TY - JOUR
AU - Horning,, Gregory
AB - Abstract In this study we investigate the origin and alteration history of serpentinites in two dredge hauls from the North Wall of the Puerto Rico Trench, where the North American plate subducts beneath the Greater Antilles island arc. Faulting of the North American plate at the North Wall of the Puerto Rico Trench exposes oceanic basement in a nearly spreading-parallel direction, offering a unique perspective on crustal accretion and alteration processes in the Cretaceous Quiet Zone. U–Pb age dating of porous zircon in an altered mafic vein in serpentinite from 20°00’00”N, 66°31’57”W yields an age of 114·8 Ma, suggesting that it originated at the Cretaceous Mid-Atlantic Ridge, c. 2200 km east of its current location. After localized high-temperature (>380 °C) hydrothermal alteration to chlorite, antigorite, talc and tremolite, peridotites underwent pervasive serpentinization at lower temperatures (∼310–240 °C). In addition to lizardite-rich serpentinite, which is prevalent in dredge D10, dredge D2 contains serpentinite dominated by antigorite. Pervasive serpentinization was followed by extensive talc alteration of several antigorite-rich serpentinites from dredge D2, which was accompanied by Fe loss. As the newly formed crust moved off-axis, open-system alteration at low temperatures by seawater (i.e. weathering) modified the mineralogy, chemical composition, and physical properties of the serpentinite. Where present, brucite dissolved and magnetite oxidized to hematite and goethite, which caused a decrease in magnetic susceptibility and an increase in porosity. Weathering led to complete oxidation of several lizardite-serpentinites in dredge D10, whereas talc-altered antigorite serpentinites in dredge D2 were less affected by oxidation. Rare earth element concentrations of lizardite-rich serpentinite in dredge D10 decreased steadily with increasing Fe(III)/Fetot, reflecting simultaneous leaching and oxidation by seawater. These results have implications for our understanding of several processes, including hydrothermal alteration of peridotite at the Cretaceous Mid-Atlantic Ridge, mass transfer and redox reactions between serpentinite and seawater during off-axis weathering, and the introduction of volatiles, oxidizing agents, and antigorite beneath the Greater Antilles Arc. INTRODUCTION Mantle peridotite is unstable in the presence of aqueous fluids at zeolite-, greenschist-, blueschist- and low-grade amphibolite-facies conditions. As a result, it undergoes a series of dissolution—precipitation and redox reactions to form serpentinite, a rock that is chiefly composed of serpentine minerals. This process, referred to as serpentinization, has gained considerable appreciation as a widespread geological process with multifaceted implications for Earth’s lithosphere, hydrosphere, and biosphere (Früh-Green et al., 2004; Bach & Früh-Green, 2010; Evans et al., 2013; McCollom & Seewald, 2013; Schrenk et al., 2013). Serpentinites occur in marine environments at mid-ocean ridges and fracture zones, at passive continental margins, in the forearc settings of subduction zones and on land in ophiolites and orogenic belts. Their occurrence in Archean, Proterozoic and Phanerozoic mountain ranges (e.g. Blais & Auvray, 1990; Attoh et al., 2006; Grimmer & Greiling, 2012) indicates that serpentinization has been operating through most of Earth’s history. Much uncertainty remains regarding the abundance and spatial distribution of serpentinized peridotite in oceanic lithosphere accreted at mid-ocean ridges (Horen et al., 1996; Carlson, 2001; Cannat et al., 2010). Differences in spreading rate and magma supply influence the relative proportions and distribution of mafic and ultramafic (olivine- and/or pyroxene-rich) lithologies (Dick et al., 2003; Cannat et al., 2010). Serpentinized peridotite is particularly common along ultraslow-spreading ridges (<20 mm a−1 full rate) where oceanic crust is thin or lacking (Dick et al., 2003; Cannat et al., 2006; Sauter et al., 2013; Schlindwein & Schmid, 2016). Along slow-spreading mid-ocean ridges (20–55 mm a−1), peridotite is common where seafloor spreading is asymmetric and accommodated by tectonic extension along large-offset detachment faults that result in the formation of oceanic core complexes (Tucholke et al., 1998; Boschi et al., 2006; Ildefonse et al., 2007; Canales et al., 2008; MacLeod et al., 2009; Smith et al., 2014). Asymmetric spreading accounts for 50% of the ridge length at the Mid-Atlantic Ridge (MAR) between 12·5° and 35°N (Escartin et al., 2008) and serpentinization is particularly widespread in this area. In contrast, the seafloor along intermediate-spreading (55–80 mm a−1) to fast-spreading (80–180 mm a−1) ridges appears to consist almost exclusively of basalt, although oceanic core complexes do occur at intermediate-spreading ocean ridges, but not as commonly as at slower-spreading ocean ridges (Tucholke et al., 2008). Full spreading rates in the Atlantic varied from more than 50 mm a−1 during the mid-Cretaceous to less than 25 mm a−1 today (Müller et al., 2008; Seton et al., 2012) and hence the abundance and distribution of serpentinite in the Atlantic is probably fairly variable. In particular, off-axis basement in the Atlantic is undersampled and so it remains unclear whether mantle peridotite was exhumed to crustal levels and affected by serpentinization in the past when seafloor spreading in the Atlantic was faster (Michael & Bonatti, 1983; Staudigel et al., 1995; Müntener & Manatschal, 2006; Jagoutz et al., 2007). Serpentinization is considered by most to be quasi-isochemical with the exception of the addition of water and the loss of Ca, and changes in trace element contents (see Dresser, 1913; Bain, 1932; Hostetler et al., 1966; Thayer, 1966; Coleman, 1971; O’Hanley, 1992; Menzies et al., 1993; Paulick et al., 2006; Beinlich et al., 2010; Plümper et al., 2012; Malvoisin, 2015; Velbel et al., 2015) However, much uncertainty remains regarding the compositional changes peridotite and serpentinite experience during prolonged interaction with seawater. Serpentinites are prone to weathering reactions in surface environments, involving mineral dissolution and oxidation, as several studies of continental sites have demonstrated (Mumpton & Thompson, 1966; Cleaves et al., 1974; Venturelli et al., 1997; Caillaud et al., 2006; Alexander et al., 2007; Bucher et al., 2015). Weathering of oceanic serpentinites has received little attention, despite the widespread occurrence of serpentinite and potential ramifications for our understanding of mass transfer and redox reactions in seafloor and subduction-zone settings. For example, previous studies documented pervasive loss of MgO in serpentinites from slow- and ultraslow-spreading mid-ocean ridges and fracture zones (Snow & Dick, 1995; Niu, 2004; Ligi et al., 2013). Although weathering of serpentinites in young oceanic lithosphere may represent a significant source of Mg for seawater, there is limited information on further compositional changes during weathering as serpentinites move off-axis. Similarly, approximately two-thirds of the iron in completely serpentinized (unweathered) peridotite is ferric. Although weathering of serpentinite can cause further oxidation, associated mineralogical and chemical effects remain poorly understood. Serpentinite is believed to be one of the principal carriers of water and fluid-mobile elements entering subduction zones (Scambelluri et al., 1995; Ulmer & Trommsdorff, 1995; Schmidt & Poli, 1998; Rüpke et al., 2004; Deschamps et al., 2011). Geophysical studies suggest that serpentinite in subducting oceanic plates forms through bending faults near the outer rise of subduction zones (Ranero et al., 2003; Van Avendonk et al., 2011). Research on bending faults has focused on subduction zones in the Pacific, where the presence of serpentinite is inferred from wide-angle seismic refraction studies, passive-source seismology, and electromagnetic studies, yet the extent to which the upper mantle below Pacific oceanic crust is serpentinized is the subject of continued research (Peacock, 2001; Ranero et al., 2003; Rüpke et al., 2004; Hacker, 2008; Key et al., 2012; Lefeldt et al., 2012; Reynard, 2013). In contrast, Korenaga (2017) suggested that bending of subducting plates does not create sufficiently negative dynamic pressure to allow mantle hydration and concluded that the geophysical observations of previous studies (e.g. Ranero et al., 2003) can be interpreted in terms of crack-like porosity, instead of serpentinization. Current models imply that serpentinite in bending faults forms through reactions of mantle peridotite with hydrothermal fluids that pass through several kilometers of crustal lithologies. Such a fluid would be strongly enriched in silica, causing talc alteration (steatitization) rather than serpentinization (e.g. Bach et al., 2013). Unfortunately, the composition and water content of hydrothermally altered mantle formed through bending faults in subduction zones have remained uncertain, mainly because serpentinites are not directly accessible on the seafloor of subducting plates in the Pacific. The Antillean arc is one of only two active subduction zones worldwide where slow-spreading oceanic lithosphere is recycled (Fig. 1). Bowin et al. (1966) provided a detailed account of serpentinized peridotite and associated basement rocks dredged from the North Wall of the Puerto Rico Trench (NWPRT), and discussed apparent similarities between NWPRT serpentinites and those from other localities; however, the origin of the NWPRT serpentinites is ambiguous. Whereas foraminifera and radiolaria in sedimentary rocks overlying NWPRT basement rocks are of Cenomanian age (Bowin et al., 1966; Chase & Hersey, 1968), Nalwalk (1969) reported Pliocene ages for the basalt, and inferred that the serpentinites intruded upper Cretaceous to Eocene sediments. Shido et al. (1974) interpreted the compositional similarities between NWPRT and Mid-Atlantic Ridge basalts as an indication of Cretaceous ages. Fig. 1. Open in new tabDownload slide Maps illustrating the North Atlantic Ocean during the Early Cretaceous and in its current configuration including the Puerto Rico Trench. (a) Bathymetric map showing the faulted blocks of the subducting North American Plate [modified from Andrews et al. (2014)]. Red–white symbol indicates location of dredges D2 and D10 on the North Wall of the Puerto Rico Trench (NWPRT). Black arrow indicates plate motion of the North American Plate relative to the Caribbean Plate. (b) Map showing the seafloor age (Ma) of the present-day North Atlantic including the Puerto Rico Trench. Oceanic basement of the North American Plate accreted at the Mid-Atlantic Ridge is moving off-axis toward the Antillean arc (AA) where it subducts underneath the Caribbean Plate (CP). Red–white symbol indicates location of samples examined in this study. Following the flow line of the 15°20’ fracture zone (bold black line) and taking changes of spreading rates into account (Müller et al., 2016), it appears that the NPWRT serpentinites examined in this study originated c. 2200 km east of their current location. CQZ, Cretaceous Quiet Zone. (c) Paleogeographical reconstruction of the Atlantic Ocean during the Early Cretaceous. The reconstructed position of NWPRT serpentinites examined in this study (dredge D10 and probably D2) at 112 ± 3 Ma is c. 13°N and 29°W. Maps in (b) and (c) and flow line (bold black line) were generated using the GPlates software version 2.0 (Williams et al., 2012) with geomagnetic reversal positions (fine black lines) from Müller et al. (2016). (See text for discussion.) Fig. 1. Open in new tabDownload slide Maps illustrating the North Atlantic Ocean during the Early Cretaceous and in its current configuration including the Puerto Rico Trench. (a) Bathymetric map showing the faulted blocks of the subducting North American Plate [modified from Andrews et al. (2014)]. Red–white symbol indicates location of dredges D2 and D10 on the North Wall of the Puerto Rico Trench (NWPRT). Black arrow indicates plate motion of the North American Plate relative to the Caribbean Plate. (b) Map showing the seafloor age (Ma) of the present-day North Atlantic including the Puerto Rico Trench. Oceanic basement of the North American Plate accreted at the Mid-Atlantic Ridge is moving off-axis toward the Antillean arc (AA) where it subducts underneath the Caribbean Plate (CP). Red–white symbol indicates location of samples examined in this study. Following the flow line of the 15°20’ fracture zone (bold black line) and taking changes of spreading rates into account (Müller et al., 2016), it appears that the NPWRT serpentinites examined in this study originated c. 2200 km east of their current location. CQZ, Cretaceous Quiet Zone. (c) Paleogeographical reconstruction of the Atlantic Ocean during the Early Cretaceous. The reconstructed position of NWPRT serpentinites examined in this study (dredge D10 and probably D2) at 112 ± 3 Ma is c. 13°N and 29°W. Maps in (b) and (c) and flow line (bold black line) were generated using the GPlates software version 2.0 (Williams et al., 2012) with geomagnetic reversal positions (fine black lines) from Müller et al. (2016). (See text for discussion.) In this study we expand on several earlier studies on NWPRT serpentinite and talc-altered serpentinite (Bowin et al., 1966; Chase & Hersey, 1968; Nalwalk, 1969; Wenner & Taylor, 1973; Perfit et al., 1980) to assess further the conditions of their emplacement and hydrothermal alteration, and to provide insight into the long-term weathering processes affecting serpentinite. We compare the mineralogical and chemical composition of NWPRT serpentinites with those from the present-day Mid-Atlantic Ridge and other settings, and discuss implications that arise from low-temperature post-serpentinization ‘open-system’ weathering. In addition, we performed U–Pb dating of porous zircon in NWPRT serpentinite to assess the origin and alteration history of serpentinite in a paleogeographical context. GEOLOGICAL SETTING