Chapter 5: Formation of replacive olivine gabbros in the Oman Moho Transition Zone: Geochemical evolution during

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Chapter 5: Formation of replacive olivine gabbros in the Oman Moho Transition Zone: Geochemical evolution during

deformation-driven dunite impregnation

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Chapter 5: Formation of replacive olivine gabbros in the Oman Moho Transition Zone

5.1 Introduction

Melt percolation within the lithospheric mantle and crust has a strong influence on the dynamics of the extensional plate boundaries. At the Mid-Atlantic slow-spreading ridge, seismic and gravimetric data indicates the presence of up to 5% melt volume in the uppermost mantle (4-10 kilometres; Dunn et al., 2005). At the East Pacific Rise fast-spreading environment, tomography models, together with petrological studies of ophiolites (Boudier et al., 1996) and active ridges (Lissenberg et al., 2013), indicate that melts accumulate both in an upper crustal magma lens and in the uppermost mantle, at the crust-mantle transition (Dunn et al., 2000). Petrological and geochemical studies of ophiolites and active ridges also highlighted the extensive role of melt percolation in the structural and geochemical evolution and thinning of the subcontinental uprising lithospheric mantle (Le Roux et al., 2007; Soustelle et al., 2009). Petrological studies (Kelemen &

Dick, 1995; Dijkstra et al., 2002; Kaczmarek & Müntener, 2008; Le Roux et al., 2008; Soustelle et al., 2009; Kaczmarek & Tommasi, 2011; Higgie & Tommasi, 2012, 2014) and geophysical data demonstrated the effects of the presence of melt on the strength and viscosity of a percolated mantle rock (up to one order of magnitude for melt fractions as low as 0.1%; Kohlstedt & Zimmerman, 1996; Rosenberg & Handy, 2005; Takei & Holtzman, 2009), leading to a strain weakening as a function of melt fraction. This weakening leads to an organisation of the melt distribution in pockets oriented at low angle to the shear plane at grain-scale size, as demonstrated in laboratory melt- bearing deformation experiments (Kohlstedt & Zimmerman, 1996). A positive feedback between the presence of melt and localisation of strain leads to the development of a layered structure, between melt-rich and melt-poor bands (Holtzman et al., 2003b, 2005; Holtzman & Kohlstedt, 2007; Higgie & Tommasi, 2012, 2014), allowing for faster melt extraction in high-porosity melt- bearing layers.

Recent petrological and geochemical studies demonstrated that reactive melt percolation in the lithospheric mantle and lower oceanic crust leads to strong geochemical variations in the percolated rock (Kelemen et al., 1995b; Bédard et al., 2000; Dick et al., 2002; Piccardo et al., 2007a; Rampone et al., 2016; Basch et al., in revision; Ferrando et al., in revision) and reactive melt (Borghini & Rampone, 2007; Borghini et al., 2007; Lissenberg & Dick, 2008; Rampone et al., 2008; Collier & Kelemen, 2010; Lissenberg et al., 2013, 2017; Nicolle et al., 2016; Paquet et al., 2016; Sanfilippo et al., 2016). Moreover, the replacive formation of hybrid rocks from the interaction between a pre-existing matrix and a percolating tholeiitic melt has been increasingly invoked in the formation of the lowermost oceanic crust (Lissenberg & Dick, 2008; Suhr et al., 2008; Drouin et al., 2009, 2010; Renna & Tribuzio, 2011; Sanfilippo & Tribuzio, 2013; Higgie &

Tommasi, 2012; Sanfilippo et al., 2013, 2014, 2015, 2016; Rampone et al., 2016; Basch et al., in revision).

The Oman ophiolites display the classic layered oceanic crust sequence (Browning, 1982;

Boudier & Nicolas, 1985; Boudier et al., 1996), indicating long-lived magma chambers in a former fast-spreading ridge environment (Nicolas & Boudier, 1995; Boudier et al., 1997). High-resolution dating indicates half-spreading rates of 10 centimetres/year (Rioux et al., 2012, 2013), similar to present-day East Pacific Rise (Toomey et al., 2007; Müller et al., 2008). At the base of the crustal sequence, a hundred-metre thick Moho Transition Zone composed of a centimetre- to metre-thick interlayered dunites and variably evolved olivine gabbros (Benn et al., 1988; Boudier & Nicolas, 1995; Kelemen et al., 1997b; Korenaga & Kelemen, 1997; Higgie & Tommasi, 2012; Jousselin et

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5.2 Structural and petrologic background

harzburgites. The Moho Transition Zone associates high shear strains and melt fractions as the upwelling mantle is forced away from the ridge axis in a corner flow (Rabinowicz et al., 1987;

Ceuleneer et al., 1988; Jousselin & Mainprice, 1998; Jousselin et al., 2012), it is therefore an ideal case study to investigate the structural and geochemical evolution of the pre-existing matrix and percolating melt in a melt-focusing deforming system (Higgie & Tommasi, 2012).

Our geochemical study follows the structural investigation of the Moho Transition Zone by Higgie & Tommasi (2012). We have used the same set of samples, from dunite to olivine gabbros, to perform mineral major and trace elements analyses and investigate the geochemical variations induced by melt-rock interaction processes during the replacive formation of the olivine gabbros.

5.2 Structural and petrologic background

The Oman ophiolite, about 500 kilometres long and 50 to 100 kilometres wide, was obducted on the Oman margin during the closure of the Neo-Tethys ocean 95 million years ago (Coleman, 1981; Hacker, 1994). The continuous gabbroic sequence and classic layered oceanic crust sequence suggests a formation of the ophiolite at a fast spreading centre (Nicolas & Boudier, 1995; Rioux et al., 2012, 2013). The tectonic context of the Oman ophiolite remains controversial, between a formation at Mid-Ocean Ridge (Boudier et al., 1988; Hacker, 1994; Godard et al., 2003, 2006; Nicolas & Boudier, 2003) and a supra-subduction zone (Pearce & Cann, 1971; Pearce et al., 1981; Shervais, 2001). This study focuses on the Moho Transition Zone (MTZ) in the Maqsad massif area, about 5 kilometres away from the inferred ridge axis (Fig. 5.1a). Its thickness ranges from a few meters to hundreds of metres, its upper limit being the base of the continuous layered gabbroic section, and its base the transition from dunites to mantle harzburgites (Benn et al., 1988;

Boudier & Nicolas, 1995; Higgie & Tommasi, 2012). We have studied an 80m-long continuous section of the MTZ in the Maqsad massif (Fig. 5.1b,c), formed by interlayered dunites, wehrlites, troctolites and variably evolved olivine gabbros occurring as lenses from millimetre- to tens of metres in size (Boudier & Nicolas, 1995; Kelemen et al., 1997; Korenaga & Kelemen, 1997; Higgie

& Tommasi, 2012). The MTZ limits and internal structures (layering, foliation, lineation) are roughly parallel to the regional Moho (Fig. 5.1c; Benn et al., 1988; Jousselin et al., 2012). The internal structures were described as the result of reactive percolation and accumulation of basaltic melts in the uppermost part of the mantle beneath an active ridge (Rabinowicz et al., 1987; Boudier

& Nicolas, 1995; Korenaga & Kelemen, 1997).

