Chapter 3: Hybrid origin of the Erro Tobbio troctolites
(Ligurian Alps, Italy): Structural and geochemical evidence of
multi-stage evolution.
narrow Jurassic Ligurian Tethys basin, opened by passive lithosphere extension and breakup of the continental lithosphere, leading to slow-spreading oceanization (Rampone & Piccardo, 2000; Manatschal & Muntener, 2009). Despite its involvement in the Alpine orogeny, the Erro-Tobbio ultramafic body (Voltri Massif, Ligurian Alps, Fig. 1) exposes kilometre-scale unaltered peridotites, mostly devoid of Alpine overprint (Bezzi & Piccardo, 1971; Chiesa et al., 1975; Ernst & Piccardo, 1979; Ottonello et al., 1979; Hoogerduijn Strating et al., 1990, 1993; Piccardo et al., 1990, 1992, 2004; Scambelluri et al., 1991; Vissers et al., 1991; Borsi et al., 1996; Capponi et al., 1999; Rampone et al., 2004, 2005), allowing the study of the pre-Alpine structural and chemical mantle evolution. The Erro-Tobbio unit is mostly made of variably serpentinized spinel-bearing lherzolites to harzburgites. Previous petrologic and structural studies documented a tectono-metamorphic decompressional evolution of these mantle rocks, from deep lithospheric settings (P > 15-20 kbar) to shallow depths (P < 5 kbar), with a progressive reequilibration from spinel- to plagioclase- to amphibole-facies conditions (Hoogerduijn Strating et al., 1990, 1993; Vissers et al., 1991; Rampone et al., 2005), and the development of extensional shear zones forming spinel tectonites, plagioclase-, hornblende-, chlorite-bearing mylonites and serpentinite mylonites (Hoogerduijn Strating et al., 1993). This extension-related exhumation was accompanied by multiple episodes of melt percolation and intrusion, namely: 1) a first open-system olivine-saturated reactive porous flow at spinel-facies conditions, leading to the dissolution of mantle clinopyroxene and orthopyroxene, and crystallization of olivine; 2) at plagioclase-facies conditions (< 8-10 kbar), a melt-rock reaction leading to the formation of plagioclase-bearing impregnated peridotites, by dissolution of olivine and crystallization of plagioclase (± opx ± cpx); 3) multiple episodes of gabbroic intrusions at shallow depths (P < 5kbar) (Piccardo et al., 2004; Rampone et al., 2004, 2005, 2014, 2016; Borghini et al., 2007; Borghini & Rampone, 2007; Piccardo & Vissers, 2007; Rampone & Borghini, 2008). Geochronological studies on gabbroic rocks from the Alpine-Apennine ophiolites indicate a large time span of gabbroic intrusions (~20 Ma) in the Alpine Tethys (Rampone et al., 2014 and reference therein). The Erro-Tobbio olivine gabbros yield the oldest Sm-Nd age of the crustal gabbroic sequences within the Alpine-Apennine ophiolites with an age of 178 ± 5 Ma (Rampone et al., 2014), therefore representing early melt intrusions in thinned lithospheric mantle exhumed at ocean-continent transition domains (Rampone & Piccardo, 2000; Manatschal & Muntener, 2009).
In the South-Eastern part of the Erro-Tobbio peridotite, the impregnated mantle peridotites are in irregular contact with a hectometre-size troctolitic body, previously described as a primitive
cumulate body (Fig. 3.1; Borghini & Rampone, 2008; Borghini et al., 2007; Rampone & Borghini, 2008; Rampone et al., 2016). Late gabbroic dykes crosscut all mantle structures, as well as the troctolitic body-impregnated peridotite contact (Borghini et al., 2007). Rampone et al. (2016) recently demonstrated the important effect of the olivine-dissolving, plagioclase-crystallizing melt-rock interaction in the Erro-Tobbio troctolitic body mineral compositions. It leads to significant enrichments in specific trace elements (Zr, Hf, Ti, HREE), coupled with strong HFSE/REE fractionation in olivine.
Figure 3.1: A: Sketch map of the Northern Apennines and Western Alps. The red square indicates the
location of the Voltri Massif, in the Ligurian Alps; B: Map of the Voltri Massif and location of the studied area within the Erro Tobbio peridotites; C: Geological map of the Mt.Foscallo area, in the Erro-Tobbio peridotites. This structural map merges new data measured on the field with previously published data from
consistent with single depleted melt increments produced by near-fractional melting of a MORB-type asthenospheric mantle source (Piccardo et al., 2004; Borghini et al., 2007; Rampone & Borghini, 2008), whereas parental melts to the troctolitic body and late gabbroic discrete intrusions resemble N-MORB-type aggregated melts (Rampone et al., 1998, 2014, 2016; Borghini & Rampone, 2007; Borghini et al., 2007; Rampone & Borghini, 2008).
3.2 Geology of the studied area
The investigated area exposes a 500-metre wide ultramafic body surrounded by serpentinized high-pressure, low-temperature Alpine shear zones. The ultramafic body preserves a pre-Alpine mantle history, displaying the association between mantle peridotites and ultramafic-mafic bodies and intrusions (from plagioclase wehrlites to troctolites to olivine gabbros) (Fig. 3.1; Borghini & Rampone, 2007; Borghini et al., 2007; Rampone & Borghini, 2008; Rampone et al., 2016). Mantle peridotites are Plagioclase-bearing Lherzolites showing in places a weak tectonic foliation defined by ortho- and clinopyroxene shape-preferred orientation. They are primarily associated to metre-size dunitic pods and centimetre-size pyroxenite layers showing a constant NNE-SSW orientation and strongly dipping to the East (Fig. 3.1). In the northernmost part of the ultramafic body, the plagioclase lherzolites are in irregular contact with a hectometre-size troctolitic body. The contact is marked by the occurrence of troctolitic and plagioclase-bearing wehrlite apophyses into the mantle peridotites, crosscutting the pyroxenite banding (Borghini & Rampone, 2007; Borghini et al., 2007; Rampone & Borghini, 2008; Rampone et al., 2016). Detailed mapping and sampling in selected outcrops revealed that the inner troctolitic body is characterized by a high modal composition variability, from plagioclase wehrlite to troctolite to dunite, and a structural complexity characterized by different generations of troctolites showing crosscutting relationships and highly variable olivine texture. In the following, based on these structural criteria, we distinguish different types of troctolites within the mafic body.
The Troctolite A is in irregular contact with the mantle peridotites and develops decimetre-thick troctolite and plagioclase-bearing wehrlite apophyses into the plagioclase lherzolites (Fig. 3.2a) in a transition zone where it is difficult to distinguish both lithotypes (Fig. 3.1). The
Troctolites A show variable olivine modal contents (from 55 to 74 vol%; Table 1, Fig. 3.2b,c) and interstitial plagioclase ± clinopyroxene, and it includes decimetre-size dunitic pods (Fig. 3.2d). The modal composition variability between olivine-rich and plagioclase-rich troctolite forms a local sub-vertical decimetre-size layering showing a NNW-SSW orientation, dipping to the East (Figs. 3.1, 3.2a).
Figure 3.2: Troctolite A structures. A: Plagioclase-rich layering within the host Troctolite A; B:
Crosscutting relationship between Troctolite A and B; C: Dunitic pod associated with the Troctolite A; D: Troctolite apophysis within the mantle peridotites at the contact between the troctolitic body and the peridotites, and gabbroic dike crosscutting the association between peridotites and troctolites.
variations at the scale of a few centimetres, from millimetre-size euhedral olivine crystals to centimetre- and decimetre-size hopper and dendritic olivine crystals (Fig. 3.3c,d,e). The olivine textural layering observed in Troctolite B, between granular and dendritic portions of the pseudo-tabular bodies (Fig. 3.3e) shows NNE-SSW orientations strongly dipping to the East, similar to the plagioclase enrichment layering in the Troctolite A (Fig. 3.1).
Figure 3.3: Troctolite B structures. A: Troctolite B crosscutting the layering of plagioclase enrichment in
Troctolite A (in red); B: Irregular contact between Troctolite A and crosscutting Troctolite B; C: Textural complexity within the Troctolite B; The white square indicates the location of (d); D: Dendritic “fishbone” olivine crystal; E: Textural variability of olivine crystals at centimetre-scale within the Troctolite B; The dashed line separates granular areas from hopper and dendritic areas.
