Geochemistry, Geophysics, Geosystems

Melt transport and deformation history in a nonvolcanic ophiolitic section, northern Apennines, Italy: Implications for crustal accretion at slow spreading settings

Authors


Abstract

[1] Field observations and petrological and geochemical data are used to constrain a conceptual model for the formation of a gabbro-peridotite section from Ligurian ophiolites (Italy). The studied section is attributed to an intraoceanic domain of the Jurassic Ligurian-Piedmontese basin and is characterized by the lack of a basalt layer, similar to nonvolcanic segments from (ultra)slow spreading ridges. The proposed model shows a “hot” lithospheric evolution in which melt transport in the mantle under spinel to plagioclase facies conditions occurred mostly in the form of grain-scale porous flow. We recognize a series of melt/peridotite interaction events, either diffuse or channeled, which modified the composition of the moderately depleted precursor mantle. In particular, localized infiltrations of MORB-type melts gave rise to formation of spinel websterite layers close to the lithosphere-asthenosphere boundary. The peridotite-websterite association was involved in a spinel facies deformation attributed to emplacement of asthenospheric material at the base of the lithosphere. The “hot” lithospheric evolution is followed by an evolution characterized by melt transport through fractures, which started with crystallization of melt into troctolite to olivine gabbro dikes. Both mantle structures and gabbroic dikes are locally crosscut by gabbroic sills. As the mantle section cooled significantly, the dip of the melt migration structures evolved from subvertical to subhorizontal. The growth of a gabbroic pluton (up to ∼400 m thick) that is intruded into the mantle sequence is attributed to accretion of gabbroic sills. The tectonomagmatic history recorded by the gabbroic pluton after its solidification is characterized by ductile shearing developed from near-solidus to amphibolite facies conditions.

1. Introduction

[2] Ophiolites offer the opportunity for 3-D investigations of ancient oceanic lithosphere. To unravel the architecture of the oceanic lithosphere and to constrain the tectonomagmatic processes through which the oceanic lithosphere is generated, the study of ophiolites provides complementary information to that acquired by ocean floor drilling and dredging. For instance, the processes of melt extraction, melt-peridotite reactions and building of gabbroic crust inferred from studies of the Oman ophiolite are considered to be relevant to the formation of modern fast spreading ridges [e.g., Nicolas and Boudier, 1995; Kelemen et al., 1995, 1997]. However, models of the creation of oceanic lithosphere at fast spreading ridges are not appropriate for (ultra)slow spreading ridges [e.g., Cannat, 1993; Cannat et al., 1997; Kelemen et al., 2007] and ophiolites representing fossil analogs of (ultra)slow spreading ridges have not been clearly identified yet.

[3] Most Jurassic ophiolites from the Alpine-Apennine belt are considered to be either remnants of an embryonic ocean, or analogs of an ocean-continent transition developed in a magma poor system [e.g., Marroni et al., 1998; Manatschal and Müntener, 2009; Piccardo and Guarnieri, 2009]. However, a few ophiolitic sequences from the Alpine-Apennine belt, considered to represent oceanward paleogeographic domains of the Jurassic basin, bear lithostratigraphic, structural and compositional similarities to modern (ultra)slow spreading ridges. These successions are essentially represented by the Chenaillet ophiolite from the western Alps [Lagabrielle and Cannat, 1990; Caby, 1995; Manatschal et al., 2011], the M. Maggiore mantle section from Corsica [Rampone et al., 2008, 2009] and by the Internal Ligurian ophiolites from the northern Apennines [Tribuzio et al., 2000; Principi et al., 2004].

[4] In this study, we wish to understand the tectonomagmatic evolution experienced by a lesser known gabbro-peridotite association from the Internal Ligurian ophiolites, exposed near the Scogna and Rocchetta Vara localities. The Scogna-Rocchetta Vara ophiolite is characterized by the absence of a basalt layer, similar to nonvolcanic segments from (ultra)slow spreading ridges [e.g., Kelemen et al., 2007], and by preservation of primary relationships between the gabbro-peridotite basement and overlying sedimentary cover [Barret and Friedrichsen, 1989; Principi et al., 2004]. On the basis of new field data, analysis of microstructures and major and trace element microanalyses of minerals, we propose a composite conceptual model for the formation and evolution of the studied ophiolite. In particular, we show (1) the compositional and structural modifications that occurred in the mantle section in conjunction with its exhumation from spinel to plagioclase facies conditions, (2) the formation of oceanic gabbroic crust, and (3) the tectonomagmatic evolution leading to exposure of the gabbro-peridotite association at the seafloor. In addition, we emphasize the similarities between the studied ophiolite and the melt-poor sections from modern (ultra)slow spreading settings.

2. Geological Setting

[5] Jurassic ophiolitic sequences representing different paleogeographic domains of the Ligurian-Piedmontese basin are exposed in the northern Apennines. The ophiolites from the External Ligurian units contain mantle sequences of subcontinental lithospheric origin [Beccaluva et al., 1984; Rampone et al., 1995; Montanini et al., 2006] and are locally associated with continental crust rocks [Molli, 1996; Montanini and Tribuzio, 2001; Renna and Tribuzio, 2009]. The association of ophiolites and continental crust material from the External Ligurian units is considered to represent a fossil ocean-continent transition such as the magma-poor continental margin of western Iberia [Marroni et al., 1998; Tribuzio et al., 2004]. The ophiolites from the Internal Ligurian units do not show relationships with continental material and are attributed to a distal domain of the basin. They are characterized by a gabbro-peridotite basement discontinuously covered by a volcanosedimentary sequence that commonly exhibits interlayering among MORB-type lava flows, sedimentary breccias and Middle-Upper Jurassic radiolarian cherts [Cortesogno et al., 1987; Principi et al., 2004].

[6] Three major ophiolite bodies are present in the Internal Ligurian units (Figure 1a), the Bracco-Levanto, Val Graveglia-Bargonasco and Scogna-Rocchetta Vara. The Bracco-Levanto ophiolite provides evidence for the occurrence of a morphological high in the Ligurian-Piedmontese basin [Cortesogno et al., 1987]. This is shown by a sequence in which the Jurassic volcanosedimentary cover is absent, apart from the local occurrence of thin breccia levels, and the gabbro-peridotite association is commonly directly overlain by Cretaceous shaly pelagites [see also Principi et al., 2004]. This paleomorphological high is made up of a gabbroic pluton intruded into mantle peridotites, as observed for oceanic core complexes from the Mid Atlantic Ridge, such as the Atlantis Massif [e.g., Cann et al., 1997] and the Kane Megamullion [Dick et al., 2008]. In the Val Graveglia-Bargonasco ophiolite, the gabbro-peridotite basement is covered by a volcanosedimentary sequence characterized by an overall continuous basalt flow layer [Principi et al., 2004]. Conversely, in Scogna-Rocchetta Vara ophiolite, the basalt layer is absent [see also Barret and Friedrichsen, 1989].

Figure 1.

(a) Location of ophiolites from the northern Apennines. The major ophiolitic bodies of the Internal Ligurian ophiolites are indicated as follows: SRV, Scogna-Rocchetta Vara; VGB, Val Graveglia-Bargonasco; BL, Bracco-Levanto. (b) Geological sketch map of Scogna-Rocchetta Vara ophiolite and simplified tectonic scheme of the northern Apennines. IL, Internal Ligurian units (Middle-Upper Jurassic ophiolitic sequences); EL, External Ligurian units (Upper Cretaceous turbitites and breccias containing ophiolitic blocks); GO, Gottero Unit (Maastrichtian-early Paleocene sandstones); TU, Tuscany nappe (late Oligocene–early Miocene turbiditic sandstones). (c) Detail of the geological sketch map of Figure 1b and cross section of Rocchetta Vara succession. The orientation of the tectonitic foliation in the mantle peridotites is N95°–70°, 30°–45°; the orientation of the magmatic layering in the gabbroic pluton is N250°–290°, 40°–75°. Another ophiolitic tectonic slice (TS) is exposed to the west of Rocchetta Vara succession and consists mostly of serpentinized peridotites of mantle origin that are overlain by radiolarian cherts. The absence of gabbroic breccias in the sedimentary cover makes the paleogeographic attribution of this tectonic slice uncertain, which has thus not been considered in this work. Maps of Figure 1b and Figure 1c are compiled after the 1:50,000 geological map scale from the ISPRA Web site (http://www.apat.gov.it/MEDIA/carg/233_PONTREMOLI/Foglio.html) and this work. (d) Detail of cross section in Figure 1c. The contact between the mantle sequence and the underlying gabbroic pluton is characterized by the occurrence of olivine-rich troctolites. Gabbroic sills show sharp contacts with respect to the mantle peridotites, crosscutting at high angles their tectonitic foliation, and the olivine-rich troctolites. A gabbroic sill within the mantle peridotites crosscuts an olivine gabbro dike at high angle.

[7] The mantle sequence from the Val Graveglia-Bargonasco ophiolite consists mainly of depleted spinel peridotites showing reequilibration under plagioclase facies conditions [Beccaluva et al., 1984; Rampone et al., 1996]. These peridotites are isotopically variably depleted relative to typical depleted mantle reservoirs [see also Rampone et al., 1998], similar to what is observed for modern oceanic lithosphere [Salters and Dick, 2002; Cipriani et al., 2004]. The Val Graveglia-Bargonasco peridotites represent either asthenospheric material that ascended in response to oceanic spreading, or exhumed subcontinental lithosphere that underwent thermochemical erosion by the upwelling asthenosphere during the rifting [see also Tribuzio et al., 2004; Piccardo and Guarnieri, 2009].

[8] Gabbros and basalts from the Bracco-Levanto ophiolite have trace element and initial Nd isotopic compositions similar to those of modern NMORB [Tribuzio et al., 1995, 2000; Rampone et al., 1998]. The gabbro-peridotite association of Bracco-Levanto ophiolite records a polyphase tectonic evolution in ductile to brittle shear zones, which was correlated with its exhumation at the seafloor [Treves and Harper, 1994; Molli, 1995, 1996; Tribuzio et al., 1995, 2000; Menna, 2009]. The gabbroic rocks from Bracco-Levanto and Scogna-Rocchetta Vara ophiolites are associated with olivine-rich troctolites [Bezzi and Piccardo, 1971]. These olivine-rich troctolites are texturally and compositionally similar to those from oceanic core complexes at the Mid Atlantic Ridge [Suhr et al., 2008; Drouin et al., 2009; Dick et al., 2010]. The olivine-rich troctolites from Internal Ligurian ophiolites have been recently shown to have formed by interaction of an olivine-rich matrix with infiltrating MORB-type melts [Renna and Tribuzio, 2011]. This interaction most likely occurred in mantle melt conduits of replacive origin, which were subsequently impregnated by MORB-type melts saturated in plagioclase + clinopyroxene to produce the olivine-rich troctolites [Renna and Tribuzio, 2011].

3. Field Relationships

[9] The studied ophiolite is exposed for ∼15 km2 and consists mainly of two tectonic successions, namely the Scogna and Rocchetta Vara (Figure 1b). The Scogna succession consists mostly of an altered gabbroic section that is locally overlain by gabbroic breccias. The Rocchetta Vara succession occurs in an orogenic-related tight syncline. In particular, the Rocchetta Vara exposure constitutes the overturned limb of this fold and stretches along the NW-SE direction for ∼6 km. The Rocchetta Vara succession is characterized by a gabbro-peridotite basement overlain by a sedimentary cover made up of gabbroic breccias, Middle-Upper Jurassic radiolarian cherts and Cretaceous shaly pelagites [Barret and Friedrichsen, 1989; Chiari et al., 2000; Principi et al., 2004].

[10] The gabbroic breccias from Scogna and Rocchetta Vara successions are poorly sorted and clast supported. The clasts are commonly nearly angular and up to meter scale; they consist of medium to coarse-grained olivine- to clinopyroxene-rich gabbro, in places showing porphyroclastic fabric. Olivine and clinopyroxene from the gabbro clasts are locally replaced by reddish patches that are made up of fine-grained hematite + calcite ± quartz. Near the contact with the gabbroic breccias, the mantle peridotites are transformed into calcite-veined hematite-bearing serpentinites (ophicalcites). Calcite- and hematite-bearing brittle structures are also found in the gabbroic rocks along the contact with overlying breccias. The calcite- and hematite-bearing brittle structures probably formed when the gabbro-peridotite basement was exposed at the seafloor. The contacts between the gabbro-peridotite association and gabbroic breccias are frequently tectonically reworked.

[11] The Rocchetta Vara mantle sequence mostly consists of peridotites commonly showing extensive serpentinization. The original fabric is generally preserved and the peridotites show porphyroclastic to tectonitic foliation (Figures 2a and 2b), characterized by alignment of porphyroclastic orthopyroxene and spinel. The peridotites locally include up to 3 cm thick pyroxene-rich layers that are generally boudinaged and elongated nearly concordantly with respect to the foliation of the host rocks. Porphyroclastic pyroxene and spinel from the pyroxenite layers and the host peridotites show the same alignment. The peridotites frequently contain subparallel plagioclase-rich veinlets (up to 2 mm thick and up to 2 cm long) that broadly follow the foliation of the host rocks (Figure 2c).

