditions 309 and 312 (data from Neo et al., 2009). The reference data for gabbro are from Neo et al. (2009) and Yamazaki et al. (2009), excluding low-MgO and high-SiO2 intrusive rocks, which have elevated REE abundances and strong negative Eu anomalies (as in Fig. 9). The amp-rich diorites and oxide-rich diorites show similar flat REE patterns in general, enriched by a factor of 1–4 relative to N-MORB, except for Eu, which shows a marked negative anomaly (EuN/Eu* = 0·5–0·8) (Fig. 8). The two lithologies are also similar in their contents of many other trace elements, including high abundances of Th, U, Nb, Ta and Ti. Notably, the amp-rich diorites have up to five times higher Zr and Hf concentrations than N-MORB, which are higher than those of the oxide-rich diorites (Fig. 8). In addition, Sr shows a striking negative anomaly for both lithologies. The tonalites have REE concentrations generally 1–3 times enriched relative to N-MORB, showing flat LREE patterns with an absence of a Eu anomaly (Fig. 8). All three samples, collected at different depths (Table 1), show smooth convergent trends from variable Gd concentrations to very similar Lu concentrations, with GdN/LuN ranging from 0·5 to 1·1. In the trace element patterns (Fig. 8), Rb and Ba are strongly depleted relative to Th and U, and Sr and Ti are strongly depleted relative to the neighboring REE, which are similar in incompatibility. More strikingly, Zr and Hf exhibit a marked enrichment relative to Nd and Sm, and the grade of enrichment is negatively correlated with the total REE content. The albitites show strongly varied REE abundances. Overall depletion is enhanced with increasing Na2O content, indicating that high-Na fluid influx effectively leaches out most REE. A positive Eu anomaly occurs for all samples and is very striking for the most depleted one (Fig. 8). Despite the strong variations in many trace elements, Zr, Hf, Nb, Ta and Ti show relative similarity (Fig. 8), consistent with their limited mobility during hydrothermal alteration. Whereas these extremely altered rocks form their own rock group, other felsic lithologies may also have undergone hydrothermal alteration to various degrees, especially the tonalite group (see below). MINERAL COMPOSITIONS Clinopyroxene The compositions of clinopyroxene from lavas and dikes, gabbro (both sub-ophitic and granular domains), cpx-rich diorite and oxide-rich diorites are plotted together for comparison in terms of Mg#, Al2O3 and TiO2 (Fig. 9) (data in Supplementary Data Table 2). The clinopyroxenes from lavas and dikes and sub-ophitic gabbros generally overlap with each other, having similarly high Mg# (70–85) and Al2O3 (2–3·5 wt %). In contrast, the clinopyroxenes from the granular gabbro, cpx-rich diorite and oxide-rich diorite have lower Mg# (65–80) and Al2O3 contents (<2 wt %), implying that they have crystallized from more evolved magmas compared with the clinopyroxenes from the sub-ophitic gabbros. This is also expressed in the TiO2 contents of the clinopyroxenes from lavas and dikes and sub-ophitic gabbro, which are restricted to 0·4–1·0 wt %, whereas the values for granular gabbro, cpx-rich diorite and oxide-rich diorite extend from c. 0·8 wt % to a lower limit close to zero. Fig. 9. Open in new tabDownload slide Clinopyroxene compositions from the various lithologies in Hole 1256D. (a) Mg# vs Al2O3. (b) Al2O3 vs TiO2. Data points for lava/dike (from Dziony et al., 2008) and for the sub-ophitic and granular domains in gabbros (from Koepke et al., 2011) are averages for a given sample. The data for clinopyroxene-rich diorite and Fe–Ti oxide-rich diorite correspond to single analytical points in various samples. Fig. 9. Open in new tabDownload slide Clinopyroxene compositions from the various lithologies in Hole 1256D. (a) Mg# vs Al2O3. (b) Al2O3 vs TiO2. Data points for lava/dike (from Dziony et al., 2008) and for the sub-ophitic and granular domains in gabbros (from Koepke et al., 2011) are averages for a given sample. The data for clinopyroxene-rich diorite and Fe–Ti oxide-rich diorite correspond to single analytical points in various samples. The REE patterns of clinopyroxenes from gabbro, cpx-rich diorite and oxide-rich diorite are illustrated in Fig. 10. Also plotted for comparison are data from Koepke et al. (2011) for two types of clinopyroxene in the gabbros, one from sub-ophitic domains regarded as relatively primitive and one from granular domains regarded as more evolved as a consequence of in situ fractionation (Koepke et al., 2011). These two types exhibit contrasting REE characteristics: the latter is highly enriched relative to N-MORB with a strong negative Eu anomaly, whereas the former is depleted with a negligible Eu anomaly. The patterns of the newly analyzed clinopyroxenes in the gabbros are parallel to those in the granular domains, with slightly decreased abundances and EuN/Eu* of 0·2–0·4. In contrast, the clinopyroxenes from the cpx-rich diorite have REE patterns intermediate between those of the sub-ophitic and granular domains, showing an enhanced negative Eu anomaly, with higher REE abundances. Compared with the REE patterns for the cpx-rich diorite, the clinopyroxenes from the oxide-rich diorite show similar LREE patterns and strong Eu anomaly, but distinctively higher HREE, especially Yb and Lu, with GdN/LuN of 0·4–0·6. Fig. 10. Open in new tabDownload slide Clinopyroxene (cpx) REE patterns normalized to N-MORB (using the data of Gale et al., 2013). Cpx data from two different textural domains in ‘spotty gabbros’ reported by Koepke et al. (2011) are plotted for comparison; Koepke et al. concluded that the ‘sub-ophitic’ domain represents an early crystallization stage and the ‘granular’ domain a late crystallization stage formed by the processes of ‘in situ fractionation’. Fig. 10. Open in new tabDownload slide Clinopyroxene (cpx) REE patterns normalized to N-MORB (using the data of Gale et al., 2013). Cpx data from two different textural domains in ‘spotty gabbros’ reported by Koepke et al. (2011) are plotted for comparison; Koepke et al. concluded that the ‘sub-ophitic’ domain represents an early crystallization stage and the ‘granular’ domain a late crystallization stage formed by the processes of ‘in situ fractionation’. Amphibole Amphiboles are present in nearly all felsic rocks, but whether they are igneous or formed by hydrothermal alteration needs close attention. Primary, magmatic amphiboles occur in the more evolved domains (granular domains) within the gabbros, evidenced by their composition in combination with textural features (Fig. 11a; Koepke et al., 2011). Magmatic amphibole is also observed in some of the felsic rocks, but only as actinolite pseudomorphs, which replace the magmatic amphibole as coherent crystals or fibrous aggregates, often containing abundant Fe–Ti oxide inclusions (Fig. 11b). Another rare type of amphibole is microcrystalline actinolite grown within quartz (Fig. 11c); this formed during an advanced stage of hydrothermal alteration at relatively low temperatures. Based on mineral chemistry, except for pargasite and hornblende in the gabbro, all of the analyzed amphiboles are actinolite with Mg/(Mg + Fe2+) between 0·5 and 0·9 (Fig. 11d). The amphiboles in gabbro have wide ranges of Al2O3 (up to 13 wt %) and TiO2 (up to 4 wt %), whereas the amphiboles in felsic lithologies have much lower and restricted Al2O3 (<4 wt %) and TiO2 (<1 wt %). Fig. 11. Open in new tabDownload slide Amphibole from various lithologies in Hole 1256D. (a) Euhedral pargasite in a granular gabbro (from Teagle et al., 2006, fig. AF3D). (b) Pseudomorph of magmatic hornblende by an actinolite aggregate containing numerous tiny Fe–Ti oxides in the amp-rich diorite 214–1-P10. (c) Small, spindly amphibole crystals embedded in quartz in the tonalite 214–1-P14. (d) Classification diagram for calcium amphiboles based on stoichiometry (apfu: atom per formula unit) at A and C sites (after Hawthorne et al., 2012). Each name can be further modified according to the Mg/(Mg + Fe2+) ratio; most amphibole compositions plotted here have Mg/(Mg + Fe2+) between 0·5 and 0·9. (e) TiO2 vs Al2O3. Data for gabbro are from Koepke et al. (2011). Fig. 11. Open in new tabDownload slide Amphibole from various lithologies