I will start by reviewing the petrological and structural results of Higgie & Tommasi (2012), before describing the mineral major and trace elements variability in the MTZ interlayered dunites, wehrlites, troctolites and olivine gabbros (Fig. 5.2a,b).

The observed modal composition layering (Fig. 5.2a,b) is parallel to a pervasive subhorizontal foliation, containing a lineation trending N110°E. The foliation is observed within altered dunitic layers (Fig. 5.2b), and the lineation is marked by a shape-preferred orientation of plagioclase and clinopyroxene in the gabbroic levels (Fig. 5.2c). The observed layering is discontinuous, with lens-shaped layers that cannot be followed laterally over more than a few metres and showing diffuse to sharp termination (Fig. 5.2d). At the scale of the 80-metre studied section (Fig. 5.1b), no systematic vertical variation of the modal composition, grain size or layer thickness was observed.

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Figure 5.1: A: Geological map of the MTZ in the Maqsad massif (based on Nicolas et al., 2000; Jousselin et al., 2012; redrawn after Higgie & Tommasi, 2012) showing the location of the studied area; B: Photograph of the studied section indicating the sampled outcrop, after Higgie & Tommasi (2012); C: Conceptual 2D cross-section (not to scale) of a fast-spreading ridge, showing the context of formation of the studied section (circle zooming on the MTZ), redrawn after Higgie & Tommasi (2012).

Higgie & Tommasi (2012) collected 34 oriented samples of all lithotypes forming the MTZ section. Some of the collected samples are characterized by a fine-scale layering (Figs. 5.1c, 5.2e,f).

They classified these samples into three groups, based on the olivine modal composition analyzed in EBSD maps: 1) Dunites (sensu lato), containing more than 70 vol% modal olivine, therefore including samples of dunites, olivine-rich wehrlites and olivine-rich troctolites (Fig. 5.2f); 2) Olivine-rich gabbros, containing between 30 vol% and 70 vol% modal olivine; and 3) Olivine gabbros, formed by less than 30 vol% of olivine (Fig. 5.2e).

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Figure 5.2: Field occurrences, after Higgie & Tommasi (2012). A: Sub-horizontal layering in the central part of the section characterized by thick gabbroic lenses (up to 10 cm thick) intercalated with thin olivine- rich layers that grade upwards into a more olivine-rich domain containing thin gabbroic layers; B:

Parallelism between the compositional layering and the foliation; C: Subhorizontal lineation (dashed lines) on the foliation plane; D-F: Detail photographs illustrating the variation in scale and structure of the layering; D: Lateral pinch-like termination of a centimetre-thick gabbroic layer; E: Diffuse transition between olivine-rich and gabbroic layers; F: Fresh-cut section showing gradational transitions between dunitic and olivine-rich gabbros; note the preferred orientation of the plagioclase aggregates parallel to the cm-scale layering.

Dunite samples are characterized by partly serpentinized plurimillimetre-size olivine grains, forming a foliation characterized by a slight elongation of the olivine crystals (aspect ratios 2:1;

Figs. 5.3a, 5.4a). Olivines are deformed with the common occurrence of kink bands normal to the crystal elongation. In most dunitic samples, clinopyroxene is more frequent than plagioclase, and occurs as rounded subidiomorphic crystals, slightly elongated within the foliation defined by olivine

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(Fig. 5.3a). In wehrlitic and troctolitic samples, both plagioclase and clinopyroxene occurs as interstitial minerals, often including small rounded olivine grains (Figs. 5.3b, 5.4a). Small anhedral spinels are common in dunitic samples, either interstitial or included in olivine (Fig. 5.4a). They form clear trails of corroded spinel grains into the dunite samples (>95 vol% olivine).

Figure 5.3: Microstructures characterizing the transition from dunitic to olivine gabbro samples. White contours outline the contact between olivines and interstitial minerals. The black arrow in the upper right corner is perpendicular to the sample foliation A: Dunite; Rounded undeformed clinopyroxene developing embayments on olivine; B: Dunite; Poikilitic plagioclase developing embayments and including olivine; C- F: Co-existing plagioclase and clinopyroxene corroding and disrupting coarse olivine crystals in elongated thin olivine grains. Both clinopyroxene and plagioclase include small rounded olivines.

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Olivine-rich gabbros show an evolution of the texture of the olivine matrix with decreasing modal composition of olivine. In the most olivine-rich gabbros, elongated olivine aggregate chains form the foliation (Figs. 5.3c,d, 5.4b), and a “network” of connected olivine crystals is often preserved (Fig. 5.3c,d). With increasing modal composition of interstitial plagioclase and clinopyroxene, the olivine is progressively more corroded and disrupted into several olivine grains (Fig. 5.3c). Single elongated olivines are anhedral and corroded by the interstitial minerals, and in places still show subgrain boundaries (Fig. 5.3c). The interstitial minerals are undeformed and show a preferred elongation within the foliation, parallel to the olivine chains and single olivines long axis (Figs. 5.3c,d, 5.4b). Interstitial and poikilitic plagioclase and clinopyroxene include small rounded clinopyroxene and olivine, and plagioclase and olivine, respectively (Fig. 5.3c,d). Few small interstitial undeformed orthopyroxenes are observed within olivine-rich gabbros (Fig. 5.4b)

Olivine gabbros range from samples composed of about 30 vol% of modal olivine (Fig.

5.4c), showing textures similar to the olivine-rich gabbros (i.e. olivine aggregates elongated within the foliation), to samples almost free of olivine (5 vol% modal olivine), where olivines are disaggregated and embedded in interstitial to poikilitic plagioclase and clinopyroxene (Figs. 5.3e,f, 5.4d). Progressive olivine corrosion by the interstitial minerals leads to the disruption of the olivine grains and therefore to a reduction of the grain size. Olivines are anhedral and corroded, showing cusped grain boundaries, and still show a shape preferred elongation within the foliation plane (Figs. 5.3e,f, 5.4d). Subgrain boundaries are rarer than in olivine-rich gabbros, but they are still observed locally. Plagioclase and clinopyroxene are undeformed, show irregular shapes, and include small rounded olivines (Fig. 5.3e,f). They also show a shape preferred elongation within the foliation plane (Figs. 5.3f, 5.4d), marking the lineation. As described in the olivine-rich gabbros, few small interstitial undeformed orthopyroxenes are observed.