The peridotites and troctolitic body are both intruded by decametre-size Olivine Gabbro bodies, centimetre- to metre-thick troctolitic to olivine gabbro dykes and centimetre-thick dykelets, all oriented NNW-SSE, and dipping to the East (40-50°; Figs. 3.1, 3.2a; Borghini & Rampone, 2007; Borghini et al., 2007) although in places dykelets occur as conjugate pairs. Dykes and dykelets are in straight contact with the host rock and show no chilled margins. They display a grain-size variability, from finer grains towards the margin of the intrusion (millimetre-size crystals), to coarse grains (centimetre-size crystals) in the core of the dyke. Figure 3.4 summarizes the field relationships mapped in the Erro-Tobbio ultramafic body, between Plagioclase lherzolites,
Host Troctolite A, Troctolite B, and Gabbroic Intrusions.
Figure 3.4: Representative sketch of the field associations and complexities observed between the peridotitic
troctolitic body, as well as within the troctolitic body.
3.3 Sampling and analytical methods
In this study, we used samples of Spinel Lherzolites, Plagioclase Lherzolites, Troctolites and
Olivine Gabbros collected during previous petrological investigations of the studied area (Fig. 3.1; Rampone et al., 2004, 2005, 2014, 2016; Borghini & Rampone, 2007; Borghini et al., 2007), as well as newly collected samples of Troctolite and Olivine Gabbros. The Spinel and Plagioclase
Lherzolites have been sampled in a nearby location in respect to the troctolitic body, within the
Erro-Tobbio peridotites, as the alteration is much less developed. These samples are used as a structural and chemical reference of the mantle protolith, prior to the formation of the troctolitic body and gabbroic dykes. Table 3.1 reports the modal composition of the 40 studied samples, namely 3 Spinel Lherzolites, 4 Plagioclase Lherzolites, 11 Troctolites A, 1 Dunite pod, 5 Wehrlite and Troctolite Apophyses, 10 Troctolite B, and 6 Troctolitic to Olivine Gabbro intrusions. We performed structural EBSD mapping of all samples, and mineral major (EPMA) element chemical analyses of 24 samples, namely 2 Spinel Lherzolites, 2 Plagioclase Lherzolites, 7 Troctolites A, 1 Dunite pod, 2 Wehrlite and Troctolite apophyses, 5 Troctolites B, and 5 Troctolitic to Olivine
Detailed methodologies for EBSD, major and trace elements analyses can be found in Supplementary Material of the Chapter 4 of this thesis.
3.4 Petrography
Spinel lherzolites show protogranular to porphyroclastic textures of olivine, orthopyroxene, clinopyroxene and spinel grains. Olivine and pyroxenes (orthopyroxene + clinopyroxene) are deformed, and display kink bands and undulatory extinctions, respectively. Clinopyroxenes and orthopyroxenes both show thin exsolution lamellaes of the complementary pyroxene. Spinels are found both as granular grains in the lherzolitic matrix and in orthopyroxene + spinel symplectites at the rim of orthopyroxene porphyroclasts, previously described as an effect of cooling of the peridotites and equilibration at lithospheric temperatures (⁓1000°C; Rampone et al., 2005; Rampone & Borghini, 2008). The spinel lherzolites display melt-rock interaction microstructures with the development of olivine embayments replacing mantle pyroxenes (i.e. pyroxene dissolution and olivine crystallization). These microstructures, associated to an increase of olivine modal compositions, have been extensively described in the Alpine-Apennine ophiolites (Piccardo et al., 2004; Rampone et al., 2005, 2008; Piccardo & Vissers, 2007; Rampone & Borghini, 2008; Basch et al., submitted) and in the Othris Massif (Dijkstra et al., 2003), and interpreted as the result of a pyroxene-dissolving, olivine crystallizing reactive melt percolation at spinel facies.
Plagioclase lherzolites have been previously described as the replacive product of melt impregnation of the spinel lherzolites (Borghini et al., 2007). They show similar textures and microstructures to the spinel facies protolith but are characterized by an enrichment in undeformed interstitial plagioclase and orthopyroxene (Table 3.1), developing embayments on kinked olivine and exsolved clinopyroxene. These melt-rock reaction microstructures are indicative of an orthopyroxene-saturated composition of the impregnated melt, as previously described in the Alpine-Apennine ophiolitic peridotites (Rampone et al., 1997, 2005, 2008, 2016; Muntener & Piccardo, 2003; Piccardo et al., 2004; Borghini & Rampone, 2007; Borghini et al., 2007; Rampone & Borghini, 2008; Basch et al., submitted) and in the Othris Massif (Dijkstra et al., 2003).
The Troctolite A shows a hypidiomorphic texture and variable grain size, from centimetre-size anhedral to millimetre-centimetre-size euhedral olivine crystals. Olivine occurs in two distinct textures, i) as fine-grained undeformed euhedral crystals embedded in interstitial to poikilitic plagioclase and clinopyroxene (Fig. 3.5a,b), and ii) as coarse (up to centimetre-size) deformed corroded grains , showing the occurrence of kink bands (Fig. 3.5c,d). These two types of olivine are commonly found together and in places, fine-grained euhedral crystals of olivine embedded in poikilitic plagioclase or clinopyroxene show the same crystallographic orientation as a neighbouring coarse corroded grain of olivine, therefore possibly indicating a disruption of coarse-grained corroded olivines by the percolating melt (Fig. 3.5c) (see later discussion).
Table 3.1: Studied samples, lithotype, modal composition, and PfJ Olivine
Modal compositions PfJ Olivine
Sample Lithotype Olivine Plagio Cpx Opx [100] [010] [001]
ETR2* Spinel Lherzolite 78 0 7 15 2.33 2.84 2.02
ETR4B* Spinel Lherzolite 75 0 5 20 2.24 1.64 1.56
ETR4A* Spinel Lherzolite 71 0 14 15 2.08 1.65 1.32
P1B* Plagio. Lherzolite 57 10 11 22 2.12 3.61 1.5
P1A* Plagio. Lherzolite 56 10 4 30 1.9 2.67 1.4
P1* Plagio. Lherzolite 55 12 11 22 2.53 4.17 2.23 MF40* Plagio. Lherzolite 53 5 12 30 2.07 3.13 1.63 MF104A Dunite 97 0 3 0 1.61 1.66 1.44 MF21* Troctolite A 74 15 11 0 1.08 1.12 1.19 MF15* Troctolite A 68 31 1 0 1.22 1.5 1.54 MF97 Troctolite A 68 30 2 0 1.19 1.33 1.4 MF51* Troctolite Apophysis 67 32 1 0 1.77 1.68 1.88 MF7A1* Troctolite A 66 23 11 0 1.52 1.58 1.23 MF7A2* Troctolite A 65 26 9 0 1.39 1.77 1.21 MF7C1* Troctolite A 65 27 8 0 1.24 1.4 1.12
MF51A1* Troctolite Apophysis 65 32 3 0 1.37 1.24 1.39
MF51A2* Troctolite Apophysis 64 35 1 0 1.25 1.48 1.41
MF96A Troctolite A 60 39 1 0 1.14 1.28 1.17
MF96B Troctolite A 61 37 2 0 1.25 1.18 1.13
MF102A1 Troctolite A 60 36 4 0 1.52 1.42 1.65
MF102B1 Troctolite A 60 36 4 0 1.19 1.34 1.42
MF47A* Wehrlite Apophysis 60 19 21 0 1.16 1.44 1.17
MF47B* Wehrlite Apophysis 59 13 28 0 1.22 1.47 1.14 MF102A2 Troctolite A 55 40 5 0 2.58 1.9 2.06 MF94B Troctolite B 60 38 2 0 1.11 1.09 1.07 MF95A Troctolite B 60 34 6 0 1.21 1.19 1.23 MF72Ga* Troctolite B 59 31 10 0 1.41 1.31 1.37 MF95B Troctolite B 59 36 5 0 1.37 1.28 1.36 MF72I* Troctolite B 57 41 2 0 1.15 1.12 1.21 MF46A* Troctolite B 55 44 1 0 1.04 1.23 1.3 MF73V2 Troctolite B 55 39 6 0 MF94A Troctolite B 55 42 3 0 2.5 1.85 1.85 MF73V1 Troctolite B 50 48 2 0 1.43 1.4 1.44 MF100 Troctolite B 45 51 4 0 1.64 1.88 1.71
MF11A1* Troctolitic gabbro 30 69 1 0 1.08 1.1 1.3
MF99 Troctolitic gabbro 30 66 4 0 1.06 1.18 1.17
MF24* Olivine gabbro 27 59 13 1 1.02 1.13 1.07
MF20* Olivine gabbro 21 60 18 1 1.02 1.04 1.05
MF2B* Olivine gabbro 16 63 19 2
MF2A* Olivine gabbro 15 68 16 1
Plagio = Plagioclase; Cpx = Clinopyroxene; Opx = Orthopyroxene.