in Hole 1256D. (a) Euhedral pargasite in a granular gabbro (from Teagle et al., 2006, fig. AF3D). (b) Pseudomorph of magmatic hornblende by an actinolite aggregate containing numerous tiny Fe–Ti oxides in the amp-rich diorite 214–1-P10. (c) Small, spindly amphibole crystals embedded in quartz in the tonalite 214–1-P14. (d) Classification diagram for calcium amphiboles based on stoichiometry (apfu: atom per formula unit) at A and C sites (after Hawthorne et al., 2012). Each name can be further modified according to the Mg/(Mg + Fe2+) ratio; most amphibole compositions plotted here have Mg/(Mg + Fe2+) between 0·5 and 0·9. (e) TiO2 vs Al2O3. Data for gabbro are from Koepke et al. (2011). Plagioclase Plagioclase An contents and REE patterns for different lithologies are illustrated in Fig. 12 (data in Supplementary Data Table 3). The data from Koepke et al. (2011) show that the plagioclases in the sub-ophitic gabbro domain generally have higher An contents (60–80 mol %) relative to those in the granular gabbro domain (50–65 mol %). However, plagioclase cores with high An contents of ∼80 mol % were also observed in the granular domains (see fig. A3 of Koepke et al. 2011). Trace element analyses of these plagioclase cores show that LREE are slightly depleted in the sub-ophitic domains compared with the granular ones. Analyses for the felsic lithologies show that the plagioclases in the cpx-rich diorite vary in An content from 45 to 65 mol %, similar to the plagioclases of the granular domain in gabbros; the similarity is also evident in terms of the LREE, which are relatively enriched with patterns similar to those of the granular domains of the gabbros. The plagioclases from the oxide-rich diorite typically have An contents of c. 50 mol %, which decrease in amount continuously to lower values of 10 mol %. The plagioclases in the amp-rich diorite are in the same compositional range as the oxide-rich diorite, but without any clear distribution. Fig. 12. Open in new tabDownload slide Plagioclase (plg) anorthite (An) content histograms for the various lithologies and REE patterns normalized to N-MORB (using data of Gale et al., 2013). An content and REE for the subophitic domain (s.d.) and granular domain (g.d.) of gabbros are from Koepke et al. (2011). Fig. 12. Open in new tabDownload slide Plagioclase (plg) anorthite (An) content histograms for the various lithologies and REE patterns normalized to N-MORB (using data of Gale et al., 2013). An content and REE for the subophitic domain (s.d.) and granular domain (g.d.) of gabbros are from Koepke et al. (2011). All plagioclases from the various lithologies studied herein display different An contents and relatively similar REE patterns with strong Eu anomalies. The plagioclases in the tonalite have a compositional range, with An contents up to ∼65 mol %. They have relatively enriched LREE patterns and stronger depletion of Sm (LaN/SmN = 3·7–8·6) than those from the other lithologies. Fe–Ti oxides Fe–Ti oxides are present in nearly all samples, both as a primary igneous phase and as secondary alteration products. The fractionation of igneous Fe–Ti oxide probably plays an important role in magma compositional evolution (see below). Analyses of coexisting ilmenite–titanomagnetite pairs can yield equilibration temperature and oxygen fugacity (fO2), using the geothermometer and oxybarometer of Andersen & Lindsley (1985), which are embedded in the program QUILF (Andersen et al., 1993). The Fe3+/∑Fe ratios of MORB glasses indicate primary fO2 between FMQ – 0·4 (0·4 log unit below the fayalite–magnetite–quartz buffer) and FMQ – 1 (Christie et al., 1986; Bézos & Humler, 2005). Studies on the basalt–ferrobasalt–andesite series from the Galapagos Rift suggest that extensive crystallization occurred along the FMQ buffer (Juster et al., 1989), with temperatures as low as 910°C for generating the Fe–Ti oxides (Perfit & Fornari, 1983). One analysis of ilmenite–titanomagnetite from the lava pond at IODP Hole 1256D, performed by Dziony et al. (2008), yields an equilibration temperature of c. 780°C and fO2 of FMQ – 0·4, probably suggesting re-equilibration during cooling in a closed system in terms of fO2. In contrast, the data for intrusive rocks in Fig. 13a (data in Supplementary Data Table 4) show that all the intrusive rocks analysed in this study were equilibrated at oxidizing conditions within a temperature range of 980–620°C, at fO2 between FMQ + 1 and FMQ + 3, not in accord with the crystallization temperatures and fO2 conditions of primary igneous processes at MORs as recorded in lavas. There are two possible explanations for the discrepancy. The high fO2 given by Fe–Ti oxides in intrusive rocks might indicate that the magma evolution of the intrusive system occurred under more oxidized conditions than those for the parental magmas. We test this hypothesis by thermodynamic modeling of magma evolution using high fO2 values (see below). Alternatively, the intrusive rocks might have re-equilibrated in the subsolidus regime in an open system for fO2, probably induced by the activity of hydrothermal fluids, which were also responsible for the moderate to strong hydrothermal alteration visible in each sample. This explanation is consistent with the strong intra-crystal heterogeneities of Fe isotopes of magnetite–ilmenite pairs in some lavas, which are most probably caused by interaction with hydrothermal fluids (Dziony et al., 2014). Fig. 13. Open in new tabDownload slide Fe–Ti oxide compositions. (a) Temperature and oxygen fugacity calculated from ilmenite–magnetite pairs using QUILF (Andersen et al., 1993). FMQ is the fayalite–magnetite–quartz buffer. The data for intrusive rock are from Dziony et al. (2014). (b, c) Trace element patterns for ilmenite normalized to N-MORB (using the data of Gale et al., 2013). Fig. 13. Open in new tabDownload slide Fe–Ti oxide compositions. (a) Temperature and oxygen fugacity calculated from ilmenite–magnetite pairs using QUILF (Andersen et al., 1993). FMQ is the fayalite–magnetite–quartz buffer. The data for intrusive rock are from Dziony et al. (2014). (b, c) Trace element patterns for ilmenite normalized to N-MORB (using the data of Gale et al., 2013). LA-ICP-MS analyses show that ilmenite preferentially contains HFSE (Nb, Ta, Zr and Hf), which can be readily incorporated by replacing Ti in the crystal structure, as well as high Yb and Lu relative to the other REE, which are usually below detection limit (Fig. 13b). The Nb and Ta concentrations in ilmenite show a gradual increase from the gabbros to the tonalites, with intermediate abundances in the cpx-diorites and oxide-rich diorites. We propose that this characteristic is inherited from the original igneous ilmenites, rather than a result of subsolidus re-equilibration or alteration, which is consistent with the high immobility of the HFSE during alteration processes. Thus, we interpret the continuous increase in Nb and Ta from the gabbros to the tonalites as a consequence of their high incompatibility during fractional crystallization of silicate mineral phases. Some attempts were made to analyze the trace element contents of titanomagnetite from the intrusive rocks, but most results are below detection limit except for some weak signals for Nb, Zr and Hf (not shown). DISCUSSION Hydrothermal alteration Because most of the felsic rocks have been affected by hydrothermal alteration, it is crucial to evaluate to what extent their primary compositions have been influenced by secondary processes. The albitites are highly altered rocks characterized by abundant albite, epidote and quartz (Fig. 5). They have extremely high Na and Al, and low Fe and Mg contents (Fig. 7). These observations support our suggestion that the original rocks have been metasomatized by high-Na fluids. Important modifications in mobile trace element budgets can also be expected, and therefore only the most immobile trace elements such as Zr, Hf and Ti (Figs 6b and 8) should be used to evaluate their petrogenesis. The albitites have Zr, Hf and Ti contents restricted to a narrow range, with concentrations similar to those of gabbroic rocks from IODP Hole 1256D. We posit that the protoliths of the albitites were gabbroic rocks. A notable feature