The described textural evolution, from dunite to wehrlite and troctolite, to olivine-rich gabbro, to olivine gabbro, is indicative of a melt-rock interaction process occurring during the reactive crystallization of the interstitial phases (Higgie & Tommasi, 2012). This reactive porous flow led to the progressive dissolution of the pre-existing olivine matrix (i.e. the MTZ dunite) by the percolating melt. Similar olivine-dissolving melt-rock interactions have been extensively described during impregnation processes and the replacive formation of plagioclase peridotites (Dijkstra et al., 2001, 2003; Borghini et al., 2007; Rampone & Borghini, 2008; Tursack & Liang, 2012; Saper & Liang, 2014; Basch et al., in revision), and more recently hybrid olivine-rich troctolites (Suhr et al., 2008; Drouin et al., 2009, 2010; Renna & Tribuzio, 2011; Sanfilippo &

Tribuzio, 2013; Sanfilippo et al., 2013, 2014, 2015, 2016; Rampone et al., 2016; Basch et al., in revision). Plagioclase and clinopyroxene are crystallized together, as indicated by the occurrence of inclusions of the co-existing mineral in both plagioclase and clinopyroxene. The corrosion and disruption of olivine during reactive crystallization of the melt and replacive formation of the olivine gabbros is well indicated by the progressively decreasing grain size at increasing tortuosity of the olivine grains, and the presence of olivine small rounded inclusions within the poikilitic minerals (Suhr et al., 2008; Drouin et al., 2010; Higgie & Tommasi, 2012). The occurrence of orthopyroxene in olivine-rich gabbros and olivine gabbros is indicative of the chemical evolution of percolating melt during late-stage crystallization and closure of the porosity at decreasing melt mass (Kelemen et al., 1997; Korenaga & Kelemen, 1997).

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Figure 5.4: EBSD phase maps; White areas are non-indexed, red is olivine, blue is plagioclase, green is clinopyroxene and yellow is orthopyroxene. The black arrow is perpendicular to the sample foliation. A:

Dunite 10OK27, 90 vol% olivine; B: Layered olivine-rich gabbro 10OK15, 25-45 vol% olivine; C: Layered olivine gabbro 10OK20B, 30 vol% olivine; D: Olivine gabbro 10OK16A, 12 vol% olivine; Samples after

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5.3 Crystallographic Preferred Orientation

5.3 Crystallographic Preferred Orientation

Higgie & Tommasi (2012) performed EBSD analyses on samples ranging from dunite (95 vol% modal olivine) to olivine gabbro (5 vol% modal olivine) and measured the Crystallographic Preferred Orientations (CPO) of olivine, clinopyroxene and plagioclase. The rather small grain size characterizing the samples from the Oman Moho Transition Zone ensured good statistics and therefore reliable measurements of the rock-forming mineral CPO (Mainprice et al., 2014) and J- index, which is the volume averaged integral of the squared orientation densities (Bunge, 1982), representative of the fabric strength (olivine J-index usually ranges from 2 to 25 in mantle peridotites; Ben Ismail & Mainprice, 1998; Tommasi et al., 2000). The CPO of olivine, clinopyroxene, and plagioclase, when present, are shown in Figures 5.5, 5.6, and 5.7, for dunite, olivine-rich gabbro, and olivine gabbro samples, respectively (data from Higgie & Tommasi, 2012). All samples have a well-developed olivine CPO. Clinopyroxene and plagioclase, when significantly present, also show clear CPO, although more dispersed than olivine.

Figure 5.5: Modal compositions, olivine, clinopyroxene and plagioclase Crystallographic Preferred Orientations of the dunitic samples. Lower hemisphere equal-area projections, colour scale fit to the maximum concentration observed in the three crystallographic axes. J-index is indicative of the fabric strength (Bunge, 1982). When insufficient amounts of grains were analyzed, the CPOs are represented without contouring (Mainprice et al., 2014). Samples after Higgie & Tommasi (2012). In all samples, the foliation is E-W.

Olivine is characterized by a strong axial-[100] CPO in the dunites (Fig. 5.5), with [100]

olivine axes parallel to the lineation, within the foliation plane, and girdle orientation of the [010]

axes, with a maximum normal to the foliation plane. [001] olivine axes shows more dispersed orientations than [100] and [010], but form a girdle normal to the lineation with a maximum in the foliation plane. In olivine gabbros (Fig. 5.6), the olivine [100] axes show more dispersed orientations, forming a girdle in the foliation plane, and the olivine [010] axes concentrate normal to

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the foliation plane, resulting in an axial-[010] pattern. Olivine-rich gabbros (Fig. 5.7) show transitional olivine CPO patterns, characterized by [100] axes forming a girdle orientation in the foliation plane, but preserving the maximum orientation parallel to the lineation.

Figure 5.6: Modal compositions, olivine, clinopyroxene and plagioclase Crystallographic Preferred Orientations of the olivine gabbro samples. Lower hemisphere equal-area projections, colour scale fit to the maximum concentration observed in the three crystallographic axes. J-index is indicative of the fabric strength (Bunge, 1982). When insufficient amounts of grains were analyzed, the CPOs are represented without contouring (Mainprice et al., 2014). Samples after Higgie & Tommasi (2012). In all samples, the foliation is E-W.

The analysis of the CPO symmetry of the olivine [100] and [010] axes (Fig. 5.8; Vollmer, 1990; Mainprice et al., 2014) shows that the transition from an axial-[100] to an axial-[010] olivine CPO is progressive, over a range of olivine modal composition from 70 vol% to 40 vol% modal olivine (progressive decrease in P/G ratio of [100] olivine axis and increase in P/G ratio of [010]

axis at decreasing modal olivine contents; Fig. 5.8). This progressive variation is however observed in all samples, regardless of the thickness of the layer (Fig. 5.8). No correlation was found between modal composition and the CPO strength (J-index = 4-12).

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Figure 5.7: Modal compositions, olivine, clinopyroxene and plagioclase Crystallographic Preferred Orientations of the olivine-rich gabbro samples. Lower hemisphere equal-area projections, colour scale fit to the maximum concentration observed in the three crystallographic axes. J-index is indicative of the fabric strength (Bunge, 1982). When insufficient amounts of grains were analyzed, the CPOs are represented without contouring (Mainprice et al., 2014). Samples after Higgie & Tommasi (2012). In all samples, the foliation is E-W.

Clinopyroxene CPO are characterized by a concentration of [001] axes parallel to the lineation, of (010) planes parallel to the foliation, and [100] axes more dispersed, both in olivine- rich gabbros (Fig. 5.7) and olivine gabbros (Fig. 5.6). The fabric strength (J-index; Bunge, 1982) does not show any correlation with the olivine or clinopyroxene modal contents, with J-indexes showing relatively low and constant values (J-index = 2-5).

Plagioclase in both olivine-rich gabbros (Fig. 5.7) and olivine gabbros (Fig. 5.6) shows a clear CPO, characterized by [100] axes orientation forming a girdle within the foliation plane, with a clear maximum parallel to the lineation. [010] axes show strong orientations normal to the

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foliation plane, with in some cases dispersion in a plane normal to the lineation. The plagioclase fabric strength (J-index; Bunge, 1982) does not show any correlation with the olivine or plagioclase modal contents, with J-indexes showing relatively low and constant values (J-index = 3-6).