Figure 3.5: Textural variability in the Troctolite A. A: Granular olivine matrix embedded in undeformed
poikilitic plagioclase; B: Granular olivine matrix embedded in poikilitic clinopyroxene. Larger olivine crystal shows the occurrence of kink bands, highlighted by the red dashed line. Interstitial plagioclase has been replaced by low-grade alteration products; C: Corroded olivine grain prior to disruption into several smaller crystals. Interstitial plagioclase has been replaced by low-grade alteration products; D: Highly corroded centimetre-size olivine, embedded in poikilitic plagioclase.
Within the Troctolite A, the textural variability includes the occurrence of dunitic aggregates (olivine > 90 vol%). In such dunitic areas (Fig. 3.6a,b,c), interstitial plagioclase surrounds millimetre-size to centimetre-size zones free of interstitial minerals, whereas in plagioclase-rich poikilitic samples (Fig. 3.6d), single olivines are entirely embedded in poikilitic plagioclase ± clinopyroxene. Clinopyroxene, orthopyroxene, and amphibole are found as thin “vermicular” crystals at the contact between olivine and poikilitic minerals and have been previously interpreted as post-cumulus crystallization of trapped melts (progressively evolving during closed system crystallization; Borghini & Rampone, 2007; Borghini et al., 2007). Spinels are found in the matrix both associated to olivine as millimetre-size corroded grains (Fig. 3.6a,d), and associated to poikilitic plagioclase and clinopyroxene, as subhedral to euhedral smaller grains (~100-200 µm, Fig. 3.6d). Troctolite apophyses (part of the Troctolite A) are very rich in coarse deformed corroded grains of olivine (Fig. 3.6b), and undeformed fine-grained olivines are rare.
Figure 3.6: EBSD phase (left column) and misorientation (right column) maps showing the textural
variation of the olivine matrix within the troctolitic body. A: Troctolite A MF96B; B: Troctolite Apophysis MF51A1; C: Troctolite A MF7A1; D: Troctolite A disaggregated MF102B1; E: Troctolite B MF101A. White areas are non-indexed pixels, corresponding to altered areas.
dunitic pods, Table 3.1). Moreover, troctolite B shows an extreme olivine textural variation, from millimetre-size euhedral crystals (Fig. 3.7a) to centimetre-size hopper (Fig. 3.7b), to decimetre-size dendritic and skeletal olivine (Figs. 3.3b,c,d,e, 3.7c,d). This textural variability leads to the formation of a layering in places (Fig. 3.4), but all olivine morphologies can also be found together at the centimetre-scale (Fig. 3.6e).
Figure 3.7: Textural variability observed into the Troctolite B pseudo-tabular bodies. A: Fine-grained
granular undeformed olivines. Plagioclase is completely replaced by low-grade alteration products, with the presence of a chlorite rim around the granular grains of olivine; B: Partially corroded coarse hopper crystal of olivine, associated to poikilitic plagioclase, here completely replaced by low-grade alteration products, and interstitial clinopyroxene; C: Coarse skeletal olivine showing the inner “branches” of olivine, associated to interstitial plagioclase and clinopyroxene; D: Single coarse skeletal olivine associated to interstitial plagioclase, partially replaced by low-grade alteration products.
The Gabbroic lenses, dykes and dykelets are mostly constituted by olivine gabbros and minor troctolites, displaying hypidiomorphic texture and fine- to coarse-grained olivine size. Subhedral plagioclase is the main rock-forming mineral (from 59 to 69 vol% modal content of plagioclase, Table 3.1). Clinopyroxene is mostly found as large anhedral crystals including pre-existing euhedral plagioclase ± olivine. Olivines (from 15 to 30 vol% modal olivine, Table 3.1) are found both as euhedral grains included in plagioclase ± clinopyroxene, and anhedral interstitial crystals in plagioclase-clinopyroxene-olivine aggregates, indicative of a eutectic crystallization of the melt. These textural features in the gabbroic intrusions are indicative of an olivine – plagioclase – clinopyroxene crystallization sequence.
3.5 Crystallographic preferred orientations
In Spinel and Plagioclase Lherzolites and Troctolite, only olivine Crystallographic Preferred Orientation (CPO) could be quantified, because of large grain size of plagioclase and pyroxenes (Fig. 3.8; Bunge, 1982; Ben Ismail & Mainprice, 1998). In the Olivine Gabbros, fine-grained euhedral plagioclase crystals also allow a representative quantification of the plagioclase CPO (Fig. 3.9).
Olivines in Spinel Lherzolites (ETR2, ETR4A, ETR4B in Table 3.1) are characterized by an axial-[100] CPO, with [100] axis showing the strongest preferred orientation in the foliation plane, parallel to the lineation, [010] axis maximum oriented normal to the foliation plane, and [001] maximum within the foliation plane, normal to the lineation (Fig. 3.8). The J-Index, representative of the fabric strength (Bunge, 1982; Ben Ismail & Mainprice, 1998), ranges from 3.64 to 5.59 in spinel peridotites. Most natural peridotites show J-Index values of olivine CPO between 2 and 20 (Tommasi et al., 2000; Soustelle et al., 2009). This axial-[100] CPO indicates that plastic deformation of olivine crystals was related to dislocation creep with activation of (010)[100] and (001)[100] slip systems, the easiest at high temperature conditions (1100-1200°C) (Ben Ismail & Mainprice, 1998; Tommasi et al., 2000; Karato et al., 2008; Drouin et al., 2010; Higgie & Tommasi, 2012). A joint activation of (010)[100] and (001)[100] slip systems in olivine is a common explanation to account for the CPO in natural peridotites deformed under asthenospheric conditions (Tommasi et al., 2000; Le Roux et al., 2008; Soustelle et al., 2009).
Olivines in Plagioclase Lherzolites (P1A, P1B in Table 3.1) are characterized by a strong axial-[010] CPO (J-Index = 5.5-7), with the strongest axis orientation being [010] normal to the foliation, and a girdle orientation of [100] and [001] within the foliation plane, showing a maximum parallel and normal to the lineation, respectively (Fig. 3.8).
In the Troctolite A with dunitic aggregates (Fig. 3.6a,b,c; MF7A1, MF7A2, MF7C1, MF96A, MF96B in Table 3.1) and in the Dunite pod associated to the Troctolite A (MF104A in Table 3.1, Fig. 3.2c), all samples are characterized by a weak (J-Index = 2.04-3.83) but clear axial-[100] olivine CPO, with strongly oriented axial-[100] and [010] axes within the foliation plane, parallel to the lineation, and normal to the foliation, respectively, and a scatter of the [001] olivine axis orientation (Fig. 3.8). This olivine CPO is similar to the one observed in the Spinel Lherzolites (Fig. 3.8). The Troctolite Apophyses show a range of weak olivine CPOs from axial-[100] to axial-[010] (J-Index = 1.86-2.1), similar to the CPO observed in the Spinel Lherzolites and Plagioclase
Figure 3.8: Modal compositions and olivine Crystallographic Preferred Orientation of Spinel lherzolite,
Plagioclase lherzolite, Troctolite apophysis, Troctolite A, and Troctolite A disaggregated samples. One-point-per-grain equal-area, lower hemisphere stereographic projections. The colour bar is scaled to the minimum concentration of the three crystallographic axes. J-index refers to the fabric strength.
The Troctolite A without dunitic aggregates (Fig. 3.6d; MF21, MF15, MF97, MF102B1 in Table 3.1) shows a very weak to random orientation of the [100] and [010] axes, and increased concentrations of the [001] olivine axis (J-Index = 2.17-3.06).
Figure 3.8: Modal composition, olivine and plagioclase Crystallographic Preferred Orientation of Olivine
gabbro and Troctolite B samples. One-point-per-grain equal-area, lower hemisphere stereographic projections. The colour bar is scaled to the minimum concentration of the three crystallographic axes. J-index refers to the fabric strength.