of the albitites is a significant positive Eu anomaly visible in REE patterns, which is, however, absent for all the other felsic rocks. This feature is well known as a record of low-temperature hydrothermal interactions (e.g. Bau, 1991), but is not related to the crystallization of plagioclase during magmatic evolution. All the other samples from IODP Hole 1256D contain only very small amounts of albite, epidote and chlorite, and display coherent major and trace element features. However, local hydrous alteration is visible, as evidenced by the partial replacement of clinopyroxene by actinolite and chlorite in some cpx-rich diorites (Fig. 4b), saussuritization of plagioclase in the oxide-rich diorite (Fig. 4d) and the tonalites (Fig. 4h), as well as actinolite plus Fe–Ti oxide replacing original high-Al hornblende (Fig. 11b). Hydrothermal alteration may have significantly modified the bulk concentrations of volatile elements (e.g. CO2, S, Cl; Zhang et al., 2017a) as well as some other trace elements (e.g. see below). Petrographic constraints on the geochemical variations The felsic intrusive rocks analyzed in this study are parts of the uppermost AML, which cools as heat is periodically extracted from the top by heat transfer to the overlying sheeted dikes (Zhang et al., 2014). The petrography suggests inhomogeneous distribution of the rock-forming minerals, which apparently determines the geochemical variations in the first order. In a EuN/Eu* vs La/Yb diagram (Fig. 14a), the cpx-rich diorites have both lower La/Yb and lower EuN/Eu* than the gabbros. This can be fully explained by fractionation of plagioclase from a parental magma with no initial Eu anomaly. In contrast, the gabbros are enriched in cumulus plagioclase. The gabbros and cpx-rich diorites are therefore complementary to each other, the gabbros representing a cumulative assemblage whereas the cpx-rich diorites represent the fractionated melts after the removal of plagioclase. The same trend is revealed in a La vs Sm diagram (Fig. 14b). It is also in agreement with the complementary REE and trace element patterns of these two lithological groups (Fig. 8). In addition, one of the gabbros shows a nearly identical trace element pattern to the AvLD, and might therefore represent a frozen melt. A subsequent fractionation stage probably occurred before the final crystallization of the cpx-rich diorites, accounting for their low concentrations of LREE (La and Ce). Fig. 14. Open in new tabDownload slide Variations in trace element concentrations. (a) La/Yb vs EuN/Eu*. It is notable that the various lithologies can be divided into four fields in this diagram. (b) La vs Sm diagram. The evolution of gabbro and cpx-rich diorite indicates a decrease of La/Sm ratio (continuous line), whereas the evolution of the other lithologies indicates an increase of La/Sm ratio (dashed line). (c) V vs Zr diagram. Also shown are mineral fractionation vectors (arrowed lines); that is, the evolution trends of fractionated melts induced by fractionation of the minerals, including plagioclase (plg), clinopyroxene (cpx), amphibole (amp), ilmenite (ilm), titanite (ttn), apatite (ap) and zircon (zir). It should be noted that accumulation of a mineral phase in the analyzed whole-rock would produce the opposite trends to those shown here for fractional crystallization. The vectors indicated are based on partition coefficients for basalt–andesite magmatic systems compiled in EarthRef.org (http://earthref.org/). Fig. 14. Open in new tabDownload slide Variations in trace element concentrations. (a) La/Yb vs EuN/Eu*. It is notable that the various lithologies can be divided into four fields in this diagram. (b) La vs Sm diagram. The evolution of gabbro and cpx-rich diorite indicates a decrease of La/Sm ratio (continuous line), whereas the evolution of the other lithologies indicates an increase of La/Sm ratio (dashed line). (c) V vs Zr diagram. Also shown are mineral fractionation vectors (arrowed lines); that is, the evolution trends of fractionated melts induced by fractionation of the minerals, including plagioclase (plg), clinopyroxene (cpx), amphibole (amp), ilmenite (ilm), titanite (ttn), apatite (ap) and zircon (zir). It should be noted that accumulation of a mineral phase in the analyzed whole-rock would produce the opposite trends to those shown here for fractional crystallization. The vectors indicated are based on partition coefficients for basalt–andesite magmatic systems compiled in EarthRef.org (http://earthref.org/). The oxide-rich diorites and amp-rich diorites are indistinguishable in EuN/Eu* vs La/Yb and Sm vs La diagrams (Fig. 14a and b). Both can be explained by combined fractionation of clinopyroxene and/or amphibole plus plagioclase from a parental magma similar to that of the cpx-rich diorites. In the Zr vs V diagram (Fig. 14c), the large variation in V of the oxide-rich diorites and amp-rich diorites can be accounted for only by an accumulation of ilmenite. The ubiquitous presence of apatite, titanite and zircon in tonalites raises the question of their implication in tonalite genesis. Their fractionation would significantly decrease the abundance of REE, Zr and the La/Yb ratio in the residual melts; such features are not observed among the tonalites studied here (Fig. 14a–c), and therefore no fractionation of such accessory minerals is involved in tonalite genesis; they are thus only products of the crystallization of the tonalitic melt itself. The composition of the tonalites can be explained by fractionation of clinopyroxene, with a little plagioclase, from a parental magma similar to that of the cpx-rich diorites. This is supported by plots of EuN/Eu* vs La/Yb, Sm vs La and Zr vs V (Fig. 15a–c), which also highlight that the tonalites cannot be derived from oxide-rich or amp-rich dioritic magmas. Alternatively, they might have formed by partial melting of the overlying sheeted dikes when heated by the AML (see discussion below). Negligible or untraceable assimilation? Fractional crystallization from parental magmas similar to MORB has been proposed to account for the generation of most of the geochemical features of MOR andesitic–dacitic lavas (Byerly, 1980; Perfit et al., 1983; Juster et al., 1989; Beier et al., 2015), and assimilation of altered crustal materials may be also important in producing some characteristic petrographic, mineralogical, whole-rock trace element and isotopic features (Coogan et al., 2003; Wanless et al., 2010, 2011; France et al., 2013; Freund et al., 2013; Grimes et al., 2013). In contrast to the present study, most of the above-cited studies relied on fresh MORB glasses; here, to characterize and quantify the magmatic processes occurring within the AML, we study plutonic samples that have been overprinted by hydrothermal circulation. Special care must thus be taken when comparing melt, cumulate or mixed compositions from models and results on natural rocks. It has been proposed that the assimilation of crustal rocks or partial melts derived from hydrothermally altered protoliths can be traced by using some ‘diagnostic’ elements, such as those fluid-mobile elements enriched by seawater-derived hydrothermal processes (e.g. Cl), or those incompatible elements highly enriched in partial melts (e.g. U, Th). Assimilation of altered oceanic crust or its partial melts is probably not a dominant process in the genesis of the samples studied herein. First, the Cl concentrations of the plutonic samples (0·05–0·1 wt %; Zhang et al., 2017b) and of the lavas and sheeted dikes (<0·05 wt %; Sano et al., 2008) are significantly lower in comparison with those felsic rocks recovered from MORs influenced significantly by assimilation [0·2–0·7 wt % Cl in the dacites of Wanless et al. (2010); 0·1–1·2 wt % Cl in the andesites of Freund et al. (2013)]. Second, assimilation is also unsupported by the consistently low U concentrations compared with partial melts of altered oceanic crust (Fig. 15), which have been estimated by experimental melting (France et al., 2014) and modeling of batch melting (Wanless et al., 2010). However, some high-Cl apatite cores (up to 5·8 wt % Cl) observed from the intrusive rocks (Zhang et al., 2017a) demonstrate local (but probably not