AND SAMPLING OF THE PUERTO RICO TRENCH Oceanic lithosphere created at the MAR is subducted beneath the Antillean arc (Fig. 1). The Puerto Rico Trench (PRT) is located north of the Antillean arc, and represents the plate boundary between the North American Plate (NAP) and the Caribbean Plate. The movement of the NAP relative to the Caribbean plate shows a transition from subduction at the Lesser Antillean arc and oblique subduction (N70°E) at the Greater Antillean arc, to sinistral strike-slip faulting, which extends from the PRT through Hispaniola to the Cayman Trough. This plate motion tears the NAP open through a complex tectonic mechanism that Molnar & Sykes (1969) described as ‘hinge-faulting’. The combination of hinge faulting and large offset normal faulting along the NWPRT has created a tectonic window into the NAP in an orientation that is almost parallel to that of seafloor spreading at the MAR. Assuming an average half spreading rate of 25 mm a−1 (Müller et al., 2008), the ∼500 km wide NWPRT exposes almost 20 Myr of Cretaceous spreading history at the MAR. However, between 120 Ma (magnetic anomaly M0) and 84 Ma (C34), the spreading rates are difficult to quantify owing to the lack of strong magnetic reversals during the Cretaceous Normal Superchron. The PRT is the deepest part of the Atlantic Ocean and it has the greatest negative gravity anomaly worldwide (ten Brink, 2005). In the 1960s and 1970s, the R.V. Chain, R.V. Robert Conrad, and R.V. Eastward recovered several dredges containing completely serpentinized peridotite and talc-altered serpentinite, diabase, basalt, and sedimentary rocks from the NWPRT (Bowin et al., 1966; Chase & Hersey, 1968; Nalwalk, 1969; Hart & Nalwalk, 1970; Shido et al., 1974; Perfit et al., 1980). The present study focuses on serpentinized peridotites and talc-altered serpentinites that were collected during cruise 19 of the R.V. Chain (Fig. 1) at water depths of ∼6850–6700 m (dredge D10, 20°00’00“N, 66°31’57“W) and ∼6950–6875 m (dredge D2, 20°00’00“N, 66°24’33“W) in the NWPRT (Bowin et al., 1966; Nalwalk, 1969). Shallower dredges in the same area recovered basalt and sedimentary units of Cretaceous age (Bowin et al., 1966). Another dredge (R.V. Chain cruise 34, dredge D2) from 7300 m depth at 19°57’00”N, 66°28’00”W recovered completely talc-altered serpentinite and hydrothermally altered diabase (Bowin et al., 1966; Nalwalk, 1969). Several dredges during cruises of the R.V. Chain (CH-34), R.V. Robert Conrad (RC8), and R.V. Eastward (E-1F-75) sampled basaltic lithologies but no serpentinite in an area between 65°45’W and 64°07’W at depths greater than 6500 m (Shido et al., 1974; Perfit et al., 1980). Shido et al. (1974) speculated that the recovered tholeiitic basalt originally formed at the MAR and was subsequently moved to its present position by seafloor spreading. Because serpentinites were not found in the eastern part of the NWPRT, Perfit et al. (1980) concluded that serpentinite does not form a continuous layer and was probably tectonically emplaced. METHODS All analyses were performed at the Woods Hole Oceanographic Institution (WHOI), Woods Hole, Massachusetts, unless indicated otherwise. Samples were inspected in transmitted and reflected light using a Zeiss AxioImager 2 microscope. Complemented by scanning electron microscopy in low-vacuum mode (SEM; Hitachi TM3000 equipped with a Bruker Quantax 70 energy-dispersive X-ray spectrometer; EDS), confocal Raman spectroscopy and electron microprobe analysis (EMPA), these efforts allowed in most instances unequivocal identification of primary (where present) and secondary minerals. Major element compositions of Cr-spinel, olivine, and secondary minerals were quantified with a JEOL JXA-8530F ‘HyperProbe’ field-emission electron probe microanalyzer (Yale University, New Haven, CT). Microprobe analyses were conducted using a 15 kV acceleration voltage and 20 nA beam current. The beam was fully focused for Cr-spinel and olivine, but to avoid beam damage we used a 5–15 μm diameter for hydrous phases. The counting times for all elements were 20 s except for F, which was measured for 120 s. The raw data were processed using the CITZAF method (Armstrong, 1995). The accuracy and precision of electronprobe microanalysis is approximately 1%. Raman spectroscopy was carried out using a Horiba LabRam HR confocal Raman spectrometer equipped with three lasers (633 nm, 532 nm, 473 nm), two gratings (600 and 1800 grooves per mm), a spectrometer with 800 mm focal length, and a Peltier-cooled (–70°C) 1024 × 256 pixel charge coupled device (CCD) detector. The use of a 100× objective with a numerical aperture of 0·9, 473 nm laser and a 100 μm confocal hole diameter allowed an effective spatial resolution better than 1 μm. Spectra were collected between 100 and 1200 cm−1 (5 s and five accumulations for most analyses), and 3500 and 3800 cm−1 (10 s and five accumulations) and the 600 grooves per mm grating was used together with a slit size of 100 μm. Where a higher spectral resolution was necessary for mineral identification, the 1800 grooves per mm grating was used. Hyperspectral maps were acquired with a step size of 0·8 μm using a computer-controlled motorized stage. Spectra and hyperspectral maps were processed with the LabSpec 6 software including the multivariate spectral analysis module. Collected Raman spectra were compared with reference spectra for mineral identification (Downs, 2006; Petriglieri et al., 2015). Variations in the composition of lizardite and chrysotile can make an unequivocal identification challenging if only the Raman bands in the 200–700 cm−1 spectral range are considered. Because the Raman bands of OH stretching vibrations of chrysotile, lizardite, antigorite, and polygonal serpentine are distinct in the spectral range of 3600–3750 cm−1, an unambiguous discrimination of these phases is possible (Petriglieri et al., 2015). Loss on ignition and major and trace element analyses of whole-rock samples were determined at the Peter Hooper GeoAnalytical Laboratory at Washington State University, Pullman, Washington, using X-ray fluorescence analysis (XRF; ThermoARL Advant’XP+) and inductively coupled plasma mass spectrometry (ICP-MS; Agilent 7700) following procedures described previously (Lichte et al., 1987; Jarvis, 1988; Longerich et al., 1990; Johnson et al., 1999). For comparison with samples from the NWPRT, the whole-rock major element compositions of additional samples recovered by seafloor drilling (to minimize weathering effects) were analyzed using XRF (GeoAnalytical Laboratory), including Deep Sea Drilling Project (DSDP) Leg 82 (Site 558 located south of the Pico Fracture Zone in 35 Ma crust), Ocean Drilling Program (ODP) Legs 107 (Tyrrhenian Sea), 125 (Conical Seamount), 147 (Hess Deep), 149 and 173 (Iberia Margin), and 153 (Kane Fracture Zone area). Concentrations of FeO were determined by mass balance from XRF data and titration of acid digested whole-rock samples conducted by Activation Laboratories (Ancaster, Canada) using techniques adapted from Wilson (1955). Helium-pycnometry (Micromeritics AccuPyc II) was conducted to measure the true (‘skeletal’) volume of samples as He occupies the internal pore space. ‘Envelope’ volumes (the volume occupied by rock + internal pore space) were approximated by performing He-pycnometry of water-saturated samples. Porosities were then approximated by subtracting the measured skeletal volume of dry samples from the measured volume of water-saturated samples. Porosities were determined for all samples of dredge D2 and for samples of dredge D10 where sufficient sample material was available. Magnetic susceptibilities (K) of serpentinites were measured with a Bartington MS2B sensor and a Bartington MS2 magnetic susceptibility meter. Thin section microscopy, SEM, Raman, and cathodoluminescence analysis revealed zircon in sample Chain 19 D10-9. Zircon was separated from the bulk-rock sample by standard crushing and magnetic separation techniques. Individual grains were subsequently handpicked under a binocular microscope, based on clarity and crystal morphology. To overcome the effects of radioactive decay induced crystal defects and associated lead loss resulting in discordant analyses, zircon grains were pretreated by the method of thermal annealing and chemical leaching for chemical abrasion thermal ionization mass spectrometry (CA-TIMS) at the Massachusetts Institute of Technology (Mattinson, 2005). This method involves heating zircons inside a furnace at 900°C for 60 h. The annealed grains are subsequently loaded into FEP Teflon® microcapsules and leached in concentrated HF at 210°C within high-pressure vessels for 12 h. The partially dissolved sample is then transferred into Savillex® FEP beakers for rinsing. The leached material is decanted with several milliliters of ultra-pure water and fluxed successively with 4N HNO3 and 6N HCl on a hot plate and/or in an ultrasonic bath. After final rinsing with ultra-pure water, zircon grains are loaded back into their microcapsules, spiked with a mixed 205Pb–233U–235U tracer solution and dissolved completely in concentrated HF at 210°C for 48 h. Essentially the high-U parts of the zircon crystals that are associated with Pb loss are preferentially removed, leaving a residue of relatively low U content. After extensive testing, it was concluded that this method is the best possible way to obtain the most concordant analyses. The total procedural blank was less than 0·1 pg for U. All ages were calculated using the decay constant λ238 = 1·55125E – 10 and λ235 = 9·8485E – 10 (Jaffey et al., 1971). RESULTS Petrography Bowin et al. (1966) provided detailed petrographic descriptions of serpentinites, which are complemented with our own petrographic observations of additional thin sections of the same samples from dredges Chain 19 D2 and D10. Owing to the fine intergrowths of secondary minerals, petrographic observations were matched with results from confocal Raman spectroscopy, electron microprobe analysis, and scanning electron microscopy. Thin section mosaics, photomicrographs, hyperspectral Raman maps, and backscattered electron images of representative mineral assemblages and textures are presented in Figs 2 and 3 and Supplementary Data Figs 1–4 (supplementary data are available for downloading at http://www.petrology.oxfordjournals.org). With the exception of relict Cr-spinel, rare orthopyroxene, and one olivine grain (sample D10-3), all rocks are completely altered. Clinopyroxene is not preserved in NWPRT serpentinites. Serpentine is by far the most abundant mineral group in all samples, being dominantly lizardite in dredge D10 and antigorite in dredge D2 (see Bowin et al., 1966). Primary modes were not explicitly quantified because bastite after orthopyroxene and clinopyroxene can be indistinguishable optically. However, under crossed polarizers almost all bastite textures reveal extinction parallel to exsolution lamellae, suggesting that the protolith was clinopyroxene poor. On the basis of thin section mosaic image analysis of pseudomorphic textures it appears that the protoliths were most probably harzburgitic, containing approximately 75–85 vol. % olivine, 10–20 vol. % orthopyroxene, 1–4 vol. % clinopyroxene, and 1 vol. % spinel (Supplementary Data Fig. 1). Fig. 2. Open in new tabDownload slide Alteration features of sample D10-9. (a) Thin section photomicrograph (crossed polarizers) showing a chlorite (Chl)–antigorite (Atg) vein that hosts zircon (Zrn). The vein cuts through lizardite (Lz) in a pseudomorphic mesh and bastite texture. A late chrysotile (Ctl) vein cuts the chlorite–antigorite vein. Magnetite is altered to hematite (Hem). (b) Backscattered electron (BSE) image of the chlorite–antigorite vein surrounded by lizardite. (c) Cathodoluminescence (CL) image of zircon. Its mottled appearance, which is unlike that of igneous zircon, should be noted. (d) Hyperspectral Raman map of lizardite, chrysotile, goethite (Gth), and hematite in mesh texture. The step size is 0·8 μm. (e) Thin section mosaic in transmitted, cross-polarized light. Several vein generations can be distinguished. Foliation of mesh and bastite texture created shear phacoids. Late chrysotile veins did not experience extensive shearing. White dashed lines outline the antigorite–chlorite vein that hosts zircon. (f) BSE image illustrating an antigorite–chlorite vein cutting through a quartz–hematite assemblage. Chrysotile vein cuts the antigorite–chlorite vein. Void spaces reflect mineral dissolution. (g) BSE image of quartz (gray) outlined by hematite (light gray) as in (f). (h) BSE image showing lizardite and hematite in mesh texture. (i) BSE image showing lizardite (dark grey) and hematite (light grey) in bastite texture. (j) BSE image of lizardite, hematite, and goethite in mesh texture. (k) Raman spectra of selected minerals in sample D10-9. Antigorite, chrysotile, and lizardite have characteristic OH-stretching bands, which can be used to distinguish them from each other (see Petriglieri et al., 2015). The OH stretching bands of chrysotile are slightly shifted toward lower wavenumbers, possibly owing to compositional differences from the ideal endmember. The Raman spectrum of zircon is characteristic of well-crystallized, non-metamict zircon. (See text for discussion). Fig. 2. Open in new tabDownload slide Alteration features of sample D10-9. (a) Thin section photomicrograph (crossed polarizers) showing a chlorite (Chl)–antigorite (Atg) vein that hosts zircon (Zrn). The vein cuts through lizardite (Lz) in a pseudomorphic mesh and bastite texture. A late chrysotile (Ctl) vein cuts the chlorite–antigorite vein. Magnetite is altered to hematite (Hem). (b) Backscattered electron (BSE) image of the chlorite–antigorite vein surrounded by lizardite. (c) Cathodoluminescence (CL) image of zircon. Its mottled appearance, which is unlike that of igneous zircon, should be noted. (d) Hyperspectral Raman map of lizardite, chrysotile, goethite (Gth), and hematite in mesh texture. The step size is 0·8 μm. (e) Thin section mosaic in transmitted, cross-polarized light. Several vein