Olivines in the Olivine Gabbros (MF20II, MF24, MF11A1, MF99 in Table 3.1) show very weak (J-Index = 1.21-1.83) (010)[001] CPO characterized by [010] and [001] showing clear maximum normal and within the foliation plane, respectively (Fig. 3.9). Plagioclases show a weak (J-Index = 1.79-4.60) (010)[100] CPO characterized by a strong orientation of the [010] axis normal to the foliation plane (Fig. 3.9). Benn & Allard (1989) and Jousselin et al. (2012) described similar CPOs of olivine and plagioclase in ophiolitic layered gabbros, as a result of the physical orientation of the rock-forming crystals during magmatic flow.
ones observed in the Olivine Gabbro intrusions (Benn & Allard, 1989, Jousselin et al., 2012). The coarse poikilitic minerals characterizing the samples of Troctolite B do not allow a reliable characterization of the Plagioclase CPO (Fig. 3.9).
3.6 Major elements mineral compositions
Major elements compositions of olivine, clinopyroxene, plagioclase, orthopyroxene and spinel analyzed in the Spinel lherzolites, Plagioclase lherzolites, Troctolites A, Dunite, Troctolites B and Olivine Gabbros are reported in the Appendix 1. Overall our data show consistency with mineral compositions reported in previous geochemical investigations of the Erro-Tobbio peridotites and associated gabbroic rocks (troctolitic body and gabbroic lenses and dykes) (Rampone et al., 1993, 1998, 2004, 2005, 2016; Borghini & Rampone, 2007; Borghini et al., 2007).
Olivines in Spinel lherzolites and Plagioclase lherzolites show rather homogeneous high
Forsterite contents (Fo = 89.5-90.5 and Fo = 89.6-90.3, respectively; Fig. 3.10a) (Table S2). Olivines in Troctolites A and Troctolites B show a slightly larger range of variation at lower Forsterite contents in olivine (Fo = 87.3-89.2). Within the Troctolitic body, the main variations are observed between samples rather than within a single sample (Fig. 3.10a), and no correlation is observed between the different olivine morphologies described in Troctolite A and Troctolite B lithotypes (granular undeformed, hopper, dendritic, and corroded deformed olivine) and Forsterite contents, as previously described by Borghini et al. (2007). Olivines within the Dunite pod associated to the Troctolite A show similar contents of Forsterite = 88.2-89.1 (Fig. 3.10a). It is noteworthy that the Wehrlite Apophysis MF47A (Table 3.1) shows the lowest Forsterite content analyzed in the Troctolite A (Fo = 87.3-87.7). Olivine gabbros show a wide range of variation of Forsterite contents in olivine from primitive compositions in the troctolitic intrusions (up to Fo = 89.2) to more evolved compositions in olivine gabbros (Fo = 81.3) (Fig. 3.10b).
Clinopyroxenes cores in Spinel Lherzolites show high Mg-values (Mg# = 90.0-91.6), high
Cr2O3 = 0.82-1.33 wt% and Al2O3 = 5.2-7.4 wt%, and low TiO2 = 0.30-0.58 wt% (Table S3; Fig.
3.11) contents. Impregnated Plagioclase Lherzolites show similar Mg-value (Mg# = 89.6-91.1) and TiO2 = 0.4-0.53 wt% contents, higher Cr2O3 = 1.02-1.40 wt%, and lower Al2O3 = 2.83-5.27 wt% concentrations. Olivine gabbros exhibit clinopyroxene compositions consistent with olivine gabbros and troctolites from the South-West Indian Ridge (Dick et al., 2002), with a positive correlation between Mg-number (Mg# = 83.5-90.8), Cr (Cr2O3 = 0.18-1.15 wt%), and Al (Al2O3 = 2.4-3.7 wt%), and negative correlation with Ti (TiO2 = 0.42-1.41 wt%) (Fig. 3.11). Clinopyroxenes in Troctolite A (and associated Dunite) and Troctolite B show high Cr (Cr2O3 = 1.17-1.67 wt%) and low Al (Al2O3 = 3.1-5.0 wt%) and Ti contents (TiO2 = 0.12-0.92 wt%) (Fig. 3.11).
Figure 3.12 shows the correlation between the clinopyroxene composition and its microstructural site. As previously documented by Borghini & Rampone (2007), clinopyroxenes in
Troctolite A show progressively decreasing Cr2O3 (Cr2O3 = 0.78-1.67 wt%) and increasing TiO2 (TiO2 = 0.12-1.24 wt%) contents from core to rim to interstitial to vermicular microstructural sites, at constant Mg-value (Mg# = 87.7-91.0).
Figure 3.10: Range of Forsterite content in olivines in A: Olivine Gabbros and B: Spinel Lherzolites,
Plagioclase Lherzolite, Dunite, Troctolites A and Troctolites B. Olivine morphology is divided into Granular undeformed and Corroded deformed within the Troctolite A, and Granular undeformed and Hopper-Dendritic within the Troctolites B.
The Cr2O3,Al2O3 and TiO2 compositional variability in clinopyroxene is well observed in major elements core-rim profiles within single clinopyroxene grains (Fig. 3.13). A progressive decrease in Cr2O3 (from 1.5 to 1.0 wt%) and Al2O3 (from 4 to 3 wt%), coupled with an increase in TiO2 (from 0.4 to 1 wt%) is observed in the profiles, from the inner core towards the contact between clinopyroxene and olivine (Rampone et al., 2005). As documented by Borghini et al. (2007), the strong heterogeneity of Cr2O3, TiO2 and Al2O3 in clinopyroxenes of Troctolites A (Fig.
3.12) is thus related to within-sample variations clearly correlated with microstructural site. Geochemical variations in the profiles are observed from ~200µm to the contact with the olivine (Fig. 3.13).
Figure 3.11: Major elements compositions of clinopyroxene cores in all studied samples, plotted against the Mg-number = Mg/(Mg+Fe). A: Cr2O3 (wt%); B: Al2O3
(wt%); C: TiO2 (wt%). Compositional
fields represent compositions of Olivine gabbros and Troctolites from the South-West Indian Ridge, after Dick et al. (2002), and Olivine-rich troctolites from the Erro-Tobbio, after Borghini & Rampone (2007).
Figure 3.12: Clinopyroxene major elements compositional variability with microstructural site in Troctolites A, plotted against the Mg-number = Mg/(Mg+Fe). A: Cr2O3 (wt%); B:
TiO2 (wt%). Compositional fields
represent compositions of Olivine gabbros and Troctolites from the South-West Indian Ridge, after Dick et al. (2002), and Olivine-rich troctolites from the Erro-Tobbio, after Borghini & Rampone (2007).
In Troctolite A, plagioclases (Table S4) are characterized by low and variable Anorthite contents (An = 52.9-66.8 mol%) (Fig. 3.14), as previously documented by Borghini et al. (2007). The same variability is observed in Troctolites B, with Anorthite contents = 55.1-66.1. Olivine
Gabbros show lower Anorthite = 51.6-62.7 mol%. In all samples of Troctolite A and Troctolite B, a
correlation is observed between the microstructural site and the Anorthite content of the analyzed plagioclase crystal. Interstitial plagioclase and crystal rims systematically show lower Anorthite contents than the plagioclase cores (Fig. 3.14), leading to a variation of Anorthite content up to 10 mol% within a single sample, in both Troctolites A and Troctolites B.
Figure 3.13: Clinopyroxene-Olivine major elements profiles in Troctolite A. Step size is 19µm in the left
profile, and 40µm for the right profile. A: Cr2O3 (wt%); B: TiO2 (wt%); C: Al2O3 (wt%). Total lengths of the
profiles are 456µm (left profile) and 920µm (right profile).
Again, these geochemical variations are well observed in major elements profiles from core to rim of plagioclase crystals, at the contact with olivine. A progressive decrease in Anorthite content (from 66 to 56 mol%), CaO (from 14 to 12 wt%), and Al2O3 (from 31 to 30 wt%) is observed in the profiles towards the rim and the contact with olivine (Fig. 3.15), as was previously suggested by Borghini & Rampone (2007). Therefore, as observed for clinopyroxene, the strong compositional variation reported in single samples of Troctolite A and Troctolite B (up to 10% Anorthite content, Figs. 3.14) is not due to variations between different crystals but to the zonation observed at the scale of a single grain (Fig. 3.15). As documented in the clinopyroxene-olivine profiles, the chemical zoning in plagioclase is observed from ~200µm to the contact with the olivine, whether it is a coarse deformed corroded grain of olivine or a small undeformed granular olivine crystal.
Figure 3.14: Range of Anorthite content in plagioclase in Troctolites A, Troctolites B and Olivine Gabbros.
Distinction has been made between interstitial and rims of plagioclase (white symbols), and cores of coarse crystals (coloured symbols).