pervasive) assimilation of seawater-derived high-salinity brine by AML magmas. Fig. 15. Open in new tabDownload slide Variation of Th vs U. Residual melt of fractional crystallization (FC) is modeled at fO2 = FMQ using a starting composition with 0·2 wt % H2O, 0·07 ppm U and 0·21 Th (see text and Fig. 17 for FC modelling details). Also shown are compositional trends of experimental partial melts (France et al., 2014) and modelled partial melts of batch melting (Wanless et al., 2010). The percentages are melt proportions. Fig. 15. Open in new tabDownload slide Variation of Th vs U. Residual melt of fractional crystallization (FC) is modeled at fO2 = FMQ using a starting composition with 0·2 wt % H2O, 0·07 ppm U and 0·21 Th (see text and Fig. 17 for FC modelling details). Also shown are compositional trends of experimental partial melts (France et al., 2014) and modelled partial melts of batch melting (Wanless et al., 2010). The percentages are melt proportions. It is worth noting that assimilation of crustal rocks or partial melts derived from hydrothermally altered protoliths is difficult to recognize in the present case for two reasons: (1) all the studied samples are plutonic rocks rather than quenched melts, and thus possible characteristic geochemical features resulting from incorporation of assimilated materials might be lost during extraction of late-magmatic fractionated melts or fluids; (2) late hydrothermal alteration involving seawater-derived fluids at intermediate and low temperatures after the cooling of the gabbroic and felsic rocks is pervasive (Alt et al., 2010; Zhang et al., 2017a), which can thus exert strong influences on the original magmatic signals of geochemical tracers such as oxygen isotopes, Cl, sulfur and other mobile elements (Sano et al., 2008; Alt & Shanks, 2011; Gao et al., 2012). Our recent study on the bulk concentrations of volatile elements (CO2, S, F, Cl, Br) for the same samples indicates that both fluid extraction and hydrothermal alteration might have occurred that significantly obscured the original magmatic imprints (Zhang et al., 2017b). Potential assimilation within the AML magma body and secondary interaction with the overlying sheeted dikes are thus difficult to track in this case, as no data on quenched glasses, melt inclusions or in situ measurement of fresh minerals exist. Major element model for gabbros and diorites To model the petrogenesis of the plutonic rocks and their felsic derivatives by means of fractional crystallization and crystal accumulation (without assimilation process), we have followed a thermodynamic approach using the MELTS program (Ghiorso & Sack, 1995). The lavas and dikes recovered from Hole 1256D have a wide compositional range, from 5 to 10 wt % MgO and Mg# of 40–65, indicating extensive differentiation from a primary melt (Fig. 3). The most primitive dikes (Mg# = 60–65) drilled from Hole 1256D are similar in major element composition to the average composition of East Pacific Rise melt inclusions and glasses reported by Wanless & Shaw (2012), indicating that they should be very close to the parental magmas that fed the AML. Therefore, we used the composition of a high-Mg# (62·3) basaltic dike (sample 309–1256D-161R-1, piece 9; see Teagle et al., 2006; Table 3) as the starting composition to model the effects of fractional crystallization and crystal accumulation. For the MELTS models, different initial water contents were considered at a pressure of 500 bars to account for the overlying rocks and water. Table 3 Starting compositions for modeling Element . SS . Element . S1 . SD . S2 . SD . (wt %) . . (ppm) . (n = 8) . . (n = 130) . . SiO2 (wt %) 49·66 U 0·047 0·017 0·072 0·035 TiO2 1·02 Th 0·126 0·028 0·215 0·091 Al2O3 15·60 Nb 2·10 0·55 3·26 1·31 FeOT 9·62 Ta 0·129 0·036 0·200 0·078 MnO 0·17 La 2·22 0·68 3·00 0·86 MgO 8·92 Ce 6·96 1·99 8·82 2·28 CaO 12·53 Sr 85·24 11·32 83·46 12·83 Na2O 2·29 Nd 6·74 1·47 8·21 1·80 K2O 0·13 Zr 63·84 14·57 80·67 19·42 P2O5 0·06 Hf 1·68 0·31 2·11 0·47 Total 100·00 Sm 2·53 0·38 3·00 0·59 Mg# 62·3 Eu 0·91 0·12 1·08 0·19 Dy 4·45 0·46 5·29 0·96 Y 29·49 3·05 34·24 6·04 Yb 2·81 0·29 3·34 0·59 Lu 0·43 0·05 0·51 0·09 Element . SS . Element . S1 . SD . S2 . SD . (wt %) . . (ppm) . (n = 8) . . (n = 130) . . SiO2 (wt %) 49·66 U 0·047 0·017 0·072 0·035 TiO2 1·02 Th 0·126 0·028 0·215 0·091 Al2O3 15·60 Nb 2·10 0·55 3·26 1·31 FeOT 9·62 Ta 0·129 0·036 0·200 0·078 MnO 0·17 La 2·22 0·68 3·00 0·86 MgO 8·92 Ce 6·96 1·99 8·82 2·28 CaO 12·53 Sr 85·24 11·32 83·46 12·83 Na2O 2·29 Nd 6·74 1·47 8·21 1·80 K2O 0·13 Zr 63·84 14·57 80·67 19·42 P2O5 0·06 Hf 1·68 0·31 2·11 0·47 Total 100·00 Sm 2·53 0·38 3·00 0·59 Mg# 62·3 Eu 0·91 0·12 1·08 0·19 Dy 4·45 0·46 5·29 0·96 Y 29·49 3·05 34·24 6·04 Yb 2·81 0·29 3·34 0·59 Lu 0·43 0·05 0·51 0·09 SS is the starting major element composition (sample 309–1256D-161R-1, piece 9; Teagle et al., 2006) used for modelling fractional crystallization and crystal accumulation. For trace elements, compositions of lavas and dikes from the IODP Hole 1256D reported by Neo et al. (2009) are used. S1 is the average value for samples with MgO 8·5–9·8 wt % and used as starting composition for modelling fractional crystallization. S2 is the average value of all lavas and dikes and used as starting composition for modelling partial melting. Table 3 Starting compositions for modeling Element . SS . Element . S1 . SD . S2 . SD . (wt %) . . (ppm) . (n = 8) . . (n = 130) . . SiO2 (wt %) 49·66 U 0·047 0·017 0·072 0·035 TiO2 1·02 Th 0·126 0·028 0·215 0·091 Al2O3 15·60 Nb 2·10 0·55 3·26 1·31 FeOT 9·62 Ta 0·129 0·036 0·200 0·078 MnO 0·17 La 2·22 0·68 3·00 0·86 MgO 8·92 Ce 6·96 1·99 8·82 2·28 CaO 12·53 Sr 85·24 11·32 83·46 12·83 Na2O 2·29 Nd 6·74 1·47 8·21 1·80 K2O 0·13 Zr 63·84 14·57 80·67 19·42 P2O5 0·06 Hf 1·68 0·31 2·11 0·47 Total 100·00 Sm 2·53 0·38 3·00 0·59 Mg# 62·3 Eu 0·91 0·12 1·08 0·19 Dy 4·45 0·46 5·29 0·96 Y 29·49 3·05 34·24 6·04 Yb 2·81 0·29 3·34 0·59 Lu 0·43 0·05 0·51 0·09 Element . SS . Element . S1 . SD . S2 . SD . (wt %) . . (ppm) . (n = 8) . . (n = 130) . . SiO2 (wt %) 49·66 U 0·047 0·017 0·072 0·035 TiO2 1·02 Th 0·126 0·028 0·215 0·091 Al2O3 15·60 Nb 2·10 0·55 3·26 1·31 FeOT 9·62 Ta 0·129 0·036 0·200 0·078 MnO 0·17 La 2·22 0·68 3·00 0·86 MgO 8·92 Ce 6·96 1·99 8·82 2·28 CaO 12·53 Sr 85·24 11·32 83·46 12·83 Na2O 2·29 Nd 6·74 1·47 8·21 1·80 K2O 0·13 Zr 63·84 14·57 80·67 19·42 P2O5 0·06 Hf 1·68 0·31 2·11 0·47 Total 100·00 Sm 2·53 0·38 3·00 0·59 Mg# 62·3 Eu 0·91 0·12 1·08 0·19 Dy 4·45 0·46 5·29 0·96 Y 29·49 3·05 34·24 6·04 Yb 2·81 0·29 3·34 0·59 Lu 0·43 0·05 0·51 0·09 SS is the starting major element composition (sample 309–1256D-161R-1, piece 9; Teagle et al., 2006) used for modelling fractional crystallization and crystal accumulation. For trace elements, compositions of lavas and dikes from the IODP Hole 1256D reported by Neo et al. (2009) are used. S1 is the average value for samples with MgO 8·5–9·8 wt % and used as starting composition for modelling fractional crystallization. S2 is the average value of all lavas and dikes and used as starting composition for modelling partial melting. Fractional crystallization has been modeled using different initial H2O contents (0 and 0·2 wt %) over a temperature interval of 5°C with fO2 buffered at FMQ and FMQ + 1. The fractionated melt and the instantaneous cumulate generated by fractional crystallization are shown in temperature steps of 20°C in Fig. 16. The instantaneous cumulate refers to the sum of the crystallized phases formed within a temperature interval of 20°C, which are not necessarily totally fractionated from the evolving magmas. As indicated by the plutonic nature and cumulate textures in the gabbroic and dioritic rocks (see above; also Teagle et al., 2006), we anticipate that both evolved melts and cumulates that formed in situ are important components of the intrusive rocks. Bulk-rock compositions can vary owing to either different magma differentiation degrees or different melt–cumulate proportions. The crystallized phases are mainly olivine, plagioclase, clinopyroxene, orthopyroxene and Fe–Ti oxide. Comparison between the modeling results indicates that dry systems tend to generate high TiO2 and Fe2O3T, but lower Al2O3, contents in both fractionated melt and instantaneous cumulate