generations can be distinguished. Foliation of mesh and bastite texture created shear phacoids. Late chrysotile veins did not experience extensive shearing. White dashed lines outline the antigorite–chlorite vein that hosts zircon. (f) BSE image illustrating an antigorite–chlorite vein cutting through a quartz–hematite assemblage. Chrysotile vein cuts the antigorite–chlorite vein. Void spaces reflect mineral dissolution. (g) BSE image of quartz (gray) outlined by hematite (light gray) as in (f). (h) BSE image showing lizardite and hematite in mesh texture. (i) BSE image showing lizardite (dark grey) and hematite (light grey) in bastite texture. (j) BSE image of lizardite, hematite, and goethite in mesh texture. (k) Raman spectra of selected minerals in sample D10-9. Antigorite, chrysotile, and lizardite have characteristic OH-stretching bands, which can be used to distinguish them from each other (see Petriglieri et al., 2015). The OH stretching bands of chrysotile are slightly shifted toward lower wavenumbers, possibly owing to compositional differences from the ideal endmember. The Raman spectrum of zircon is characteristic of well-crystallized, non-metamict zircon. (See text for discussion). Fig. 3. Open in new tabDownload slide Mineral assemblages of samples from dredges D2 and D10. (a) BSE image of weathered mesh texture in sample D10-8. Mesh centers (dark grey) are porous and consist of lizardite that has a slightly lower iron content than lizardite in mesh rims (medium grey). Magnetite is altered to hematite (white). (b) BSE image of antigorite and minor lizardite in non-pseudomorphic texture of sample D2-2. (c) Hyperspectral Raman map of lizardite, hematite, and goethite in mesh texture of sample D10-4. Step size is 0·8 μm. (d) BSE image of relict olivine armored by Cr-spinel in sample D10-3. (e) BSE image of antigorite and accessory hematite in porous non-pseudomorphic texture of sample D2-2. Inset shows hematite (medium grey) together with goethite (dark grey) in antigorite matrix. (f) BSE image of partially altered orthopyroxene, lizardite and accessory hematite in sample D10-4. Goethite forms millimeter-sized concretions. (g) BSE image of lizardite and minor hematite in mesh and bastite texture of sample D10-4. Abbreviations are the same as in Fig. 2. Fig. 3. Open in new tabDownload slide Mineral assemblages of samples from dredges D2 and D10. (a) BSE image of weathered mesh texture in sample D10-8. Mesh centers (dark grey) are porous and consist of lizardite that has a slightly lower iron content than lizardite in mesh rims (medium grey). Magnetite is altered to hematite (white). (b) BSE image of antigorite and minor lizardite in non-pseudomorphic texture of sample D2-2. (c) Hyperspectral Raman map of lizardite, hematite, and goethite in mesh texture of sample D10-4. Step size is 0·8 μm. (d) BSE image of relict olivine armored by Cr-spinel in sample D10-3. (e) BSE image of antigorite and accessory hematite in porous non-pseudomorphic texture of sample D2-2. Inset shows hematite (medium grey) together with goethite (dark grey) in antigorite matrix. (f) BSE image of partially altered orthopyroxene, lizardite and accessory hematite in sample D10-4. Goethite forms millimeter-sized concretions. (g) BSE image of lizardite and minor hematite in mesh and bastite texture of sample D10-4. Abbreviations are the same as in Fig. 2. Dredge D10 Serpentinites from dredge D10 display pseudomorphic mesh and bastite textures after olivine and pyroxene, respectively, in addition to transitional hourglass and ribbon textures, and minor non-pseudomorphic interpenetrating and interlocking textures (Figs 2 and 3; Supplementary Data Figs 1 and 3). Mesh and bastite textures resemble those of serpentinites from the MAR, with lizardite-rich mesh cells that are outlined by magnetite. Unlike serpentinite from the MAR, magnetite in NWPRT serpentinite is, in most instances, completely altered to hematite (Figs 2 and 3, Supplementary Data Figs 1, 3 and 4). The latter can be partially to completely altered to goethite (Figs 2d, j and 3c). Porous mesh centers appear darker than mesh rims in backscattered electron mode (Fig. 3a). Mesh centers do not contain any brucite or Ni—Fe sulfides. Bastite is composed of lizardite, minor chrysotile ± chlorite, and accessory hematite (Fig. 3f and g). Chlorite can be locally abundant in bastite textures (Supplementary Data Fig. 1, D10-12). At least two generations of serpentine veins can be distinguished, with earlier veins chiefly composed of lizardite and magnetite that are partially to completely replaced by hematite and goethite, and a later generation of monomineralic chrysotile veins (Supplementary Data Fig. 1). Except for sample D10-25, which is brecciated and cemented by calcite, serpentinites from dredge D10 are carbonate-poor. Antigorite occurs together with chlorite in veins crossing the mesh-textured matrix, mainly composed of lizardite in sample D10-9 (Fig. 2a, b and f). Bastite after orthopyroxene in sample D10-9 is talc-rich, and tremolite occurs in bastite after clinopyroxene (and probably clinopyroxene lamellae exsolved from orthopyroxene). Phacoid mesh cells and bastite with kink bands and undulose extinction are common in sample D10-9 (Fig. 2e). Late chrysotile/lizardite veins have remained largely unsheared and cut sheared antigorite veins, as well as sheared mesh and bastite textures. Antigorite—chlorite veins in sample D10-9 contain accessory zircon, which displays a complex irregular internal texture that appears porous in transmitted light and cloudy when examined in cathodoluminescence (Fig. 2a and c). Serpentinite is partially altered to poorly crystalline clay (Supplementary Data Fig. 4c) and, in a few cases, to quartz intergrown with goethite (Fig. 3f and g). Spinel is generally only partially altered to ferrichromite along grain boundaries and intra-grain fractures; however, in some cases, spinel is completely altered to ferrichromite. Dredge D2 Samples from dredge D2 are chiefly composed of antigorite and display mostly non-pseudomorphic and transitional textures (Fig. 3b and e;Supplementary Data Fig. 2). Pseudomorphic textures are apparent in parts of samples D2-2 and D2-4. Individual ‘blades’ of antigorite several hundred micrometers across occur in interpenetrating and interlocking textures. In sample D2-6, antigorite is, in places, associated with lizardite, but in most areas antigorite is the only serpentine mineral present (Fig. 3; Supplementary Data Fig. 2). Minor amounts of chlorite are intergrown with antigorite in most samples and in some cases chrysotile occurs together with antigorite. Samples D2-1, D2-3, and D2-6 are strongly talc-altered, whereas talc is only a minor phase in samples D2-2, D2-4, and D2-5 (Supplementary Data Fig. 2). In the latter samples, magnetite is altered to hematite and goethite. In the former, alteration to monomineralic talc in veins and along reaction fronts has completely obliterated precursor minerals and textures, but where talc alteration is incomplete hematite and goethite are also present. In contrast to samples from dredge D10, there is no unaltered Cr-spinel in samples from dredge D2; it is completely altered to ferrichromite, magnetite, and chlorite. Mineral compositions Electron microprobe analyses of serpentine, talc, and chlorite are given in Supplementary Data Table 1. The one single olivine grain in sample D10-3 is armored by unaltered Cr-spinel (Fig. 3d) and has a magnesium number [Mg# = 100 × Mg/(Mg + Fe)] of 92. Subsolidus exchange between the spinel host and olivine possibly led to an increase in the forsterite content of olivine by 1–1·5 Mg#s. Unusually high measured Cr2O3 contents of up to 0·7 wt % may be due to an analytical interference with the Cr-spinel host. Spinel surrounding the olivine has a chromium number [Cr# = 100 × Cr/(Cr + Al)] of 45 and a Mg# of 62, with a TiO2 content of ∼0·3 wt %. Electron microprobe analyses of spinel were complemented by semi-quantitative SEM-EDS analyses. Calculated Cr#s using both methods show reasonable agreement. The Cr#s of spinel in other samples from dredge D10 vary from 36 to 46, with Mg#s of 61–68. The Cr contents of spinel vary by ∼1·5 Cr# (EMPA) and TiO2 concentrations range from <0·01 wt % to 0·3 wt % in individual samples. We analyzed serpentine in pseudomorphic mesh and bastite textures, in veins and other non-pseudomorphic textures. Major element compositions of lizardite, chrysotile, and antigorite are illustrated in the ternary chemograph for Mg—Fe—Si (Fig. 4). In addition to variations in the relative proportions of Mg, Fe and Si in serpentine, spread is attributed to variations in the Fe(III)/Fetot. Most compositions fall between the tie-line connecting lizardite [Mg3Si2O5(OH)4] with the hypothetical dioctahedral endmember ferric lizardite [Fe2Si2O5(OH)4] and the one connecting antigorite [Mg48Si34O85(OH)62] with the hypothetical dioctahedral endmember ferric antigorite [Fe32Si34O85(OH)62]. Raman spectroscopy of samples from dredge D10 confirms that lizardite is the dominant serpentine mineral, but a number of lizardite analyses fall above the lizardite–ferric-lizardite tie-line, with (Mg + Fe)/Si similar to that of antigorite. Lizardite in mesh texture has an Mg# of 94–98. Mesh centers can be more magnesian (Mg# = 97–98) than mesh rims (Mg# = 94–95), such as in sample D10-8 [see Figs 3a and 4; see also Supplementary Data (SD) Fig. 5 and SD Table 1]. Lizardite in the bastite texture of samples D10-3, D10-4, and D10-9 has a lower Mg# (90–93) than bastite of samples D10-7 and D10-8 (95–97). Antigorite tends to have a lower Mg# than lizardite and chrysotile, although there is considerable overlap. Apparent trends towards Si in samples D10-9 and D10-4, and in samples from dredge D2, can be explained by mixed analyses with talc (Fig. 4). Outliers toward Fe are probably due to mixed analysis with submicrometer-sized hematite or goethite. Fig. 4. Open in new tabDownload slide Electron microprobe data for lizardite, chrysotile, and antigorite illustrated in a Mg–Fe–Si chemograph. Most analyses fall above the tie-line between lizardite [Mg3Si2O5(OH)4] and hypothetical endmember dioctahedral ferric lizardite [Fe2Si2O5(OH)4] and below the tie-line between antigorite [Mg48Si34O85(OH)62] and hypothetical endmember dioctahedral ferric antigorite [Fe32Si34O85(OH)62]. In addition to variations in the Fe valence and relative proportions of Mg, Fe, and Si, spread is influenced by mixed analyses with minor talc and chlorite. (See text for discussion). Fig. 4. Open in new tabDownload slide Electron microprobe data for lizardite, chrysotile, and antigorite illustrated in a Mg–Fe–Si chemograph. Most analyses fall above the tie-line between lizardite [Mg3Si2O5(OH)4] and hypothetical endmember dioctahedral ferric lizardite [Fe2Si2O5(OH)4] and below the tie-line between antigorite [Mg48Si34O85(OH)62] and hypothetical endmember dioctahedral ferric antigorite [Fe32Si34O85(OH)62]. In addition to variations in the Fe valence and relative proportions of Mg, Fe, and Si, spread is influenced by mixed analyses with minor talc and chlorite. (See text for discussion). Lizardite in mesh texture has lower Cr2O3 and Al2O3 and higher NiO concentrations than lizardite in bastite texture, though overlapping Mg# (Fig. 5). However, with concentrations of up to 1·1 wt % in sample D10-7, Al2O3 can be strongly enriched in lizardite in mesh texture relative to its olivine precursor, which is generally Al-poor. Fig. 5. Open in new tabDownload slide U–Pb concordia diagram for zircon single grain dates (sample D10-9). The 2σ uncertainties on the concordia are illustrated by grey bands (Jaffey et al., 1971). The derived age is consistent with accretion and hydrothermal alteration of peridotite and mafic veins at the Cretaceous Mid-Atlantic Ridge. The diagram was constructed with the U–Pb_Redux software package (Bowring et al., 2011; McLean et al., 2011). Ages on concordia are in million years. MSWD, mean square weighted deviation. Fig. 5. Open in new tabDownload slide U–Pb concordia diagram for zircon single grain dates (sample D10-9). The 2σ uncertainties on the concordia are illustrated by grey bands (Jaffey et al., 1971). The derived age is consistent with accretion and hydrothermal alteration of peridotite and mafic veins at the Cretaceous Mid-Atlantic Ridge. The diagram was constructed with the U–Pb_Redux software package (Bowring et al., 2011; McLean et al., 2011). Ages on concordia are in million years. MSWD, mean square weighted deviation. Measured concentrations of Cl do not seem to vary in a systematic manner with Na or K, suggesting that Cl is not dominantly hosted by salts, but instead is structurally bound in phyllosilicates (SD Table 1). However, some of the samples with the highest Cl contents also have the highest Na concentrations, which is probably due to mixed analysis with halite. Concentrations are highest in lizardite (up to ∼0·3 wt % in sample D10-8) and near or below the limit of detection (∼0·01 wt %) in chlorite, whereas concentrations of Cl in antigorite range from 0·05 to 0·1 wt %. Table 1: U–Pb data Sample . Concentrations . Isotopic ratios . Age (Ma) . fraction . Pb(c) (pg) . Pb/Pb . Th/U . 206Pb/ 204Pb . 208Pb/ 206Pb . 206Pb/ 238U . % err. . 207Pb/ 235U . % err. . 207Pb/ 206Pb . % err. . 206Pb/ 238U . abs. err. . 207Pb/ 235U . abs. err. . 207Pb/ 206Pb . abs. err. . Corr. coef. . (a) . (b) . (c) . . (c) . (d) . (e) . (f) . (e) . (f) . (e) . (f) . . (f) . . (f) . . (f) . . z1 0·15 6·4 0·48 403·7 0·152 0·018046 (·27) 0·12267 (3·28) 0·04932 (3·16) 115·29 0·31 117·5 3·6 162 74 0·49 z3 0·15 12·0 0·49 742·6 0·157 0·017950 (·16) 0·11955 (1·70) 0·04833 (1·64) 114·68 0·18 114·7 1·8 114 39 0·44 z4 0·15 10·1 0·52 622·8 0·165 0·018015 (·18) 0·12208 (1·88) 0·04917 (1·80) 115·10 0·20 117·0 2·1 155 42 0·47 z5 0·22 10·0 0·46 628·5 0·147 0·017985 (·16) 0·11999 (1·96) 0·04841 (1·90) 114·91 0·18 115·1 2·1 118 45 0·45 z6 0·19 17·5 0·48 1072·5 0·153 0·017967 (·10) 0·11976 (1·11) 0·04837 (1·07) 114·79 0·11 114·9 1·2 116 25 0·39 Sample . Concentrations . Isotopic ratios . Age (Ma) . fraction . Pb(c) (pg) . Pb/Pb . Th/U . 