Figure 2.

Main mesostructures of Scogna-Rocchetta Vara ophiolite: (a) Porphyroclastic peridotite with pyroxenite banding (M. Sovrani area). (b) Peridotite-pyroxenite association showing a tectonitic foliation (M. Gruzzo area). (c) Tectonized peridotite showing plagioclase-rich veinlets elongated concordantly with the foliation of the host rock. At the top of the photograph, an olivine gabbro dike displays diffuse contacts with the host peridotite (M. Gruzzo area). (d) Replacive dunite with spinel trails showing diffuse contacts with the host tectonized peridotite. The orientation of these spinel trails is nearly concordant with the contact with the host rock and its foliation (M. Gruzzo area). (e) Ductile shear foliation in clinopyroxene-rich gabbros forming a low angle with respect to modal layering (M. Sovrani area).

[12] The Rocchetta Vara mantle sequence locally also includes dunite bodies, that are up to meter scale in thickness. These bodies are elongated nearly parallel to the tectonitic foliation and include mm-scale aggregates made up of euhedral spinel, which generally form trails reaching up to 0.3 m in length (Figure 2d). The orientation of the spinel trails is geometrically nearly concordant with the contact with the host rocks and their foliation. The contacts between the tectonized peridotites and the dunites are characterized by a gradual inward decrease in modal proportions of orthopyroxene, thus suggesting a replacive origin for the dunites. The contact zone commonly also displays a gradual inward decrease of the subparallel plagioclase-rich veinlets, thus indicating that the replacive dunites formed after the plagioclase-rich veinlets. Plagioclase rarely occurs within the dunites and develops irregular and thin veinlets (<1 mm) containing, in places, grains of pyroxene.

[13] The peridotite foliation and the dunite bodies are locally crosscut by troctolite to olivine gabbro dikes (Figure 2c). These dikes are up to meter scale in thickness and commonly show diffuse contacts with the host rocks; the angle between the peridotite foliation and the gabbroic dikes is <10°. In the troctolite to olivine gabbro dikes, euhedral olivine and plagioclase grains show parallel alignment, thus leading to a magmatic foliation that is subparallel to the contacts with the host mantle rocks and their foliation. The thickest dikes also display subparallel modal and/or grain size layering.

[14] In the NW sector of the Rocchetta Vara succession, a gabbroic intrusion (up to 400 m thick) is exposed below a ∼150 m thick mantle sequence (Figure 1c). In particular, the mantle sequence overlying the gabbroic intrusion contains sills up to 3 m thick made up of coarse-grained clinopyroxene-rich gabbro. These gabbroic sills are elongated nearly perpendicular to the peridotite foliation. The gabbroic sills display sharp planar boundaries to the host peridotites, without grain size reduction. We also found one gabbroic sill crosscutting an olivine-rich gabbro dike at a high angle (∼70°; Figure 1d). The thickest sills exhibit modal and/or grain size layering that is subparallel to the contacts with the host mantle rocks.

[15] The Rocchetta Vara gabbroic intrusion consists broadly of coarse-grained clinopyroxene-rich gabbros, associated with minor amounts of medium to coarse-grained olivine gabbros to troctolites. Poor exposure conditions, and locally the marked alteration of rocks, obscure the contacts between the gabbros, olivine gabbros and troctolites. Note, however, that the troctolites form lenticular bodies (up to a few tens of meters in thickness) occurring at different distances from the contact with overlying sediments. Troctolites to olivine gabbros and gabbros are locally characterized by a weak modal and/or grain size layering, which is subparallel to the orientation of the gabbroic sills. In addition, the troctolites commonly show a magmatic foliation produced by alignment of euhedral olivine and plagioclase grains, which is at a low angle to the modal/grain size layering. The clinopyroxene-rich gabbros from the Rocchetta Vara gabbroic intrusion are in places characterized by a ductile shear foliation that forms a low angle with respect to the magmatic layering (Figure 2e). The shear zones have a width of several meters and were found near the contacts with the olivine-rich troctolites and the mantle peridotite lenses. Sheared gabbros exhibit porphyroclastic and, rarely, mylonitic fabric.

[16] Within the Rocchetta Vara gabbroic intrusion, there are a few mantle bodies (up to 50 m in thickness) that are elongated subparallel to the magmatic layering of the host gabbros. The mantle peridotites from these lenses retain a tectonitic to porphyroclastic foliation and show subparallel plagioclase-rich veinlets. The mantle lenses within the gabbroic pluton also contain meter-scale dunite bodies as well as gabbroic dikes showing diffuse contacts with the host peridotites. The structures of these mantle lenses and those of the mantle sequence enclosing the gabbroic intrusion are geometrically concordant. In the Rocchetta Vara succession, the contact between the gabbroic intrusion and overlying mantle sequence is characterized by occurrence of olivine-rich troctolites (Figure 1d). These rocks are exposed for a thickness of ∼75 m and contain sills (up to meter scale in thickness) made up of clinopyroxene-rich gabbros, which show sharp planar contacts with the host olivine-rich troctolites.

[17] The Scogna gabbroic section is mostly made up of clinopyroxene-rich gabbros, olivine gabbros and troctolites, similar to that from Rocchetta Vara succession. In addition, the Scogna gabbroic section contains two bodies of olivine-rich troctolite (∼50 m thick), lying at different distances from the contact with the overlying gabbroic breccias. Furthermore, the Scogna gabbroic section locally include basalt dikes displaying chilled margins and reaching a couple of meters in thickness. The basalt dikes are commonly porphyritic, with up to 10 vol % plagioclase phenocrysts, and crosscut the fabric of the host gabbros (i.e., modal/grain size layering and troctolite foliation) at a high angle.

[18] Figure 3 displays a paleotectonic reconstruction of the Scogna-Rocchetta Vara ophiolite in Upper Jurassic. The mantle peridotites display a foliation that is at a high angle with respect to the magmatic layering in the gabbroic pluton. These peridotites include diffuse plagioclase-rich veinlets and, locally, dunite bodies that are both elongated concordantly with the mantle foliation. In addition, the peridotites are crosscut by olivine-rich gabbroic dikes that are subparallel to the mantle structures and commonly display diffuse contacts with the host peridotites. Near the contact with the gabbroic pluton, the gabbroic dikes are postdated by clinopyroxene-rich gabbroic sills, displaying sharp planar boundaries with the host peridotites. The gabbroic pluton consists of clinopyroxene-rich gabbros, olivine gabbros and troctolites and locally contains mantle peridotite lenses as well as bodies made up of olivine-rich troctolites. The intrusive fabric of the gabbroic pluton is subparallel to the orientation of the gabbroic sills. The gabbroic pluton locally includes ductile shear zones and basalt dikes at low and high angles, respectively, relative to the intrusive fabric.

Figure 3.

Paleotectonic reconstruction of Scogna-Rocchetta Vara ophiolite in the Middle Jurassic. The scale of replacive dunite bodies, gabbroic dikes, gabbroic sills, and basalt dikes is exaggerated. The contact between the gabbro-peridotite basement and overlying gabbroic breccias is arbitrarily depicted as subhorizontal. However, the original geometric relationships between the fabrics of the gabbro-peridotite basement and the overlying sediments are most likely obliterated by the orogenic tectonics (see text for further details).

4. Petrography and Major Element Mineral Compositions

[19] The main petrographic features and localities of Scogna-Rocchetta Vara samples selected for chemical analyses are reported in Tables 1a1c. Major element analyses of mineral cores (Tables 27) were carried out using a JEOL JXA-8200 electron microprobe located at Dipartimento di Scienze della Terra, Università degli Studi di Milano (Italy); conditions of analyses were 15 kV and 15 nA, and natural standards were utilized. For comparative purposes, new chemical analyses were also carried out on dunite, gabbro and basalt samples from other successions of the Internal Ligurian ophiolites.

Table 1a. Location and Main Petrographic Features of Ultramafic Samples Selected for Chemical Analysesa
SampleUnitLocalityLatitude and LongitudeRock TypebOlivineOrthopyroxeneClinopyroxeneSpinelPlagioclaseRemarks
Modal Amounts (vol %)Degree of Alteration (vol %)Modal Amounts (vol %)Degree of Alteration (vol %)Modal Amounts (vol %)Degree of Alteration (vol %)Modal Amounts (vol %)Degree of Alteration (vol %)Modal Amounts (vol %)Degree of Alteration (vol %)
  • a

    Mineral abbreviations after Kretz [1983].

  • b

    Modes are estimated using point counting.

  • c

    One thousand points in each standard size section.

  • d

    Five hundred points in each standard size section.

  • e

    Modes are visually estimated.

MG3Scogna–R.VaraM. Gruzzo44°13′34.59″N, 9°47′48.66″Eperidotitec59702210610251390 
MG4Scogna–R.VaraM. Gruzzo44°13′33.65″N, 9°47′49.92″Eperidotitec646019155152510100no Spl porph
CC2Scogna–R.VaraM. Sovrani44°15′6.29″N, 9°44′53.02″Eperidotitec688020304501157100mantle lens within gabbroic pluton; no Cpx porph
MG2aScogna–R.VaraM. Gruzzo44°13′33.96″N, 9°47′51.31″Eperidotited63601610410251590with 20 mm thick pyroxenite layer
MG2aScogna–R.VaraM. Gruzzo44°13′33.96″N, 9°47′51.31″Ewebsterited570530280707100 
MG20Scogna–R.VaraM. Gruzzo44°13′33.76″N, 9°47′50.15″Eperidotited64701810410151390with 15 mm thick pyroxenite layer
MG20Scogna–R.VaraM. Gruzzo44°13′33.76″N, 9°47′50.15″Ewebsterited770480350505100 
MG5aScogna–R.VaraM. Gruzzo44°13′34.36″N, 9°47′50.05″Edunitee90100____105__spinel trails
MG5bScogna–R.VaraM. Gruzzo44°13′34.34″N, 9°47′50.35″Edunitee85100__<5100<55<5100altered Pxs within Pl veinlets
VF102Bracco-LevantoBonassola44°10′32.60″N, 9°35′08.65″Edunitee90100____1010__spinel trails
Table 1b. Location and Main Petrographic Features of Gabbroic Samples Selected for Chemical Analysesa
SampleUnitLocalityLatitude and LongitudeRock TypeOl (vol %)Degree of Ol Alteration (vol %)Pl IgneousPl NeoblastsDegree of Pl Alteration (vol %)Cpx IgneousCpx NeoblastsAmphiboleRemarks
Volume %Grain (mm)Volume %Grain (mm)Volume %Grain (mm)Volume %Grain (mm)Volume %Grain (mm)
  • a

    Mineral abbreviations after Kretz [1983]; modes are visually estimated.

MG2bScogna–R.VaraM. Gruzzo44°13′33.95″N, 9°47′51.30″Etroctolite dike45100503.5–5.0_ 3052.5–5.0_ _ foliation in Ol and Pl grains, 30 mm thick
MG12Scogna–R.VaraM. Gruzzo44°14′52.99″N, 9°44′55.78″EOl gabbro dike25100452–8.5.0_ 40303.5–10.0_ _ 2 m thick
RV37aScogna–R.VaraM. Sovrani46°85′186″N, 9°48′138″Egabbroic sill_ 502.5–10.0_ 60503.5–15.0_ _ 50 mm thick
CC11Scogna–R.VaraM. Sovrani44°15′26.85″N, 9°44′33.53″Esheared gabbro5100302.5–10300.5–1.030204.0–10.0100.1–0.2accessory Ti-prg neoblastic<0.05 
CC12aScogna–R.VaraM. Sovrani44°15′26.82″N, 9°44′33.46″Esheared gabbro5100201.2–7.5300.1–0.230354.0–8.0< 50.05–0.110, Hbl neoblastic0.2–0.7with relics of neoblastic Cpx
CC12bScogna–R.VaraM. Sovrani44°15′26.82″N, 9°44′33.46″Esheared gabbro5100101.0–3.6400.1–0.220<51.0–2.0< 50.01–0.240, Hbl neoblastic0.2–0.8with relics of neoblastic Cpx
FG1/1Internal Ligurian ophiolitesBonassola44°10′50.01″N, 9°34′27.73″ECpx-rich gabbro_ 503.5–15.0_ 100502.5–20.5_ accessory Ti-prg interstitial0.05 
Table 1c. Location and Main Petrographic Features of Basaltic Samples Selected for Chemical Analysesa
SampleUnitLocalityLatitude and LongitudeRock TypeGroundmassPl PhenocrystalsDegree of Pl Alteration (vol %)Cpx Phenocrystals
Volume %Grain (mm)Volume %Grain (mm)Volume %Grain (mm)
  • a

    Mineral abbreviations after Kretz [1983]; modes are visually estimated.