than slightly hydrous systems, whereas the SiO2 content is not so strongly influenced by aH2O. More oxidized conditions tend to generate relatively higher SiO2 contents, but lower Fe2O3T in both fractionated melt and instantaneous cumulate at a late stage (i.e. after the crystallization of Fe–Ti oxides). The various plutonic samples studied in this work are also plotted in Fig. 16 for comparison. For clarity we discuss below each of the lithologies based on selective diagrams. Fig. 16. Open in new tabDownload slide Fractional crystallization modelling using MELTS for SiO2, TiO2, Fe2O3T, Al2O3 and MgO contents. The starting composition is listed in Table 6. The modelling is performed with a temperature interval of 5°C. Different fO2 (FMQ, FMQ + 1) and initial H2O contents (0, 0·2 wt %) have been applied. Fig. 16. Open in new tabDownload slide Fractional crystallization modelling using MELTS for SiO2, TiO2, Fe2O3T, Al2O3 and MgO contents. The starting composition is listed in Table 6. The modelling is performed with a temperature interval of 5°C. Different fO2 (FMQ, FMQ + 1) and initial H2O contents (0, 0·2 wt %) have been applied. As shown in Fig. 16, the gabbros and the cpx-rich diorites have slightly lower MgO and Fe2O3T contents compared with the starting composition, and have distinctive Al2O3 contents. As indicated by comparison with the modeled evolution trends of the fractionated melt and instantaneous cumulate, the gabbros and the cpx-rich diorites do not correspond to pure fractionated melts or cumulates, but rather to a mixture of the two components. For example, as shown in Fig. 17a, most of the samples have lower TiO2 contents than both the fractionated melt and instantaneous cumulate for a given MgO content, but can be formed by mixing between the low-TiO2 early stage instantaneous cumulate and intermediate-stage fractionated melts. Therefore, the compositional variations of gabbros and the cpx-rich diorites can be interpreted by varying mixing degrees involving fractionated melt and instantaneous cumulate. Even though a starting composition with a lower TiO2 content (e.g. 0·5 wt %) might produce fractionated melts directly matching the gabbros and diorites in terms of TiO2 versus MgO in Fig. 17a, the fact that the East Pacific Rise (EPR) MORB glasses (Wanless & Shaw, 2012) and the lavas from IODP Hole 1256D (Neo et al., 2009; Sano et al., 2011) all have TiO2 ≥1·0 wt % does not support the existence of such a low-TiO2 primitive melt. Fig. 17. Open in new tabDownload slide Fractional crystallization modelling results for petrogenetic interpretation. (a) Gabbros and cpx-rich diorites. (b) Oxide-rich diorites. (c) Amp-rich diorites. (d) Tonalites. Experimental data of France et al. (2010) and Erdmann et al. (2015) are also plotted in (d), which demonstrate that generation of low-Al2O3 partial melts requires very low aH2O. (See text for details.) Fig. 17. Open in new tabDownload slide Fractional crystallization modelling results for petrogenetic interpretation. (a) Gabbros and cpx-rich diorites. (b) Oxide-rich diorites. (c) Amp-rich diorites. (d) Tonalites. Experimental data of France et al. (2010) and Erdmann et al. (2015) are also plotted in (d), which demonstrate that generation of low-Al2O3 partial melts requires very low aH2O. (See text for details.) The oxide-rich diorites have the highest Fe2O3T and TiO2 contents among the studied lithologies (Fig. 16), implying significant contribution from an Fe–Ti oxide-rich cumulate. The MgO and Fe2O3T contents of the oxide-rich diorites can be best reproduced by mixing fractionated melt and oxide-bearing instantaneous cumulate (Fig. 17b). Both fractionated melt and instantaneous cumulate are derived from a parental magma containing 0·2 wt % initial H2O with fO2 buffered at FMQ. More oxidizing fO2 conditions would shift the stability of Fe–Ti oxide to a higher temperature, and the instantaneous cumulate consequently evolves to lower Fe2O3T. Therefore, the oxide-rich diorites were most probably formed at fO2 close to FMQ, whereas more oxidized conditions cannot produce high-Fe and high-Ti diorites (Fig. 16). We conclude that fO2 during magma differentiation in the uppermost AML is relatively reduced. This is consistent with the redox condition for extensive MORB crystallization at the Galapagos Rift that is roughly along the FMQ buffer (Perfit & Fornari, 1983; Juster et al., 1989). The amp-rich diorites have distinctively lower TiO2 and Fe2O3T but higher SiO2 contents compared with the oxide-rich diorites (Fig. 16), implying that cumulates rich in Fe–Ti oxides are not incorporated in the rock itself. As shown in the Al2O3 vs MgO diagram (Fig. 17c), most amp-rich diorites can be interpreted as pure fractionated melts that evolved from a parental magma containing 0·2 wt % H2O, whereas some samples might have formed in an H2O-depleted system. The difference in Al2O3 content is most probably controlled by the stability of plagioclase, because increasing aH2O would strongly depress its stability field (Michael & Chase, 1987; Danyushevsky, 2001; Almeev et al., 2012). However, the oxidation condition cannot be inferred because it has little effect on the evolution of the pure fractionated melt (see Fig. 16). Alternative models implying hydrous partial melting of previously hydrothermally altered rocks can also account for the genesis of tonalites; these will be discussed further in the trace element section below. Based on major element contents and associated models, an fO2 close to the FQM buffer is applicable for the generation of all studied lithologies. More oxidizing conditions can clearly be ruled out, for the oxide-rich diorites (Fig. 17b). A dry system fails in reproducing most amp-rich diorites and tonalites (Fig. 17c); however, a slightly wet system containing c. 0·2 wt % initial H2O content can account for their petrogenesis. The small amounts of water that are needed to reproduce the natural rock compositions and to crystallize amphibole and apatite, which are ubiquitous phases in the studied plutonic rocks, could be partly derived from the breakdown of alteration phases that are present in previously altered sheeted dikes within the roof of the AML. There is ample evidence for this process from both the petrography (France et al., 2009; Teagle et al., 2010; Koepke et al., 2011) and the geochemistry (France et al., 2014; Erdmann et al., 2015). Trace element model for gabbros and diorites To model trace element variations during fractional crystallization, the average composition of the most primary lavas and dikes with MgO 8·5–9·8 wt % from Hole 1256D (Neo et al., 2009) is used as a proxy for the primary melt (S1 in Table 3). Because major elemental modeling with MELTS shows that the gabbros and diorites can be best interpreted as products of fractional crystallization of a primitive MORB melt (with initial 0·2 wt % H2O) at fO2 close to FMQ (see discussion above and Fig. 17a–c), modeling of trace elements has been performed with the phase relations determined at such conditions. The partition coefficients for modelling fractional crystallization are from Wanless et al. (2010), except that Eu for plagioclase is assumed to be 1·7 because Eu is expected to be a compatible element for plagioclase in basaltic systems at relatively reduced conditions (Wilke & Behrens, 1999). The trace element features of the gabbros and most diorites can be fully reproduced by mixing fractional crystallization products (melt and cumulate) at intermediate crystallization degrees (F = 40%, where F is the melt proportion) and a melt:cumulate ratio ranging from 3:2 to 1:4 (Fig. 18a and b). This result is consistent with the study of Koepke et al. (2011), who highlighted that the IODP Hole 1256D gabbros are heterogeneous rocks that have formed through in situ crystallization, with primitive clinopyroxene in the sub-ophitic domains and highly evolved ones in the granular domains of the same thin section. Clinopyroxenes in the granular gabbros studied herein have highly fractionated REE patterns (Fig. 10) that are very similar to those in the granular domains in the patchy gabbros described by Koepke et al. (2011). This similarity therefore supports our suggestion that the granular gabbros correspond to