Figure 5.8:Variation of symmetry of the olivine CPO, redrawn after Higgie & Tommasi (2012), described by the ratio between the Point (P), Girdle (G) components of the distribution of the olivine [100] and [010]

axes, function of the olivine modal content. Point(P), Girdle(G) and Random (R) components of the distribution of the olivine [100] and [010] axes are calculated from the eigen values (λ1, λ2, λ3) of the normalized orientation matrix for each crystallographic axis as P = (λ1 - λ3), G = 2x (λ2 - λ3), and R = 3x λ3

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The CPO analyses performed by Higgie & Tommasi (2012) highlighted good correlations with the modal composition of the various layers of dunite, wehrlite, troctolite, olivine-rich gabbro and olivine gabbro. The olivine axial-[100] patterns described in dunite samples (>70 vol% modal olivine; Fig. 5.5) are typically described in mantle rocks deformed in simple shear at high temperature (>1000°C) and moderate pressures (e.g. Zhang & Karato, 1995; Ben Ismail &

Mainprice, 1998; Bystricky et al., 2000; Tommasi et al., 2000; Karato et al., 2008). In contrast, the olivine axial-[010] patterns analyzed in olivine-rich gabbros (30-70 vol% modal olivine; Fig. 5.7) and olivine gabbros (<30 vol% modal olivine; Fig. 5.6) are consistent with CPO obtained in olivine-melt aggregates in melt-bearing simple shear experiments (Zimmerman et al., 1999, Holtzman et al., 2003b). Such patterns have also been described as a transition of preferential slip system at high pressure (>3GPa) and/or water contents (Durham & Goetze, 1977; Jung & Karato, 2001; Mainprice et al., 2005; Jung et al., 2006, 2009). At the Moho Transition Zone, the millimetre- scale at which this transition is observed, and the lack of hydrous minerals in the rock mineralogical assemblage indicate that it is unlikely that the transition from olivine axial-[100] to axial-[010] is related to high pressure or water contents. Moreover, the parallelism between the compositional layer and deformation structures (Fig. 5.2a,b), and the correlation between the olivine crystals elongation and CPO, and the plagioclase and clinopyroxene CPO in the olivine-rich gabbros and olivine gabbros, points to a deformation-controlled melt distribution (Fig. 5.9; Higgie & Tommasi, 2012).

Figure 5.9: Representative sketch of the formation process of the dunite-olivine gabbro layering observed at the Oman Moho Transition Zone. Positive feedback between melt focusing and deformation leads to the formation of melt-rich layers, and therefore to the replacive formation of olivine gabbros.

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These observations point to a positive feedback between melt distribution and deformation at the Moho Transition Zone (Fig. 5.9; Higgie & Tommasi, 2012), as was observed in melt-bearing shearing experiments (Holtzman et al., 2003b). The focussed melt percolation in deforming layers allowed extensive reaction between the olivine matrix (dunite) and the percolating melt, progressively dissolving olivine and crystallizing clinopyroxene and plagioclase (Higgie &

Tommasi, 2012), leading to the replacive formation of olivine-poor gabbros (5 vol% modal olivine). The preservation of a clear olivine CPO in all layers, independently of their modal composition (Figs. 5.5, 5.6, 5.7, 5.8), implies that the instantaneous melt fraction always remained below the critical value allowing for the loss of cohesion of the solid matrix (20-40% melt/rock ratio; Rosenberg & Handy, 2005). The progressive olivine dissolution and the CPO transition from olivine axial-[100] to axial-[010] are therefore a process integrated over time, indicative of an open system layer-parallel melt percolation rather than a melt impregnation process (Kelemen et al., 1997; Korenaga & Kelemen, 1997; Higgie & Tommasi, 2012).

5.4 Major and trace elements mineral compositions

We selected a set of samples representative of the lithological variation observed from dunite (94 vol% modal olivine) to olivine-rich wehrlite to olivine-rich troctolite to olivine-rich gabbros to olivine gabbros (12 vol% modal olivine) (Table 5.1), and characterized by a well- defined CPO (Figs. 5.5, 5.6, 5.7; Higgie & Tommasi, 2012). Mineral major and trace elements of olivine, clinopyroxene, plagioclase and spinel, together with trace elements analyses standard deviations are given in Appendix 2. Overall our data shows consistency with mineral major and trace elements compositions reported in previous geochemical investigations of the Oman Moho Transition Zone layering between dunites and variably evolved olivine gabbroic lenses (Kelemen et al., 1997; Korenaga & Kelemen, 1997; Koga et al., 2001; Higgie & Tommasi, 2012; Nicolle et al., 2016), except for the olivine trace elements that represent a first dataset in the Oman MTZ.

Table 5.1: Modal compositions and J-index of the studied samples of dunite, olivine-rich gabbro and olivine gabbro; OR = Olivine-rich layer; PR = Plagioclase/Pyroxene-rich layer

Modal compositions J-index

Sample Lithotype Ol % Cpx % Plg % Total J-index Ol J-index Cpx J-index Plg

10OK16A Olivine Gabbro 12 32 56 100 3.07 2.07 3.00

10OK25A - PR Olivine Gabbro 17 43 40 100 5.10 3.11 4.56

10OK10 Olivine Gabbro 27 28 45 100 3.86 2.16 3.61

09OAT10A1 - PR Ol-rich Gabbro 34 3 63 100 3.08 - 5.40

10OK25A - OR Ol-rich Gabbro 45 26 29 100 4.61 3.86 4.76

09OAT10A1 - OR2 Ol-rich Gabbro 48 16 36 100 5.51 5.07 5.82

09OAT10A1 - OR1 Ol-rich Gabbro 65 0 35 100 2.90 - 5.62

10OK34B Dunite 76 15 9 100 4.37 4.47 4.63

10OK27 Dunite 90 10 0 100 3.94 2.75 -

10OK11B Dunite 94 6 0 100 6.46 5.78 -

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5.4 Major and trace elements mineral compositions

Olivines show a significant range of Forsterite content from dunites to olivine gabbros (Fo

= 79.59-88.99 mol%), largely correlated with the olivine modal content (Fig. 5.10). Dunites (>70%

modal olivine) show the highest olivine Forsterite values (Fo = 87.30-88.99 mol%), and olivine gabbros the lowest (Fo = 79.59-84.50 mol%). Olivine-rich gabbros show intermediate Fo contents (Fo = 84.02-88.36 mol%) and strong variations between samples.

Figure 5.10: Modal composition of olivine (determined by EBSD) vs Forsterite content. Layered samples have been separated and each layer considered as single samples.

Moderately incompatible elements in olivine show good correlations with the Forsterite contents (Fig. 5.11). In the dunites, olivines show the lowest Mn (1447-1653 ppm), Co (140.5-171.8 ppm), Zn (28.8-54.7 ppm) and highest Ni (1814-2651 ppm) concentrations. Consistently, olivines in the olivine gabbros show the highest Mn (1778-2407 ppm), Co (169.5-182.9 ppm), Zn (62.4-99.3 ppm) and lowest Ni (1136-1625 ppm) concentrations, and olivines in the olivine-rich gabbros show intermediate compositions of Mn (1540-2052 ppm), Co (155.9-174.3 ppm), Zn (46.1-75.3 ppm) and Ni (1292-1837 ppm). Overall, olivine compositions in the studied samples are intermediate between the compositions of olivine in spinel peridotites (De Hoog et al., 2010) and in the Oman lower crustal gabbros (Browning, 1982), and similar to those reported by Korenaga & Kelemen (1997) in the Moho Transition Zone gabbros (Fig. 5.11). They are also consistent with the olivine compositional trend analyzed in variably evolved olivine gabbros from the Erro Tobbio peridotitic body (see chapter 3 of this thesis).