SC1Scogna–R.VaraScogna44°17′38.55″N, 9°42′11.66″Ebasalt dike850.5–1.5151.6–5.06051
BR100Bracco-LevantoBracco44°16′23.81″N, 9°33′46.35″Ebasalt dike950.5–2.052.0–8.090_ 
BR101Bracco-LevantoBracco44°15′00.97″N, 9°34′06.43″Ebasalt dike950.1–2.554.5–10.880_ 
VF100Bracco-LevantoBonassola44°10′34.06″N, 9°35′06.08″Ebasalt dike1000.01–0.75_ 90_ 
Table 2. Major Element Olivine Compositionsa
 Sample
MG3MG4CC2MG2aMG20
  • a

    Average values (wt %); perid, peridotite. Dash indicates not analyzed. SD values are in parentheses.

Rock typeperidperidperidperidperid
MineralOl porphOl porphOl porphOl porphOl porph
Number53657
SiO241.22 (0.20)41.49 (0.09)41.48 (0.27)41.29 (0.38)40.68 (0.20)
TiO20.01 (0.01)0.01 (<0.01)0.01 (0.02)0.01 (0.01)0.01 (0.01)
Al2O30.01 (0.01)0.02 (<0.01)<0.01 (<0.01)0.02 (0.01)<0.01 (<0.01)
Cr2O30.02 (0.02)0.01 (0.0100.01 (0.02)<0.01 (<0.01)<0.01 (<0.01)
FeO9.42 (0.37)9.66 (0.02)10.22 (0.08)10.53 (0.03)10.29 (0.09)
MnO0.13 (0.02)0.14 (0.03)0.15 (0.02)0.15 (0.01)0.14 (0.03)
NiO– (–)0.37 (0.03)0.37 (0.04)0.33 (0.03)0.34 (0.04)
MgO49.54 (0.08)49.89 (0.03)49.58 (0.15)49.11 (0.38)48.89 (0.48)
CaO0.05 (0.02)0.03 (0.01)0.06 (0.01)0.05 (0.02)0.05 (0.01)
Na2O<0.01 (0.01)0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)
K2O<0.01 (<0.01)0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)
Sum100.42 (0.51)101.61 (0.04)101.89 (0.40)101.49 (0.70)100.42 (0.56)
Forsterite (mol %)90.4 (0.3)90.2 (0.0)89.6 (0.1)89.3 (0.1)89.4 (0.2)
Table 3. Major Element Orthopyroxene Compositionsa
 Sample and Mineral
MG3MG4CC2MG2aMg20
Opx porphOpx neoOpx porphOpx neoOpx porphOpx neoOpx porphOpx neoOpxOpx neoOpx
  • a

    Average values (wt %); perid, peridotite; web, websterite. Mg # = Mg/(Mg + Fe3+ + Fe2+) * 100; Cr # = Cr/(Cr + Al)*100. SD values are in parentheses.

Rock typeperidperidperidperidperidperidwebwebhost lherzwebhost lherz
Number3831261335634
SiO256.23 (0.76)55.29 (1.15)57.37 (0.23)55.64 (1.05)56.63 (0.32)55.85 (0.34)56.35 (0.31)56.68 (0.83)56.58 (0.61)54.95 (0.57)55.47 (0.21)
TiO20.11 (0.01)0.12 (0.04)0.12 (0.01)0.13 (0.05)0.11 (0.03)0.13 (0.03)0.18 (0.02)0.20 (0.01)0.22 (0.22)0.24 (0.03)0.20 (0.02)
Al2O33.23 (0.1702.99 (0.45)2.96 (0.21)2.78 (0.30)2.90 (0.20)2.75 (0.17)2.73 (0.12)2.62 (0.17)2.63 (0.23)2.73 (0.15)2.60 (0.24)
Cr2O30.89 (0.07)0.78 (0.05)0.84 (0.02)0.78 (0.06)0.83 (0.03)0.85 (0.06)0.84 (0.01)0.79 (0.06)0.83 (0.09)0.85 (0.05)0.76 (0.06)
FeO6.23 (0.06)6.23 (0.13)6.26 (0.10)6.20 (0.14)6.61 (0.09)6.41 (0.10)6.75 (0.03)6.69 (0.09)6.77 (0.15)6.60 (0.09)6.57 (0.09)
MnO0.15 (0.01)0.15 (0.03)0.14 (0.00)0.14 (0.02)0.15 (0.03)0.13 (0.03)0.14 (<0.01)0.17 (0.02)0.15 (0.02)0.18 (0.01)0.16 (0.02)
NiO0.02 (0.03)0.06 (0.05)0.11 (0.02)0.09 (0.02)0.09 (0.03)0.10 (0.04)0.11 (0.01)0.09 (0.02)0.09 (0.03)0.06 (0.04)0.06 (0.06)
MgO32.70 (0.05)32.66 (0.42)32.65 (0.32)32.70 (0.51)32.88 (0.18)32.72 (0.31)32.65 (0.15)32.57 (0.32)32.69 (0.61)33.01 (0.66)32.33 (0.34)
CaO1.32 (0.11)1.44 (0.11)1.39 (0.15)1.38 (0.07)1.29 (0.14)1.29 (0.18)1.27 (0.04)1.53 (0.28)1.24 (0.12)1.53 (0.07)1.26 (0.28)
Na2O0.01 (<0.01)0.02 (0.02)0.01 (0.01)0.02 (0.01)0.01 (<0.01)0.01 (0.02)0.02 (<0.01)0.03 (0.01)0.04 (0.04)0.03 (0.01)0.05 (0.05)
K2O0.01 (<0.01)0.01 (<0.01)0.01 (0.01)0.00 (<0.01)<0.01 (<0.01)<0.01 (<0.01)0.01 (0.01)<0.01 (<0.01)<0.01 (<0.01)0.01 (<0.01)0.02 (0.01)
Sum100.88 (0.90)99.74 (0.93)101.86 (0.23)99.86 (1.41)101.49 (0.33)100.25 (0.35)101.04 (0.32)101.51 (0.83)101.25 (0.99)100.19 (0.19)99.47 (0.38)
Cr #15.6 (0.7)15.1 (1.9)15.9 (0.7)15.9 (1.4)16.2 (1.1)17.1 (1.1)17.1 (0.5)16.8 (0.9)17.5 (0.4)17.3 (0.4)16.3 (0.2)
Mg #90.3 (0.1)90.3 (0.3)90.3 (0.1)90.4 (0.2)89.9 (0.1)90.1 (0.1)89.6 (0.0)89.7 (0.1)89.6 (0.1)89.9 (0.2)89.8 (0.1)
Table 4. Major Element Clinopyroxene Compositionsa
 Sample and Mineral
MG3MG4CC2: Cpx neoMG2aMG20MG2b: Cpx CoreMG12: Cpx CoreRV37a: Cpx CoreCC11CC12a: Cpx porphFG1/1: Cpx porphSC1BR100: G Mass
Cpx porphCpx neoCpx porphCpx neoCpx neoCpx dissCpx neoCpx dissCpx porphCpx neoPhenoxG Mass
  • a

    Average values (wt %). SRV-IL, Scogna-Rocchetta Vara ophiolite; BL, other bodies of Internal Ligurian ophiolites; perid, peridotite; web, websterite; troc-dike, troctolitic dike; gabb-sill, gabbroic sill; Ol-gab dike, olivine gabbro dike; cpx-gab, Cpx-rich gabbro; HT1-gab, plagioclase + clinopyroxene assemblage; HT2-gab, plagioclase + hornblende assemblage; B-dike, basalt dike. Mg # = Mg/(Mg + Fe3+ + Fe2+) * 100; Cr # = Cr/(Cr + Al) * 100. SD values are in parentheses.

UnitSRVSRVSRVSRVSRVSRVSRVSRVSRVSRVSRVSRVSRVSRVSRVBLSRVSRVBL
Rock typeperidperidperidperidperidwebhost peridwebhost peridtroc-dikeOl-gab-dikegab-sillHT1-gabHT1-gabHT2-gabCpx-gabB-dikeB-dikeB-dike
Number6115881036263946652106
SiO253.04 (0.71)52.51 (0.80)52.99 (0.18)52.39 (0.18)52.69 (0.94)52.77 (0.55)52.56 (0.65)51.93 (0.97)51.89 (0.44)52.78 (0.68)51.88 (0.57)54.08 (0.24)53.09 (0.35)53.07 (0.43)53.41 (0.33)53.04 (0.17)52.85 (0.40)51.87 (0.39)51.08 (0.33)
TiO20.25 (0.03)0.28 (0.03)0.28 (0.02)0.29 (0.02)0.31 (0.02)0.42 (0.04)0.50 (0.01)0.48 (0.07)0.58 (0.04)0.58 (0.05)0.39 (0.02)0.43 (0.11)0.54 (0.07)0.55 (0.07)0.44 (0.03)0.55 (0.0200.48 (0.02)1.06 (0.13)0.94 (0.12)
Al2O34.16 (0.11)3.79 (0.69)5.00 (0.28)4.13 (0.28)4.08 (0.70)4.03 (0.16)3.58 (0.44)3.69 (0.38)3.95 (0.04)3.55 (0.46)3.37 (0.42)3.37 (0.26)2.95 (0.15)2.87 (0.31)3.73 (0.23)2.82 (0.17)3.74 (0.09)3.20 (0.44)3.60 (0.28)
Cr2O31.33 (0.01)1.23 (0.21)1.38 (0.02)1.36 (0.02)1.26 (0.09)1.47 (0.04)1.40 (0.15)1.41 (0.09)1.44 (0.11)1.21 (0.17)0.31 (0.04)0.93 (0.20)0.23 (0.02)0.21 (0.02)1.18 (0.06)0.18 (0.04)1.16 (0.03)0.21 (0.10)0.31 (0.08)
FeO3.08 (0.09)2.93 (0.21)3.16 (0.25)2.84 (0.25)2.88 (0.14)3.06 (0.13)3.08 (0.04)3.17 (0.51)3.17 (0.02)3.30 (0.53)3.96 (0.03)3.16 (0.17)5.03 (0.49)4.63 (0.16)3.77 (0.17)4.94 (0.30)4.53 (0.11)7.52 (1.24)6.30 (0.36)
MnO0.09 (0.03)0.09 (0.01)0.08 (0.02)0.09 (0.02)0.09 (0.02)0.10 (0.02)0.10 (0.02)0.10 (0.02)0.10 (0.01)0.11 (0.02)0.14 (0.02)0.10 (0.01)0.15 (0.02)0.17 (0.02)0.12 (0.01)0.16 (0.02)0.13 (0.03)0.20 (0.05)0.16 (0.02)
NiO0.06 (0.03)0.06 (0.03)0.04 (0.04)0.04 (0.04)0.05 (0.02)0.05 (0.02)0.07 (0.01)0.05 (0.03)0.07 (0.01)0.04 (0.02)0.01 (0.01)0.03 (0.03)<0.01 (<0.01)<0.01 (<0.01)0.04 (0.03)0.03 (0.03)0.02 (0.02)0.02 (0.02)<0.01 (0.01)
MgO17.22 (0.26)17.36 (0.63)16.68 (0.83)16.63 (0.83)16.40 (0.57)16.58 (0.48)16.42 (0.58)17.52 (1.78)16.80 (0.04)17.30 (1.70)15.97 (0.09)16.83 (0.44)16.97 (0.82)16.49 (0.38)15.82 (0.34)16.22 (0.11)17.14 (0.07)15.65 (0.45)15.77 (0.16)
CaO21.52 (0.51)22.27 (0.77)21.26 (1.08)22.90 (1.08)22.97 (0.36)21.99 (0.61)22.32 (0.07)20.79 (2.18)20.85 (0.07)21.37 (2.41)21.25 (0.36)21.83 (0.66)21.06 (1.44)21.91 (0.25)21.86 (0.28)22.10 (0.21)19.86 (0.20)19.88 (0.64)20.24 (0.28)
Na2O0.19 (0.01)0.19 (0.02)0.21 (0.01)0.22 (0.01)0.16 (0.01)0.48 (0.03)0.52 (0.04)0.45 (0.06)0.49 (0.01)0.46 (0.07)0.45 (0.03)0.41 (0.04)0.43 (0.0200.45 (0.05)0.48 (0.05)0.45 (0.01)0.43 (0.01)0.44 (0.04)0.41 (0.02)
K2O<0.01 (0.01)0.01 (0.01)0.00 (<0.01)0.00 (<0.01)<0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)0.01 (0.01)0.02 (0.01)<0.01 (<0.01)<0.01 (0.00)0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)0.00 (<0.01)<0.01 (<0.01)<0.01 (<0.01)
Sum100.95 (0.39)100.70 (0.43)101.09 (0.22)100.90 (0.22)100.89 (0.62)100.95 (0.50)100.56 (0.53)99.59 (0.62)99.35 (0.60)100.70 (0.50)97.73 (0.63)101.19 (0.22)100.46 (0.36)100.34 (0.29)100.83 (0.21)100.48 (0.20)100.33 (0.24)100.03 (0.31)98.83 (0.33)
Cr #17.7 (0.4)17.9 (0.8)15.6 (0.8)18.1 (0.8)17.4 (1.7)19.6 (0.6)20.8 (0.3)20.4 (0.9)19.7 (1.0)          
Mg #90.9 (0.2)91.4 (0.4)90.4 (0.3)91.3 (0.3)91.0 (0.5)90.6 (0.2)90.5 (0.3)90.8 (0.5)90.4 (<0.1)90.4 (0.6)87.8 (0.1)90.5 (0.3)85.8 (0.6)86.4 (0.3)88.2 (0.5)85.4 (0.7)87.1 (0.5)78.8 (0.3)81.7 (0.9)
Table 5. Major Element Spinel Compositionsa
 Sample, Rock Type, and Mineral
MG3: peridMG4: perid, spl neoCC2: peridMG2a: Spl porphMG20: webMG20: Host perid, spl porphMG5a: Dun, Spl TrailMG5b: Dun, Spl TrailVF102: Dun, Spl Trail
Spl porphspl neoSpl Rimmed by PlSpl porphSpl Rimmed by PlHost peridwebspl porphspl neo
  • a

    Average values (wt %). SRV-IL, scogna rocchetta Vara ophiolite; BL, other bodies of Internal Ligurian Ophiolites; perid, peridotite; web, websterite; dun, dunite. Mg # = Mg/(Mg + Fe2+) * 100; Cr # = Cr/(Cr + Al) * 100. SD values are in parentheses.