206Pb/ 204Pb . 208Pb/ 206Pb . 206Pb/ 238U . % err. . 207Pb/ 235U . % err. . 207Pb/ 206Pb . % err. . 206Pb/ 238U . abs. err. . 207Pb/ 235U . abs. err. . 207Pb/ 206Pb . abs. err. . Corr. coef. . (a) . (b) . (c) . . (c) . (d) . (e) . (f) . (e) . (f) . (e) . (f) . . (f) . . (f) . . (f) . . z1 0·15 6·4 0·48 403·7 0·152 0·018046 (·27) 0·12267 (3·28) 0·04932 (3·16) 115·29 0·31 117·5 3·6 162 74 0·49 z3 0·15 12·0 0·49 742·6 0·157 0·017950 (·16) 0·11955 (1·70) 0·04833 (1·64) 114·68 0·18 114·7 1·8 114 39 0·44 z4 0·15 10·1 0·52 622·8 0·165 0·018015 (·18) 0·12208 (1·88) 0·04917 (1·80) 115·10 0·20 117·0 2·1 155 42 0·47 z5 0·22 10·0 0·46 628·5 0·147 0·017985 (·16) 0·11999 (1·96) 0·04841 (1·90) 114·91 0·18 115·1 2·1 118 45 0·45 z6 0·19 17·5 0·48 1072·5 0·153 0·017967 (·10) 0·11976 (1·11) 0·04837 (1·07) 114·79 0·11 114·9 1·2 116 25 0·39 (a) Thermally annealed and pretreated single zircon. (b) Total common Pb in analyses. (c) Measured ratio corrected for spike and fractionation only. (d) Radiogenic Pb. (e) Corrected for fractionation, spike and blank. All common Pb assumed to be laboratory blank. Also corrected for initial Th/U disequilibrium using radiogenic 208Pb and Th/U[magma] = 2·8. (f) Errors (err.) are 2σ; abs. err., absolute error. Mass fractionation correction of 0·25%/a.m.u. ± 0·04%/a.m.u. (atomic mass unit) was applied to single-collector Daly analyses. Total procedural blank less than 0·1 pg for U. Blank isotopic composition: 206Pb/204Pb = 18·31 ± 0·53, 207Pb/204Pb = 15·38 ± 0·35, 208Pb/204Pb = 37·45 ± 1·1. Corr. coef., correlation coefficient. Ages calculated using the decay constants λ238 = 1·55125E – 10 and λ235 = 9·8485E – 10 (Jaffey et al., 1971). Table 1: U–Pb data Sample . Concentrations . Isotopic ratios . Age (Ma) . fraction . Pb(c) (pg) . Pb/Pb . Th/U . 206Pb/ 204Pb . 208Pb/ 206Pb . 206Pb/ 238U . % err. . 207Pb/ 235U . % err. . 207Pb/ 206Pb . % err. . 206Pb/ 238U . abs. err. . 207Pb/ 235U . abs. err. . 207Pb/ 206Pb . abs. err. . Corr. coef. . (a) . (b) . (c) . . (c) . (d) . (e) . (f) . (e) . (f) . (e) . (f) . . (f) . . (f) . . (f) . . z1 0·15 6·4 0·48 403·7 0·152 0·018046 (·27) 0·12267 (3·28) 0·04932 (3·16) 115·29 0·31 117·5 3·6 162 74 0·49 z3 0·15 12·0 0·49 742·6 0·157 0·017950 (·16) 0·11955 (1·70) 0·04833 (1·64) 114·68 0·18 114·7 1·8 114 39 0·44 z4 0·15 10·1 0·52 622·8 0·165 0·018015 (·18) 0·12208 (1·88) 0·04917 (1·80) 115·10 0·20 117·0 2·1 155 42 0·47 z5 0·22 10·0 0·46 628·5 0·147 0·017985 (·16) 0·11999 (1·96) 0·04841 (1·90) 114·91 0·18 115·1 2·1 118 45 0·45 z6 0·19 17·5 0·48 1072·5 0·153 0·017967 (·10) 0·11976 (1·11) 0·04837 (1·07) 114·79 0·11 114·9 1·2 116 25 0·39 Sample . Concentrations . Isotopic ratios . Age (Ma) . fraction . Pb(c) (pg) . Pb/Pb . Th/U . 206Pb/ 204Pb . 208Pb/ 206Pb . 206Pb/ 238U . % err. . 207Pb/ 235U . % err. . 207Pb/ 206Pb . % err. . 206Pb/ 238U . abs. err. . 207Pb/ 235U . abs. err. . 207Pb/ 206Pb . abs. err. . Corr. coef. . (a) . (b) . (c) . . (c) . (d) . (e) . (f) . (e) . (f) . (e) . (f) . . (f) . . (f) . . (f) . . z1 0·15 6·4 0·48 403·7 0·152 0·018046 (·27) 0·12267 (3·28) 0·04932 (3·16) 115·29 0·31 117·5 3·6 162 74 0·49 z3 0·15 12·0 0·49 742·6 0·157 0·017950 (·16) 0·11955 (1·70) 0·04833 (1·64) 114·68 0·18 114·7 1·8 114 39 0·44 z4 0·15 10·1 0·52 622·8 0·165 0·018015 (·18) 0·12208 (1·88) 0·04917 (1·80) 115·10 0·20 117·0 2·1 155 42 0·47 z5 0·22 10·0 0·46 628·5 0·147 0·017985 (·16) 0·11999 (1·96) 0·04841 (1·90) 114·91 0·18 115·1 2·1 118 45 0·45 z6 0·19 17·5 0·48 1072·5 0·153 0·017967 (·10) 0·11976 (1·11) 0·04837 (1·07) 114·79 0·11 114·9 1·2 116 25 0·39 (a) Thermally annealed and pretreated single zircon. (b) Total common Pb in analyses. (c) Measured ratio corrected for spike and fractionation only. (d) Radiogenic Pb. (e) Corrected for fractionation, spike and blank. All common Pb assumed to be laboratory blank. Also corrected for initial Th/U disequilibrium using radiogenic 208Pb and Th/U[magma] = 2·8. (f) Errors (err.) are 2σ; abs. err., absolute error. Mass fractionation correction of 0·25%/a.m.u. ± 0·04%/a.m.u. (atomic mass unit) was applied to single-collector Daly analyses. Total procedural blank less than 0·1 pg for U. Blank isotopic composition: 206Pb/204Pb = 18·31 ± 0·53, 207Pb/204Pb = 15·38 ± 0·35, 208Pb/204Pb = 37·45 ± 1·1. Corr. coef., correlation coefficient. Ages calculated using the decay constants λ238 = 1·55125E – 10 and λ235 = 9·8485E – 10 (Jaffey et al., 1971). Talc contains 5–6 wt % of structurally bound water. It has FeO contents c. 2–3 wt %, lower than those of antigorite, which it replaces in samples from dredge D2 (SD Table 1). Concentrations of CaO, TiO2, Cr2O3, CoO, MnO, and K2O are at or below the detection limit of the electron probe for these elements. Chlorite compositions are close to that of clinochlore, although with variable Fe contents (SD Fig. 5 and SD Table 1). In sample D10-9, chlorite is enriched in Fe and has an Mg# of 89–91 relative to chlorite in samples from D2, which have Mg# ranging from 94 to 96. U–Pb chronology and Raman spectroscopy of zircon Dating of porous zircon separated from sample D10-9 yielded ages ranging from 114·68 to 115·29 Ma. Three of the five ages produced a concordant 206Pb/238U age of 114·793 ± 0·085/0·12/0·17 Ma with a mean square weighted deviation of 1·5 (see Fig. 5 and Table 1; uncertainties refer to the three ages that produced the concordant age). Raman spectroscopy can be deployed to assess the crystallinity of zircon, which deteriorates during metamictization (Nasdala et al., 1995). Zircon from sample D10-9 shows well-defined Raman bands (Fig. 2k) between 200 and 230 cm−1. The antisymmetric stretching band of SiO4 (B1g) in zircon has a full width at half maximum of less than 5 cm−1 and its frequency ranges from 1006·5 to 1008·0 cm−1. These results indicate that zircon is well crystallized and was not affected by radiation damage and metamictization to a significant extent (see Nasdala et al., 1995). Major element compositions of whole-rock samples Major element compositions of whole-rock samples are given in Table 2 and illustrated in Figs 6–8, which, for comparison, also include the major element compositions of serpentinites from the MAR and other settings. It is important to bear in mind that whole-rock major and trace element data for sample D10-9 represent a mixed analysis of an ultramafic host rock and mafic veins. Table 2: Major element compositions of whole-rock samples (wt %), loss on ignition (LOI, wt %), titrated Fe(II) contents (wt %), skeletal densities (g cm-3), and magnetic susceptibilities (SI 10-2) from dredges D2 and D10 Sample no.: . D10-1 . D10-3 . D10-4 . D10-7 . D10-8 . D10-9 . D10-10 . D10-12 . D10-15 . D10-16 . D2-1 . D2-2 . D2-3 . D2-4 . D2-5 . D2-6 . Rock type: . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . SiO2 39·741 39·543 39·632 39·480 38·560 40·102 37·876 39·275 40·416 39·182 49·406 43·194 49·369 42·515 42·511 51·190 TiO2 0·014 0·024 0·018 0·025 0·014 0·067 0·026 0·028 0·034 0·021 0·004 0·004 0·003 0·002 0·003 0·002 Al2O3 1·055 0·692 0·989 1·173 1·032 1·275 1·362 0·551 0·744 1·420 0·698 0·608 0·712 0·367 0·413 0·520 FeO* 7·934 8·319 8·871 7·947 9·509 10·067 8·240 9·413 9·215 9·204 4·607 7·482 4·510 7·214 7·161 4·466 MnO 0·060 0·084 0·076 0·092 0·098 0·106 0·098 0·087 0·085 0·087 0·097 0·065 0·119 0·080 0·086 0·107 MgO 36·443 36·707 36·066 37·396 36·485 34·597 35·249 36·737 36·572 36·369 34·443 35·771 34·124 36·652 37·249 32·994 CaO 0·044 0·024 0·091 0·030 0·033 0·173 0·456 0·026 0·080 0·036 0·045 0·017 0·022 0·022 0·031 0·026 Na2O 0·179 0·171 0·167 0·176 0·222 0·168 0·433 0·072 0·109 0·286 0·089 0·117 <0·001 <0·001 <0·001 0·219 K2O 0·045 0·021 0·043 0·027 0·024 0·028 0·052 0·009 0·024 0·043 0·003 0·003 <0·001 <0·001 <0·001 0·003 P2O5 0·015 0·012 0·016 0·009 0·009 0·024 0·010 0·019 0·014 0·011 0·005 b.d.l. <0·001 <0·001 <0·001 0·004 LOI (%) 13·88 13·35 13·17 12·71 13·14 12·31 15·96 12·56 12·52 12·71 9·50 11·59 9·67 12·08 11·93 10·47 Sum 99·41 98·95 99·14 99·06 99·12 98·91 99·76 98·77 99·81 99·37 98·90 98·85 98·52 98·96 99·40 100·00 FeO† 0·4 0·6 1·5 0·2 b.d.l. 1·4 b.d.l. 0·9 0·9 0·3 3·2 2·5 3·1 3·2 3·3 3 Fe2O3‡ 8·37 8·58 8·19 8·61 10·46 9·63 9·05 9·46 9·24 9·90 1·56 5·54 1·57 4·46 4·29 1·63 Sum§ 100·2 99·8 100·0 99·9 100·2 99·9 100·7 99·7 100·7 100·4 99·1 99·4 98·7 99·4 99·8 100·2 Fe(III)/Fetot 0·95 0·93 0·85 0·98 0·99 0·87 0·99 0·91 0·91 0·97 0·33 0·69 0·34 0·58 0·57 0·35 Skeletal density 2·66 2·72 2·71 2·79 2·69 2·69 2·70 2·67 2·66 2·71 2·69 2·67 2·70 2·66 2·66 2·73 Envelope density 2·46 2·33 n.a. 2·41 n.a. n.a. n.a. n.a. 2·55 2·03 2·38 2·28 2·32 2·41 2·40 2·52 Porosity 6·52 11·75 n.c. 9·70 n.c. n.c. n.c. n.c. 4·05 23·32 11·64 14·55 13·90 9·54 9·89 7·73 Magn. susc. 0·278 0·319 0·849 1·618 1·261 1·844 1·822 2·123 1·521 1·966 0·031 2·259 0·061 0·115 0·143 0·035 Av. Cr# of spinel n.a. 44 37 34 35 43 n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. Sample no.: . D10-1 . D10-3 . D10-4 . D10-7 . D10-8 . D10-9 . D10-10 . D10-12 . D10-15 . D10-16 . D2-1 . D2-2 . D2-3 . D2-4 . D2-5 . D2-6 . Rock type: . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . SiO2 39·741 39·543 39·632 39·480 38·560 40·102 37·876 39·275 40·416 39·182 49·406 43·194 49·369 42·515 42·511 51·190 TiO2 0·014 0·024 0·018 0·025 0·014 0·067 0·026 0·028 0·034 0·021 0·004 0·004 0·003 0·002 0·003 0·002 Al2O3 1·055 0·692 0·989 1·173 1·032 1·275 1·362 0·551 0·744 1·420 0·698 0·608 0·712 0·367 0·413 0·520 FeO* 7·934 8·319 8·871 7·947 9·509 10·067 8·240 9·413 9·215 9·204 4·607 7·482 4·510 7·214 7·161 4·466 MnO 0·060 0·084 0·076 0·092 0·098 0·106 0·098 0·087 0·085 0·087 0·097 0·065 0·119 0·080 0·086 0·107 MgO 36·443 36·707 36·066 37·396 36·485 34·597 35·249 36·737 36·572 36·369 34·443 35·771 34·124 36·652 37·249 32·994 CaO 0·044 0·024 0·091 0·030 0·033 0·173 0·456 0·026 0·080 0·036 0·045 0·017 0·022 0·022 0·031 0·026 Na2O 0·179 0·171 0·167 0·176 0·222 0·168 0·433 0·072 0·109 0·286 0·089 0·117 <0·001 <0·001 <0·001 0·219 K2O 0·045 0·021 0·043 0·027 0·024 0·028 0·052 0·009 0·024 0·043 0·003 0·003 <0·001 <0·001 <0·001 0·003 P2O5 0·015 0·012 0·016 0·009 0·009 0·024 0·010 0·019 0·014 0·011 0·005 b.d.l. <0·001 <0·001 <0·001 0·004 LOI (%) 13·88 13·35 13·17 12·71 13·14 12·31 15·96 12·56 12·52 12·71 9·50 11·59 9·67 12·08 11·93 10·47 Sum 99·41 98·95 99·14 99·06 99·12 98·91 99·76 98·77 99·81 99·37 98·90 98·85 98·52 98·96 99·40 100·00 FeO† 0·4 0·6 1·5 0·2 b.d.l. 1·4 b.d.l. 0·9 0·9 0·3 3·2 2·5 3·1 3·2 3·3 3 Fe2O3‡ 8·37 8·58 8·19 8·61 10·46 9·63 9·05 9·46 9·24 9·90 1·56 5·54 1·57 4·46 4·29 1·63 Sum§ 100·2 99·8 100·0 99·9 100·2 99·9 100·7 99·7 100·7 100·4 99·1 99·4 98·7 99·4 99·8 100·2 Fe(III)/Fetot 0·95 0·93 0·85 0·98 0·99 0·87 0·99 0·91 0·91 0·97 0·33 0·69 0·34 0·58 0·57 0·35 Skeletal density 2·66 2·72 2·71 2·79 2·69 2·69 2·70 2·67 2·66 2·71 2·69 2·67 2·70 2·66 2·66 2·73 Envelope density 2·46 2·33 n.a. 2·41 n.a. n.a. n.a. n.a. 2·55 2·03 2·38 2·28 2·32 2·41 2·40 2·52 Porosity 6·52 11·75 n.c. 9·70 n.c. n.c. n.c. n.c. 4·05 23·32 11·64 14·55 13·90 9·54 9·89 7·73 Magn. susc. 0·278 0·319 0·849 1·618 1·261 1·844 1·822 2·123 1·521 1·966 0·031 2·259 0·061 0·115 0·143 0·035 Av. Cr# of spinel n.a. 44 37 34 35 43 n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. * All Fe calculated as FeO. † FeO determined by titration. ‡ Calculated by mass balance from XRF and titration data. § Sum taking whole-rock FeO and Fe2O3 contents into account. n.a., not analyzed; n.c., not calculated; b.d.l., below detection limit. Samples of dredge D2 do not contain fresh spinel. Table 2: Major element compositions of whole-rock samples (wt %), loss on ignition (LOI, wt %), titrated Fe(II) contents (wt %), skeletal densities (g cm-3), and magnetic susceptibilities (SI 10-2) from dredges D2 and D10 Sample no.: . D10-1 . D10-3 . D10-4 . D10-7 . D10-8 . D10-9 . D10-10 . D10-12 . D10-15 . D10-16 . D2-1 . D2-2 . D2-3 . D2-4 . D2-5 . D2-6 . Rock type: . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . SiO2 39·741 39·543 39·632 39·480 38·560 40·102 37·876 39·275 40·416 39·182 49·406 43·194 49·369 42·515 42·511 51·190 TiO2 0·014 0·024 0·018 0·025 0·014 0·067 0·026 0·028 0·034 0·021 0·004 0·004 0·003 0·002 0·003 0·002 Al2O3 1·055 0·692 0·989 1·173 1·032 1·275 1·362 0·551 0·744 1·420 0·698 0·608 0·712 0·367 0·413 0·520 FeO* 7·934 8·319 8·871 7·947 9·509 10·067 8·240 9·413 9·215 9·204 4·607 7·482 4·510 7·214 7·161 4·466 MnO 0·060 0·084 0·076 0·092 0·098 0·106 0·098 0·087 0·085 0·087 0·097 0·065 0·119 0·080 0·086 0·107 MgO 36·443 36·707 36·066 37·396 36·485 34·597 35·249 36·737 36·572 36·369 34·443 35·771 34·124 36·652 37·249 32·994 CaO 0·044 0·024 0·091 0·030 0·033 0·173 0·456 0·026 0·080 0·036 0·045 0·017 0·022 0·022 0·031 0·026 Na2O 0·179 0·171 0·167 0·176 0·222 0·168 0·433 0·072 0·109 0·286 0·089 0·117 <0·001 <0·001 <0·001 0·219 K2O 0·045 0·021 0·043 0·027 0·024 0·028 0·052 0·009 0·024 0·043 0·003 0·003 <0·001 <0·001 <0·001 0·003 P2O5 0·015 0·012 0·016 0·009 0·009 0·024 0·010 0·019 0·014 0·011 0·005 b.d.l. <0·001 <0·001 <0·001 0·004 LOI (%) 13·88 13·35 13·17 12·71 13·14 12·31 15·96 12·56 12·52 12·71 9·50 11·59 9·67 12·08 11·93 10·47 Sum 99·41 98·95 99·14 99·06 99·12 98·91 99·76 98·77 99·81 99·37 98·90 98·85 98·52 98·96 99·40 100·00 FeO† 0·4 0·6 1·5 0·2 b.d.l. 1·4 b.d.l. 0·9 0·9 0·3 3·2 2·5 3·1 3·2 3·3 3 Fe2O3‡ 8·37 8·58 8·19 8·61 10·46 9·63 9·05 9·46 9·24 9·90 1·56 5·54 1·57 4·46 4·29 1·63 Sum§ 100·2 99·8 100·0 99·9 100·2 99·9 100·7 99·7 100·7 100·4 99·1 99·4 98·7 99·4 99·8 100·2 Fe(III)/Fetot 0·95 0·93 0·85 0·98 0·99 0·87 0·99 0·91 0·91 0·97 0·33 0·69 0·34 0·58 0·57 0·35 Skeletal density 2·66 2·72 2·71 2·79 2·69 2·69 2·70 2·67 2·66 2·71 2·69 2·67 2·70 2·66 2·66 2·73 Envelope density 2·46 2·33 n.a. 2·41 n.a. n.a. n.a. n.a. 2·55 2·03 2·38 2·28 2·32 2·41 2·40 2·52 Porosity 6·52 11·75 n.c. 9·70 n.c. n.c. n.c. n.c. 4·05 23·32 11·64 14·55 13·90 9·54 9·89 7·73 Magn. susc. 0·278 0·319 0·849 1·618 1·261 1·844 1·822 2·123 1·521 1·966 0·031 2·259 0·061 0·115 0·143 0·035 Av. Cr# of spinel n.a. 44 37 34 35 43 n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. Sample no.: . D10-1 . D10-3 . D10-4 . D10-7 . D10-8 . D10-9 . D10-10 . D10-12 . D10-15 . D10-16 . D2-1 . D2-2 . D2-3 . D2-4 . D2-5 . D2-6 . Rock type: . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . SiO2 39·741 39·543 39·632 39·480 38·560 40·102 37·876 39·275 40·416 39·182 49·406 43·194 49·369 42·515 42·511 51·190 TiO2 0·014 0·024 0·018 0·025 0·014 0·067 0·026 0·028 0·034 0·021 0·004 0·004 0·003 0·002 0·003 0·002 Al2O3 1·055 0·692 0·989 1·173 1·032 1·275 1·362 0·551 0·744 1·420 0·698 0·608 0·712 0·367 0·413 0·520 FeO* 7·934 8·319 8·871 7·947 9·509 10·067 8·240 9·413 9·215 9·204 4·607 7·482 4·510 7·214 7·161 4·466 MnO 0·060 0·084 0·076 0·092 0·098 0·106 0·098 0·087 0·085 0·087 0·097 0·065 0·119 0·080 0·086 0·107 MgO 36·443 36·707 36·066 37·396 36·485 34·597 35·249 36·737 36·572 36·369 34·443 35·771 34·124 36·652 37·249 32·994 CaO 0·044 0·024 0·091 0·030 0·033 0·173 0·456 0·026 0·080 0·036 0·045 0·017 0·022 0·022 0·031 0·026 Na2O 0·179 0·171 0·167 0·176 0·222 0·168 0·433 0·072 0·109 0·286 0·089 0·117 <0·001 <0·001 <0·001 0·219 K2O 0·045 0·021 0·043 0·027 0·024 0·028 0·052 0·009 0·024 0·043 0·003 0·003 <0·001 <0·001 <0·001 0·003 P2O5 0·015 0·012 0·016 0·009 0·009 0·024 0·010 0·019 0·014 0·011 0·005 b.d.l. <0·001 <0·001 <0·001 0·004 LOI (%) 13·88 13·35 13·17 12·71 13·14 12·31 15·96 12·56 12·52 12·71 9·50 11·59 9·67 12·08 11·93 10·47 Sum 99·41 98·95 99·14 99·06 99·12 98·91 99·76 98·77 99·81 99·37 98·90 98·85 98·52 98·96 99·40 100·00 FeO† 0·4 0·6 1·5 0·2 b.d.l. 1·4 b.d.l. 0·9 0·9 0·3 3·2 2·5 3·1 3·2 3·3 3 Fe2O3‡ 8·37 8·58 8·19 8·61 10·46 9·63 9·05 9·46 9·24 9·90 1·56 5·54 1·57 4·46 4·29 1·63 Sum§ 100·2 99·8 100·0 99·9 100·2 99·9 100·7 99·7 100·7 100·4 99·1 99·4 98·7 99·4 99·8 100·2 Fe(III)/Fetot 0·95 0·93 0·85 0·98 0·99 0·87 0·99 0·91 0·91 0·97 0·33 0·69 0·34 0·58 0·57 0·35 Skeletal density 2·66 2·72 2·71 2·79 2·69 2·69 2·70 2·67 2·66 2·71 2·69 2·67 2·70 2·66 2·66 2·73 Envelope density 2·46 2·33 n.a. 2·41 n.a. n.a. n.a. n.a. 2·55 2·03 2·38 2·28 2·32 2·41 2·40 2·52 Porosity 6·52 11·75 n.c. 9·70 n.c. n.c. n.c. n.c. 4·05 23·32 11·64 14·55 13·90 9·54 9·89 7·73 Magn. susc. 0·278 0·319 0·849 1·618 1·261 1·844 1·822 2·123 1·521 1·966 0·031 2·259 0·061 0·115 0·143 0·035 Av. Cr# of spinel n.a. 44 37 34 35 43 n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. * All Fe calculated as FeO. † FeO determined by titration. ‡ Calculated by mass balance from XRF and titration data. § Sum taking whole-rock FeO and Fe2O3 contents into account. n.a., not analyzed; n.c., not calculated; b.d.l., below detection limit. Samples of dredge D2 do not contain fresh spinel. Fig. 6. Open in new tabDownload slide Whole-rock major element data illustrating the effects of weathering and talc alteration on MgO/SiO2 and Al2O3/SiO2 ratios. Weathering leads to a decrease in MgO/SiO2 owing to MgO loss, which is most apparent in samples from dredge D10. Talc-altered samples of dredge D2 are less weathered but have low Al2O3/SiO2 and MgO/SiO2 owing to addition of SiO2. The terrestrial melting array of peridotite and samples from ODP Leg 209 (Mid-Atlantic Ridge) are shown for comparison. [Adapted from Paulick et al. (2006) with additional data from Kodolanyi et al. (2012).] Fig. 6. Open in new tabDownload slide Whole-rock major element data illustrating the effects of weathering and talc alteration on MgO/SiO2 and Al2O3/SiO2 ratios. Weathering leads to a decrease in MgO/SiO2 owing to MgO loss, which is most apparent in samples from dredge D10. Talc-altered samples of dredge D2 are less weathered but have low Al2O3/SiO2 and MgO/SiO2 owing to addition of SiO2. The terrestrial melting array of peridotite and samples from ODP Leg 209 (Mid-Atlantic Ridge) are shown for comparison. [Adapted from Paulick et al. (2006) with additional data from Kodolanyi et al. (2012).] Fig. 7. Open in new tabDownload slide Variation diagrams showing skeletal density-major oxide relationships of NWPRT serpentinites from dredges D2 and D10 in comparison with serpentinites recovered by the DSDP and ODP from mid-ocean ridge (ODP Leg 147 and 153), MAR off-axis (DSDP Leg 82), continental margins (ODP Legs 107, 149, and 173), and subduction-zone (ODP Leg 125) settings. See text for discussion. Fig. 7. Open in new tabDownload slide Variation diagrams showing skeletal density-major oxide relationships of NWPRT serpentinites from dredges D2 and D10 in comparison with serpentinites recovered by the DSDP and ODP from mid-ocean ridge (ODP Leg 147 and 153), MAR off-axis (DSDP Leg 82), continental margins (ODP Legs 107, 149, and 173), and subduction-zone (ODP Leg 125) settings. See text for discussion. Fig. 8. Open in new tabDownload slide Variation diagram illustrating Fe(III)/Fetot vs SiO2 (wt %) for lizardite-serpentinites from dredge D10 and variably talc-altered antigorite-serpentinites from dredge D2. Talc-altered antigorite-serpentinites have lower Fe(III)/Fetot values than serpentinites that did not experience talc alteration. Additional data from Paulick et al. (2006). Fig. 8. Open in new tabDownload slide Variation diagram illustrating Fe(III)/Fetot vs SiO2 (wt %) for lizardite-serpentinites from dredge D10 and variably talc-altered antigorite-serpentinites from dredge D2. Talc-altered antigorite-serpentinites have lower Fe(III)/Fetot values than serpentinites that did not experience talc alteration. Additional data from Paulick et al. (2006). Dredge D10 Lizardite-serpentinites from dredge D10 have low concentrations of MgO (∼35–37 wt %) and CaO (<0·5 wt %), whereas SiO2 (∼39–40 wt %) and Al2O3 (∼0·5–1·4 wt %) concentrations are comparable with those of serpentinites from the MAR (ODP Leg 153) (Fig. 7). The MgO/SiO2 ratios of NWPRT serpentinites fall below the terrestrial melting array (Jagoutz et al., 1979) (Fig. 6). Concentrations of FeOtot (8–10 wt %) and TiO2 (∼0·01–0·03 wt % for most samples) are relatively high, and do not show any apparent correlation with MgO and SiO2. Because sample D10-9 hosts altered mafic veins that contain zircon, it has TiO2 and P2O5 concentrations that are significantly higher than those of other serpentinites from dredges D10 and serpentinites from other settings (Table 2). Fe(III)/Fetot ratios range from 0·85 to 0·99 (Fig. 8). In samples D10-8 and D10-10, the concentration of FeO was below the detection limit of the employed titration technique (0·1 wt %), suggesting that these samples are completely oxidized [Fe(III)/Fetot > 0·99]. Loss on ignition (LOI) values for dredge 10 lizardite-serpentinite range from 12 to 16 wt %. Dredge D2 Concentrations of SiO2 in samples from dredge D2 are variable (∼42–51 wt %, Fig. 7) and increase with increasing extent of talc alteration as MgO/SiO2 and Al2O3/SiO2 ratios decrease (Fig. 6). Concentrations of FeOtot and Fe(III)/Fetot ratios are significantly lower than those of samples from dredge D10 and decrease with increasing extent of talc alteration (Figs 7 and 8). Al2O3, K2O, and P2O5 concentrations in samples from dredge D2 are consistently lower than those of samples from dredge D10, whereas concentrations of MnO, Na2O, and CaO are indistinguishable in the two dredge hauls (Fig. 7; Table 2). Trace and rare earth element compositions of whole-rock samples Trace and rare earth element (REE) compositions are presented in Table 3 and plotted in Figs 9 and 10. Table 3: ICP-MS analyses of trace elements in whole-rocks (μg g-1) from dredges D2 and D10 Sample no.: . D10-1 . D10-3 . D10-4 . D10-7 . D10-8 . D10-9 . D10-10 . D10-12 . D10-15 . D10-16 . D2-1 . D2-2 . D2-3 . D2-4 . D2-5 . D2-6 . D10-4* . RSD % . Rock type: . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Lz-srp . . La 0·141 0·194 0·388 0·101 0·081 0·403 0·383 0·205 0·154 0·201 0·157 0·095 0·114 0·085 0·094 0·123 0·385 0·49 Ce 0·272 0·616 1·141 0·260 0·286 1·328 1·014 0·756 0·501 0·484 0·260 0·137 0·197 0·131 0·134 0·164 1·143 0·08 Pr 0·039 0·072 0·233 0·021 0·022 0·221 0·051 0·090 0·078 0·044 0·039 0·024 0·027 0·021 0·027 0·032 0·224 2·79 Nd 0·185 0·317 1·245 0·096 0·098 1·146 0·196 0·440 0·409 0·158 0·176 0·100 0·134 0·076 0·101 0·147 1·223 1·24 Sm 0·052 0·126 0·423 0·024 0·031 0·441 0·040 0·155 0·172 0·042 0·037 0·028 0·026 0·032 0·034 0·025 0·410 2·22 Eu 0·040 0·087 0·254 0·007 0·011 0·143 0·012 0·074 0·102 0·012 0·022 0·016 0·025 0·020 0·029 0·022 0·242 3·41 Gd 0·075 0·175 0·488 0·037 0·036 0·567 0·043 0·176 0·237 0·056 0·041 0·029 0·036 0·042 0·045 0·030 0·490 0·38 Tb 0·017 0·040 0·089 0·007 0·007 0·113 0·008 0·032 0·048 0·012 0·007 0·005 0·006 0·009 0·010 0·005 0·087 1·47 Dy 0·139 0·285 0·594 0·052 0·046 0·753 0·056 0·211 0·328 0·093 0·046 0·039 0·041 0·065 0·064 0·037 0·571 2·77 Ho 0·031 0·062 0·115 0·011 0·011 0·161 0·014 0·045 0·070 0·021 0·011 0·009 0·006 0·016 0·017 0·008 0·118 1·67 Er 0·107 0·187 0·330 0·037 0·034 0·467 0·042 0·137 0·196 0·067 0·034 0·024 0·020 0·049 0·051 0·025 0·332 0·37 Tm 0·019 0·029 0·053 0·006 0·006 0·068 0·007 0·021 0·030 0·013 0·005 0·004 0·003 0·009 0·009 0·004 0·052 2·01 Yb 0·145 0·202 0·356 0·045 0·045 0·486 0·055 0·134 0·201 0·089 0·029 0·026 0·020 0·073 0·071 0·025 0·351 1·04 Lu 0·025 0·037 0·062 0·007 0·009 0·076 0·010 0·023 0·034 0·015 0·004 0·005 0·002 0·012 0·015 0·003 0·057 5·04 Ba 1·75 1·74 2·42 1·55 1·25 1·17 7·40 2·13 0·97 0·78 0·78 0·32 0·17 <0·1 <0·1 <0·1 2·02 12·99 Th 0·04 0·07 0·04 0·05 0·03 0·07 0·13 0·07 0·03 0·04 0·05 0·02 0·03 0·03 0·03 0·02 0·03 21·51 Nb 0·055 0·265 0·563 0·151 0·041 1·182 0·261 0·386 0·266 0·086 0·061 0·027 0·028 0·049 0·035 0·042 0·531 4·17 Y 0·791 1·495 2·758 0·259 0·271 3·688 0·269 1·173 1·567 0·443 0·352 0·253 0·232 0·370 0·441 0·317 2·677 2·09 Hf 0·02 0·03 0·04 0·06 0·02 1·43 0·07 0·08 0·04 0·03 0·03 0·02 0·02 0·02 0·03 0·03 0·03 14·95 Ta 0·003 0·011 0·021 0·006 0·002 0·085 0·004 0·022 0·011 0·004 0·004 0·003 0·002 0·006 0·004 0·002 0·020 2·47 U 0·481 0·519 0·529 0·145 0·211 0·710 0·551 0·824 0·769 0·248 0·246 0·587 0·292 0·462 0·533 0·231 0·529 0·06 Pb 0·089 0·203 0·308 3·312 2·879 0·130 17·480 0·234 0·138 2·246 2·146 0·616 1·367 0·283 0·261 1·689 0·303 1·32 Rb 1·19 0·46 1·39 0·36 0·26 0·47 0·55 0·30 0·34 0·53 0·18 0·12 0·14 0·08 0·09 0·11 1·28 6·09 Cs 0·04 0·02 0·03 0·01 0·01 0·01 0·01 0·01 0·02 0·01 0·01 0·00 0·01 0·01 0·00 0·00 0·03 18·10 Sr 6·82 5·25 7·71 4·35 4·82 5·52 8·63 4·41 4·21 5·92 3·08 1·19 2·18 1·60 1·53 2·43 7·55 1·43 Sc 10·48 7·26 12·16 10·03 8·13 9·11 8·70 5·78 7·86 9·80 6·68 7·57 6·72 5·55 5·96 7·61 11·71 2·66 Zr 0·68 1·28 1·74 3·36 1·98 68·04 7·50 2·49 1·17 3·35 1·01 0·85 0·78 1·68 1·18 0·84 1·50 10·21 Tl 0·02 0·03 0·03 0·03 0·02 0·07 0·02 0·06 0·36 0·03 n.a. n.a. n.a. n.a. n.a. n.a. 0·03 14·8 Ni† 2221 2262 2049 1967 2087 2089 1606 2205 2169 2403 1569 1727 1483 1937 2065 1548 n.a. — Cr† 2517 1982 2320 2178 1647 2827 2924 1532 2118 2735 1383 2239 2358 943 1075 1572 n.a. — Sample no.: . D10-1 . D10-3 . D10-4 . D10-7 . D10-8 . D10-9 . D10-10 . D10-12 . D10-15 . D10-16 . D2-1 . D2-2 . D2-3 . D2-4 . D2-5 . D2-6 . D10-4* . RSD % . Rock type: . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Lz-srp . . La 0·141 0·194 0·388 0·101 0·081 0·403 0·383 0·205 0·154 0·201 0·157 0·095 0·114 0·085 0·094 0·123 0·385 0·49 Ce 0·272 0·616 1·141 0·260 0·286 1·328 1·014 0·756 0·501 0·484 0·260 0·137 0·197 0·131 0·134 0·164 1·143 0·08 Pr 0·039 0·072 0·233 0·021 0·022 0·221 0·051 0·090 0·078 0·044 0·039 0·024 0·027 0·021 0·027 0·032 0·224 2·79 Nd 0·185 0·317 1·245 0·096 0·098 1·146 0·196 0·440 0·409 0·158 0·176 0·100 0·134 0·076 0·101 0·147 1·223 1·24 Sm 0·052 0·126 0·423 0·024 0·031 0·441 0·040 0·155 0·172 0·042 0·037 0·028 0·026 0·032 0·034 0·025 0·410 2·22 Eu 0·040 0·087 0·254 0·007 0·011 0·143 0·012 0·074 0·102 0·012 0·022 0·016 0·025 0·020 0·029 0·022 0·242 3·41 Gd 0·075 0·175 0·488 0·037 0·036 0·567 0·043 0·176 0·237 0·056 0·041 0·029 0·036 0·042 0·045 0·030 0·490 0·38 Tb 0·017 0·040 0·089 0·007 0·007 0·113 0·008 0·032 0·048 0·012 0·007 0·005 0·006 0·009 0·010 0·005 0·087 1·47 Dy 0·139 0·285 0·594 0·052 0·046 0·753 0·056 0·211 0·328 0·093 0·046 0·039 0·041 0·065 0·064 0·037 0·571 2·77 Ho 0·031 0·062 0·115 0·011 0·011 0·161 0·014 0·045 0·070 0·021 0·011 0·009 0·006 0·016 0·017 0·008 0·118 1·67 Er 0·107 0·187 0·330 0·037 0·034 0·467 0·042 0·137 0·196 0·067 0·034 0·024 0·020 0·049 0·051 0·025 0·332 0·37 Tm 0·019 0·029 0·053 0·006 0·006 0·068 0·007 0·021 0·030 0·013 0·005 0·004 0·003 0·009 0·009 0·004 0·052 2·01 Yb 0·145 0·202 0·356 0·045 0·045 0·486 0·055 0·134 0·201 0·089 0·029 0·026 0·020 0·073 0·071 0·025 0·351 1·04 Lu 0·025 0·037 0·062 0·007 0·009 0·076 0·010 0·023 0·034 0·015 0·004 0·005 0·002 0·012 0·015 0·003 0·057 5·04 Ba 1·75 1·74 2·42 1·55 1·25 1·17 7·40 2·13 0·97 0·78 0·78 0·32 0·17 <0·1 <0·1 <0·1 2·02 12·99 Th 0·04 0·07 0·04 0·05 0·03 0·07 0·13 0·07 0·03 0·04 0·05 0·02 0·03 0·03 0·03 0·02 0·03 21·51 Nb 0·055 0·265 0·563 0·151 0·041 1·182 0·261 0·386 0·266 0·086 0·061 0·027 0·028 0·049 0·035 0·042 0·531 4·17 Y 0·791 1·495 2·758 0·259 0·271 3·688 0·269 1·173 1·567 0·443 0·352 0·253 0·232 0·370 0·441 0·317 2·677 2·09 Hf 0·02 0·03 0·04 0·06 0·02 1·43 0·07 0·08 0·04 0·03 0·03 0·02 0·02 0·02 0·03 0·03 0·03 14·95 Ta 0·003 0·011 0·021 0·006 0·002 0·085 0·004 0·022 0·011 0·004 0·004 0·003 0·002 0·006 0·004 0·002 0·020 2·47 U 0·481 0·519 0·529 0·145 0·211 0·710 0·551 0·824 0·769 0·248 0·246 0·587 0·292 0·462 0·533 0·231 0·529 0·06 Pb 0·089 0·203 0·308 3·312 2·879 0·130 17·480 0·234 0·138 2·246 2·146 0·616 1·367 0·283 0·261 1·689 0·303 1·32 Rb 1·19 0·46 1·39 0·36 0·26 0·47 0·55 0·30 0·34 0·53 0·18 0·12 0·14 0·08 0·09 0·11 1·28 6·09 Cs 0·04 0·02 0·03 0·01 0·01 0·01 0·01 0·01 0·02 0·01 0·01 0·00 0·01 0·01 0·00 0·00 0·03 18·10 Sr 6·82 5·25 7·71 4·35 4·82 5·52 8·63 4·41 4·21 5·92 3·08 1·19 2·18 1·60 1·53 2·43 7·55 1·43 Sc 10·48 7·26 12·16 10·03 8·13 9·11 8·70 5·78 7·86 9·80 6·68 7·57 6·72 5·55 5·96 7·61 11·71 2·66 Zr 0·68 1·28 1·74 3·36 1·98 68·04 7·50 2·49 1·17 3·35 1·01 0·85 0·78 1·68 1·18 0·84 1·50 10·21 Tl 0·02 0·03 0·03 0·03 0·02 0·07 0·02 0·06 0·36 0·03 n.a. n.a. n.a. n.a. n.a. n.a. 0·03 14·8 Ni† 2221 2262 2049 1967 2087 2089 1606 2205 2169 2403 1569 1727 1483 1937 2065 1548 n.a. — Cr† 2517 1982 2320 2178 1647 2827 2924 1532 2118 2735 1383 2239 2358 943 1075 1572 n.a. — * Repeat analysis. RSD is relative standard deviation for duplicate analyses of sample D10-4. † X-ray fluorescence analyses (XRF). Table 3: ICP-MS analyses of trace elements in whole-rocks (μg g-1) from dredges D2 and D10 Sample no.: . D10-1 . D10-3 . D10-4 . D10-7 . D10-8 . D10-9 . D10-10 . D10-12 . D10-15 . D10-16 . D2-1 . D2-2 . D2-3 . D2-4 . D2-5 . D2-6 . D10-4* . RSD % . Rock type: . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Lz-srp . . La 0·141 0·194 0·388 0·101 0·081 0·403 0·383 0·205 0·154 0·201 0·157 0·095 0·114 0·085 0·094 0·123 0·385 0·49 Ce 0·272 0·616 1·141 0·260 0·286 1·328 1·014 0·756 0·501 0·484 0·260 0·137 0·197 0·131 0·134 0·164 1·143 0·08 Pr 0·039 0·072 0·233 0·021 0·022 0·221 0·051 0·090 0·078 0·044 0·039 0·024 0·027 0·021 0·027 0·032 0·224 2·79 Nd 0·185 0·317 1·245 0·096 0·098 1·146 0·196 0·440 0·409 0·158 0·176 0·100 0·134 0·076 0·101 0·147 1·223 1·24 Sm 0·052 0·126 0·423 0·024 0·031 0·441 0·040 0·155 0·172 0·042 0·037 0·028 0·026 0·032 0·034 0·025 0·410 2·22 Eu 0·040 0·087 0·254 0·007 0·011 0·143 0·012 0·074 0·102 0·012 0·022 0·016 0·025 0·020 0·029 0·022 0·242 3·41 Gd 0·075 0·175 0·488 0·037 0·036 0·567 0·043 0·176 0·237 0·056 0·041 0·029 0·036 0·042 0·045 0·030 0·490 0·38 Tb 0·017 0·040 0·089 0·007 0·007 0·113 0·008 0·032 0·048 0·012 0·007 0·005 0·006 0·009 0·010 0·005 0·087 1·47 Dy 0·139 0·285 0·594 0·052 0·046 0·753 0·056 0·211 0·328 0·093 0·046 0·039 0·041 0·065 0·064 0·037 0·571 2·77 Ho 0·031 0·062 0·115 0·011 0·011 0·161 0·014 0·045 0·070 0·021 0·011 0·009 0·006 0·016 0·017 0·008 0·118 1·67 Er 0·107 0·187 0·330 0·037 0·034 0·467 0·042 0·137 0·196 0·067 0·034 0·024 0·020 0·049 0·051 0·025 0·332 0·37 Tm 0·019 0·029 0·053 0·006 0·006 0·068 0·007 0·021 0·030 0·013 0·005 0·004 0·003 0·009 0·009 0·004 0·052 2·01 Yb 0·145 0·202 0·356 0·045 0·045 0·486 0·055 0·134 0·201 0·089 0·029 0·026 0·020 0·073 0·071 0·025 0·351 1·04 Lu 0·025 0·037 0·062 0·007 0·009 0·076 0·010 0·023 0·034 0·015 0·004 0·005 0·002 0·012 0·015 0·003 0·057 5·04 Ba 1·75 1·74 2·42 1·55 1·25 1·17 7·40 2·13 0·97 0·78 0·78 0·32 0·17 <0·1 <0·1 <0·1 2·02 12·99 Th 0·04 0·07 0·04 0·05 0·03 0·07 0·13 0·07 0·03 0·04 0·05 0·02 0·03 0·03 0·03 0·02 0·03 21·51 Nb 0·055 0·265 0·563 0·151 0·041 1·182 0·261 0·386 0·266 0·086 0·061 0·027 0·028 0·049 0·035 0·042 0·531 4·17 Y 0·791 1·495 2·758 0·259 0·271 3·688 0·269 1·173 1·567 0·443 0·352 0·253 0·232 0·370 0·441 0·317 2·677 2·09 Hf 0·02 0·03 0·04 0·06 0·02 1·43 0·07 0·08 0·04 0·03 0·03 0·02 0·02 0·02 0·03 0·03 0·03 14·95 Ta 0·003 0·011 0·021 0·006 0·002 0·085 0·004 0·022 0·011 0·004 0·004 0·003 0·002 0·006 0·004 0·002 0·020 2·47 U 0·481 0·519 0·529 0·145 0·211 0·710 0·551 0·824 0·769 0·248 0·246 0·587 0·292 0·462 0·533 0·231 0·529 0·06 Pb 0·089 0·203 0·308 3·312 2·879 0·130 17·480 0·234 0·138 2·246 2·146 0·616 1·367 0·283 0·261 1·689 0·303 1·32 Rb 1·19 0·46 1·39 0·36 0·26 0·47 0·55 0·30 0·34 0·53 0·18 0·12 0·14 0·08 0·09 0·11 1·28 6·09 Cs 0·04 0·02 0·03 0·01 0·01 0·01 0·01 0·01 0·02 0·01 0·01 0·00 0·01 0·01 0·00 0·00 0·03 18·10 Sr 6·82 5·25 7·71 4·35 4·82 5·52 8·63 4·41 4·21 5·92 3·08 1·19 2·18 1·60 1·53 2·43 7·55 1·43 Sc 10·48 7·26 12·16 10·03 8·13 9·11 8·70 5·78 7·86 9·80 6·68 7·57 6·72 5·55 5·96 7·61 11·71 2·66 Zr 0·68 1·28 1·74 3·36 1·98 68·04 7·50 2·49 1·17 3·35 1·01 0·85 0·78 1·68 1·18 0·84 1·50 10·21 Tl 0·02 0·03 0·03 0·03 0·02 0·07 0·02 0·06 0·36 0·03 n.a. n.a. n.a. n.a. n.a. n.a. 0·03 14·8 Ni† 2221 2262 2049 1967 2087 2089 1606 2205 2169 2403 1569 1727 1483 1937 2065 1548 n.a. — Cr† 2517 1982 2320 2178 1647 2827 2924 1532 2118 2735 1383 2239 2358 943 1075 1572 n.a. — Sample no.: . D10-1 . D10-3 . D10-4 . D10-7 . D10-8 . D10-9 . D10-10 . D10-12 . D10-15 . D10-16 . D2-1 . D2-2 . D2-3 . D2-4 . D2-5 . D2-6 . D10-4* . RSD % . Rock type: . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Lz-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Atg-srp . Lz-srp . . La 0·141 0·194 0·388 0·101 0·081 0·403 0·383 0·205 0·154 0·201 0·157 0·095 0·114 0·085 0·094 0·123 0·385 0·49 Ce 0·272 0·616 1·141 0·260 0·286 1·328 1·014 0·756 0·501 0·484 0·260 0·137 0·197 0·131 0·134 0·164 1·143 0·08 Pr 0·039 0·072 0·233 0·021 0·022 0·221 0·051 0·090 0·078 0·044 0·039 0·024 0·027 0·021 0·027 0·032 0·224 2·79 Nd 0·185 0·317 1·245 0·096 0·098 1·146 0·196 0·440 0·409 0·158 0·176 0·100 0·134 0·076 0·101 0·147 1·223 1·24 Sm 0·052 0·126 0·423 0·024 0·031 0·441 0·040 0·155 0·172 0·042 0·037 0·028 0·026 0·032 0·034 0·025 0·410 2·22 Eu 0·040 0·087 0·254 0·007 0·011 0·143 0·012 0·074 0·102 0·012 0·022 0·016 0·025 0·020 0·029 0·022 0·242 3·41 Gd 0·075 0·175 0·488 0·037 0·036 0·567 0·043 0·176 0·237 0·056 0·041 0·029 0·036 0·042 0·045 0·030 0·490 0·38 Tb 0·017 0·040 0·089 0·007 0·007 0·113 0·008 0·032 0·048 0·012 0·007 0·005 0·006 0·009 0·010 0·005 0·087 1·47 Dy 0·139 0·285 0·594 0·052 0·046 0·753 0·056 0·211 0·328 0·093 0·046 0·039 0·041 0·065 0·064 0·037 0·571 2·77 Ho 0·031 0·062 0·115 0·011 0·011 0·161 0·014 0·045 0·070 0·021 0·011 0·009 0·006 0·016 0·017 0·008 0·118 1·67 Er 0·107 0·187 0·330 0·037 0·034 0·467 0·042 0·137 0·196 0·067 0·034 0·024 0·020 0·049 0·051 0·025 0·332 0·37 Tm 0·019 0·029 0·053 0·006 0·006 0·068 0·007 0·021 0·030 0·013 0·005 0·004 0·003 0·009 0·009 0·004 0·052 2·01 Yb 0·145 0·202 0·356 0·045 0·045 0·486 0·055 0·134 0·201 0·089 0·029 0·026 0·020 0·073 0·071 0·025 0·351 1·04 Lu 0·025 0·037 0·062 0·007 0·009 0·076 0·010 0·023 0·034 0·015 0·004 0·005 0·002 0·012 0·015 0·003 0·057 5·04 Ba 1·75 1·74 2·42 1·55 1·25 1·17 7·40 2·13 0·97 0·78 0·78 0·32 0·17 <0·1 <0·1 <0·1 2·02 12·99 Th 0·04 0·07 0·04 0·05 0·03 0·07 0·13 0·07 0·03 0·04 0·05 0·02 0·03 0·03 0·03 0·02 0·03 21·51 Nb 0·055 0·265 0·563 0·151 0·041 1·182 0·261 0·386 0·266 0·086 0·061 0·027 0·028 0·049 0·035 0·042 0·531 4·17 Y 0·791 1·495 2·758 0·259 0·271 3·688 0·269 1·173 1·567 0·443 0·352 0·253 0·232 0·370 0·441 0·317 2·677 2·09 Hf 0·02 0·03 0·04 0·06 0·02 1·43 0·07 0·08 0·04 0·03 0·03 0·02 0·02 0·02 0·03 0·03 0·03 14·95 Ta 0·003 0·011 0·021 0·006 0·002 0·085 0·004 0·022 0·011 0·004 0·004 0·003 0·002 0·006 0·004 0·002 0·020 2·47 U 0·481 0·519 0·529 0·145 0·211 0·710 0·551 0·824 0·769 0·248 0·246 0·587 0·292 0·462 0·533 0·231 0·529 0·06 Pb 0·089 0·203 0·308 3·312 2·879 0·130 17·480 0·234 0·138 2·246 2·146 0·616 1·367 0·283 0·261 1·689 0·303 1·32 Rb 1·19 0·46 1·39 0·36 0·26 0·47 0·55 0·30 0·34 0·53 0·18 0·12 0·14 0·08 0·09 0·11 1·28 6·09 Cs 0·04 0·02 0·03 0·01 0·01 0·01 0·01 0·01 0·02 0·01 0·01 0·00 0·01 0·01 0·00 0·00 0·03 18·10 Sr 6·82 5·25 7·71 4·35 4·82 5·52 8·63 4·41 4·21 5·92 3·08 1·19 2·18 1·60 1·53 2·43 7·55 1·43 Sc 10·48 7·26 12·16 10·03 8·13 9·11 8·70 5·78 7·86 9·80 6·68 7·57 6·72 5·55 5·96 7·61 11·71 2·66 Zr 0·68 1·28 1·74 3·36 1·98 68·04 7·50 2·49 1·17 3·35 1·01 0·85 0·78 1·68 1·18 0·84 1·50 10·21 Tl 0·02 0·03 0·03 0·03 0·02 0·07 0·02 0·06 0·36 0·03 n.a. n.a. n.a. n.a. n.a. n.a. 0·03 14·8 Ni† 2221 2262 2049 1967 2087 2089 1606 2205 2169 2403 1569 1727 1483 1937 2065 1548 n.a. — Cr† 2517 1982 2320 2178 1647 2827 2924 1532 2118 2735 1383 2239 2358 943 1075 1572 n.a. — * Repeat analysis. RSD is relative standard deviation for duplicate analyses of sample D10-4. † X-ray fluorescence analyses (XRF). Dredge D10 Samples from dredge D10 show a statistically significant (0·85 ≤ R2 ≤ 0·94) inverse linear correlation between REE (and Y) contents and Fe(III)/Fetot ratios, except for La and Ce (Fig. 9). Sample D10-9 is treated separately owing to the addition of REE in altered mafic veins. Negative slopes and intercepts ≥0·98 suggest near-complete removal of REE from completely oxidized samples (Fig. 9). The least oxidized samples have the highest REE concentrations and relatively flat or concave REE patterns (Fig. 10a). Moreover, the least oxidized samples have a positive Eu anomaly, except for sample D10-9, which has a slightly negative Eu anomaly [samples with Eu or Ce concentrations that are enriched (or depleted) relative to their neighboring REE are referred to as having positive (or negative) Eu or Ce anomalies]. More oxidized samples have U-shaped REE patterns and the most oxidized samples have a negative Eu anomaly and a positive Ce anomaly. Sample D10-9, which contains porous zircon, has the highest concentrations of Zr, Hf, Y, and Nb. Sample D10-10 is one of the most oxidized samples and has the highest concentrations of Ba, Th, Pb, and Sr. Fig. 9. Open in new tabDownload slide Variation diagrams illustrating the changes in REE content with increasing extent of oxidation [Fe(III)/Fetot] during weathering of serpentinite. An inverse linear relationship between REE (+Y) heavier than Ce and Fe(III)/Fetot is apparent. Loss of REE (>Ce) is probably due to complexing with hard ligands during open-system alteration of serpentinite by seawater. (See text for discussion) Fig. 9. Open in new tabDownload slide Variation diagrams illustrating the changes in REE content with increasing extent of oxidation [Fe(III)/Fetot] during weathering of serpentinite. An inverse linear relationship between REE (+Y) heavier than Ce and Fe(III)/Fetot is apparent. Loss of REE (>Ce) is probably due to complexing with hard ligands during open-system alteration of serpentinite by seawater. (See text for discussion) Fig. 10. Open in new tabDownload slide Chondrite-normalized REE patterns of samples from dredges D2 and D10. (a) REE patterns of lizardite serpentinite from dredge D10 vary with Fe(III)/Fetot. Increased Fe(III)/Fetot reflects increased interaction with seawater under open-system conditions. Least oxidized samples preserve concave-upward REE patterns with positive Eu anomalies inherited from refertilized peridotite that interacted with black-smoker type hydrothermal fluids. Positive Ce anomalies are probably due to increased resistance of Ce4+ to weathering. Concave-upward patterns become U-shaped patterns as samples become more oxidized. (b) REE patterns of antigorite-serpentinite from dredge D2 do not vary systematically with Fe(III)/Fetot. Patterns are similar to those of the less oxidized samples of dredge D10, but concentrations are about one order of magnitude lower. Fig. 10. Open in new tabDownload slide Chondrite-normalized REE patterns of samples from dredges D2 and D10. (a) REE patterns of lizardite serpentinite from dredge D10 vary with Fe(III)/Fetot. Increased Fe(III)/Fetot reflects increased interaction with seawater under open-system conditions. Least oxidized samples preserve concave-upward REE patterns with positive Eu anomalies inherited from refertilized peridotite that interacted with black-smoker type hydrothermal fluids. Positive Ce anomalies are probably due to increased resistance of Ce4+ to weathering. Concave-upward patterns become U-shaped patterns as samples become more oxidized. (b) REE patterns of antigorite-serpentinite from dredge D2 do not vary systematically with Fe(III)/Fetot. Patterns are similar to those of the less oxidized samples of dredge D10, but concentrations are about one order of magnitude lower. Dredge D2 There is no clear correlation between Fe(III)/Fetot and REE content in samples from dredge D2 (Fig. 10b). REE concentrations are comparable with those of the most oxidized samples from dredge D10. The less oxidized and more strongly talc-altered samples from dredge D2 (samples D2-1, D2-3, D2-6) exhibit REE patterns that are enriched in light REE (LREE) relative to heavy REE (HREE). All samples display a negative Ce anomaly and a positive Eu anomaly. Strongly weathered antigorite serpentinite samples (D2-4 and D2-5) have U-shaped REE patterns similar to the more oxidized samples from dredge D10. Physical properties of whole-rock samples Skeletal densities range from 2·66 to 2·79 g cm−3 whereas envelope densities range from 2·03 to 2·55 g cm−3 (Table 2). There is significant overlap between the densities of samples from dredges D2 and D10 despite significantly lower LOI values in some samples from dredge D2 (owing to the presence of talc). Porosities range from 4·0 vol. % in sample D10-15 to 14·6 vol. % in sample D2-2. Sample D10-16 has an apparent porosity of 23·3 vol. %, but this sample is characterized by open fractures, probably owing to relaxation after sampling. High porosities within the weathered mesh texture are also evident from backscattered electron images of this sample (Supplementary Data Fig. 4c). There is no apparent correlation between skeletal density and porosity. Major element oxide—density relationships are discussed in the context of other DSDP and ODP samples. Dredge D10 serpentinites have magnetic susceptibilities ranging from 0·28 to 2·1 SI 10−2 (Table 2, Fig. 11). With the exception of sample D2-2, magnetic susceptibilities of samples from dredge D2 are lower than those from dredge D10 (Fig. 11, Table 2). Fig. 11. Open in new tabDownload slide Density vs magnetic susceptibility of NWPRT samples. Additional data for serpentinites from the MAR (ODP Legs 153 and 209) and off-axis in c. 35 Ma basement (DSDP Leg 82) are shown for comparison [data from Klein et al. (2014)]. The oxidation of magnetite to hematite and goethite during weathering, as well as the dissolution of magnetite during talc alteration, causes a decrease in magnetic susceptibility. Fig. 11. Open in new tabDownload slide Density vs magnetic susceptibility of NWPRT samples. Additional data for serpentinites from the MAR (ODP Legs 153 and 209) and off-axis in c. 35 Ma basement (DSDP Leg 82) are shown for comparison [data from Klein et al. (2014)]. The oxidation of magnetite to hematite and goethite during weathering, as well as the dissolution of magnetite during talc alteration, causes a decrease in magnetic susceptibility. DISCUSSION Origin of NWPRT serpentinites The origin and geological history of NWPRT serpentinites have remained ambiguous since their recovery in 1961. The oldest sediments overlying basement rocks are believed to be of Late Cretaceous age (Bowin et al., 1966). Unfortunately, it is difficult to assess whether or not dredging captured the entire sedimentary stratigraphic column. A Cretaceous age for the serpentinite would be consistent with the ages of peridotites in Puerto Rico and Hispaniola (Bowin et al., 1966), but the paleogeographical and geodynamic relationships between NWPRT serpentinite in the North American Plate and serpentinized peridotite of the Caribbean Plate and the South Wall of the Puerto Rico Trench are not fully understood (Perfit et al., 1980). Bowin et al. (1966) suggested that the serpentinized peridotite may form a continuous layer of pre-Cenomanian age, whereas Nalwalk (1969) concluded that the