In Olivine Gabbros, no systematic zoning is observed in plagioclase, and the analyzed range of Anorthite content is mainly observed between samples (Fig. 3.14), with plagioclase in troctolitic dykes showing higher Anorthite contents (MF11A1, MF99, An = 53.8-62.7 mol%; Table 3.1) than plagioclases forming the olivine gabbro dykes (MF2A, MF24, An = 51.6-54.6 mol%; Table 3.1).
In Spinel Lherzolites, spinels (Table S6) exhibit high Mg-number (Mg# = 66.9-72.8 mol%), low Cr-number (Cr# = 14.2-18.6 mol%), and very low TiO2 (0.02-0.16 wt%) (Fig. 3.16), consistent with the spinel major elements compositions reported from plagioclase-free peridotites from the South-West Indian Ridge (Seyler et al., 2003). In Olivine Gabbros, spinels show low Mg-number (Mg# = 25.2-36.1 mol%), and high Cr-number (Cr# = 63.6-69.0 mol%) and TiO2 (1.22-1.49 wt%) (Fig. 3.16).
In Troctolites A, Dunites and Troctolites B, spinels show intermediate compositions of Mg-numbers (Mg# = 19.2-55.6) and Cr-Mg-numbers (Cr# = 40.5-64.5) between spinel compositions in the
Spinel lherzolites and the Olivine gabbros, and a negative correlation is observed between the
Mg-number and the Cr-Mg-number (Fig. 3.16), consistent with spinel compositions in Troctolites from the Mid-Atlantic Ridge (Miller et al., 2009). Some spinels in Troctolites A, Troctolites B, and most of
Troctolites B is consistent with the trend reported for spinels analyzed in troctolites from the
Mid-Atlantic Ridge (Fig. 3.16; Miller et al., 2009).
Figure 3.15: Plagioclase-Olivine major elements
profile in Troctolite A. Step size is 54µm. A: Anorthite content (mol%); B: CaO (wt%); C: Al2O3 (wt%). Total length of the profile is 864µm.
Figure 3.16: Spinel TiO2 and Cr-number
(Cr/(Al+Cr)) core major elements compositions plotted against the spinel
Mg-number (Mg/(Mg+Fe)) in Spinel
Lherzolites, Plagioclase Lherzolites, Troctolites A, Dunite, Troctolites B, and Olivine Gabbros. Compositional fields represent compositions of spinels in Plagioclase-free peridotites and Plagioclase peridotites from the South-West Indian Ridge (Seyler et al., 2003; Paquet et al., 2016), and in Troctolites from the Mid-Atlantic Ridge (Miller et al., 2009).
The major elements compositions of olivine, clinopyroxene and plagioclase show correlations which shed light on the magmatic processes (fractional crystallization, melt-rock interaction, subsolidus chemical reequilibration) ruling their formation.
Figure 3.17 shows the Mg-Fe partitioning between olivine and clinopyroxene in all studied lithotypes (Spinel Lherzolite, Plagioclase Lherzolite, Troctolite A, Dunite, Troctolite B, Olivine
Gabbro). Overall, the studied samples show a positive correlation between Forsterite contents in
olivine (from Fo = 81.3 in Olivine Gabbro to Fo = 90.5 in Spinel Lherzolite) and Mg-value in clinopyroxene (from Mg# = 83.5 in Olivine Gabbro to Mg# = 91.6 in Spinel Lherzolite). This correlation is consistent with the Mg-Fe equilibrium lines calculated between olivine and clinopyroxene by Kawasaki & Ita (1994) (Kd ol/cpx (Fe#) = 1.39) and Lissenberg et al. (2008) (Kd ol/cpx (Fe#) = 1.30) (Fig. 3.17). Couples of olivine and clinopyroxene cores in Troctolites A, Dunite, and Troctolites B show compositions (Fo = 87.3-89.2 mol%, Mg# = 87.7-91 mol%) that are intermediate between the Mg-rich couples analyzed in Spinel and Plagioclase Lherzolites (Fo = 89.5-90.5 mol%, Mg# = 89.6-91.6 mol%) and the most evolved compositions in Olivine Gabbros (Fo = 81.3-89.2 mol%, Mg# = 83.5-90.8 mol%).
Figure 3.17: Olivine – Clinopyroxene cores Mg# correlation in the studied samples, compared to theoretical
Fe-Mg equilibrium between olivine and clinopyroxenes, after Lissenberg et al., 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.
Figure 3.18: Olivine Forsterite content (mol%) against clinopyroxene Cr2O3 (wt%) core compositions in the
studied samples. The compositional field represents the compositions of clinopyroxene in olivine gabbros from the Southwest Indian Ridge (Dick et al., 2002)
Olivine-clinopyroxene core couples show significantly enriched clinopyroxene Cr2O3 (up to 1.75 wt%) compositions at high Forsterite contents in olivine in the Troctolite A, Dunite and
Troctolite B, in respect to Cr2O3 concentrations analyzed in the Spinel and Plagioclase Lherzolites (Cr2O3 = 0.82-1.46 wt%) (Fig. 3.18). A trend of positive correlation between Forsterite contents in olivine and Cr2O3 concentrations in clinopyroxene is observed in the olivine gabbro serie from the Southwest Indian Ridge (Dick et al., 2002), suggesting covariation of the minerals Fo (olivine) and Cr2O3 (cpx) during fractional crystallization processes (Fig. 3.18).
Figure 3.19: Anorthite content (mol%) in plagioclase cores versus Forsterite content (mol%) in olivine
cores in olivine-plagioclase couples from the studied Troctolites A, Troctolites B and Olivine Gabbros. 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), and Pineto ophiolite (Sanfilippo & Tribuzio, 2013).
Figure 3.19 shows Anorthite and Forsterite contents (mol%) in plagioclase-olivine core couples in Troctolite A, Troctolite B and Olivine Gabbro. Within the troctolitic body, plagioclase-olivine couples show significant variations in Anorthite content of plagioclase cores (An = 58.4-66.8 mol%) at constant Forsterite content in associated olivines (87.3-89.2 mol%). Olivine Gabbros define a trend of evolution characterized by a positive correlation between the Anorthite content in plagioclase cores and the Forsterite content in olivine (from An51.6-Fo81.3 to An62.7-Fo89.2). This
An-Fo trend of variation in the Olivine Gabbros shows a similar slope to the trends defined by olivine gabbros in the oceanic lower crust from the Mid-Cayman Rise (Elthon et al., 1987), 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), and Pineto ophiolite
Clinopyroxene Mg-number (mol%) shows similar correlations with plagioclase Anorthite (mol%) (Fig 3.20), with relatively constant Mg-number in Troctolite A and Troctolite B (Mg# = 87.7-91.0 mol%) at varying Anorthite content (An = 55.1-67.0 mol%). The Olivine Gabbros show a positive correlation between Mg-number in clinopyroxene and Anorthite content in plagioclase (from An51.6-Mg#83.7 to An62.7-Fo90.1). As described above for the Forsterite-Anorthite position
correlation, the slope defined by the Anorthite – Mg-value (cpx) covariation in Olivine Gabbros is consistent with the trends documented in the oceanic gabbroic suites at the East Pacific Rise (Dick & Natland, 1996; Lissenberg et al., 2013), Mid-Atlantic Ridge (Ross & Elthon, 1997;Lissenberg & Dick, 2008; Suhr et al., 2008; Drouin et al., 2009; Miller et al., 2009) and South-West Indian Ridge (Dick et al., 2002), but shifted towards higher Mg-values of clinopyroxene (at given Anorthite, ~10 mol% Mg# higher than clinopyroxenes at the East Pacific Rise, ~15 mol% Mg# higher than the Mid-Atlantic Ridge and ~20 mol% Mg# higher than the South-West Indian Ridge clinopyroxenes; Fig. 3.20).
Figure 3.20: Anorthite content (mol%) in plagioclase cores versus Mg-number (mol%) in clinopyroxene
cores in plagioclase-clinopyroxene couples from the studied Troctolites A, Troctolites B and Olivine Gabbros. Compositional trends represent plagioclase-clinopyroxene couples in olivine gabbros and troctolites from East Pacific Rise (Dick & Natland, 1996; Lissenberg et al., 2013), South-West Indian Ridge (Dick et al., 2002), and Mid-Atlantic Ridge (Ross & Elthon, 1997; Lissenberg & Dick, 2008; Suhr et al., 2008; Drouin et al., 2009; Miller et al., 2009).