evolved melts that were produced by fractional crystallization. Fig. 18. Open in new tabDownload slide Trace element patterns of modelled products in comparison with natural samples. (a) Gabbro and cpx-rich diorite. (b) Oxide-rich diorite and amp-rich diorite. The modelled products may include three parts: fractionated melt (M), instantaneous cumulate (C) and assimilated tonalite (T). The partition coefficients for modelling fractional crystallization are after Wanless et al. (2010) except that Eu for plagioclase is assumed to be 1·7 (see text for detailed discussion). (c) Tonalites. FC, fractional crystallization; PM, partial melting. The arrows indicate potential influences from hydrothermal alteration. Starting composition for FC modeling (‘primitive melt’) is the average composition of lavas and dikes with 8·5–9·8 wt % MgO, whereas starting composition for PM modeling is the average composition of all lavas and dikes (data from Neo et al., 2009). Assimilated tonalite corresponds to sample 235–1-P5. Also shown are the partial melts of experimental hydrous partial melting of an altered sheeted dike at temperatures of 955–975°C (France et al., 2014). Experimental partial melts at temperatures ≥1000°C, which have even lower REE contents and negligible Sr anomalies, are not shown here. The inset photographs (from left to right) are a CL image of quartz (sample 212–1-P6_1) showing three generations and a Cl Kα X-ray map of apatite with heterogeneous Cl distribution (sample 212–1-P14). Fig. 18. Open in new tabDownload slide Trace element patterns of modelled products in comparison with natural samples. (a) Gabbro and cpx-rich diorite. (b) Oxide-rich diorite and amp-rich diorite. The modelled products may include three parts: fractionated melt (M), instantaneous cumulate (C) and assimilated tonalite (T). The partition coefficients for modelling fractional crystallization are after Wanless et al. (2010) except that Eu for plagioclase is assumed to be 1·7 (see text for detailed discussion). (c) Tonalites. FC, fractional crystallization; PM, partial melting. The arrows indicate potential influences from hydrothermal alteration. Starting composition for FC modeling (‘primitive melt’) is the average composition of lavas and dikes with 8·5–9·8 wt % MgO, whereas starting composition for PM modeling is the average composition of all lavas and dikes (data from Neo et al., 2009). Assimilated tonalite corresponds to sample 235–1-P5. Also shown are the partial melts of experimental hydrous partial melting of an altered sheeted dike at temperatures of 955–975°C (France et al., 2014). Experimental partial melts at temperatures ≥1000°C, which have even lower REE contents and negligible Sr anomalies, are not shown here. The inset photographs (from left to right) are a CL image of quartz (sample 212–1-P6_1) showing three generations and a Cl Kα X-ray map of apatite with heterogeneous Cl distribution (sample 212–1-P14). For the cpx-rich diorites, major element modeling (Fig. 17a) indicates that they probably formed at a similar crystallization stage to the gabbro, but involving less cumulate and more evolved melt. However, the strong positive Zr and Hf anomalies cannot be produced only by fractional crystallization, but may be explained by contamination with Zr–Hf-rich tonalitic magmas (Fig. 8). Calculations based on trace elements (Fig. 18a) allow us to estimate the proportions of such a mixture; this is composed at 80% of an evolved magma produced by fractional crystallization (F = 40%, including 50% melt and 30% cumulate) and 20% of a tonalitic component (compare the trace element composition of tonalite sample 235–1-P5). The contamination of some of the diorites by an evolved tonalitic melt is also consistent with the contact relations that were observed in the drilled cores, as the tonalitic veins are not restricted to intrusions within the pyroxene hornfels, but also form intrusions within the gabbros (Fig. 2). In addition, evidence from apatite also suggests a mixing model for diorites: two contrasting endmember compositions of apatite (i.e. chlorapatite and fluorapatite) have been observed in gabbros and tonalites, respectively, whereas both of these endmember compositions exist in diorites (Zhang et al., 2017a). Extreme high-degree fractionation could partially reproduce the amp-rich and oxide-rich diorites, and similarly the enrichments in U–Th and REE; nevertheless, it fails to explain the enrichments of Zr and Hf of some samples. In addition, extreme fractionation would produce a marked negative Eu anomaly, which is not observed in these rocks. Figure 18b shows that the trace element features of the oxide-rich diorites can be reproduced by considering either a pure melt or a mixture of melt and a cumulate at a high crystallization degree (F = 25%). The cumulate proportion involved in the mixture is estimated to be >20% to account for the LREE concentrations in the natural samples. A similar mixing model can also reproduce the amp-rich diorite compositions. However, a large proportion (c. 50%) of the tonalitic component, enriched in U–Th, LREE and Zr–Hf, could alternatively have been involved. This would better reproduce the enrichments of these elements in the amp-rich diorites. Tonalites: fractional crystallization, partial melting and hydrothermal alteration High-silica magmas at the dike–gabbro transition zone of the oceanic crust have commonly been attributed to two contrasting processes (see review by Koepke et al., 2007); that is, extensive fractional crystallization of MORB (FC model), or partial melting of former crustal mafic rocks triggered by heating from the underlying AML (PM model). However, determining the exact mechanism of formation of oceanic plagiogranites in a given ridge setting is still very challenging. There have been numerous studies on oceanic plagiogranites from ophiolites (see table 1 of Koepke et al., 2007), but the complex and controversial tectonic settings of many ophiolites, involving mainly MOR or subduction zone related (e.g. Peters & Kamber, 1994; Rollinson, 2009; Dilek & Furnes, 2014), makes the petrological and tectonic significances of these plagiogranites equivocal for a normal MOR setting. Recently, Brophy (2009) proposed that the trends of REE vs SiO2 variations could discriminate between partial melting and fractional crystallization for felsic melt generation at a MOR setting. It is notable that this method could work only in conditions at which amphibole is one of the major phases in the residue after partial melting or is one of the fractionated phases. Haase et al. (2016) applied this method to propose that the plagiogranite intrusions in the Oman ophiolite might be the products (i.e. residue melts) of extensive fractional crystallization of parental mafic melts. However, the fulfillment of the requirements of the method is questionable. Alternatively, a partial melting model has been supported by studies from ophiolites (e.g. Troodos, Gillis & Coogan, 2002; Oman, France et al., 2009; Koepke et al., 2014) and drilled oceanic crust (e.g. Mid-Atlantic Ridge, Koepke et al., 2005; Southwest Indian Ridge, Koepke et al., 2007), and has been tested by experimental studies (Koepke et al., 2004, 2007; France et al., 2010, 2014; Erdmann et al., 2015). In addition, extensive fractional crystallization of MORB melts combined with partial melting and assimilation of altered crust has been proposed to explain some felsic lavas from oceanic settings (e.g. Wanless et al., 2010; Freund et al., 2013). Amphibole is lacking in the two-pyroxene hornfelses intersected at IODP Hole 1256D, which are believed to represent melting residue of previous partial melting events (Koepke et al., 2008; Zhang et al., 2014; Erdmann et al., 2015). In addition, the petrography and composition of the pyroxene hornfelses, which were observed in strong association with the intrusive felsic rocks, closely resemble the residues produced in partial melting experiments, providing important evidence for the model of partial melting. It is also indicated by experimental studies that amphibole cannot be a stable phase in the melting residue at some conditions, if temperatures are higher than 950°C (e.g. France et al., 2010; Wolff et al., 