serpentinites intruded Upper Cretaceous to Eocene sediments and are of post-Cenomanian age, because the serpentinites were found in a relatively small area and because sediments collected with some of the serpentinites are of Eocene age. Paleomagnetic age constraints place the time of accretion of the NWPRT basement in the Cretaceous Normal Superchron, from Turonian in the east to Aptian in the west (Fig. 1). The position where dredge D10 was recovered (20°00’00”N, 66°31’57”W) is in c. 112 ± 3 Ma basement according to the models of Müller et al. (2008, 2016). Paleomagnetic data do not provide conclusive evidence for a mid-ocean ridge origin of the NWPRT serpentinite, because the serpentinite could have been emplaced in Cretaceous oceanic crust at a later time; for example, during the Eocene as Nalwalk (1969) suggested. But an Early Cretaceous age of 112 ± 3 Ma is consistent with the U–Pb age of zircon in sample D10-9 (114·8 Ma), which we take as evidence that the peridotite protolith originates from the MAR. Moreover, high-temperature hydrothermal alteration of peridotite requires a heat source that was, in all likelihood, associated with the Cretaceous MAR. Taking into account this age, flow-lines of fracture zones, and changes in spreading rates, the reconstructed origin of the host-rock can be traced back to ∼13°N and 29°W at (or near) the Cretaceous Mid-Atlantic Ridge, c. 2200 km east of its current location (Fig. 1). Igneous processes Extensive to complete alteration of peridotite protoliths makes reconstruction of their igneous history at the Cretaceous MAR challenging. If peridotite does not undergo extensive melt—rock interactions there is a robust correlation between the extent of melting and the Cr# of spinel, the modal proportion of clinopyroxene, and the HREE content of clinopyroxene (Hellebrand et al., 2001; Warren, 2016). Serpentinites from dredge D10 contain unaltered spinel, whereas the spinel in samples from dredge D2 is altered to ferrite chromite and magnetite. In dredge D10, Cr#s of 36–46 suggest ∼14–16% melting, ∼0·4–0·7 μg g−1 Yb in clinopyroxene and a modal clinoyroxene content of ∼1–4% (Hellebrand et al., 2001; Warren, 2016). Thin section mosaics of samples from dredge D10 also suggest that the serpentinite precursor was a clinopyroxene-poor harzburgite. Given the approximate modal percentage of primary minerals (75–85 vol. % olivine, 10–20 vol. % orthopyroxene, 1–4 vol. % clinopyroxene, and 1 vol. % spinel), and a Yb concentration in clinopyroxene of ∼0·4–0·7 μg g−1, the bulk-rock Yb concentration is estimated to be ∼0·03–0·09 μg g−1 (Workman & Hart, 2005). Of all samples from dredge D10, only samples D10-7 and D10-8 have Yb concentrations that fall within this range. Samples with higher Yb concentrations could reflect lower degrees of melting, but that seems unlikely as these samples also have spinels with higher Cr#, which is commonly interpreted to reflect a higher degree of melting (Hellebrand et al., 2001; Warren, 2016). Symplectites of Cr-spinel and (now altered) pyroxene, altered mafic veins containing zircon (sample D10-9), and elevated concentrations of relatively incompatible and immobile elements (Zr, Th, and Nb) indicate refertilization of peridotite during melt—rock interaction (Paulick et al., 2006; Jöns et al., 2010). Although the least oxidized lizardite-serpentinites indeed have REE patterns that resemble the REE patterns of refertilized abyssal harzburgite (Deschamps et al., 2013), distinct enrichments in fluid-mobile LREE can also be a result of hydrothermal alteration (Paulick et al., 2006; Jöns et al., 2010). Hydrothermal alteration history of NWPRT serpentinites On the basis of petrographic observations, several alteration processes can be distinguished. Localized chlorite—serpentine—tremolite—talc alteration of the peridotite host and blackwall alteration of mafic veins represents the onset of high-temperature hydrothermal processes (sample D10-9). Veins of mafic origin, such as the ones in sample D10-9, now consist almost exclusively of chlorite and antigorite (Fig. 2) and accessory zircon. Zircon has a porous texture and does not show any internal zoning characteristic of igneous zircon when examined with cathodoluminescence, but instead the internal zoning is rather cloudy or nebulous (Fig. 2c). Hydrothermal zircon of similar appearance has been reported to form at temperatures as high as 600°C in magmatic—hydrothermal systems to temperatures as low as 270°C in rodingitized gabbro associated with serpentinized peridotite (Dubinska et al., 2004; Schaltegger, 2007). On the basis of these observations we conclude that zircon in sample D10-9 is most probably a hydrothermal alteration product of igneous zircon (see Tomascheck et al., 2003). Most peridotites from dredge D10 underwent pervasive serpentinization under static conditions without major tectonic shearing before, during, or after serpentinization took place. Olivine was altered to serpentine and magnetite in pseudomorphic mesh textures, whereas orthopyroxene was altered to serpentine ± chlorite in bastite textures (Fig. 3). Density–magnetic susceptibility relationships of NWPRT serpentinite overlap with those from other MAR sites (Fig. 11); however, magnetic susceptibilities are lower than those of most samples from ODP Leg 153. Magnetic susceptibilities of NWPRT serpentinites were probably higher before weathering by seawater commenced, causing alteration of magnetite to hematite and goethite (see below). Skeletal densities of completely serpentinized brucite-free NWPRT serpentinites are higher than those of completely serpentinized brucite- or iowaite-bearing peridotite from the other sites. Seawater is undersaturated with respect to brucite (Snow & Dick, 1995) and the mesh texture in samples from dredge D10 is unusually porous for serpentinite (Fig. 3). Hence, it is likely that brucite was once present in these samples, but later dissolved during open-system weathering by seawater, causing an increase in porosity and skeletal density. During serpentinization, iron originally contained in olivine is chiefly distributed among serpentine, magnetite, and brucite, in addition to traces of Ni—Fe sulfides and alloys (Klein & Bach, 2009; Klein et al., 2009). The Fe content of lizardite in mesh texture is lower than that of its olivine precursor, which allows excess iron to partition into magnetite and brucite. More Fe is partitioned into brucite and less magnetite is generated during serpentinization with decreasing temperature, because of an inverse, temperature-dependent relationship between the Fe content of brucite and the abundance of magnetite (McCollom & Bach, 2009; Klein et al., 2013, 2014; McCollom et al., 2016). Therefore, the abundance of pseudomorphic hematite after magnetite throughout the mesh texture in samples from dredge D10 suggests that brucite was probably Fe-poor or only slightly enriched in Fe. Together with the lack of talc and tremolite in bastite texture after pyroxene in most samples from dredge D10, this indicates that initial alteration of mafic veins was followed by pervasive serpentinization that took place at temperatures lower than 380°C but higher than 200°C (e.g. Klein et al., 2013). To constrain further the temperature of serpentinization we can apply the oxygen isotope thermometer of Saccocia et al. (2009) to oxygen isotope data for the NWPRT samples D2-2 and D10-2 from the literature (Wenner & Taylor, 1973). Assuming the oxygen isotopic composition of the serpentinizing fluid was close to that of modern day seawater, the δ18O data for samples D2-2 and D10-2 yield temperatures of 309 ± 3 °C and 243 ± 3 °C, respectively (Wenner & Taylor, 1973; Saccocia et al., 2009), which seems to be consistent with the presence of antigorite in the former sample and lizardite in the latter. However, the δ18O of seawater has varied significantly in the geological past (Veizer et al., 1999) and the δ18O of hydrothermal fluid can evolve during fluid—rock interaction (Alt & Bach, 2006). Lower temperatures are calculated if the serpentinizing fluid was more depleted in 18O and vice versa. It is conceivable that the oxygen isotopic signature of serpentinites was modified during post-serpentinization weathering by seawater and it is not clear whether the nature of distinct serpentine minerals has an effect on temperature estimates. However, temperatures of ∼240–310 °C are reasonable for serpentinization in a mid-ocean ridge setting and consistent with phase petrological constraints. Talc alteration is common in mid-ocean ridge serpentinite and is due to fluid-mediated diffusive or advective Si mass transfer from mafic intrusions (Boschi et al., 2006; Bach et al., 2013; Ogasawara et al., 2013). Talc-altered serpentinites from dredge D2 have lower Fe contents and Fe(III)/Fetot (Table 2; Figs 7 and 8) than serpentinites that are not affected by talc alteration. Indeed, talc has a higher Mg# than its serpentine precursor (Supplementary Data Table 1) and the significantly lower Fe content of talc cannot be solely explained by dilution owing to the addition of Si (or Mg). Moreover, the domains affected by talc alteration are free of magnetite and other opaque minerals (Supplementary Data Fig. 2). Therefore, it appears that some of the Fe originally present in serpentinite was removed during talc alteration. Mass transfer and redox reactions during weathering of serpentinites by seawater Weathering has multifaceted effects on the mineralogy and chemical composition of serpentinite. Changes in mineralogy include the formation of hematite and goethite at the expense of magnetite, and the replacement of serpentine by poorly crystalline clay minerals, quartz, and carbonate in boxwork textures (Milliken et al., 1996; Bucher et al., 2015). Clay minerals, quartz, and carbonate are only locally abundant in the NWPRT serpentinites. Snow & Dick (1995) attributed pervasive loss of Mg from serpentinized peridotite to the dissolution of brucite and incongruent dissolution of olivine and orthopyroxene, none of which are saturated in seawater at low temperatures. Brucite can be altered either to iowaite (D’Antonio & Kristensen, 2004; Klein et al., 2014) via the oxidation of iron and addition of Cl, or to hydrous magnesium carbonates (Mumpton & Thompson, 1966; Power et al., 2009), depending on the ambient aCO2(aq). Serpentinites from the NWPRT contain neither brucite nor iowaite or hydrous magnesium carbonates, and their compositions fall below the terrestrial MgO/SiO2 vs Al2O3/SiO2 array of peridotites (Fig. 6), consistent with the removal of Mg or the addition of Si to the peridotite protolith, or both. Brucite usually forms in mesh texture during serpentinization in olivine-rich domains in peridotite and dunite. Mesh texture in NWPRT serpentinite is unusually porous (Fig. 3) compared with unweathered serpentinite from the MAR, pointing to mineral dissolution and tentatively suggesting that brucite was once present in the mesh texture. The high porosity is also reflected in the very low envelope densities of 2·03 g cm−3 (see Bowin et al., 1966) that are significantly lower than those of unweathered serpentinite (Table 1). The widespread formation of hematite in NWPRT serpentinites from dredges D2 and D10 is tied to the oxidation of magnetite: 4 Fe3O4+ O2(aq)= 6 Fe2O3.(1) In turn, goethite formation at the expense of hematite is evident in most samples (Figs 2 and 3): Fe2O3+ H2O=2 FeO(OH).(2) The abundance of hematite and goethite in weathered lizardite-serpentinites is also reflected in their high Fe contents, which suggests that Fe is largely immobile during weathering. Assuming that no Fe was added during weathering, this indicates that the peridotite protoliths of NWPRT serpentinites at the Cretaceous MAR contained ≥8 wt % FeO, equivalent to ∼1·1 mol FeO per kg of rock. Considering the generalized reaction 2FeO+ H2O=Fe2O3+ H2(aq).(3) and almost complete oxidation of several strongly weathered NWPRT serpentinites, up to 0·55 mol H2(aq) can be generated per kg of protolith. Because about two-thirds of the iron in completely serpentinized (but unweathered) peridotite is ferric (Klein et al., 2009, 2014; Andreani et al., 2013), serpentinization may have generated as much as 0·37 mol H2(aq) per kg of protolith. Depending on the level of oxygenation of bottom seawater in the Atlantic, oxygen was consumed [reaction (1)] or hydrogen was generated [reaction (3)]. If seawater was oxygen depleted up to 0·18 mol H2(aq) could have been generated during weathering. However, the timescales of H2(aq) generation during weathering of serpentinite are poorly understood and it is unclear whether the H2(aq) generation rate is sufficient to support microbial ecosystems with chemical energy in off-axis environments. Several NWPRT serpentinites are almost completely oxidized (samples D10-1, D10-7, D10-8, D10-10, D10-16), which requires that in addition to ferric iron in hematite and goethite, virtually all iron in serpentine is ferric. However, the talc-altered antigorite-serpentinites from dredge D2 have lower Fe(III)/Fetot than lizardite-serpentinites from dredge D10, which suggests that talc-altered rocks may be more resistant to oxidation (Fig. 8). Electron microprobe data for lizardite from dredge D10 indicate a slight cation deficiency of Mg and Fe in the octahedral sheet, whereas Si in the tetrahedral sheet is stoichiometric (Fig. 4). This result is probably due to preferential leaching of Mg relative to Fe and Si, a behavior that has been documented previously in chrysotile and other phyllosilicates (Bales & Morgan, 1985; Nagy, 1995). Lizardite in mesh texture is enriched in Al2O3 relative to its olivine precursor (SD Fig. 4, SD Table 1), suggesting that Al2O3 originally present in pyroxene is mobile on a local scale. However, bulk-rock Al2O3 contents remained largely unchanged during serpentinization and subsequent weathering, as indicated by the lack of correlation of Al2O3 contents with density (Fig. 7). Most of the CaO was lost during serpentinization as most of the secondary minerals in serpentinite discriminate against it. Sodium, K, and Rb appear