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Figure 5.11: Forsterite content in olivine vs composition in moderately incompatible elements in ppm. A:

Mn (olivine); B: Co (olivine); C: Ni (olivine); D: Zn (olivine). Compositional fields are after De Hoog et al., 2010 for spinel peridotites; Browning, 1982 for lower crustal gabbros; Korenaga & Kelemen, 1997 for Moho Transition Zone gabbros; and Erro Tobbio olivine gabbros (see chapter 3 of this thesis). Layered samples have been separated and each layer considered as single samples.

Incompatible elements (Ti, Y, REE) in olivine also show correlations with the Forsterite content (Fig. 5.12) between the analyzed lithotypes. Olivine in the dunites show the highest Yb concentrations (YbN = 0.073-0.363 times C1), whereas olivines in the olivine gabbros exhibit the lowest Yb concentrations (YbN = 0.020-0.062 times C1) (Fig. 5.12). Again, olivines in olivine-rich gabbros display intermediate incompatible trace elements compositions (YbN = 0.036-0.111 times C1). The C1-normalised Yb concentrations (after Sun & McDonough, 1989) are slightly lower than those analyzed in olivines from the MORB-type olivine gabbros in the Erro Tobbio peridotitic body (Fig. 5.12; Rampone et al., 2016; see chapter 3 of this thesis).

Consistently with the olivines YbN contents from the Oman Moho Transition Zone (Fig.

5.12), variations are observed in olivine REE patterns, with olivines in dunites showing higher REE absolute concentrations than olivine-rich gabbros and olivine gabbros, at similar MREE/HREE

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fractionation (Fig. 5.13a). These olivine REE abundances are lower than those of olivines in olivine-rich troctolites from the Mid-Atlantic Ridge (Drouin et al., 2009) and the Erro Tobbio peridotitic body (Rampone et al., 2016; chapter 3 of this thesis). Olivines analyzed in both dunites and olivine gabbros do not show the positive Zr-Hf anomalies reported in olivine-rich troctolites from the Mid-Atlantic Ridge (Drouin et al., 2009) and the Erro Tobbio peridotitic body (Rampone et al., 2016; chapter 3 of this thesis), but show broadly similar concentrations in the most incompatible elements in olivine (Fig. 5.13b; Drouin et al., 2009).

Figure 5.12: Forsterite content in olivine vs Yb normalized to C1-chondrite. Normalisation values after Sun

& McDonough (1989). The compositional field of olivine gabbros from the Erro Tobbio (Rampone et al., 2016; see chapter 3 of this thesis) is represented for comparison.

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Figure 5.13: Average REE and trace elements compositions of olivine. A: C1-chondrite normalized REE compositions of olivine in dunite, olivine-rich gabbro, and olivine gabbro; B: Primitive Mantle normalized compositions of olivine. C1-Chondrite and Primitive Mantle normalisation values after Sun & McDonough (1989). Compositional fields represent the olivine compositions in olivine-rich troctolites from the Mid- Atlantic Ridge (Drouin et al., 2009) and from the Erro Tobbio ophiolite (Rampone et al., 2016).

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Clinopyroxenes Mg-values and Cr2O3 (wt%) concentrations (Fig. 5.14) show a positive correlation with Forsterite contents in olivine, with the highest Mg-values and Cr2O3 in clinopyroxene analyzed in dunites (Mg# = 88.19-90.49 mol%, Cr2O3 = 1.18-1.38 wt%) and the lowest in olivine gabbros (Mg# = 83.38-87.59 mol%, Cr2O3 = 0.37-0.56 wt%). Consistently with Forsterite contents in olivine, clinopyroxenes in olivine-rich gabbros show intermediate Mg-value and Cr2O3 compositions (Mg# = 86.66-90.21 mol%, Cr2O3 = 0.46-1.27 wt%). The correlation between the Mg-value of clinopyroxene and the Forsterite content of olivine follows the expected theoretical equilibrium line calculated assuming mineral-melt Fe-Mg distribution coefficients of 0.30 for olivine and 0.23 for clinopyroxene, after Lissenberg & Dick (2008). Clinopyroxenes Cr2O3

compositions in olivine gabbros are consistent with compositions of clinopyroxenes reported in Oman lower crustal gabbros (Browning, 1982) and Oman Moho Transition Zone gabbros (Korenaga & Kelemen, 1997), whereas Cr2O3 contents of clinopyroxenes in dunites are anomalously high and similar to those of clinopyroxenes in the Erro Tobbio olivine-rich troctolites (Borghini & Rampone, 2007; chapter 3 of this thesis).

Figure 5.14: Forsterite contents in olivine vs major elements compositions of clinopyroxene.

A: Clinopyroxene Mg-value (mol%), confronted to the theoretical Fe-Mg equilibrium between olivine and clinopyroxene, after Lissenberg &

Dick., 2008. Mg# = Mg/(Mg+Fe). The dashed lines represent the calculated olivine- clinopyroxene equilibrium line assuming an uncertainty of ±0.02 on the mineral-melt partition coefficients; B: Clinopyroxene Cr2O3

(wt%), compared to the compositional field of Moho Transition Zone gabbros (Korenaga &

Kelemen, 1997) and lower crustal gabbros (Browning, 1982), and to troctolites from the Erro Tobbio peridotitic body (Borghini &

Rampone, 2007; chapter 3 of this thesis).

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Figure 5.15: Average REE and trace elements compositions of clinopyroxene. A: C1-chondrite normalized REE compositions of clinopyroxene in dunite, olivine-rich gabbro, and olivine gabbro; B: Primitive Mantle normalized compositions of clinopyroxene. C1-Chondrite and Primitive Mantle normalisation values after Sun & McDonough (1989). Compositional fields represent the clinopyroxene compositions in lower crustal gabbros (Kelemen et al., 1997), Moho Transition Zone gabbroic sills (Kelemen et al., 1997; Koga et al.,

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With the exception of the olivine-rich gabbro sample 09OAT10A1 (Table 5.1) (YbN = 2.94 times C1; CeN/YbN = 0.23), REE and trace elements of clinopyroxenes analyzed in the dunites, olivine-rich gabbros and olivine gabbros (Fig. 5.15a) show homogeneous compositions (YbN = 3.93-5.88 times C1; CeN/YbN = 0.110-0.163), dunites showing slightly higher absolute concentrations (YbN = 4.64-5.88) and lower LREE fractionation (CeN/YbN = 0.110-0.124). They show similar REE compositions to clinopyroxenes analyzed in the Oman Transition Zone gabbroic sills (Kelemen et al., 1997, Koga et al., 2001; Nicolle et al., 2016) and Oman lower crustal gabbros (Kelemen et al., 1997), and slightly lower trace elements concentrations than clinopyroxenes analyzed in MORB-type olivine-rich troctolites from the Mid-Atlantic Ridge (Drouin et al., 2009).

In the olivine-rich gabbros and olivine gabbros, clinopyroxenes show slightly positive Sr anomalies (SrN/Sr*N = 1.25-1.42) (Fig. 5.15b), and slightly positive Eu anomaly (EuN/Eu*N = 0.97-1.13) (Fig.