2013), or if the water activity (aH2O) is low during partial melting (Erdmann et al., 2015). In this regard, the method of Brophy (2009) cannot be used to support or exclude a partial melting model that might not involve amphibole. The natural tonalites studied herein have low MgO contents of 1–2 wt %, which can potentially be reproduced by low-degree (most probably <20%) partial melting of basaltic protoliths (e.g. Koepke et al., 2004; France et al., 2010; Erdmann et al., 2015). However, the low Al2O3 contents of natural tonalites (∼11 wt %) cannot be reproduced by hydrous partial melting experiments at high aH2O (e.g. Koepke et al., 2004; France et al., 2010), but can be simulated only at very low aH2O (Erdmann et al., 2015) (see Fig. 17d). Low aH2O enhances the stabilization of plagioclase (rich in Al2O3) in the residues during partial melting, which thus results in the relative depletion of Al2O3 in the partial melt. Although fractional crystallization models can reproduce felsic melts, following the same evolution trend as the amp-rich diorites, with similar major element compositions to the natural tonalites (e.g. Al2O3 and MgO in Fig. 17d), the discrepancy in trace elements between them is large and excludes high-degree fractional crystallization as the main genetic process for the tonalites (see below). When normalized to N-MORB, the three tonalitic samples show variable trace element patterns, characterized by strong negative Sr anomalies, absence of Eu anomalies, slight enrichment in Ta relative to Nb, and strikingly positive Zr and Hf anomalies (Fig. 18c). A model of fractional crystallization with a melt fraction (F) of 20% (for F > 20%, melt composition has SiO2 < 60 wt % and Al2O3 > 11 wt %, which are not consistent with the tonalitic samples) can reproduce a trace pattern roughly parallel to the most REE-enriched tonalite, but striking inconsistencies exist for Ta, Eu and Zr–Hf. Fractional crystallization with F < 20% would generate higher overall REE but enhanced depletions in Nb and Ta, which do not correspond to the natural samples. As ilmenite and plagioclase strongly incorporate Ta and Eu, respectively, negative Ta and Eu anomalies in high-SiO2 fractionated melts formed by fractional crystallization are inevitable. The trace element concentrations of partial melts derived from experimental hydrous partial melting of altered sheeted dikes at 955–975°C (melting degree is 40–50%) (France et al., 2014) can roughly mimic the most REE-depleted tonalite for some trace elements (but not for Yb, Lu, Zr and Hf) (Fig. 18c). Lower-degree melting would generate partial melts with more elevated REE concentrations, but unfortunately, because of the limited size of melt pools within experimental products at low melt fractions, such data are not available. We thus model the trace element compositions of low-degree partial melts by using the average composition of the lavas and dikes from Hole 1256D as the starting composition (S2 in Table 3). According to observations on natural hornfels (Koepke et al., 2008; Erdmann et al., 2015), we assume that the residual phases include clinopyroxene (40%), orthopyroxene (10%), plagioclase (45%) and magnetite (5%) in a fixed proportion. Hence, the resultant melt trace element composition is dependent entirely on the partial melting degree. The trace element pattern of a partial melt derived by 10% melting can mimic the most REE-enriched tonalite for most elements, except for Zr and Hf (see discussion below on hydrothermal alteration effects). Compared with the fractional crystallization melt, the Ta and Eu issues are eliminated in the partial-melting model for two reasons (Fig. 18c). First, there is considerably less ilmenite in the partial melting residue than in the cumulate from fractional crystallization, which results in high Ta in the partial melt. Second, the Eu partition coefficient between plagioclase and silicate melt is highly dependent on oxygen fugacity, with a negative correlation (Wilke & Behrens, 1999), as Eu2+/Eu3+ is reduced at higher fO2 and the incorporation ability of Eu3+ in plagioclase is much less than that of Eu2+. For example, at log fO2 of c. –16 (close to FMQ – 1 at 900°C) the partition coefficient of Eu between plagioclase and silicate melt is ∼1·7, but it deceases to ∼0·3 at log fO2 of c. –9 (close to FMQ + 2 at 1000°C). Because the sheeted dikes were probably hydrothermally altered before the partial melting event (Koepke et al., 2008; Erdmann et al., 2015), it is obvious that the potential partial melting event proceeded at a more oxidized condition (ΔFMQ = +1 to +2 or higher; Koepke et al., 2008) compared with those typical for MORB crystallization (ΔFMQ = –1 to 0; Christie et al., 1986; Bézos & Humler, 2005). The partial melting modeling was performed using partition coefficients from Bédard (2006), except that for Eu the partition coefficient between plagioclase and melt it was assumed to be 0·3. We suggest that the original high-SiO2 melts, which later formed the tonalites, were generated by low-degree (F ≤ 10%) partial melting of an altered basaltic protolith, rather than by high-degree fractional crystallization from a high-MgO primary melt. This is consistent with the good match for trace elements between natural pyroxene hornfels from the roof of the AML and the experimental residue that was formed in the partial melting experiments on altered sheeted dikes (Fischer et al., 2016; France et al., 2014). In addition, two of the three tonalites studied herein occur as isolated veins within hornfels away from gabbros, supporting the conclusion that the tonalitic veins are the products of hydrous partial melting of altered sheeted dikes. Specific geochemical characteristics of the tonalites are enhanced enrichments in Zr and Hf together with enhanced depletions in REE, to a decreasing extent from La to Lu (Fig. 18c), which we propose to be related to hydrothermal alteration of the tonalitic rocks. The decreasing partition coefficients from La to Lu between aqueous fluids and melt (Reed et al., 2000; Lukanin & Dernov-Pegarev, 2010; Tsay et al., 2014) match exactly the observed variations for the three tonalites, which probably indicate the presence of multiple pulses of hydrothermal fluid flux. Microstructures of zoned quartz and apatite provide support for such repeated hydrothermal alteration events. As shown in Fig. 18c, the cathodoluminescence (CL) image of quartz in a tonalite shows three zones that might indicate two events of hydrothermal interaction subsequent to the primary crystallization of quartz from the melt. The inner zone (Q0), which shows a bright violet CL color, is eroded and overgrown by the mantle zone (Q1) which shows a dark violet CL color. A thin rim zone (Q2) with a dark red CL color further erodes the mantle zone, which is in contact with actinolite grown on the quartz grain boundaries. Chlorine mapping of apatite (Fig. 18c) also indicates two hydrothermal alteration events that modify the original crystal growth from the melt. The major zone (A0) displays a euhedral crystal outline with a more or less homogeneous intermediate Cl content. A later zone (A1), with very low Cl, intersects the original zone. Finally, a high-Cl zone (A2) was formed, probably owing to rock–brine interaction at subsolidus conditions. Detailed in situ trace element analyses of zoned apatite (Zhang et al., 2017a) have documented strong depletion in REE from core to rim, most probably reflecting the influence of hydrothermal alteration. Implications for the nature of the AML In previous sections, we have compared the modeling products of magma fractional crystallization with the natural samples to discuss how they probably formed in the AML. We suggested that the gabbros, cpx-rich diorites, oxide-rich diorites and amp-rich diorites might represent evolved magmas containing both melt and batch cumulate; the estimated crystallization degree is within the range 60–75% (F = 40–25%; Fig. 18). In contrast, the high-silica tonalites appear to be products of low-degree partial melting of altered sheeted dikes. Core drilling at Hole 1256D shows that there are no high-SiO2 or low-MgO components in extrusive rocks (Fig. 3), suggesting that the dioritic and tonalitic lithologies