UnitSRVSRVSRVSRVSRVSRVSRVSRVSRVSRVSRVSRVSRVBL
Number747325879721797
SiO20.05 (0.06)0.03 (0.00)0.05 (0.02)0.02 (0.02)0.00 (0.00)0.17 (0.32)0.05 (0.04)0.05 (0.04)0.23 (0.12)0.03 (0.03)0.02 (0.03)0.05 (0.05)0.03 (0.01)0.04 (0.01)
TiO20.35 (0.02)0.38 (0.05)0.32 (0.04)0.36 (0.11)0.43 (0.02)0.37 (0.03)0.67 (0.11)0.68 (0.03)0.65 (0.06)0.59 (0.05)0.62 (0.02)0.21 (0.02)0.25 (0.02)0.24 (0.01)
Al2O331.39 (0.63)30.33 (0.47)27.38 (1.16)29.67 (1.10)27.56 (0.09)27.44 (0.78)24.94 (2.32)25.24 (0.89)26.08 (3.51)27.40 (1.24)27.09 (0.75)43.63 (0.69)42.67 (0.46)44.80 (0.31)
Cr2O336.13 (0.94)38.41 (0.65)37.28 (0.95)39.04 (0.49)38.22 (0.06)37.63 (1.04)39.34 (1.76)39.48 (0.51)39.54 (0.10)40.21 (0.60)39.66 (0.16)22.90 (0.18)23.30 (0.19)22.63 (0.19)
FeO17.18 (0.74)18.35 (0.10)19.99 (1.44)18.19 (1.01)21.16 (0.13)20.90 (0.42)23.67 (1.27)21.39 (0.62)19.55 (0.32)21.72 (0.20)20.69 (0.23)15.09 (0.35)15.88 (0.42)13.19 (0.13)
MnO0.09 (0.02)0.09 (0.00)0.11 (0.03)0.14 (0.06)0.11 (0.11)0.12 (0.02)0.16 (0.02)0.12 (0.02)0.12 (0.02)0.13 (0.05)0.13 (<0.01)0.07 (0.03)0.08 (0.02)0.08 (0.03)
NiO0.02 (0.02)0.22 (0.01)0.11 (0.02)0.13 (0.06)0.16 (0.01)0.13 (0.02)0.09 (0.03)0.14 (0.03)0.15 (0.04)0.14 (0.04)0.13 (0.01)0.26 (0.03)0.26 (0.03)0.27 (0.03)
MgO15.10 (0.35)14.25 (0.51)13.36 (1.24)13.80 (1.18)13.32 (0.05)12.94 (0.35)10.99 (0.74)12.51 (0.37)13.20 (2.00)12.78 (0.75)13.06 (0.28)18.69 (0.27)17.67 (0.26)18.95 (0.17)
CaO0.01 (0.01)0.01 (0.00)0.01 (0.01)0.01 (0.02)0.01 (0.01)0.01 (0.01)0.01 (<0.01)0.01 (0.01)0.05 (0.04)0.03 (0.05)<0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)
Na2O0.01 (0.01)0.02 (0.02)0.05 (0.12)0.02 (0.03)0.01 (0.01)0.01 (0.01)0.02 (0.01)0.01 (0.01)0.02 (0.02)<0.01 (0.01)<0.01 (<0.01)0.01 (0.01)0.01 (0.01)<0.01 (<0.01)
K2O<0.01 (<0.01)0.01 (0.01)<0.01 (<0.01)0.01 (0.01)<0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)0.01 (0.01)<0.01 (<0.01)0.01 (0.01)<0.01 (<0.01)<0.01 (<0.01)0.01 (<0.01)
Sum100.33 (0.95)102.09 (0.23)98.66 (1.14)101.38 (0.72)100.96 (0.06)99.72 (0.47)99.89 (1.14)99.59 (1.10)99.61 (5.47)103.05 (1.95)101.41 (0.98)100.92 (0.74)100.16 (0.50)100.22 (0.33)
Mg #65.3 (1.6)0.6 (<0.1)60.5 (4.2)0.6 (<0.1)58.9 (0.1)57.8 (2.0)50.0 (2.6)56.4 (1.3)0.5 (<0.1)0.5 (<0.1)0.5 (<0.1)75.8 (0.9)72.8 (0.8)76.9 (0.5)
Cr #43.6 (1.1)45.9 (0.8)57.7 (1.5)46.9 (1.1)48.2 (<0.1)57.8 (1.3)51.5 (3.2)51.2 (1.2)50.6 (3.3)49.6 (1.2)49.5 (0.6)26.0 (0.4)26.8 (0.3)25.3 (0.3)
Table 6. Major Element Plagioclase Compositionsa
 Sample and Mineral
MG2b: Pl CoreMG12: Pl CoreCC11CC12a
Pl porphPl neoPl porphPl neo
  • a

    Average values (wt %); troc-dike, troctolitic dike; HT1-gab, plagioclase + clinopyroxene assemblage; HT2-gab, plagioclase + hornblende assemblage. SD values are in parentheses.

Rock typetroc-dikeOl-gab-dikeHT1-gabHT1-gabHT2-gabHT2-gab
Number65751010
SiO252.72 (2.00)52.07 (0.29)53.77 (0.14)54.07 (0.29)54.00 (0.29)53.93 (1.32)
TiO20.07 (0.04)0.06 (0.01)0.06 (0.02)0.04 (0.01)0.05 (0.02)0.01 (0.02)
Al2O330.39 (0.40)29.31 (0.17)29.15 (0.26)29.31 (0.47)28.52 (0.13)29.82 (1.39)
Cr2O3<0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)<0.01 (<0.01)0.01 (0.01)
FeO0.21 (0.10)0.22 (0.04)0.17 (0.03)0.14 (0.03)0.18 (0.02)0.11 (0.04)
MnO0.01 (0.01)0.01 (0.01)0.01 (0.01)0.01 (0.01)0.01 (<0.01)0.01 (0.01)
NiO0.01 (0.01)0.01 (0.02)0.01 (0.02)0.01 (0.01)0.00 (0.01)0.01 (0.02)
MgO0.16 (0.20)0.05 (0.01)0.02 (0.01)0.01 (0.00)0.02 (0.01)0.04 (0.10)
CaO12.73 (0.25)11.88 (0.16)11.83 (0.15)11.87 (0.24)12.49 (0.20)12.45 (0.73)
Na2O4.41 (0.23)4.43 (0.05)4.88 (0.06)4.91 (0.13)4.36 (0.13)4.38 (0.56)
K2O0.02 (0.01)0.02 (0.01)0.03 (0.00)0.02 (0.01)0.03 (0.00)0.11 (0.29)
Sum100.72 (1.78)98.05 (0.60)99.93 (0.35)100.40 (0.33)99.67 (0.11)100.90 (1.03)
An (mol %)61.4 (1.7)59.7 (0.5)57.2 (0.6)57.1 (1.2)61.2 (1.1)60.7 (3.8)
Table 7. Major Element Amphibole Compositionsa
 Sample
CC11CC12aCC12b
  • a

    Average values (wt %). HT1-gab, plagioclase + clinopyroxene assemblage; HT2-gab, plagioclase + hornblende assemblage; Mg # = Mg/(Mg + Fe2+) *1 00; dash indicates not analyzed. SD values are in parentheses.

Rock typeHT1-gabHT2-gabHT2-gab
MineralTi-prg neoHbl neoHbl neo
Number9910
SiO245.35 (0.54)48.21 (1.08)49.13 (0.68)
TiO22.61 (0.2300.80 (0.43)0.81 (0.29)
Al2O311.29 (0.38)9.49 (0.42)8.90 (0.46)
Cr2O30.40 (0.07)0.54 (0.48)0.58 (0.35)
FeO6.85 (0.35)6.54 (0.25)6.60 (0.21)
MnO0.09 (0.02)0.09 (0.01)0.10 (0.02)
NiO0.04 (0.03)0.07 (0.03)0.05 (0.02)
MgO17.16 (0.22)18.05 (0.42)17.80 (0.34)
CaO12.07 (0.07)12.09 (0.13)11.85 (0.14)
Na2O2.81 (0.13)2.16 (0.22)2.16 (0.22)
K2O0.03 (0.01)0.08 (0.01)0.08 (0.01)
Cl<0.01 (<0.01)0.21 (0.03)– (–)
Sum98.71 (0.23)98.34 (0.22)98.06 (0.23)
Mg #81.7 (0.9)83.1 (0.7)82.8 (0.7)

4.1. Tectonized Peridotites

[20] The sampled peridotites have ∼5 vol % clinopyroxene (Tables 1a1c). In all samples, olivine is partly replaced by serpentine and minor Fe oxide phases, and plagioclase is altered into fine-grained aggregates of prehnite ± epidote ± hydrogrossular ± chlorite. The peridotites hosting the gabbroic pluton and those within the gabbroic pluton have similar modal compositions and microstructures.

[21] The tectonitic foliation is shown by the alignment of porphyroclastic olivine, orthopyroxene, clinopyroxene and spinel. Porphyroclastic olivine is up to 3 mm in length and have 89.3 to 90.4 mol % forsterite (Table 2). Porphyroclastic orthopyroxene and clinopyroxene (up to 7 mm and 3 mm in length, respectively) are exsolved and commonly show undulose extinction and kink bands. They are mantled by polygonal aggregates made up of unstrained orthopyroxene and clinopyroxene (both 0.3–0.5 mm), which are associated with accessory amounts of spinel (Figure 4a) and olivine. In these aggregates, neoblastic olivine and spinel are smaller (both ≤0.2 mm) than the associated pyroxenes.

Figure 4.

Main microstructures of the studied samples. (a) Tectonized peridotite MG4: neoblastic assemblage of orthopyroxene (Opx II) + clinopyroxene (Cpx) + spinel (Spl) at the edge of a porphyroclastic orthopyroxene (Opx I). The peridotite matrix is crosscut by serpentine (Srp) microveins. (b) Tectonized peridotite MG3: porphyroclastic spinel (Spl) rimmed by plagioclase (Pl) and orthopyroxene (Opx). (c) Tectonized peridotite MG3: plagioclase-rich veinlet (Pl) crosscutting a porphyroclastic olivine (Ol) and clinopyroxene (Cpx). A spinel (Spl) occurs at the boundary with the host peridotite. Plagioclase is replaced by fine-grained aggregates of prehnite ± epidote ± hydrogrossular. (d) Websterite layer MG2a: 120° triple junctions between neoblastic clinopyroxene (Cpx), orthopyroxene (Opx), and spinel (Spl). (e) Sheared gabbro CC11: neoblastic clinopyroxene (Cpx II) associated with minor amounts of Ti pargasite (Ti-Prg) at the edge of porphyroclastic clinopyroxene (Cpx I). The Cpx II + Ti-Prg aggregates are in contact with polygonal aggregates made up of plagioclase (Pl). (f) Thin section of sheared gabbro CC12b: mylonitic structure made up of bands rich in neoblastic hornblende (Hbl) and plagioclase (Pl).