5.15).

Figure 5.16: A: Mg-value in clinopyroxene (Mg/(Mg+Fe)) vs C1-normalized Yb concentrations in clinopyroxene. The fractional crystallization trend of evolution (Kelemen et al, 1997) is plotted as comparison; B: Chondrite-normalized Nd/Yb ratios of calculated melts in equilibrium with clinopyroxene, using clinopyroxene/melt partition coefficients from Hart & Dunn (1993). The black dots represent Pacific MORB samples (from the RIDGE Petrological Database), and the compositional field represent melts calculated in equilibrium from Oman Moho Transition Zone clinopyroxenes, using the same set of partition coefficients as we are (Koga et al., 2001). Normalisation values after Sun & McDonough (1989).

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Clinopyroxenes analyzed in the dunites show enrichments in YbN at constant Mg-value (Fig.

5.16a), and therefore do not follow the expected trend of fractional crystallization of a primitive melt (Kelemen et al., 1997), defined roughly by the clinopyroxenes analyzed in olivine-rich gabbros and olivine gabbros (Fig. 5.16a). Melts calculated in equilibrium with clinopyroxene, using the experimental set of clinopyroxene/melt partition coefficients from Hart & Dunn (1993) show variations in YbN consistent with the variation observed in clinopyroxenes, at constant NdN/YbN

fractionation (Fig. 5.16b). These equilibrium melts are similar to melts previously calculated in equilibrium with clinopyroxenes in the olivine gabbros (Koga et al., 2001) and Pacific MORBs (RIDGE Petrological database; Fig. 5.16b).

Plagioclase shows high Anorthite contents in both dunites (An = 82.56-85.07 mol%) and olivine-rich and olivine gabbros (An = 87.43-92.12 mol%) as previously reported in the Oman Moho Transition Zone (Browning, 1982, 1984; Kelemen et al., 1997; Korenaga & Kelemen, 1997;

Koga et al., 2001; Nicolle et al., 2016). It shows higher Anorthite contents (Fig. 5.17) than analyzed in oceanic gabbroic suites at slow- (South-West Indian Ridge, Dick et al., 2002; Mid-Atlantic Ridge, Ross & Elthon, 1997; Lissenberg & Dick, 2008; Suhr et al., 2008; Drouin et al., 2009;

Miller et al., 2009; Mid-Cayman Rise, Elthon, 1987) and fast-spreading environments (East Pacific Rise, Lissenberg et al., 2013), and in the Erro Tobbio olivine gabbros (Rampone et al., 2016;

chapter 3 from this thesis). Previous studies interpreted the high Anorthite contents in plagioclase as a result of the crystallization of hydrous melts, despite the absence of hydrous phases in the mineralogical assemblage (i.e. amphibole; Fig. 5.4) (Kelemen et al., 1997; Korenaga & Kelemen, 1997; Koga et al., 2001). In a Forsterite-Anorthite covariation diagram, the studied Oman dunites and gabbros define different trends of evolution, from variable Anorthite contents in plagioclase (An = 82.56-92.12 mol%) at constant Forsterite in olivine (Fo = 87.30-88.99 mol%) in the most olivine-rich samples (dunites and olivine-rich gabbros), to variable Forsterite contents (Fo = 79.59- 88.36 mol%) at constant Anorthite (An = 87.43-92.12 mol%) in the olivine-rich gabbros and olivine gabbros (Fig. 5.17).

Plagioclase shows homogeneous REE and trace elements compositions in dunites, olivine- rich gabbros and olivine gabbros (Fig. 5.18a), at slightly lower trace elements absolute concentrations than plagioclases recently analyzed in gabbroic lenses at the Oman Moho Transition Zone (Nicolle et al., 2016), and plagioclases analyzed in the Oman lower crustal gabbros (Kelemen et al., 1997) and MORB-type olivine-rich troctolites from the Mid-Atlantic Ridge (Drouin et al., 2009). However, plagioclases in dunites show slightly higher LREE concentrations and fractionation (SmN = 0.189-0.260 times C1; CeN/SmN = 1.38-1.84) than plagioclases in olivine-rich gabbros (SmN = 0.203 times C1; CeN/SmN = 1.17) and olivine gabbros (SmN = 0.190-0.211 times C1; CeN/SmN = 0.94-1.57). Plagioclases all show strong positive Eu and Sr anomaly (Fig. 5.18b).

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Figure 5.17: Anorthite content (mol%) in plagioclase vs Forsterite content (mol%) in olivine for olivine- plagioclase couples. Compositional trends and fields represent olivine-plagioclase couples in olivine gabbros and troctolites from South-West Indian Ridge (Dick et al.,2002), Mid-Atlantic Ridge (Ross &

Elthon, 1997; Lissenberg & Dick, 2008; Suhr et al., 2008; Drouin et al., 2009; Miller et al., 2009), Mid- Cayman Rise (Elthon, 1987), East Pacific Rise (Lissenberg et al., 2013), olivine gabbros from the Erro Tobbio ophiolite (see chapter 3 from this thesis), Oman lower crustal gabbros (Browning, 1984), and Oman Moho Transition Zone olivine gabbros (Korenaga & Kelemen, 1997). The trend of fractional crystallization, assuming a primitive MORB composition, was taken from Korenaga & Kelemen, (1997).

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Figure 5.18: Average REE and trace elements compositions of plagioclase. A: C1-chondrite normalized REE compositions of plagioclase in dunite, olivine-rich gabbro, and olivine gabbro; B: Primitive Mantle normalized compositions of plagioclase. C1-Chondrite and Primitive Mantle normalisation values after Sun

& McDonough (1989). Compositional fields represent the plagioclase compositions in lower crustal gabbros (Kelemen et al., 1997), Moho Transition Zone gabbroic sills (Nicolle et al., 2016), and olivine-rich

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5.5 Discussion

5.5 Discussion

5.5.1 Replacive origin of the Oman Moho Transition Zone

Dunites, olivine-rich gabbros, and olivine gabbros show microstructures indicative of the progressive corrosion of a pre-existing olivine matrix (Fig. 5.3; Higgie & Tommasi, 2012), with the development of embayments of interstitial clinopyroxene and plagioclase on olivine, and the disruption of deformed olivine aggregates in olivine-rich gabbros and olivine gabbros. Similar textural evolution of the olivine matrix has been previously described during the replacive formation of olivine-rich troctolites in oceanic settings (Suhr et al., 2008; Drouin et al., 2010;

Sanfilippo et al., 2013, 2015, 2016; Ferrando et al., in press) and ophiolites (Renna & Tribuzio, 2011; Sanfilippo & Tribuzio, 2012; Sanfilippo et al., 2014, Rampone et al., 2016; Renna et al., 2016; Basch et al., in revision; chapters 3 and 4 of this thesis). These melt-rock reaction microstructures, indicative of an olivine-dissolving reactive melt percolation, are accompanied by variations in the olivine CPO (Basch et al., in revision), from clear axial-[100] patterns, typically observed in mantle peridotites deformed at asthenospheric conditions (high temperature, moderate pressure; Zhang & Karato, 1995; Ben Ismail & Mainprice, 1998; Bystricky et al., 2000; Tommasi et al., 2000; Karato et al., 2008) to axial-[010], described in melt-bearing simple shear experiments (Zimmerman et al., 1999, Holtzman et al., 2003b) and impregnated peridotites (Le Roux et al., 2008; Soustelle et al., 2009). In the Oman Moho Transition Zone, Higgie & Tommasi (2012) described such a variation from axial-[100] olivine CPO in the replacive dunites, to axial-[010]

CPO in the olivine gabbros (Figs. 5.5, 5.6, 5.7, 5.8), as a result of deformation in the presence of melt. The CPO symmetry transition is observed in samples showing a range from 70 vol% to 40 vol% of olivine modal composition (Fig. 5.8; Higgie & Tommasi, 2012). This progressive olivine CPO evolution from olivine-rich samples (>70 mol% modal olivine; axial-[100] CPO) to olivine- poor samples (<40 vol% modal olivine; axial-[010] CPO) is recording the variability in the amount of melt and deformation (integrated over time) involved in the reactive melt percolation process.