described in this study, regardless of whether they were generated by fractional crystallization or partial melting, correspond to un-eruptible magmas that ended their route to the surface inside the AML. The amount of evolved melt produced by high-degree fractional crystallization in the AML is significant, and is in contrast with gabbros from the lower crust, which correspond, more or less, to pure cumulates with a calculated melt porosity of a few per cent (Natland & Dick, 1996; Korenaga & Kelemen, 1998). Here we examine the liquid evolution and mobility based on MELTS models, with fO2 buffered at FMQ and an initial H2O content of 0·2 wt %. As shown in Fig. 19a, the dioritic members with Mg# of c. 30 (Fig. 7) were formed slightly below the stability of Fe–Ti oxides (at c. 1100°C), corresponding to melt fractions less than 25%. Meanwhile, as a result of increasing SiO2 in the fractionated melt, the melt viscosity increased continuously with the process of fractional crystallization and rose dramatically when Fe–Ti oxides started to crystallize (not shown). Liquid velocity is calculated using the liquid diapir model of Sparks et al. (1980) as gΔρd2/(12μ), where g is gravitational acceleration, Δρ is density contrast, d is the diameter of a liquid droplet (here assumed as 0·1 m), and μ is liquid viscosity. The variation of calculated liquid velocity as a function of temperature is shown in Fig. 19b; it is notable that the liquid velocity decreases extremely rapidly after the onset of crystallization of Fe–Ti oxides. The intense drop of liquid mobility (or increase of viscosity) can explain why the fractionated melts can hardly be separated from solid phases and do not rise further. This accounts for the lack of felsic extrusive rocks with Mg# ≤30 in the cores at Hole 1256D (Fig. 3). The tonalitic magmas have even higher SiO2 contents (i.e. higher viscosity), and usually occur as intrusive veins with centimeter-scale sizes. As liquid velocity shows a negative correlation with viscosity but a positive correlation with vein size, we expect that the tonalite liquid velocity should be lower than that of the dioritic magmas. Fig. 19. Open in new tabDownload slide Properties of fractionated melt. (a) Liquid mass and Mg# of fractionated melt. (b) Liquid mobility (m h–1). The modeling is performed using MELTS with fO2 at the FMQ buffer and an initial H2O content of 0·2 wt %. Liquid mobility (i.e. velocity) is calculated assuming the diameter of liquid droplet as 0·1 m. (Note the logarithmic scale for liquid mobility.) Fig. 19. Open in new tabDownload slide Properties of fractionated melt. (a) Liquid mass and Mg# of fractionated melt. (b) Liquid mobility (m h–1). The modeling is performed using MELTS with fO2 at the FMQ buffer and an initial H2O content of 0·2 wt %. Liquid mobility (i.e. velocity) is calculated assuming the diameter of liquid droplet as 0·1 m. (Note the logarithmic scale for liquid mobility.) The extensive occurrence of felsic intrusive rocks in the upper part of the AML supports the proposal of Natland & Dick (1996) that the shallow AML is rich in (but not composed solely of) highly fractionated iron-rich and siliceous melts that are expelled from deeper gabbroic cumulates. These high-viscosity and un-eruptible felsic magmas have densities lower than the less evolved basaltic magmas underneath, and thus can act as a density filter causing MORB melts to pool in the lower part of the AML (Fig. 20). Therefore, the eruptible zone for lavas and dikes lies in the lower part of the AML, which is composed of less evolved basaltic melts. Although the erupted lavas and dikes have a wide range of compositions spanning 5–10 wt % MgO, they are distinctive from the intrusive dioritic rocks in both major and trace elements (Figs 3 and 8). This confirms that the magmas in the upper part of the AML are not eruptible but tend to crystallize to form felsic intrusive rocks (Fig. 20). Fig. 20. Open in new tabDownload slide Schematic cross-section model illustrating the nature of the AML at two stages reflecting the periodic replenishment of magma from the mantle. (a) Final stage of replenishment. At this stage, after moving upward associated with pervasive assimilation at the top, the AML will start to cool, developing a stratified structure with a more primitive melt in the lower part, and more evolved magma with higher viscosity in the upper part, which has a very low potential to erupt. The sheeted dikes directly overlying the AML are affected by the heat of the AML and consequently are partially melted in the lower part and contact metamorphosed in the upper part, giving rise to granoblastic dikes and tonalitic veins (e.g. Koepke et al., 2008). (b) Inter-replenishment stage. At this stage, the AML is shrinking and magma solidification is enhanced to generate heterogeneously distributed evolved dioritic intrusive (frozen evolved melts) and gabbroic cumulates. The solidifying AML chamber and the overlying hornfels (and enclosed tonalitic patches and veins) are overprinted by the pervasive hydrothermal circulation of seawater-derived fluids. Fig. 20. Open in new tabDownload slide Schematic cross-section model illustrating the nature of the AML at two stages reflecting the periodic replenishment of magma from the mantle. (a) Final stage of replenishment. At this stage, after moving upward associated with pervasive assimilation at the top, the AML will start to cool, developing a stratified structure with a more primitive melt in the lower part, and more evolved magma with higher viscosity in the upper part, which has a very low potential to erupt. The sheeted dikes directly overlying the AML are affected by the heat of the AML and consequently are partially melted in the lower part and contact metamorphosed in the upper part, giving rise to granoblastic dikes and tonalitic veins (e.g. Koepke et al., 2008). (b) Inter-replenishment stage. At this stage, the AML is shrinking and magma solidification is enhanced to generate heterogeneously distributed evolved dioritic intrusive (frozen evolved melts) and gabbroic cumulates. The solidifying AML chamber and the overlying hornfels (and enclosed tonalitic patches and veins) are overprinted by the pervasive hydrothermal circulation of seawater-derived fluids. CONCLUSIONS Bulk-rock geochemistry, mineral compositions and magma evolution models for IODP Hole 1256D intrusive rocks indicate that the gabbros and diorites are formed from a mixture of fractionated melt and cumulate with a variable melt:cumulate ratio. The estimated corresponding melt fraction varies between 25 and 40%. In contrast, the high-SiO2 tonalites are unlikely to be the products of extreme fractional crystallization, but rather of low-degree hydrous partial melting of the sheeted dike complex overlying the AML. The tonalitic components have been partly incorporated in dioritic magmas by mixing. The upper part of the AML, documented herein by various plutonic rocks, contains abundant low-density and high-viscosity felsic magmas, which are not eruptible, but may act as a barrier to MORB melts pooling in the lower part of the AML. ACKNOWLEDGEMENTS We gratefully acknowledge the shipboard crew and Scientific Party of IODP Expeditions 312 and 335 for their assistance in data collection. The samples used in this study were provided by the Integrated Ocean Drilling Program. We thank Felix Genske and an anonymous reviewer for critical and insightful comments. We also thank Prof. Marjorie Wilson for editorial handling and polishing the language. FUNDING This research was funded by the DFG (Deutsche Forschungsgemeinschaft) project KO 1723/17. SUPPLEMENTARY DATA Supplementary data for this paper are available at Journal of Petrology online. REFERENCES Almeev R. R. , Holtz F., Koepke J., Parat F. ( 2012 ). 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TI - Felsic Plutonic Rocks from IODP Hole 1256D, Eastern Pacific: Implications for the Nature of the Axial Melt Lens at Fast-Spreading Mid-Ocean Ridges
JF - Journal of Petrology
DO - 10.1093/petrology/egx064
DA - 2017-08-01
UR - https://www.deepdyve.com/lp/oxford-university-press/felsic-plutonic-rocks-from-iodp-hole-1256d-eastern-pacific-5B0NMWUyMr
SP - 1535
VL - 58
IS - 8
DP - DeepDyve
ER -