[22] Porphyroclastic orthopyroxene have low TiO2 concentrations (Figure 5) and relatively high Cr2O3 abundances (0.8–0.9 wt %). Porphyroclastic clinopyroxene have low amounts of Na2O and TiO2 (Figure 6), and high Cr2O3 concentrations (1.3–1.4 wt %). Neoblastic orthopyroxene and clinopyroxene have commonly slightly lower concentrations of Al2O3 and Cr2O3 than their porphyroclastic counterparts (Tables 3 and 4), most likely as a result of formation of synkinematic spinel, which incorporates a high amount of these elements (Table 5). Figure 6 also shows that analyzed porphyroclastic clinopyroxenes are chemically similar to those from another mantle section of the Internal Ligurian ophiolites [Rampone et al., 1996].

Figure 5.

Variation of TiO2 versus Mg # [Mg/(Mg+Fe2++Fe3+) × 100] for orthopyroxene cores from tectonized peridotites and websterite layers of Scogna-Rocchetta Vara ophiolite. Data are averaged per sample; the error bars represent one standard deviation of the mean value.

Figure 6.

Variation of TiO2 and Na2O versus Mg # [Mg/(Mg+Fe2++Fe3+) × 100] for clinopyroxene cores from mantle and crustal rocks of Scogna-Rocchetta Vara ophiolite. Data are averaged per sample; the error bars represent one standard deviation of the mean value. Grey field indicates compositions of porphyroclastic clinopyroxenes from mantle peridotites of Val Graveglia-Bargonasco ophiolite (Internal Ligurian units [Rampone et al., 1996]). The compositions of clinopyroxenes from the olivine-rich troctolites of Internal Ligurian ophiolites are from Renna and Tribuzio [2011]. The compositions of clinopyroxenes from gabbroic and basaltic rocks of Bracco-Levanto ophiolite (Internal Ligurian units) are also shown [Rampone et al., 1998; this work].

[23] Porphyroclastic spinel (up to 3 mm in length) occurs in accessory modal amounts (Tables 1a1c). Spinel is commonly rimmed by a plagioclase corona reaching up to 0.2 mm in thickness (Figure 4b). Orthopyroxene grains, locally up to 0.1 mm thick, are commonly associated with the coronitic plagioclase, at the contact with porphyroclastic olivine. Neoblastic spinel from the pyroxene-dominated polygonal aggregates is chemically similar to the porphyroclastic spinel (Figure 7). We applied the Ca-in-Opx geothermometer [Brey and Köhler, 1990] to the polygonal aggregates, assuming a confining pressure of 1.0 GPa; neoblastic orthopyroxene has 1.3–1.4 wt % CaO, which correspond to temperature estimates of 1090–1130°C.

Figure 7.

Variation of TiO2 versus Cr # [Cr/(Cr+Al) × 100] for spinel cores from tectonized peridotites, websterite layers, and replacive dunites of Scogna-Rocchetta Vara ophiolite. Data are averaged per sample; the error bars represent one standard deviation of the mean value. Compositions of porphyroclastic spinels from peridotites of Val Graveglia-Bargonasco ophiolite (Internal Ligurian units [Rampone et al., 1996]) and of spinels trails from replacive dunites of another section of Alpine ophiolites [Piccardo et al., 2007] are displayed as field A and field B, respectively. The compositions of spinels from replacive dunites (this work) and basalt dikes [Cortesogno and Gaggero, 1992] from other bodies of the Internal Ligurian ophiolites are also shown.

[24] The tectonized peridotites commonly contain subparallel plagioclase-rich veinlets that are commonly 0.5 to 2 mm thick and up to 20 mm long. The plagioclase-rich veinlets commonly crosscut the porphyroclastic olivines and follow the edges of the pyroxene dominated neoblastic aggregates (Figure 4c). Within the neoblastic aggregates, there are sinuous apophyses of these veinlets (up to 150 μm thick). Note that the plagioclase-rich veinlets show concave contacts against the peridotite matrix. The plagioclase-rich veinlets sometimes contain euhedral orthopyroxene grains that are up to 0.5 mm in length. Plagioclase-rich veinlets with similar microstructures were documented for a nearly undeformed mantle section from the Internal Ligurian ophiolites and interpreted to reflect melt impregnation [Rampone et al., 1997].

4.2. Pyroxenite Layers

[25] These rocks consist of deformed orthopyroxene porphyroclasts (up to 5 mm in length) rimmed by unstrained aggregates made up of clinopyroxene, orthopyroxene and minor spinel. In particular, porphyroclastic orthopyroxene commonly shows clinopyroxene exsolution lamellae, undulose extinction and kink bands. In the neoblastic aggregates, pyroxenes (0.3–0.5 mm) show 120° triple junctions (Figure 4d); neoblastic spinel is smaller (≤0.2 mm) than associated pyroxenes, and olivine is locally present as small (≤0.1 mm) serpentinized grains. Minor amounts of porphyroclastic spinel are also commonly present and elongated parallel to porphyroclastic orthopyroxene. On the basis of the measured modes (Tables 1a1c), these rocks are hereafter referred to as websterites.

[26] Plagioclase-rich veinlets are commonly found along the contacts between the websterites and the host peridotites. These veins show sinouos apophyses (up to 0.2 mm thick) within the websterites, which also contain altered plagioclase as coronas around the porphyroclastic spinel. Orthopyroxene, clinopyroxene and spinel from the websterites have relatively high concentrations of TiO2 (Figures 57). Clinopyroxene also has high concentrations of Na2O. Equilibrium temperatures for the neoblastic mineral assemblage were calculated by applying the Ca-in-Opx geothermometer [Brey and Köhler, 1990], assuming P = 1.0 GPa; the CaO content of orthopyroxene is ∼1.5 wt %, which gave estimates of ∼1150°C.

[27] In the tectonized peridotite enclosing the websterite layer, at the thin section scale, the minerals are chemically distinct from those of the other tectonized peridotites considered in the present study. Porphyroclastic orthopyroxene and spinel are enriched in TiO2 with respect to porphyroclastic orthopyroxene and spinel from the other peridotites (Figures 5 and 7). In addition, porphyroclastic olivine, orthopyroxene and spinel from the peridotite hosting the websterite layer has slightly lower Mg # with respect to the other peridotites (see also Tables 2 and 5). Porphyroclastic orthopyroxene and spinel from the websterite layer and adjacent tectonized peridotite are therefore chemically similar. Furthermore, we analyzed a clinopyroxene grain (∼250 μm) rimming a porphyroclastic orthopyroxene at a distance of 5 mm from the websterite layer. This clinopyroxene has higher TiO2 and Na2O than clinopyroxenes from the other peridotites and chemically resembles the neoblastic clinopyroxene from the websterite (Figure 6).

4.3. Replacive Dunites

[28] In these rocks, olivine is altered into serpentine and minor Fe oxide phases, and is associated with minor amounts of euhedral spinel. Spinel is characterized by low Cr # and TiO2 concentrations, similar to spinel from replacive dunites from another mantle section of the Internal Ligurian ophiolites (Figure 7). Plagioclase films (up to 0.5 mm thick) are common among the olivine grains and locally contain euhedral pyroxene grains. Both plagioclase and pyroxene are replaced by fine-grained aggregates of prehnite ± epidote ± hydrogrossular and serpentine ± minor Fe oxide phases, respectively.

4.4. Gabbroic Rocks

[29] Troctolite to olivine gabbro dikes intruding the mantle sequence are generally medium grained (Tables 1a1c) and show nearly equigranular structure. In these rocks, plagioclase (anorthite = 60–61 mol % (Table 6)) and olivine are euhedral, and clinopyroxene (up to 20 vol %) is anhedral to subhedral. In the gabbroic sills, olivine is <10 vol %. The gabbros from the sills are commonly coarse grained; olivine and plagioclase are euhedral to subhedral, and clinopyroxene is subophitic. In these rocks, plagioclase and olivine are frequently replaced by microaggregates made up of epidote + albite and serpentine + Fe oxide phases, respectively.

[30] Clinopyroxene-rich gabbros from the pluton are modally and texturally similar to the gabbros from the sills (Tables 1a1c). Clinopyroxenes from the plutonic gabbros have slightly lower Mg # than clinopyroxenes from the dikes and sills (Figure 6), and plagioclase with anorthite component ranging from 57 to 61 mol % (Table 6). Clinopyroxenes from gabbros of other Internal Ligurian plutonic sequences have similar major element compositions [Tribuzio et al., 1995; Rampone et al., 1998; this work].

[31] The gabbroic sequences from the studied ophiolite are locally associated with olivine-rich (commonly 80–90 vol %) troctolites [Renna and Tribuzio, 2011]. These rocks have olivine (Fo = 87–88 mol %) and accessory spinel with rounded to embayed morphology. The high Mg # values and the high Cr2O3 concentrations of accessory poikilitic clinopyroxene (88–90 and 1.3–1.5 wt %, respectively) from the olivine-rich troctolites were correlated with a reaction between an olivine-spinel matrix and infiltrating MORB-type melts [Renna and Tribuzio, 2011]. Conversely, in the troctolites either from the dikes or the gabbroic sequences, olivine and plagioclase occur in nearly equal modal amounts and are both euhedral, thereby suggesting a cumulate origin.

4.5. Sheared Gabbros

[32] Two distinct mineral assemblages were recognized in sheared gabbros. The first is characterized by recrystallization of clinopyroxene and plagioclase. In porphyroclastic metagabbros, in particular, polygonal aggregates of neoblastic clinopyroxene and plagioclase occur at the edges of deformed clinopyroxene and plagioclase porphyroclasts, respectively. In addition, fine-grained neoblastic Ti pargasite (Table 7) is commonly present as an accessory within the aggregates rimming the porphyroclastic clinopyroxene (Figure 4e). Within individual samples, neoblastic clinopyroxene and plagioclase are chemically similar to the porphyroclastic counterparts (Tables 4 and 6).

[33] The second mineral assemblage is represented by the hornblende + plagioclase amphibolite facies association (Figure 4f). In the amphibolite facies sheared gabbros, medium- to fine-grained aggregates of neoblastic hornblende are present at the rims of porphyroclastic clinopyroxene and fine-grained plagioclase aggregates are found at the margin of porphyroclastic plagioclase. Neoblastic hornblende has lower TiO2, Al2O3 and Na2O than neoblastic Ti pargasite (Table 7) and contains significant amounts of Cl (∼0.2 wt %). Associated neoblastic plagioclase is chemically similar to porphyroclastic plagioclase of igneous origin (Table 6). Locally, the hornblende + plagioclase assemblage mantles the association of neoblastic clinopyroxene + plagioclase + Ti pargasite.

[34] Sheared gabbros showing crystallization of clinopyroxene + plagioclase + Ti pargasite and, in places, of hornblende + plagioclase were reported for other gabbroic bodies from the Internal Ligurian ophiolites [i.e., Cortesogno et al., 1975; Molli, 1994, 1995, 1996; Tribuzio et al., 1995, 2000; Menna, 2009]. These shear zones bear similar structural, microstructural and compositional features to those documented in the present study. Temperature conditions of the ductile shearing events recorded by the gabbros from Scogna-Rocchetta Vara ophiolite were evaluated applying the amphibole-plagioclase geothermometer of Holland and Blundy [1994], assuming pressure conditions of 0.2 GPa. Temperature estimates of ∼850°C were obtained for the plagioclase + Ti pargasite (+ clinopyroxene) assemblage. The hornblende + plagioclase pairs yielded temperature values of ∼710°C.

4.6. Basalt Dikes

[35] These rocks commonly include phenocrystic plagioclase in an aphanitic groundmass made up of euhedral plagioclase, ophitic clinopyroxene and accessory ilmenite. Phenocrystic olivine and clinopyroxene are also locally present. Plagioclase is replaced by microaggregates consisting of epidote and albite, and olivine is altered into serpentine and minor Fe oxide phases. Phenocrystic clinopyroxene shows slightly higher Mg # and Cr2O3, and slightly lower TiO2 than groundmass clinopyroxene (Figure 6 and Table 4).

5. Trace Element Compositions of Clinopyroxenes

[36] Clinopyroxene cores from tectonized peridotites, websterites, gabbroic rocks and basalts were analyzed for trace element concentrations (Table 8) by laser ablation ICP-MS at C.N.R.–Istituto di Geoscienze e Georisorse, Unità di Pavia. This instrument couples a Nd:YAG laser source (Brilliant, Quantel) operating at 266 nm with a quadrupole ICP-MS (Drc-e, Perkin Elmer). Analyses were carried out with a spot diameter of ∼40 μm. Data reduction was performed using the “Glitter” software package [Van Achterbergh et al., 2001]. Ablation signal and integration intervals were selected by inspecting the time-resolved data to ensure that no inclusions were present in the analyzed volume. NIST SRM 612 and 44Ca were used as external and internal standards, respectively. Accuracy was tested on the BCR2-g (USGS) reference glass and is estimated to be better than ±5% (1σ).