Therefore, the positive feedback between melt focusing and deformation led to an open-system extensive reactive melt percolation in the olivine-poor layers (Zimmerman et al., 1999, Holtzman et al., 2003b, 2005; Holtzman & Kohlstedt, 2007; Higgie & Tommasi, 2012, 2014), whereas the small amounts of melts percolating within the dunitic layers allowed the preservation of the mantle dunite protolith axial-[100] CPO, and behaved as a closed-system impregnation.

The constant Forsterite content in olivine (Fo = 87.30-88.99 mol%) and Mg-value in clinopyroxene (Mg# = 88.19-90.49 mol%) analyzed in the dunites and olivine-rich gabbros (Figs.

5.14a, 5.16a, 5.17) is indicative of the buffering of the melt composition by olivine dissolution, at decreasing melt/rock ratios. Therefore, dunites follow trends of reactive crystallization, showing variably evolved compositions of plagioclase (An = 82.56-92.12 mol%) at constant Forsterite contents, as previously reported in impregnated peridotites (Collier & Kelemen, 2010) and olivine- rich troctolites (Sanfilippo et al., 2016; chapter 3 of this thesis). Consistently, clinopyroxenes show enriched Cr2O3 compositions, as previously reported in melt-rock interaction dominated systems, during assimilation of a dunitic component (olivine + Cr-spinel; Borghini et al., 2007; Renna &

Tribuzio, 2011; Renna et al., 2016; chapter 3 of this thesis). In the dunites, clinopyroxene and plagioclase show slightly enriched HREE (Figs. 5.15a, 5.16) and LREE (Fig. 5.18) compositions, respectively, compared to clinopyroxenes and plagioclase analyzed in the olivine gabbros. This is

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indicative of the progressive evolution of the percolating melt during reactive percolation at decreasing melt mass.

On the contrary, in the olivine gabbro layers, the high Anorthite content in plagioclase (An

= 87.43-92.12 mol%; Fig. 5.17), REE absolute concentrations in plagioclase (SmN = 0.190-0.211 times C1; Fig. 5.18), Cr2O3 compositions in clinopyroxene (Cr2O3 = 0.37-0.56 wt%; Fig. 5.14b) and REE absolute concentrations in clinopyroxene (YbN = 3.97-4.30 times C1; Figs. 5.15a, 5.16), together with the lack of Eu negative anomaly in clinopyroxene, are indicative of the low fractionation of the melt in an open-system percolation process, as previously proposed in the Oman Moho Transition Zone (<50% fractionation, and probably <10%; Kelemen et al., 1997; Korenaga &

Kelemen, 1997; Koga et al., 2001). Consistently with the preservation of a clear axial-[010] CPO in the olivine gabbros, the low degree of fractionation of the melt implies a percolation process occurring at low instantaneous melt/rock ratio (<20-40%; Rosenberg & Handy, 2005; Higgie &

Tommasi, 2012), integrated over time (Kelemen et al., 1997; Korenaga & Kelemen, 1997; Koga et al., 2001; Higgie & Tommasi, 2012).

Within the olivine gabbro, poikilitic clinopyroxene includes rounded plagioclase crystals, and show positive Sr anomalies (SrN/SrN* = 1.25-1.42) and slightly positive Eu anomalies (EuN/Eu*N = 0.97-1.13). Within all samples of dunite, olivine-rich gabbro and olivine gabbro, clinopyroxene-plagioclase couples show constant Sr partitioning (Sr plagioclase / Sr clinopyroxene = 14- 18), consistent with the Sr partitioning calculated from experimental clinopyroxene/melt (Kd cpx-melt

Sr = 0.1283; Hart & Dunn, 1993) and plagioclase/melt (Kd plagio-melt Sr = 1.801; Laubier et al., 2014) partition coefficients. The clinopyroxene-melt partition coefficient of Sr is lower than the partition coefficients of the neighbouring Pr and Nd (Hart & Dunn, 1993), thus a negative anomaly in Sr is expected in clinopyroxene crystallized from melts showing flat REE patterns. The positive anomaly analyzed in clinopyroxenes of olivine gabbros indicates Sr enrichment in the percolating melt crystallizing the clinopyroxene. Such positive Sr, Eu anomalies in clinopyroxene, coupled with microstructural observations of rounded plagioclase crystals included into poikilitic clinopyroxene, suggest the corrosion of plagioclase, together with olivine, during the melt percolation, as previously described in the lower gabbroic crust at the Mid-Atlantic Ridge and East Pacific Rise (Blackman et al., 2006; Lissenberg & Dick, 2008; Lissenberg et al., 2013; Lissenberg & MacLeod, 2017). The corrosion of a gabbroic matrix by the percolating melt is in agreement with a process of open system melt percolation and low fractionation of the melt, involving reaction between the pre- existing matrix and the reactive melt (Kelemen et al., 1997; Korenaga & Kelemen, 1997, 1998;

Koga et al., 2001; Higgie & Tommasi, 2012).

In summary, combined microstructural and mineral chemistry evidence show that the dunite and olivine gabbro layers record processes of closed system reactive crystallization and open system melt percolation, respectively. The change of olivine CPO symmetry from axial-[100] in the dunites to axial-[010] in olivine gabbros (Higgie & Tommasi, 2012) marks the transition from the closed system impregnation to open system melt percolation. Mineral compositions in the dunite layers are controlled by the dunite-consuming melt-rock interaction process, leading to a buffering of the melt composition by the pre-existing olivine matrix, whereas mineral compositions in the olivine gabbro layers (pre-existing olivine matrix and interstitial plagioclase and clinopyroxene crystallized from the melt) are controlled by the composition of percolating melt. The structural and geochemical evolution of single layers is entirely ruled by the melt-rock ratio integrated over time, and the related process of melt transport into the Moho Transition Zone (closed vs open system)

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5.5 Discussion

peculiar features that are indicative of an anomalous melt composition. In the following, we discuss the composition of the percolating melt leading to the formation of the impregnated dunites and olivine gabbros.