3.7 Trace elements mineral compositions
The trace elements compositions and standard deviation of olivine, clinopyroxene, plagioclase and orthopyroxene analyzed in the Spinel Lherzolite, Plagioclase Lherzolite, Troctolite
A, Dunite, Troctolite B and Olivine Gabbro samples are reported in Appendix 1. Overall, the data presented in this study are consistent with mineral compositions reported in previous studies of the Erro-Tobbio Peridotites, Troctolitic body, and Olivine Gabbro intrusions (Rampone et al., 1993, 1998, 2004, 2005, 2016; Borghini & Rampone, 2007; Borghini et al., 2007).
Olivines in Spinel Lherzolites show homogeneous compositions of the moderately
incompatible elements (Ni, Mn, Zn, Co) consistent with the Spinel peridotites reported by De Hoog et al. (2010), characterized by high Ni (2815-3571 ppm), and low Mn (1110-1167 ppm), Zn (29-40 ppm) and Co (141-146 ppm) concentrations (Fig. 3.21). Plagioclase Lherzolites show slightly lower Ni (3006-3340 ppm), similar Mn (1087-1156 ppm) and higher Zn (48-61 ppm) and Co (143-158 ppm) concentrations. These olivine compositions in Plagioclase Lherzolites are consistent with the compositions reported in Alpine-Apennine impregnated peridotite by Sanfilippo et al. (2014) and Rampone et al. (2016) (Fig. 3.21). Olivines in Olivine Gabbros define a wide compositional range, well correlated with their Forsterite content (Fig. 3.21). They exhibit a positive correlation between Forsterite content (81.3-89.2 mol%) and Ni concentrations (973-2356 ppm), and a negative correlation with Mn (1329-2198 ppm), Zn (56-112 ppm) and Co (132-198 ppm) (Fig. 3.21), consistent with a Fractional Crystallization trend (Drouin et al., 2009; Rampone et al., 2016) and with Erro-Tobbio olivine gabbro compositions reported by Rampone et al., 2016. As reported by Rampone et al. (2016), olivines in the troctolitic body (Troctolite A, Dunite, Troctolite B) display Ni (1627-2689 ppm), Mn (1340-1855 ppm), Zn (35-74 ppm) and Co (122-173 ppm) compositions that are intermediate between olivines in the Olivine Gabbros and in the Spinel and Plagioclase
Lherzolites (Fig. 3.21). In Troctolites A and Troctolites B, no systematic compositional variation in moderately incompatible elements is observed, within a single sample, between olivines with different textures (i.e. granular undeformed, corroded deformed, hopper and dendritic crystals).
Overall, the moderately incompatible elements show a good correlation with the Forsterite content in olivine, defining a trend similar to the expected geochemical correlation obtained after a simple process of fractional crystallization (Drouin et al., 2009; De Hoog et al., 2010; Ferrando et al., in revision).
Highly incompatible elements in olivine (C1-normalized Yb, Y, Zr, Ti) do not follow a linear trend of correlation with the Forsterite content between all lithotypes (Fig. 3.22). Olivines in
Spinel Lherzolites show low Y (0.008-0.033 ppm), Zr (0.004 ppm) and Ti (11-26 ppm)
concentrations (Fig. 3.22), low HREE abundances (YbN = 0.015-0.13 times C1) and high HREE fractionation (DyN/YbN = 0.119-0.245) (Fig. 3.23a,b), comparable to olivines from the Gakkel Ridge lherzolites reported by D’Errico et al. (2016). Olivines in Plagioclase Lherzolites show slightly higher Y (0.047-0.08 ppm), Zr (0.016), and Ti (27-61 ppm) concentrations (Fig. 3.22), as well as higher HREE abundances (YbN = 0.169-0.303 times C1) and lower HREE fractionation (DyN/YbN = 0.062-0.067) (Fig. 3.23a,b).
Figure 3.21: Variation of the moderately incompatible
elements, as a function of Forsterite content in olivine forming the Spinel lherzolite, Plagioclase Lherzolite, Troctolite A, Dunite, Troctolite B and Olivine Gabbro. A: Mn; B: Ni; C: Zn; D: Co. Compositional field and reference data points represent olivine compositions in Spinel peridotite (De Hoog et al., 2010), Olivine Gabbros, Troctolites and Peridotite from the Erro-Tobbio peridotites (Rampone et al., 2016), Internal Ligurides and Lanzo ophiolites (Sanfilippo et al., 2014).
Figure 3.22: Variation of highly incompatible elements,
as a function of Forsterite content in olivine forming the Spinel Lherzolite, Plagioclase Lherzolite, Troctolite A, Dunite, Troctolite B and Olivine Gabbro. A: Yb normalized to C1 chondrite; B: Y; C: Zr; D: Ti. Reference data points represent olivine compositions in Olivine Gabbros, Troctolites and Peridotites from the Erro-Tobbio peridotites (Rampone et al., 2016), Internal Ligurides and Lanzo ophiolites (Sanfilippo et al., 2014). C1-chondrite normalization values after Sun & McDonough (1989).
Figure 3.23: Average REE and trace elements compositions of olivine. A-B: C1-chondrite and Primitive
Mantle normalized REE and trace elements compositions of olivine in Spinel Lherzolite, Plagioclase Lherzolite, and Olivine Gabbro; C-D: C1-chondrite and Primitive Mantle normalized REE and trace elements compositions of olivine in Troctolite A and associated Dunite. In Troctolite A, empty symbols represent Granular grains, and coloured symbols represent Corroded grains; E-F: C1-chondrite and Primitive Mantle normalized REE and trace elements compositions of olivine in Troctolite B. Empty symbols represent Granular grains, and coloured symbols represent Hopper and Dendritic grains. Compositional fields represent the olivine composition in Gakkel Ridge spinel lherzolite (D’Errico et al., 2016), Mid-Atlantic Ridge olivine-rich troctolites (Drouin et al., 2009) and Erro-Tobbio troctolites (Rampone et al., 2016). C1-chondrite and Primitive Mantle normalization values after Sun & McDonough (1989).
Olivines in the Olivine Gabbros show moderate incompatible elements variations over the large range of Forsterite content observed (81.3-89.2 mol%). They show low Y (0.018-0.128 ppm), Zr (0.001-0.229 ppm), Ti (43-172 ppm) (Fig. 3.22), and HREE abundances (YbN = 0.09-0.73 times C1) and HREE fractionation (DyN/YbN = 0.026-0.087) consistent with olivine REE compositions reported in oceanic MORB settings (Drouin et al., 2009) and from Rampone et al. (2016) in Erro-Tobbio olivine gabbros (Fig. 3.23a,b). In Troctolites A and associated Dunites, olivines show
significant trace elements variations at constant Forsterite contents (Fo = 87.3-89.2 mol%), and a correlation between the olivine texture and composition is observed, with the strongest trace elements enrichments in Y (0.019-0.329 ppm), Zr (0.001-0.694 ppm) and Ti (51-349 ppm) being observed in corroded deformed olivines (Fig. 3.22). These corroded deformed olivines also show strong enrichments in HREE abundances (YbN = 0.23-0.91 times C1) at constant HREE fractionation (DyN/YbN = 0.026-0.073) (Fig. 3.23c,d). Granular undeformed olivines in Troctolite A also show variable Y (0.023-0.196 ppm), Zr (0.001-0.381), Ti (63-224 ppm) concentrations (Fig. 3.22), although defining weaker enrichments. They also have rather heterogeneous HREE abundances (YbN = 0.23-0.67 times C1) at constant HREE fractionation (DyN/YbN = 0.017-0.049).
Olivines in the Troctolite B show similar trace elements enrichments, correlated with the olivine textures, with hopper and dendritic olivines showing stronger enrichments in Y (0.029-0.42 ppm), Zr (0.042-1.054 ppm), Ti (67-256 ppm) concentrations (Fig. 3.22), and HREE abundances (YbN = 0.56-0.73 times C1) at constant HREE fractionation (DyN/YbN = 0.031-0.053). Granular undeformed olivines show homogeneous trace elements contents, with Y (0.038-0.092 ppm), Zr (0.033-0.154 ppm), Ti (78-177 ppm) concentrations, and HREE abundances (YbN = 0.33-0.57 times C1; Fig. 3.23e,f) similar to olivines analyzed in the most primitive Olivine Gabbro intrusion (Figs. 3.22, 3.23a,b). Overall, olivines in the troctolitic body (Troctolite A, Dunite, Troctolite B) show HREE compositions consistent with MORB-type olivine compositions described in Mid-Atlantic Ridge olivine-rich troctolites (Drouin et al., 2016) and the olivine compositions in olivine-rich troctolites and troctolites from the Alpine-Apennine ophiolites (Sanfilippo et al., 2014; Rampone et al., 2016) (Figs. 3.22; 3.23c,d,e,f).