Table 8. Trace Element Clinopyroxene Compositions of Ultramafic and Mafic Rocks Obtained With LA-ICP-MSa
 Sample and Mineral
MG4CC2: Cpx neoMG2aMG20: Cpx neoMG2b: Cpx CoreRV37a: Cpx CoreCC11: Cpx CoreCC12a: Cpx CoreFG1-1: Cpx CoreSC1BR100: GMass
Cpx porphCpx neoCpx dissCpx neoPhenoxGMass
  • a

    Average values (ppm). SRV-IL, Scogna-Rocchetta Vara ophiolite; BL, other bodies of Internal Ligurian ophiolites; perid, peridotite; web, websterite; troc-dike, troctolitic dike; Eu/Eu* = EuN/√(GdN * SmN); Sr/Sr* = SrN/√(CeN * NdN); Zr/Zr* = ZrN/√(NdN * SmN); Ti/Ti* = TiN/√(GdN * DyN); normalized to C1 chondrite [Anders and Ebihara, 1982]; Cpx-gab, clinopyroxene-rich gabbro; B-dike, basalt dike. SD values (1 sigma) are in parentheses.

UnitSRVSRVSRVSRVSRVSRVSRVSRVSRVSRVBLSRVSRVBL
Rock typeperidperidperidhost peridwebwebtroc-dikegab-sillCpx-gabCpx-gabCpx-gabB-dikeB-dikeB-dike
Number3422443333112611
V323 (2)(7)358 (23)380 (13)377 (9)378 (13)411 (23)308 (24)397 (10)320 (6)336 (11)276 (7)383 (48)447 (25)
Cr11,055 (177)(248)10,686 (631)10,905 (109)11,122 (139)12,017 (634)10,340 (193)8,683 (574)1,885 (132)9,901 (339)1,351 (100)8,396 (564)1,437 (585)2,060 (836)
Co34.8 (1.5)33.8 (0.6)34.3 (6.6)26.7 (1.6)32.7 (1.8)31.7 (1.8)30.8 (5.6)33.8 (6.1)42.8 (1.2)37.1 (3.0)32.9 (0.8)31.0 (0.2)33.2 (4.3)36.3 (2.0)
Ni491 (17)480 (13)453 (32)380 (9)398 (80422 (34)424 (32)488 (34)201 (9)297 (9)158 (8)203 (2)83 (6)105 (12)
Sc65.5 (0.9)61.3 (1.3)67.9 (0.7)76.8 (1.1)74.5 (1.2)71.6 (3.3)75.4 (1.8)59.9 (5.3)75.6 (0.9)77.5 (1.0)110 (2)97.9 (3.7)118 (19)137 (9)
Ti2,085 (53)2,162 (29)2,302 (114)3,286 (30)3,059 (60)3,238 (196)4,478 (40)2,097 (237)3,350 (91)2,890 (59)3,741 (114)3,332 (30)6,234 (701)6,656 (542)
Sr0.26 (0.04)0.49 (0.24)0.29 (0.08)3.91 (0.74)1.49 (0.10)2.79 (0.23)13.9 (0.8)11.4 (0.6)14.5 (0.8)21.1 (1.0)15.0 (0.5)11.0 (0.2)11.7 (1.7)11.0 (0.4)
Zr3.29 (0.11)2.95 (0.30)2.23 (0.23)16.93 (0.11)13.26 (0.40)16.47 (0.63)16.95 (1.15)8.18 (0.91)8.20 (0.33)7.58 (0.28)11.71 (0.39)10.21 (0.72)28.41 (4.95)33.91 (5.12)
Nb0.02 (0.02)0.02 (0.01)<0.01<0.10.10 (0.01)0.09 (0.03)0.05 (0.01)0.04 (0.01)0.02 (0.01)0.01 (0.01)0.04 (0.01)0.05 (0.01)0.03 (0.03)0.03 (0.01)
Y15.8 (0.4)13.6 (0.7)16.8 (1.5)18.8 (0.1)19.9 (1.1)19.6 (0.5)17.0 (0.6)10.5 (0.5)12.0 (0.3)9.60 (0.53)15.9 (0.3)14.7 (0.7)23.7 (3.9)25.5 (2.6)
Hf0.27 (0.06)0.33 (0.07)0.26 (0.03)0.64 (0.14)0.56 (0.12)0.60 (0.14)0.66 (0.04)0.27 (0.05)0.25 (0.09)0.34 (0.09)0.54 (0.18)0.57 (0.24)1.34 (0.39)1.54 (0.25)
Ta<0.01<0.01<0.010.01 (0.01)0.02 (0.02)0.02 (<0.01)0.01 (0.01)0.01 (0.01)<0.01<0.01<0.01<0.010.01 (0.01)0.01 (0.01)
La0.01 (0.01)0.01 (0.00)<0.040.33 (0.03)0.36 (0.05)0.39 (0.05)0.32 (0.03)0.21 (0.05)0.23 (0.03)0.31 (0.04)0.32 (0.04)0.20 (0.04)0.41 (0.12)0.41 (0.04)
Ce0.17 (0.02)0.18 (0.01)0.10 (<0.01)1.86 (0.19)1.82 (0.04)1.87 (0.08)2.08 (0.05)1.19 (0.20)1.70 (0.26)1.40 (0.06)1.58 (0.11)1.17 (0.00)2.22 (0.46)2.31 (0.20)
Pr0.12 (0.01)0.09 (0.02)0.05 (<0.01)0.42 (0.01)0.41 (0.03)0.39 (0.06)0.52 (0.01)0.26 (0.05)0.37 (0.08)0.31 (0.02)0.43 (0.04)0.34 (0.03)0.66 (0.13)0.63 (0.08)
Nd1.14 (0.11)1.27 (0.14)1.16 (0.15)2.98 (0.03)2.65 (0.18)3.42 (0.42)3.76 (0.22)1.81 (0.05)2.53 (0.15)1.97 (0.09)3.01 (0.25)2.41 (0.03)4.81 (0.95)4.71 (0.51)
Sm0.86 (0.12)0.83 (0.05)0.90 (0.08)1.81 (0.11)1.56 (0.09)1.51 (0.14)1.48 (0.07)0.90 (0.06)1.22 (0.15)0.86 (0.10)1.70 (0.14)1.09 (0.09)2.05 (0.36)2.34 (0.21)
Eu0.33 (0.02)0.36 (0.06)0.37 (0.06)0.46 (0.03)0.48 (0.05)0.61 (0.13)0.56 (0.09)0.41 (0.07)0.49 (0.08)0.43 (0.02)0.56 (0.07)0.42 (0.01)0.74 (0.11)0.89 (0.09)
Gd1.79 (0.47)1.71 (0.21)1.72 (0.21)2.59 (0.37)2.58 (0.32)2.66 (0.18)2.37 (0.05)1.33 (0.37)1.88 (0.12)1.43 (0.04)2.46 (0.49)1.95 (0.35)3.61 (0.54)3.87 (0.35)
Tb0.38 (0.06)0.32 (0.04)0.38 (0.02)0.54 (0.06)0.50 (0.03)0.54 (0.09)0.44 (0.01)0.25 (0.06)0.33 (0.02)0.23 (0.03)0.40 (0.05)0.45 (0.02)0.57 (0.12)0.69 (0.05)
Dy2.94 (0.33)2.63 (0.13)3.08 (0.17)3.62 (0.23)3.82 (0.32)3.61 (0.12)3.28 (0.29)1.92 (0.17)2.12 (0.25)1.88 (0.09)2.78 (0.13)2.60 (0.01)4.45 (0.73)5.10 (0.61)
Ho0.63 (0.02)0.50 (0.09)0.61 (0.05)0.75 (0.00)0.79 (0.06)0.75 (0.04)0.70 (0.04)0.39 (0.06)0.47 (0.03)0.45 (0.01)0.59 (0.07)0.54 (0.04)0.95 (0.11)1.00 (0.14)
Er1.70 (0.16)1.50 (0.08)1.69 (0.03)2.11 (0.06)2.20 (0.24)1.92 (0.14)1.71 (0.24)1.31 (0.16)1.38 (0.23)0.99 (0.16)1.73 (0.22)1.85 (0.10)2.78 (0.51)2.94 (0.25)
Tm0.22 (0.04)0.19 (0.01)0.24 (0.01)0.32 (0.02)0.32 (0.01)0.29 (0.05)0.27 (0.02)0.16 (0.02)0.18 (0.03)0.15 (0.02)0.23 (0.06)0.24 (<0.01)0.37 (0.10)0.38 (0.05)
Yb1.77 (0.08)1.35 (0.11)1.13 (0.01)1.82 (0.04)1.58 (0.17)1.96 (0.19)1.67 (0.18)1.13 (0.10)1.40 (0.03)1.04 (0.13)1.65 (0.03)1.48 (0.37)2.46 (0.51)2.60 (0.24)
Lu0.20 (0.03)0.17 (0.02)0.23 (0.05)0.27 (0.03)0.27 (0.03)0.19 (0.03)0.25 (0.02)0.17 (0.03)0.15 (0.01)0.14 (0.01)0.20 (0.04)0.23 (0.00)0.35 (0.07)0.38 (0.06)
LaN/SmN0.010.010.020.120.150.160.140.140.120.230.120.120.130.11
Eu/Eu*0.810.910.910.640.730.920.911.150.981.190.830.880.830.89
Sr/Sr*0.040.070.060.110.050.070.330.520.470.850.460.440.240.22
Zr/Zr*0.220.190.140.480.430.470.470.420.310.380.340.410.590.67
Ti/Ti*0.450.510.500.540.490.520.800.660.840.880.720.740.780.75
ZrN/HfN0.320.230.230.690.630.720.680.800.870.580.580.470.560.58

5.1. Tectonized Peridotites

[37] Porphyroclastic and neoblastic clinopyroxenes from the peridotites of Scogna-Rocchetta Vara ophiolite have similar trace element compositions (Figure 8a). Their chondrite-normalized REE patterns are characterized by a marked LREE depletion (LaN/SmN = 0.01–0.02) relative to MREE and HREE, which are nearly flat at ∼9 times chondrite. Incompatible elements normalized to chondrite show depletions of Sr, Zr and Ti relative to neighboring REE. Zr is also depleted relative to Hf (ZrN/HfN = 0.2–0.3). Note that clinopyroxene from the peridotites within the gabbroic pluton has the same trace element signature of clinopyroxene from the peridotites hosting the gabbroic pluton. A similar geochemical fingerprint was observed for porphyroclastic clinopyroxenes from peridotites of another mantle sequence of the Internal Ligurian ophiolites [Rampone et al., 1996; Piccardo et al., 2004].

Figure 8.

REE and incompatible element compositions of clinopyroxene cores from mantle and crustal rocks of Scogna-Rocchetta Vara ophiolite, normalized to C1 chondrite [Anders and Ebihara, 1982]. (a) Clinopyroxenes from mantle rocks. Note that clinopyroxene from the peridotites enclosed within the gabbroic pluton has the same trace element signature as clinopyroxene from the peridotites hosting the gabbroic pluton. Clinopyroxenes from peridotites of Val Graveglia-Bargonasco ophiolite (Internal Ligurian units [Piccardo et al., 2004]) are also reported. Compositions of clinopyroxenes from peridotites of Mid Atlantic Ridge [Brunelli et al., 2006] and from westerites of Southeast Indian Ridge [Dantas et al., 2007] are displayed as field A and field B, respectively. (b) Clinopyroxenes from gabbroic dikes and sills. Clinopyroxenes from the olivine-rich troctolites of Internal Ligurian ophiolites [Renna and Tribuzio, 2011] and the gabbroic dikes intruding the mantle sequences of Val Graveglia-Bargonasco [Piccardo et al., 2004] are also reported. (c) Clinopyroxenes from plutonic gabbroic rocks. Clinopyroxene of a gabbro from Bracco-Levanto ophiolite (Internal Ligurian units (this work)) is also reported. (d) Clinopyroxenes from basalt dikes. Clinopyroxenes from basalt dikes of Bracco-Levanto ophiolite (this work) are also reported. The incompatible element compositions of clinopyroxene at the equilibrium with NMORB were calculated through average NMORB compositions of Hofmann [1988] and experimentally derived coefficients for clinopyroxene/basalt partitioning of Hart and Dunn [1993].

5.2. Websterite Layers

[38] Clinopyroxene from the selected websterite is less depleted in LREE (LaN/SmN = 0.14) than clinopyroxene from the tectonized peridotites (Figure 8a). The REE pattern of clinopyroxene from the websterite is characterized by nearly flat MREE and HREE at about 13 times chondrite and a slight negative Eu anomaly (Eu/Eu* = 0.7). Sr is depleted relative to LREE, with higher absolute concentrations than in clinopyroxenes from the peridotites. Zr and Ti are depleted with respect to neighboring REE. In particular, the Zr depletion is less marked than in clinopyroxenes from the peridotites. This is also shown by the relatively high ZrN/HfN value (0.6) of clinopyroxene from the websterite. In the thin section including the websterite layer, we analyzed a clinopyroxene grain from the host peridotite; the clinopyroxene from the peridotite has trace element compositions resembling those of clinopyroxene from the websterite.