5.5.2 MORB-type melt signature

The REE compositions of clinopyroxene analyzed in the olivine gabbros (Fig. 5.15a) show patterns consistent with crystallization from a primitive MORB-type melt, as was described in previous studies of the Oman Moho Transition Zone (Kelemen et al., 1997; Koga et al., 2001;

Nicolle et al., 2016). Clinopyroxenes analyzed in the dunites show an enrichment in YbN at constant Mg-value (Fig. 5.16a), consistent with an evolution of the melt controlled by concomitant olivine- consuming melt-rock interaction and trapped melt crystallization. In the olivine-rich gabbros and olivine gabbros, clinopyroxenes show a rough correlation between the Mg-value and the YbN

concentration (Fig. 5.16a), indicative of an evolution of the melt by fractionation (Kelemen et al., 1997). Although different processes are involved in the geochemical evolution of the melts percolating in the dunites and olivine gabbros, the NdN/YbN ratio in calculated melts in equilibrium with the clinopyroxenes are constant (NdN/YbN = 0.81-1.54; Fig. 5.16b), indicative of a common MORB-type parental melt composition in all samples, similar, although more primitive to equilibrium melts calculated in Pacific MORBs (RIDGE petrological database; Koga et al., 2001)

Moreover, the composition of clinopyroxene in the Oman Moho Transition Zone (Koga et al., 2001; this study) is similar in REE concentrations and L-HREE fractionation to clinopyroxenes analyzed in the Oman lower crustal gabbros (Kelemen et al., 1997; Koga et al., 2001). This indicates that the melts percolating within the Moho Transition Zone represent the parental melts feeding the oceanic crust. The occurrence of MORB melts percolating the dunites and leading to the replacive formation of the olivine gabbros confirms that the aggregation of polybaric melts leading to the formation of MORB-type chemical compositions occurs prior to their percolation in the Moho Transition Zone, possibly in the dunite channels crosscutting the mantle harzburgites, during focused melt percolation (Kelemen et al., 1995b).

5.5.3 Controlling factor of the mineral major elements compositions

The open system percolation and low melt fractionation (<50%, probably less than 10%;

Kelemen et al., 1997; Korenaga & Kelemen, 1997; Koga et al., 2001) described in the olivine gabbro layers implies that the melt composition does not evolve significantly, and therefore that the crystallized minerals (interstitial plagioclase and clinopyroxene) represent the primary composition of the incoming melt. This in turn implies that the percolating melt is in equilibrium with minerals forming the olivine gabbro, namely olivine with Forsterite = 79.59-84.86 mol% (Fig. 5.10), clinopyroxene with Mg# = 83.4-87.59 mol% (Fig. 5.14) and plagioclase with Anorthite = 87.43- 91.77 mol% (Fig. 5.17). The high Anorthite content observed in all layers of the Oman Moho Transition Zone, coupled with Forsterite contents and Mg-values in clinopyroxene indicative of a melt Mg-value = 58-60 mol% (Gale et al., 2013; Ferrando et al., in press), is not consistent with the mineral compositional trends reported in oceanic gabbroic suites at the South-West Indian Ridge (Dick et al.,2002), Mid-Atlantic Ridge (Ross & Elthon, 1997; Lissenberg & Dick, 2008; Suhr et al., 2008; Drouin et al., 2009; Miller et al., 2009), Mid-Cayman Rise (Elthon, 1987) and East Pacific Rise (Lissenberg et al., 2013, 2017) (Fig. 5.17).

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High-Anorthite plagioclase crystals have been previously described in MORB lavas (Neilsen et al., 1993), in lower oceanic crust gabbros at the East Pacific Rise (Dick & Natland, 1996; Lissenberg et al., 2013), in the Oman lower crust and Moho Transition Zone (Browning, 1984; Kelemen et al., 1997; Korenaga & Kelemen, 1997; Benoit et al., 1999; Koga et al., 2001;

Koepke et al., 2009; Abily et al., 2011). The dunite samples show that crystallization of clinopyroxene prior to plagioclase accompanies the high Anorthite contents in plagioclase (Fig.

5.3a; Browning, 1984; Kelemen et al., 1997; Korenaga & Kelemen, 1997; Koga et al., 2001). An [olivine – clinopyroxene – plagioclase] crystallization sequence is not consistent with the fractional crystallization of a MORB-type melt at low pressures (2kbar; Koga et al., 2001). Previous studies of the Moho Transition Zone interpreted the high Anorthite contents in plagioclase and reversed clinopyroxene-plagioclase saturation order as due to crystallization from hydrous melts (Kelemen et al., 1997; Korenaga & Kelemen, 1997; Koga et al., 2001), known to favour the early appearance of clinopyroxene on the liquidus of basaltic melts, and drastically increase Anorthite contents in plagioclase at a given temperature (Berndt et al., 2005; Koepke et al., 2009).

Many recent studies of the Oman lower crust and Moho Transition Zone describe the mixture between percolating MORB melts and high-temperature hydrothermal fluids (Benoit et al., 1999; Koepke et al., 2009), possibly brought by synmagmatic normal faults reaching the lower crust (Python et al., 2007, 2011; Abily et al., 2011; Rospabé et al., 2017), to explain the occurrence of hydrous products (amphibole, hydrogarnet, diopsidite, enstatite) and peculiar mineral compositions in dunites, wehrlites, olivine gabbros, and gabbronorites. However, in the studied samples of dunites, olivine-rich gabbros and olivine gabbros from the Moho Transition Zone, no interstitial hydrous minerals have been observed (Figs. 5.3, 5.4).

To investigate the effects of increased water contents in the crystallization order and the composition of minerals crystallized from the percolating melt, we performed a geochemical modelling using the pMELTS thermodynamic software (Ghiorso et al., 2002). We selected as initial melt an average MORB composition characterized by Mg#(melt) ranging from 58 to 60, and rather high Cao contents (i.e. CaO/Na2O > 5.5, and CaO/Al2O3 > 0.85), representing the average of 77 MORB melts after the database of Gale et al. (2013). The calculated average melt composition is given in Table 5.2. We performed the crystallization of the dry average MORB-type melt at 2kbar, as well as after addition of 1 wt% to 4 wt% H2O (Table 5.2). Consistently with the results of experimental crystallization of hydrous melts (Berndt et al., 2005; Husen et al., 2016), the Anorthite content of plagioclase progressively increases with H2O contents in the melt (from An = 78.86 mol% in dry melts to An = 83.91 mol% for H2O = 4 wt%; Table 5.2). The crystallization order of the MORB-type melt also reverses from [olivine – plagioclase – clinopyroxene] to [olivine – clinopyroxene – plagioclase] in melts containing more than 2 wt% H2O (Table 5.2; Berndt et al., 2005; Husen et al., 2016). However, hydrous melts are characterized by a saturation temperature of plagioclase and clinopyroxene progressively decreasing to temperatures lower than 1100°C for melts containing more than 2wt% H2O (Table 5.2; Husen et al., 2016).

In order to get information about the crystallization temperatures recorded by minerals in the studied MTZ olivine gabbros, we calculated equilibrium temperatures of olivine-clinopyroxene and plagioclase-clinopyroxene couples, using the REE-geothermometers from Sun & Liang (2012, 2013a,b, 2014) and Sun et al. (2017), respectively.