Clinopyroxenes in the Spinel Lherzolites show strong LREE depletion (CeN/YbN = 0.065-0.068), coupled to flat M-HREE patterns (YbN = 10.08-11.32 times C1) (Fig. 3.24a,b), at higher abundances than depleted spinel lherzolites from the Gakkel Ridge (D’Errico et al., 2016) and comparable to clinopyroxene compositions that were reported in the spinel peridotites from the Erro-Tobbio ophiolite (Rampone et al., 2005). In Plagioclase Lherzolites, clinopyroxenes show higher REE absolute concentrations (YbN = 11.94-14.8 times C1) at constant LREE fractionation (CeN/YbN = 0.092-0.136), and the development of an Eu negative anomaly (Fig. 3.24a,b), similar to what was previously described in the Erro-Tobbio impregnated plagioclase peridotites by Borghini et al. (2007). In Olivine Gabbros, clinopyroxenes show enriched M-HREE concentrations (YbN = 14.65-28.85 times C1), Eu negative anomaly and modest LREE depletion (CeN/YbN = 0.163-0.311) (Fig. 3.24a,b), indicative of a MORB-like signature of parental melts, as described in previous studies (Borghini & Rampone, 2007; Borghini et al., 2007; Piccardo & Vissers, 2007; Rampone et al., 2016). Clinopyroxene cores in Troctolite A and associated Dunite show homogeneous trace elements absolute concentration, constant LREE fractionation, and no Eu anomaly (YbN = 6.66 times C1; CeN/YbN = 0.228-0.342) (Fig. 3.24c,d), consistent with the clinopyroxene reported in olivine-rich troctolites from the Mid-Atlantic Ridge (Drouin et al., 2009). In the Troctolite B, clinopyroxene cores show compositions characterized by relatively high REE abundances (YbN = 13.6-19.35 times C1; CeN/YbN = 0.263-0.329) and negative Eu anomalies (Fig. 3.24e,f).
Figure 3.24: Average REE and trace elements compositions of clinopyroxene. A-B: C1-chondrite and
Primitive Mantle normalized REE and trace elements compositions of clinopyroxene in Spinel Lherzolite, Plagioclase Lherzolite, and Olivine Gabbro; C-D: C1-chondrite and Primitive Mantle normalized REE and trace elements compositions of clinopyroxene in Troctolite A and associated Dunite. In Troctolite A, empty symbols represent Interstitial grains and rims, and coloured symbols represent Coarse grains and cores; E-F: C1-chondrite and Primitive Mantle normalized REE and trace elements compositions of clinopyroxene in Troctolite B. Empty symbols represent Interstitial grains and rims, and coloured symbols represent Coarse grains and cores. Compositional fields represent the clinopyroxene composition in Gakkel Ridge spinel lherzolite (D’Errico et al., 2016), Erro-Tobbio spinel and plagioclase peridotites (Rampone et al., 2005; Borghini et al., 2007), Mid-Atlantic Ridge olivine-rich troctolites (Drouin et al., 2009) and Erro-Tobbio troctolites (Borghini et al., 2007; Rampone et al., 2016). C1-chondrite and Primitive Mantle normalization values after Sun & McDonough (1989).
As previously described by Borghini et al. (2007), in both Troctolite A and Troctolite B, significant REE enrichments have been observed from core to rim to interstitial to vermicular microstructural sites in clinopyroxene crystals. Figure 3.24c,d,e,f shows the strong enrichment analyzed in the outermost rims of clinopyroxene, up to YbN = 37.8 times C1, the development of a
significant Eu negative anomaly and the disappearance of the Zr-Hf negative anomaly usually observed in clinopyroxenes. This correlation between geochemical enrichments and microstructural site is well observed in trace elements profiles from core to rim of a clinopyroxene crystal (Fig. 3.25a), where we clearly see the progressive REE enrichment at constant LREE fractionation, development of a negative Eu anomaly, and decrease in the Zr-Hf negative anomaly.
Figure 3.25: C1-chondrite normalized REE and Primitive Mantle normalized trace elements compositions
along a profile from core (blue) to rim (yellow) of A: clinopyroxene; and B: plagioclase. The analyses points are located on the reflected light photograph.
Plagioclase cores in Olivine Gabbros show REE compositions (SmN = 0.48-0.85 times C1; CeN/SmN = 1.90-3.21; Fig. 3.26a,b) consistent with the MORB-type plagioclase compositions analyzed in olivine-rich troctolites at the Mid-Atlantic Ridge (Drouin et al., 2009), and with compositions of plagioclase previously reported in the Erro-Tobbio troctolitic body (Borghini et al., 2007; Rampone et al., 2016). In the Troctolite A and Troctolite B, plagioclase cores show slightly more enriched REE abundances (SmN = 0.5-1.42 times C1) than plagioclases analyzed in the Olivine Gabbros and in Mid-Atlantic Ridge olivine-rich troctolites (Drouin et al., 2009), at constant
As described above for clinopyroxene, plagioclase also displays core-rim trace elements variations, with progressive REE enrichment towards the crystal outer rim, at the contact with olivine. Figure 3.25b shows a trace elements profile from core to rim, showing the progressive REE enrichment at constant LREE fractionation, during melt-rock interactions processes.
Figure 3.26: Average REE and trace elements compositions of plagioclase and orthopyroxene. A-B:
C1-chondrite and Primitive Mantle normalized REE and trace elements compositions of plagioclase in Olivine Gabbro; C-D: C1-chondrite and Primitive Mantle normalized REE and trace elements compositions of plagioclase in Troctolite A and Troctolite B; E-F: C1-chondrite and Primitive Mantle normalized REE and trace elements compositions of orthopyroxene in Spinel Lherzolite, Plagioclase Lherzolite and Olivine Gabbro. Compositional fields represent the plagioclase composition in Mid-Atlantic Ridge olivine-rich troctolites (Drouin et al., 2009) and Erro-Tobbio troctolites (Borghini et al., 2007; Rampone et al., 2016), and orthopyroxene composition in Gakkel Ridge spinel lherzolite (D’Errico et al., 2016), and Erro-Tobbio plagioclase peridotites. C1-chondrite and Primitive Mantle normalization values after Sun & McDonough (1989).
Orthopyroxene is only found in Spinel Lherzolites, Plagioclase Lherzolites and the most
evolved Olivine Gabbros. In Spinel Lherzolites, as for clinopyroxene, the analyzed orthopyroxene REE absolute concentrations (YbN = 2.48-2.62 times C1; CeN/YbN = 0.006; Fig. 3.26e,f) are slightly higher than those of orthopyroxenes in depleted lherzolites from the Gakkel Ridge (D’Errico et al., 2016). Orthopyroxenes analyzed in Plagioclase Lherzolites (YbN = 3.75-4.85 times C1; CeN/YbN = 0.005) are comparable to the compositions previously reported in the Erro-Tobbio impregnated plagioclase peridotites (Rampone et al., 2016). Orthopyroxenes analyzed in the Olivine
Gabbro show higher REE absolute concentrations (YbN = 6.54 times C1) and CeN/YbN fractionation = 0.017, indicative of the MORB-type affinity of the parental melt (Fig. 3.26e,f).
3.8 Discussion
3.8.1. Replacive origin of the Troctolite A
As reported in previous studies, the Erro-Tobbio troctolitic body is in irregular contact with the host impregnated Plagioclase lherzolites (Borghini & Rampone, 2007; Borghini et al., 2007; Rampone et al., 2016). The host Troctolite A crosscuts the pyroxenite banding associated to the mantle peridotites (Fig. 3.4) and develops wehrlite and troctolite apophyses into the mantle
Plagioclase Lherzolites (Fig. 3.2a). The Troctolite A shows a strong textural complexity with the occurrence of two distinct types of olivines within single samples (Fig. 3.6a,b,c,d), i.e. millimetre-size undeformed granular olivine grains (Fig. 3.5a,b) and coarse (up to centimetre-size), deformed and corroded crystals (Fig. 3.5c,d). As inferred in oceanic settings during formation of olivine-rich troctolites (Suhr et al., 2008; Drouin et al., 2010), Rampone et al. (2016) interpreted the textural complexity of the Erro-Tobbio troctolites as the result of melt-rock interactions leading to the dissolution of the olivine matrix and crystallization of interstitial plagioclase. Although they were not able to distinguish two olivine generations in a specific troctolite sample, they inferred that the millimetre-size undeformed granular olivine grains could represent a second generation of “olivine
2”, whether of magmatic origin or representing disrupted coarse olivine grains. Detailed EBSD
analyses (size, shape, misorientation; Fig. 3.6) allow us to interpret the coarse deformed and corroded olivine as the pre-existing, possibly mantle relitic “olivine 1”. The occurrence of coarse corroded grains almost disrupted into several granular olivines (Fig. 3.5c) suggests that most of the small undeformed olivine grains are formed after extensive corrosion and disruption of the coarse relitic pre-existing olivines. This process of textural evolution of the olivine matrix during progressive melt-rock interaction and replacive formation of hybrid olivine-rich troctolites has been previously inferred in oceanic settings (Suhr et al., 2008; Drouin et al., 2010) and recently demonstrated in ophiolitic settings at the Mt.Maggiore peridotitic body (Basch et al., in revision, see Chapter 4).