5.3. Gabbroic Rocks

[39] Igneous clinopyroxenes from the gabbroic dikes, gabbroic sills and clinopyroxene-rich gabbros of the gabbroic pluton display nearly parallel REE and incompatible element patterns (Figures 8b and 8c). In particular, the LREE are depleted relative to MREE and HREE (LaN/SmN = 0.12–0.23), which are nearly flat and range from about 7 to 11 times chondrite. The incompatible element patterns display negative Sr, Zr and Ti anomalies relative to neighboring REE and show relatively high ZrN/HfN values (0.6–0.9). These patterns mimic the pattern of clinopyroxene at equilibrium with NMORB, calculated using average NMORB compositions [Hofmann, 1988] and experimentally derived partition coefficients [Hart and Dunn, 1993]. Note that the incompatible element signature of clinopyroxene from the gabbroic rocks is similar to that of clinopyroxene from the websterite layers. A similar incompatible element fingerprint was also found for clinopyroxenes from the olivine-rich troctolites [Renna and Tribuzio, 2011] and other gabbroic bodies of the Internal Ligurian ophiolites [Tribuzio et al., 1995; Rampone et al., 1996, 1997; Piccardo et al., 2004; this work].

5.4. Basalt Dikes

[40] Phenocrystic and groundmass clinopyroxenes display nearly parallel REE and incompatible element patterns (Figure 8d). The REE patterns are characterized by LREE depletion (LaN/SmN = 0.11–0.12) and the incompatible element patterns show negative Sr, Zr and Ti anomalies relative to neighboring REE, similar to clinopyroxenes from the gabbroic rocks; the ZrN/HfN value of clinopyroxenes from the basalts is relatively high 0.5–0.6. Phenocrystic clinopyroxene has lower concentrations of incompatible elements than groundmass clinopyroxene. For instance, HREE are ∼10 and ∼16 times chondrite for phenocrystic and groundmass clinopyroxene, respectively. Groundmass clinopyroxene also differs in a barely appreciable negative Eu anomaly (Eu/Eu* = 0.8) and the most marked negative Sr anomaly. Figure 8d also shows that groundmass clinopyroxenes from basalt dikes intruding other gabbroic bodies of the Internal Ligurian ophiolites have a similar incompatible element signature.

6. Whole-Rock Compositions of Basalts

[41] One basalt dike from Scogna gabbroic section was selected for whole-rock major and trace element analyses (Table 9). For comparative purposes, new analyses were also carried out on two basalt dikes intruding other gabbroic bodies from the Internal Ligurian ophiolites. The chemical analyses were carried out by ICP-MS spectrometry at “Activation Laboratories” (Ancaster, Ontario). Precision and accuracy of ICP-MS analyses are commonly within 10%.

Table 9. Whole-Rock Major and Trace Element Compositions of Basaltsa
 Unit
SRVBLBL
  • a

    B-dike, basalt dike; SRV, Scogna-Rocchetta Vara ophiolite; IL, other bodies of Internal Ligurian ophiolites; LOI, loss on ignition. Mg # = molar Mg/(Mg + Fetot2+).

SampleSC1VF100BR101
Rock typeb-dikeb-dikeb-dike
 
Major Elements (wt %)
SiO249.7150.449.16
TiO21.281.421.514
Al2O316.3315.915.89
Fe2O37.717.078.7
MnO0.1410.1720.143
MgO8.817.68.78
CaO7.38.937.95
Na2O3.764.153.75
K2O0.820.170.5
P2O50.120.130.16
L.O.I3.602.913.09
Total99.698.999.6
Mg #69.468.066.7
 
Trace Elements (ppm)
V220236250
Cr260290210
Co323934
Ni10011090
Cu208060
Zn<30110<30
Ga131914
Rb554
Sr191208188
Ba61710
Zr99136102
Nb1.63.42.5
Y24.134.329.8
Hf2.432.6
Ta0.090.210.13
Pb<5<4<5
Th0.210.260.21
U0.070.20.08
La3.014.253.96
Ce10.013.412.7
Pr1.752.342.19
Nd9.0611.311.4
Sm3.013.343.75
Eu1.061.231.3
Gd3.924.424.87
Tb0.720.80.9
Dy4.445.045.46
Ho0.971.061.16
Er2.73.273.27
Tm0.4160.4840.512
Yb2.712.963.3
Lu0.370.410.48
Zr/Y4.123.923.42

[42] The selected sample has a Mg # value [molar Mg/(Mg+Fetot2+) × 100] of 69, relatively high Al2O3 and TiO2 concentrations (16.3 and 1.3 wt %, respectively). The concentrations of Cr and Ni are 260 and 100 ppm, respectively. Its chondrite normalized REE pattern is characterized by slight LREE depletion (Figure 9) relative to MREE and HREE (LaN/SmN = 0.6, for SmN = 20). With respect to average NMORB compositions [Hofmann, 1988], the basalt dike is slightly LREE enriched and HREE depleted. In addition, normalization of incompatible trace elements to average NMORB compositions [Hofmann, 1988] reveals that Zr is slightly enriched over Y (Zr/Y = 4.1) and LREE. Basalt dikes intruding other gabbroic bodies of the Internal Ligurian ophiolites [Rampone et al., 1998; this work] display a similar incompatible element signature (Figure 9).

Figure 9.

REE and incompatible compositions of the basalt dikes Scogna-Rocchetta Vara and Bracco-Levanto ophiolites (Internal Ligurian units), normalized to C1 chondrite [Anders and Ebihara, 1982] and average NMORB compositions [Hofmann, 1988], respectively. One of the basalt dikes from Bracco-Levanto ophiolite is from Rampone et al. [1998].

7. Discussion

7.1. Compositional and Structural Modifications in the Mantle Section

[43] The data presented allowed us to reconstruct a composite tectonomagmatic evolution for the spinel plagioclase mantle section of Scogna-Rocchetta Vara ophiolite before the intrusion of gabbroic rocks, which commenced under spinel facies conditions. In particular, this section deals with the exhumation of the mantle section from spinel to plagioclase facies conditions, which is associated with repeated events of interaction with migrating melts.

7.1.1. The Peridotite Precursor

[44] The studied peridotites contain 4–6 vol % of clinopyroxene (Tables 1a1c), which is characterized by marked depletions in LREE and Zr and nearly flat MREE and HREE (Figure 8a). These features imply a moderately depleted geochemical signature of the spinel facies mantle precursor, which was similarly documented for another mantle sequence of the Internal Ligurian ophiolites [Rampone et al., 1996; Piccardo et al., 2004] and for other mantle bodies of the Alpine-Apennine ophiolites [Rampone et al., 1997, 2008; Piccardo et al., 2007; Tribuzio et al., 2004; Müntener et al., 2004, 2010]. Taken as a whole, these peridotites share many compositional similarities with moderately depleted abyssal peridotites from modern ocean lithosphere [e.g., Johnson et al., 1990; Hellebrand et al., 2001; Brunelli et al., 2006].

[45] The mantle peridotites from Scogna-Rocchetta Vara ophiolite could represent residues after a low degree partial melting of an asthenospheric source [see also Rampone et al., 1996; Tribuzio et al., 2004]. Alternatively, their geochemical signature could have been produced by interaction of the peridotite protoliths with highly depleted, olivine-saturated melts derived from the underlying asthenosphere, as recently proposed for other depleted peridotites from the Alpine-Apennine ophiolites [Piccardo et al., 2007; Rampone et al., 2008]. Note that percolation and crystallization of highly depleted melts in mantle sequences from slow and ultraslow spreading settings have been also recognized on the basis of compositions of clinopyroxenes from peridotites [Seyler et al., 2001] and websterite layers [Dantas et al., 2007].

7.1.2. Origin of the Websterite Layers

[46] The depleted tectonized peridotites from Scogna-Rocchetta Vara ophiolite locally include thin layers made up of spinel websterites. Pyroxenite layering has been documented in other mantle sections from the Alpine-Apennine ophiolites [e.g., Montanini et al., 2006; Rampone and Borghini, 2008]. In particular, Montanini et al. [2006] analyzed pyroxenite layers from the External Ligurian mantle section, which retains a subcontinental origin [Beccaluva et al., 1984; Rampone et al., 1995]. The External Ligurian pyroxenites in places show relics of garnet facies assemblages, thus providing evidence for equilibration at deep lithospheric levels, most likely in pre-Jurassic times [Montanini et al., 2006]. Conversely, the websterites from Scogna-Rocchetta Vara ophiolites do not have garnet relics and do not retain a geochemical signature indicating a formation by precursor garnet-bearing assemblages.

[47] Clinopyroxene from the studied websterites has relatively high concentrations of Na2O and TiO2 and a REE pattern characterized by slight LREE fractionation and nearly flat MREE and HREE, similar to clinopyroxene from the associated gabbroic rocks (Figures 5 and 8a). Clinopyroxene from the websterites is therefore geochemically enriched with respect to clinopyroxene from the tectonized peridotites. In addition, the clinopyroxenes disseminated in the peridotite close to the websterite layer are chemically similar to the clinopyroxenes from the websterite layer. Furthermore, in the peridotite samples containing the websterite layer, porphyroclastic orthopyroxene and spinel are enriched in TiO2 and depleted in Mg # with respect to the other peridotites. The MORB-type melts producing the websterites therefore led to local refertilization of the host peridotite, thus implying that the websterites formed after the depletion event recorded by the tectonized peridotites.

[48] Pyroxenites are also locally found in mantle sections from slow and ultraslow spreading ridges [Fujii, 1990; Juteau et al., 1990; Kempton and Stephens, 1997; Dantas et al., 2007]. In particular, the main petrological features of websterites from Scogna-Rocchetta Vara ophiolite are similar to those of the spinel websterite layers dredged from the Southwest Indian Ridge Oblique Supersegment [Dantas et al., 2007]. The websterites from Southwest Indian Ridge Oblique Supersegment differ in the highly depleted incompatible element signature of the clinopyroxenes (Figure 8). These websterites were interpreted as cumulates produced by melts segregated into veins under spinel facies conditions, although an origin through melt-peridotite reaction could not be excluded [Dantas et al., 2007]. Sobolev et al. [2007] envisaged that olivine-free pyroxenite layers in the upper mantle are produced by reaction between peridotites and silica-rich melts, in turn formed by melting of recycled basalts and gabbros of the oceanic crust transformed to eclogite. However, this petrogenetic process does not match the MORB-type incompatible element signature of clinopyroxenes from the studied websterites, which argues against a high proportion of garnet component in the hypothetical melts reacting with the peridotites.

[49] Experimental determinations shows that crystallization of clinopyroxene + orthopyroxene ± olivine ± spinel from MORB-type melts occur at P ≥ 0.9 GPa [e.g., Elthon, 1993], i.e., under pressure conditions that are compatible with the presence of spinel in the host peridotites [Gasparik, 1984]. The formation of Scogna-Rocchetta Vara websterite layers is therefore attributed to infiltration of MORB-type melts under spinel facies mantle conditions, thus implying the existence of a thick conductive lid. These melts led to local refertilization of the host tectonized peridotites and were most likely produced by the underlying asthenosphere.

7.1.3. Spinel Facies Deformation

[50] The peridotites and websterite layers from Scogna-Rocchetta Vara ophiolite show a foliation produced by the alignment of porphyroclastic orthopyroxene and spinel. In particular, porphyroclastic orthopyroxenes are rimmed by polygonal aggregates consisting of orthopyroxene and clinopyroxene plus minor spinel and olivine (Figures 4a and 4d). CaO compositions of neoblastic orthopyroxene from both peridotites and websterites (1.3–1.5 wt % (Table 3)) yield temperature evaluations of 1090–1150°C on the basis of the Ca-in-Opx geothermometer [Brey and Köhler, 1990]. Therefore, we propose that the spinel facies deformation occurred close to the asthenosphere-lithosphere boundary, which is classically fixed at 1200°C [e.g., Ceuleneer and Rabinowicz, 1992].

[51] Spinel tectonized peridotites are reported for other mantle sections of the Alpine-Apennine ophiolites attributed to ocean-continent transitions [e.g., Vissers et al., 1991; Montanini et al., 2006; Kaczmarek and Müntener, 2008]. They locally constitute up to km-scale thick extensional shear zones that were correlated with uplift of subcontinental mantle in response to the lithospheric extension that led to opening of the Ligurian-Piedmontese basin [see also Vissers et al., 1995]. In particular, the spinel facies deformation in these mantle shear zones was inferred to occur before the interaction of the peridotites with melts derived from the asthenosphere [Piccardo and Vissers, 2007]. In the Scogna-Rocchetta Vara mantle sequence, the spinel facies deformation affected both the depleted peridotites and the included websterite layers, thus implying that it followed the injection of the MORB-type melts that gave rise to the websterite layers (i.e., a depletion event in the underlying asthenosphere).