to be enriched in NWPRT samples compared with unweathered serpentinites from the present-day MAR (see Paulick et al., 2006). In addition to possible uptake during weathering, their enrichment may simply reflect an artifact owing to evaporation of seawater in pore spaces during sample storage. Leaching experiments would solve this question but limited sample supply precluded these measurements. REE usually occur in the trivalent oxidation state, the major exceptions being the redox-sensitive Ce and Eu, which can also occur as Ce4+ and Eu2+, respectively. This behavior can lead to enrichments or depletions of these elements relative to the other REE (Ludden & Thompson, 1979). Although less common, some of the other REE can also be divalent (Sm, Tm, Yb) or tetravalent (Pr, Tb, Dy) (e.g. Beck, 1978; Hu et al., 1997). With the exception of La and Ce, weathered lizardite-serpentinites from the NWPRT (dredge 10) display a strong inverse linear relationship between REE concentration and Fe(III)/Fetot (Fig. 9). However, the increase in Fe(III)/Fetot is unlikely to be the cause for the removal of REE. The increase of Fe(III)/Fetot is due to the oxidation of ferrous iron by seawater, and Fe oxidation therefore monitors the progress of seawater alteration. The correlation of REE depletion with Fe oxidation thus suggests that REE were removed from the serpentinite as a result of complexing with hard ligands such as CO32− ⁠, OH−, F−, SO42− ⁠, and PO43− (Wood, 1990) during prolonged interaction with seawater under open-system conditions. The negative Ce anomaly of seawater (Elderfield & Greaves, 1982) has been explained by oxidation of Ce3+ to Ce4+ and subsequent fractionation from seawater during the formation of ferromanganese deposits through direct adsorption or scavenging by colloidal iron oxide (Glasby et al., 1978). The oxidation state of Ce has not been determined for NWPRT serpentinites, but fractionation of Ce from seawater via oxidation of trivalent to tetravalent Ce may explain its positive anomaly in the NWPRT serpentinite. Alternatively, tetravalent Ce is not enriched, but less depleted than other REE owing to the increased stability of tetravalent Ce on mineral surfaces relative to complexation with hard ligands in seawater. Like Ce, La also shows no strong correlation with Fe(III)/Fetot. The reason for this behavior is unclear, although analytical interferences during ICP-MS measurements may have caused some of the scatter. The U-shaped REE patterns of the most oxidized NWPRT serpentinites from dredge D10 (Fig. 10; samples D10-7, D10-8, D10-10 and D10-16) may be inherited from melt refertilization or fluid-dominated serpentinization (Niu, 2004; Paulick et al., 2006; Deschamps et al., 2013). However, our data clearly indicate that weathering overprinted pre-existing REE patterns formed by igneous and hydrothermal processes. The resulting effects are particularly apparent in weathered serpentinites from dredge D10, whereas talc-altered antigorite serpentinites from dredge D2 appear to have been more weathering resistant. The linearly decreasing concentrations of middle REE (MREE) and HREE with increasing Fe(III)/Fetot suggests a more or less uniform depletion during weathering with limited fractionation between the MREE and the HREE. In contrast to the MREE and HREE, depletion of La and Ce is less apparent. Antigorite in mid-ocean ridge serpentinite Lizardite appears to be the prevalent serpentine mineral in serpentinized peridotite and other olivine-rich rocks formed at mid-ocean ridges. Whereas chrysotile is common, though less abundant than lizardite, few studies report antigorite in mid-ocean ridge serpentinites. Aumento & Loubat (1971) detected antigorite as a minor phase by X-ray diffraction in three dredged serpentinites from the MAR at 45°N. Grobéty (2003) described intimate intergrowths of antigorite with lizardite and chrysotile in serpentinite from Hess Deep, a tectonic window into fast-spreading crust in the Pacific where serpentinization is estimated to have taken place at 400 °C ± 50 °C (Früh-Green et al., 1996). Antigorite forms together with Fe-rich brucite at the expense of olivine in troctolite from the Atlantis Massif, an oceanic core complex near the MAR (Beard et al., 2009). Antigorite at the Atlantis Massif is thought to have formed at temperatures higher than 300°C, possibly as high as 500°C, although serpentine phase relations cannot be considered a reliable temperature indicator. In fact, sample D10-9 contains lizardite, chrysotile, and antigorite, clearly indicative of disequilibrium on a centimeter scale rather than steep temperature gradients. Ribeiro da Costa et al. (2008) analyzed antigorite in fault zones from the Rainbow and Menez Hom hydrothermal areas at the MAR in deformed, non-pseudomorphic serpentinites. Oxygen-isotope thermometry of these samples yields temperatures lower than 300°C, which those researchers explained by dissolution of chrysotile and recrystallization to antigorite under dynamic conditions. Antigorite is the dominant serpentine mineral in samples from dredge D2 where it may have formed at the expense of chrysotile (or lizardite), possibly together with brucite or via addition of Si: 7 Chr=Atg+3 Brc(4) 16Chr+2 SiO2(aq)=Atg.(5) Although none of the serpentinites from dredge D2 contain brucite, its former presence cannot be completely ruled out, as it would have dissolved during weathering or reacted to serpentine during subsequent talc alteration. Because three of six antigorite-serpentinites examined in this study are strongly talc altered (samples D2-1, D2-3, D2-6), the addition of Si to form antigorite is conceivable. However, petrographic constraints cannot unequivocally resolve whether antigorite formation was linked to Si-metasomatism, owing to prograde metamorphism, or if antigorite formed directly at the expense of primary minerals. In sample D10-9, lizardite in pseudomorphic mesh and bastite textures is cut by veins that are chiefly composed of antigorite and chlorite (Fig. 2). These veins are cut by a later generation of chrysotile veins. Pervasive antigorite formation, as in samples from dredge D2, is not recorded in sample D10-9, possibly suggesting that two distinct modes of antigorite formation were at work in samples from dredges D2 and D10. Geophysical studies suggest that serpentinization can affect the oceanic mantle to several kilometers below the seafloor, which has led some to revisit Hess’ original hypothesis (Hess, 1962) that the seismic Moho at slow- and ultraslow-spreading mid-ocean ridges may be a serpentinization front (Minshull et al., 1998; Schlindwein & Schmid, 2016). The deepest drill hole into oceanic serpentinite (ODP Leg 153, Hole 920D) penetrated only 200·8 m below the seafloor (Shipboard Scientific Party, 1995), so seafloor serpentinites sampled to date probably do not represent the full range of serpentinite expected to be present deeper in the basement. Our samples give us reason to speculate that antigorite may be more common or even the dominant serpentine mineral at greater depths. Petrological phase relations certainly support the idea that antigorite is the stable phase at high temperatures and pressures (Evans, 2004), and tectonic windows such as the NWPRT and Hess Deep, oceanic core complexes such as the Atlantis Massif, and ophiolites that are unaffected by any metamorphic overprint during obduction may represent suitable natural laboratories to access these deeper levels of the serpentinized oceanic lithosphere. CONCLUSIONS On the basis of paleomagnetic constraints (Müller et al., 2008, 2016) and the zircon age presented in this study, we conclude that serpentinized peridotite exposed on the NWPRT originates from the Mid-Atlantic Ridge. Alternative models for the origin of NWPRT serpentinite, such as a late-stage intrusion into Upper Cretaceous–Eocene sediments as suggested by Nalwalk (1969), are inconsistent with the age constraints. Mantle melting, melt—rock interaction, and high-temperature hydrothermal alteration including serpentinization and, notably, the formation of antigorite, as well as talc alteration of antigorite-serpentinite, took place during the Early Cretaceous, most probably at or near the MAR. Continued serpentinization farther off-axis cannot be ruled out, but petrological phase relations and oxygen isotope thermometry indicate that serpentinization took place at elevated temperatures characteristic of hydrothermal systems found at or near the MAR, such as Rainbow, Logatchev, and Lost City (Kelley et al., 2001; Allen & Seyfried, 2004). Antigorite is found in tectonic windows that expose exhumed basement such as the Hess Deep (Früh-Green et al., 1996; Grobéty, 2003) and the Atlantis Massif (Beard et al., 2009), in fault zones from the Rainbow and Menez Hom areas on the MAR (Ribeiro da Costa et al., 2008), in the Guatemala forearc (Kodolanyi & Pettke, 2011), and in the Mariana forearc (Murata et al., 2009) where serpentinite mud from the décollement ‘erupts’ to form mud ‘volcanoes’ (Fryer et al., 1985; Fryer, 2012). The presence of antigorite in samples from dredge D2 and sample D10-9 reinforces the question of how common and widespread antigorite is in slow-spreading oceanic basement and magma-poor continental margins. Conceivably, antigorite may be found at pressure—temperature conditions exceeding those that have been accessed by seafloor drilling in the Atlantic and it may thus be more widespread then previous studies suggest. Serpentinites at the present-day MAR show similar mineral assemblages and whole-rock compositions and signs of weathering, but are typically not completely oxidized, suggesting that weathering continues off-axis, possibly to or beyond the estimated sealing age of 65 Ma (Stein & Stein, 1994; Stein et al., 1995). Weathering modifies the mineral contents of serpentinites with potentially important consequences for the geophysical properties (specific gravity, porosity, magnetic susceptibility, etc.) of the uppermost basement entering the Puerto Rico Trench. Chemical changes during fluid-dominated serpentinization (Paulick et al., 2006) and weathering of serpentinites overprint patterns typically associated with asthenospheric melting and melt—rock interaction. It can thus be suggested that weathering effects must be considered when REE patterns or trace element ratios of bulk-rock seafloor serpentinites are used to infer possible magmatic processes. Subduction of water-rich oxidized lithologies would enhance melt production and affect the redox budget of the magmas produced in subduction zones. The obliquity of subduction prevents melting beneath the Greater Antillean arc today. Yet, if serpentinite is also a significant component of the subducting slab beneath the Lesser Antillean arc south of the Puerto Rico Trench, the effects of serpentinite subduction should be more obvious as subduction is not oblique there and volcanism is very active. The remainder of the NWPRT serpentinite is entrained in the sinistral strike-slip system that separates the North American and Caribbean plates. The low shear strength of serpentinite, and in particular talc-altered serpentinite (Reinen et al., 1994; Moore & Rymer, 2007) may affect the seismicity of this system. The formation of magnetite during serpentinization at the Cretaceous MAR generated H2, which, by analogy to present-day hydrothermal systems at the MAR, can be oxidized by microorganisms to gain metabolic energy (Amend et al., 2011; Brazelton et al., 2012). Using thermodynamic and kinetic constraints, Bach (2016) suggested that weathering of olivine to hematite and goethite can support microbial ecosystems with chemical energy in ridge flank environments. Similarly, oxidation of already serpentinized peridotite may generate additional H2 in low-temperature off-axis environments if the oxygenation levels of bottom seawater are low. This process may proceed until the rock is completely oxidized. However, it is unclear whether the rates of H2(aq) generation during weathering of serpentinite would actually suffice to support microbial organisms with chemical energy. The Puerto Rico Trench is a fascinating and conveniently located hadal environment offering numerous opportunities to gain insight into igneous, geodynamic, hydrothermal, and probably microbiological, processes in the Atlantic. The present paper was written on the basis of a limited number and amount of samples that were collected more than 50 years ago. We hope the results and interpretations presented here encourage future expeditions to the Puerto Rico Trench. ACKNOWLEDGEMENTS We are grateful for discussions with Colin Devey and Chris German during the early stages of this study. Carl Bowin is thanked for providing access to samples from cruise 19 of the R.V. Chain. We thank James Eckert for his assistance with electron microprobe analysis at Yale University, and Karmina Aquino for porosity measurements of samples from dredge D10. We are grateful to IODP and the curators of the core repositories in Bremen, Kochi, and College Station for providing access to drill core samples. This paper benefited from comments and suggestions by Othmar Müntener, Eric Hellebrand and an anonymous reviewer. Editorial handling by James Beard is much appreciated. FUNDING F.K. and H.R.M. acknowledge financial support from the Ocean Exploration Institute at the Woods Hole Oceanographic Institution. This research also benefited from financial support by the US National Science Foundation grant OCE-1059534 (to F.K. and S.E.H.). SUPPLEMENTARY DATA Supplementary data for this paper are available at Journal of Petrology online. REFERENCES Alexander E. B. , Coleman R. G., Keeler-Wolf T., Harrison S. P. ( 2007 ). 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TI - Mid-ocean Ridge Serpentinite in the Puerto Rico Trench: from Seafloor Spreading to Subduction
JF - Journal of Petrology
DO - 10.1093/petrology/egx071
DA - 2017-09-01
UR - https://www.deepdyve.com/lp/oxford-university-press/mid-ocean-ridge-serpentinite-in-the-puerto-rico-trench-from-seafloor-pempPUtlwC
SP - 1729
VL - 58
IS - 9
DP - DeepDyve
ER -