At the scale of the sample, the Troctolite A is also characterized by variations in the texture of the olivine matrix (taken as a whole, olivines 1 + olivines 2), between samples characterized by plagioclase-free dunitic aggregates surrounded by interstitial phases (Fig. 3.6a,b,c), and disaggregated samples where single olivines are completely embedded in poikilitic plagioclase (Fig. 3.6d). This textural variability is well correlated with a change in the Crystallographic Preferred Orientations of olivine. The olivine matrix of Troctolite A characterized by plagioclase-free dunitic aggregates shows an axial-[100] fabric (Fig. 3.8), similar to the Spinel lherzolites and Dunite pods. This axial-[100] CPO is typically reported in natural peridotites deformed under asthenospheric
1998; Tommasi et al., 2000; Karato et al., 2008; Drouin et al., 2010; Higgie & Tommasi, 2012). The samples characterized by a disaggregated olivine matrix, embedded in poikilitic plagioclase, show random orientation of [100] and [010] olivine axes, and a stronger concentration of the [001] axis (Fig. 3.8). Randoming of the olivine CPO have been previously reported in zones of melt accumulation in the Oman Moho Transition Zone (Ceuleneer & Rabinowicz, 1992; Boudier & Nicolas, 1995; Jousselin et al., 1998; Dijkstra et al., 2002; Higgie & Tommasi, 2012) and during the replacive formation of olivine-rich troctolites at the Atlantis Massif (Drouin et al., 2010). It has been interpreted as a loss of cohesion of the solid matrix during impregnation at high melt/rock ratios (20-40% melt fraction; Rosenberg & Handy, 2005). Melt-rock interaction microstructures, indicating the corrosion of the pre-existing olivine matrix, together with the preservation of dunitic pods within the host Troctolite A (Figs. 3.2d, 3.4) and the correlation between the observed texture of the olivine matrix and its CPO (Fig. 3.8), suggest a replacive formation of Troctolites A. We infer that they formed from a mantle Dunite protolith (itself preserving the mantle precursor axial-[100] CPO), after an olivine-undersaturated melt reactive percolation at variable melt-rock ratios (Fig. 3.27). The disaggregation of the olivine matrix associated to the loss of the olivine CPO are indicative of high instantaneous melt-rock ratios (>20-40%; Rosenberg & Handy, 2005), whereas the samples preserving the mantle olivine CPO indicate a reactive percolation at low melt-rock ratios (Fig. 3.27). Texture and CPO analyses thus indicate that Troctolites A are likely the hybrid product of reactive percolation and impregnation of a pristine olivine matrix by melts crystallizing plagioclase and minor clinopyroxene.
Peculiar geochemical compositional trends of the rock-forming minerals, not consistent with a simple fractional crystallization process, support the replacive origin of the Troctolites A. Despite strong variations in olivine modal compositions (from 55 vol% in troctolites to 97 vol% in dunitic pods), olivines and clinopyroxenes in the Dunite and the Troctolite A show a narrow range of composition (Fo = 88.2-89.1 mol%; Fig. 3.10; Mg# = 89-91 mol%; Fig. 3.20). These constant compositions of the mafic minerals are coupled with significant within-sample variations in plagioclase Anorthite contents (An = 52.9-66.8 mol%; Figs. 3.19, 3.20), and therefore do not follow compositional trends of fractional crystallization defined by the oceanic gabbroic sequences (East Pacific Rise, Dick & Natland, 1996; Lissenberg et al., 2013; Mid-Cayman Rise, Elthon, 1987; 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). Rather, these peculiar compositional trends (Figs. 3.19, 3.20) are indicative of the buffering of the melt Mg-value by an olivine-dissolving reactive porous flow (Mg# melt ≈ 68-70; Collier & Kelemen, 2010; Sanfilippo et
al., 2016).
To better constrain and quantify the role of reactive crystallization in the formation of these peculiar An-Fo and An-Mg# compositional trends (Figs. 3.19, 3.20), we performed a geochemical modelling of equilibrium crystallization, assuming variable assimilated mass of olivine Fo89 (olivine composition in the dunite), using the pMELTS thermodynamic program (Ghiorso et al., 2002; Collier & Kelemen, 2010).
Figure 3.27: Sketch of the evolution of the olivine textures and associated CPOs during progressive
olivine-dissolving, plagioclase-crystallizing melt-rock interaction and replacive formation of the Troctolite A. A: Coarse-grained dunite protolith showing an axial-[100] olivine CPO; B: Troctolite A impregnated at low melt-rock ratios, and thus preserving dunitic aggregates and axial-[100] olivine CPO; C: Disaggregated troctolite A impregnated at high instantaneous melt-rock ratios. The arrows within small olivine grains represent the loss of cohesion of the solid matrix leading to the randoming of the olivine CPO.
As an initial melt, we calculated the average of 66 MORB compositions from the compilation of Gale et al. (2013). In order to reproduce the low-Anorthite compositions in plagioclase (low CaO/Na2O) coupled to high Fo and Mg-value in olivine and clinopyroxene, respectively (high melt Mg-value), we arbitrarily selected in the database primitive MORBs characterized by an Mg#>67, CaO/Al2O3>0.6, and CaO/Na2O<5,. The resulting MORB composition is given in Table 3.2.
Ini.Melt 49.83 1.13 7.95 9.25 17.08 11.81 2.75 0.20 0.00 0.00 100 0.68 0.69 4.33
Figure 3.28: pMELTs numerical simulations of the major elements compositions of olivine (Forsterite
content), plagioclase (Anorthite content) and clinopyroxene (Mg-value) during equilibrium crystallization and reactive crystallization of a calculated primitive MORB, after Gale et al. (2013). Varying assimilation rates of a dunite (100% olivine) from 0.5g/°C to 3g/°C of cooling, are modelled, and compared to the core compositions of olivine-plagioclase, and clinopyroxene-plagioclase couples analyzed in the Troctolite A and Troctolite B. Compositional trends of oceanic gabbroic suites are plotted as comparison for East Pacific Rise (Dick & Natland, 1996; Lissenberg et al., 2013), 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) and Mid-Cayman Rise (Elthon et al., 1987).
We modelled isobaric (P = 3kbar) reactive equilibrium crystallization of the average MORB melt, cooling at steps of 5°C after assimilation of fixed mass of olivine (0g, 0.5g, 1g, 3g) per 1°C of cooling (Fig. 3.28). Similar models of reactive crystallization using the pMELTS thermodynamic program (Ghiorso et al., 2002) have been previously performed by Collier & Kelemen (2010) and Sanfilippo et al. (2016), involving the assimilation of mantle lherzolites. Based on field and microstructural observations, we decided to assimilate 100% olivine Fo89, similar to the Forsterite composition of olivine in the Dunite pods. Figure 3.28 shows the crystal line of descent of olivine, plagioclase and clinopyroxene after equilibrium reactive crystallization involving 0g/°C, 0.5g/°C, 1g/°C and 3g/°C of olivine Fo89 assimilation. Before the evidence of low melt/rock ratios (preservation of clear CPO patterns; Fig. 3.8) involved in the closed system reactive crystallization of the troctolitic body (see later discussion 3.9.3. Closed-system crystallization of the troctolitic body), we preferred an Equilibrium crystallization to the common Fractional crystallization. The crystallization order is [olivine-plagioclase-clinopyroxene], as expected from the crystallization of a MORB melt at low pressures (<7 kbar; Bender et al., 1978; Husen et al., 2016). With an increase in assimilation rates (from 0g/°C to 3g/°C of cooling), the Forsterite content in olivine and the Mg-value in clinopyroxene are progressively buffered by the composition of the assimilated olivine (Fo = 89 mol%), while the Anorthite contents in plagioclase evolve freely towards lower Anorthitic