[52] Mantle sections characterized by a widespread spinel tectonitic foliation are uncommon at slow and ultraslow spreading ridges [e.g., Kelemen et al., 2007; Achenbach et al., 2011]. Nevertheless, the occurrence of spinel facies deformation events in the mantle sections from these settings is documented by grain size reduction of pyroxenes in association with neoblastic spinel, similar to the peridotites of the present study [Jaroslow et al., 1996; Ceuleneer and Cannat, 1997; Brunelli et al., 2006; Dick et al., 2010]. In addition, the websterite layers from Southwest Indian Ridge Oblique Supersegment [Dantas et al., 2007] and those from Scogna-Rocchetta Vara ophiolite show similar spinel-bearing deformation structures. In the oceanic settings, these structures were interpreted to reflect recrystallization and static recovery in conjunction with the mantle flow that followed the accretion of the asthenosphere into the oceanic lithosphere [Jaroslow et al., 1996; Ceuleneer and Cannat, 1997].

[53] The development of the spinel tectonitic foliation in the Scogna-Rocchetta Vara mantle section is attributed to an exhumation process that involved the lower lithospheric mantle. We propose that the spinel facies deformation followed the emplacement of asthenospheric material at the base of the lithosphere, most likely in the Jurassic. We cannot define, however, whether the studied mantle was accreted to a thick oceanic lithosphere or it was incorporated beneath an extending subcontinental lithosphere.

7.1.4. Melt Impregnation Under Plagioclase Facies Conditions

[54] The tectonized peridotites from Scogna-Rocchetta Vara ophiolite contain a high modal percentage of plagioclase (7–15 vol % (Tables 1a1c)), which suggests the addition of a Al2O3-rich component (i.e., a melt) to the peridotite. The plagioclase commonly develop orthopyroxene-bearing veinlets that are subparallel to the spinel facies foliation planes (Figure 2c). These veinlets frequently have a length exceeding the length of the porphyroclastic minerals. It is therefore unlikely that the plagioclase formed by subsolidus reaction between mineral grains (i.e., orthopyroxene, clinopyroxene and spinel) in mutual contact. Similar veinlets were reported for a nearly undeformed mantle section from the Internal Ligurian ophiolites and were interpreted as crystallization products of orthopyroxene-saturated melts percolating through the peridotites [Rampone et al., 1997]. Infiltration and impregnation of peridotites by orthopyroxene-saturated melts in the plagioclase stability field was also proposed for other mantle sections of the Alpine-Apennine ophiolites [e.g., Piccardo et al., 2007; Rampone et al., 2008; Müntener et al., 2010]. A similar petrogenetic process explains the formation of the plagioclase-rich veinlets in the Scogna-Rocchetta Vara mantle section. In particular, the studied tectonized peridotites show that the spinel facies foliation planes favored the infiltration of these melts.

[55] The process of melt impregnation under plagioclase facies conditions is also shown by a few geochemical features of the studied peridotites. For instance, the high Cr # of porphyroclastic and neoblastic spinel (Figure 7) and of associated pyroxenes (Tables 35) hypothetically correspond to a highly refractory mantle peridotite [Hellebrand et al., 2001; Brunelli et al., 2006], thus arguing against the moderately depleted geochemical signature indicated by the clinopyroxene. We attribute the high Cr # of spinel, orthopyroxene and clinopyroxene from Scogna-Rocchetta Vara peridotites to chemical modifications occurring in response to a reaction with infiltrating melts in the plagioclase stability field, as typically documented for plagioclase-bearing peridotites from (ultra)slow spreading ridges [e.g., Kelemen et al., 2007; Dick et al., 2010]. For instance, the moderately depleted plagioclase-bearing peridotites from the Romanche Fracture Zone (equatorial Atlantic) have spinel and pyroxenes with high Cr # [Tartarotti et al., 2002], similar to the peridotites of the present study.

[56] The plagioclase-rich veinlets are common along the contacts between the tectonized peridotites and the websterite layers, with frequent apophyses within the websterites. We thus also argue that also the websterites were subjected to melt infiltration in the plagioclase stability field and led to chemical reequilibration of both porphyroclastic and neoblastic minerals. A chemical reequilibration of clinopyroxene from the peridotites and included websterites is consistent with its marked negative Sr anomaly, and with the fact that the deepest Sr negative anomaly is associated with the development of a negative Eu anomaly (Figure 8a). In addition, the low concentrations of Na2O in the clinopyroxene from the peridotites (0.2 wt %), which hypothetically imply a highly refractory nature for these mantle rocks [Hellebrand and Snow, 2003], are most likely partly related to the incorporation of Na2O into newly formed plagioclase [see also Tribuzio et al., 2004; Müntener et al., 2010].

[57] A thorough geochemical study of plagioclase-rich veinlets from the tectonized peridotites of Scogna-Rocchetta Vara ophiolite is hampered by the extensive alteration of primary minerals. Rampone et al. [1997] showed that the orthopyroxene-saturated melts impregnating another mantle section of the Internal Ligurian ophiolites were depleted in incompatible elements with respect to typical NMORB compositions [Rampone et al., 1997]. We propose that the melts impregnating Scogna-Rocchetta Vara mantle section had a similar geochemical fingerprint, which is consistent with the depleted signature retained by the spinel facies clinopyroxenes. If the impregnating melts had a MORB-type signature, the depleted clinopyroxenes would record a significant increase in the concentrations of most incompatible trace elements (e.g., the LREE), thus leading to a chemical refertilization of the host peridotite [Müntener et al., 2010]. The melts impregnating the studied peridotites could be genetically related to silica-undersaturated, highly depleted melts that had previously reacted with the mantle sequence, thus producing residual melts enriched in silica by preferential dissolution of pyroxenes. In particular, the geochemical depletion of the percolating melts may be attributed to low degree (5%–7%) fractional melting of a slightly depleted spinel-bearing mantle source [Piccardo et al., 2007; Rampone et al., 2008].

7.1.5. Formation of Dunitic Conduits

[58] The tectonized plagioclase peridotites are locally replaced by meter-scale dunite bodies with spinel trails. These bodies are subparallel to the spinel facies foliation planes and the plagioclase-rich veinlets in the host rocks (Figures 2d and 3). There is a general consensus that replacive dunites form by pyroxene reactive dissolution by olivine-saturated melts migrating in the peridotites [e.g., Kelemen et al., 1997]. The spinel concentrations in replacive dunites are commonly attributed to precipitation from hybrid melts formed by mixing between primitive basaltic melts and secondary melts enriched in SiO2 and Cr2O3 [Irvine, 1977; Dick and Bullen, 1984; Arai and Yurimoto, 1994; Zhou et al., 1994]. The secondary melts are produced in response to the formation of the replacive dunite, when primitive basalts migrate upward and dissolve pyroxenes from the peridotites. The dunite bodies from Scogna-Rocchetta Vara ophiolite may thus be considered as conduits of olivine-saturated melt.

[59] Dunite bodies crosscutting plagioclase-impregnated peridotites were found in other depleted mantle sections from the Alpine-Apennine ophiolites and related to channeled porous flow of olivine-saturated melts [Müntener and Piccardo, 2003; Piccardo et al., 2007; Rampone et al. 2008]. In particular, on the basis of clinopyroxene incompatible element compositions, Piccardo et al. [2007] showed that the replacive dunites from one of these sections formed by infiltration of melts with MORB-type geochemical signature. These MORB-derived dunites are characterized by spinels with relatively high Cr # and TiO2 concentrations (Figure 7), as also reported for the replacive dunites in equilibrium with MORB-type melts from the Oman ophiolite [Kelemen et al., 1997]. The spinels from the dunites considered in this study differ in the lower Cr # and TiO2 concentrations, with respect to the spinels from the MORB-type dikes of the Internal Ligurian ophiolites [Cortesogno and Gaggero, 1992]. These data argue against the involvement of typical MORB-type melts for the formation of the replacive dunites from Scogna-Rocchetta Vara ophiolite. In particular, the low concentrations of TiO2 in the spinels from the replacive dunites indicate an equilibration with melts depleted in incompatible trace elements relative to MORB.

[60] In the Scogna-Rocchetta Vara mantle section, the pyroxene-dissolving melts forming the dunites and the orthopyroxene-saturated, highly depleted melts that led to plagioclase crystallization in the tectonized peridotites could be attributed to the same event. However, the replacive structural relationships indicate that the plagioclase-forming melts in the tectonized peridotites are not generated by the physically associated dunite bodies. Impregnation of the observed peridotites is attributed to the occurrence of dunitic conduits deeper in the mantle section. These high-permeability conduits may have prevented new infiltrating melts from reacting with the host tectonized peridotites [see also Kelemen et al., 1995]. The prolonged migration of these olivine-saturated melts presumably allowed the interaction with the plagioclase-impregnated peridotites.

7.2. Formation of the Gabbroic Rocks

[61] We recognized a composite evolution for the migration of melts that led to the building of the gabbroic crust from Scogna-Rocchetta Vara ophiolite, which started with the intrusion of troctolite to olivine gabbro dikes into the mantle sequence.

7.2.1. Intrusion of Gabbroic Dikes and Sills Into the Mantle Sequence

[62] The gabbroic dikes crosscut at a low angle the foliation planes in mantle peridotites and range in composition from troctolites to olivine gabbros. Locally, the dikes also crosscut the replacive dunite bodies. The troctolite to olivine gabbro dikes commonly show diffuse boundaries with respect to the host peridotites, thus indicating that the melt injections occurred when the mantle section was under high-temperature conditions, presumably slightly lower than the crystallization temperature of the troctolites (∼1200°C) [e.g., Grove et al., 1992].

[63] The gabbroic sills occur near the contact between the mantle sequence and the gabbroic pluton (Figure 1d) and are olivine free to olivine poor. The sills show sharp contacts against the host tectonized peridotites, thus attesting that they formed within a mantle sequence that was cooler than their crystallization temperature. In addition, the gabbroic sills in places crosscut the troctolite to olivine gabbro dikes. The temperature of the mantle sequence hence progressively decreased from the intrusion of the troctolitic dikes to the formation of the gabbroic sills.

[64] We propose that the dikes and sills document the progressive transfer from a “hot lithospheric regime,” where melt migration and crystallization is controlled by reaction with the mantle peridotites, to a “cold lithospheric regime,” in which melt migration occurs in fractures and melts evolve through fractional crystallization. This change was associated with a rotation of the dip of the melt migration structures, which evolved from subvertical to subhorizontal.

7.2.2. Growth of the Gabbroic Pluton

[65] The gabbroic pluton consists mostly of coarse-grained clinopyroxene-rich gabbros, associated with minor, medium grained olivine gabbros to troctolites. The different gabbro types constituting the pluton do not show systematic modal layering. However, they locally display a weak modal and/or grain size layering, which is at a high angle with respect to the mantle structures and subparallel to the orientation of the gabbroic sills. In addition, the troctolites show a magmatic foliation produced by alignment of olivine and plagioclase grains, which is nearly concordant with the modal/grain size layering. Trace element compositions of clinopyroxene from the gabbroic rocks show formation from MORB-type melts, similar to the gabbroic rocks from the dikes and sills (Figures 8b and 8c). The Scogna-Rocchetta Vara gabbroic pluton also contains up to 75 m thick bodies of olivine-rich troctolites, which were interpreted to result from a reaction process driven by the infiltration of MORB-type melts saturated in plagioclase + clinopyroxene into an olivine-rich, spinel-bearing matrix [Renna and Tribuzio, 2011]. In addition, we found up to 50 m thick lens-like mantle bodies within the gabbroic pluton.

[66] The main structural and compositional features of Scogna-Rocchetta Vara gabbroic pluton are similar to those of other gabbroic bodies from the Internal Ligurian ophiolites [Cortesogno et al., 1987; Rampone et al., 1998; Tribuzio et al., 2000; Menna, 2009], as well as of those of many gabbroic sequences exposed along slow and ultraslow spreading ridges. Notable examples are the gabbroic sequences from (1) Kane Fracture Zone of Mid Atlantic Ridge 23°N [Ross and Elthon, 1997; Dick et al., 2008; Lissenberg and Dick, 2008], (2) 15°N section of Mid Atlantic Ridge [Kelemen et al., 2007], (3) Atlantis Massif (Mid Atlantic Ridge, 30°N) [Blackman et al., 2006], (4) Atlantis Bank at Southwest Indian Ridge [Robinson et al., 1989; Dick et al., 2000], and (5) Mid Cayman Rise [Elthon, 1987]. In particular, we wish to emphasize that the gabbroic complexes of Atlantis Massif [Suhr et al., 2008; Tamura et al., 2008; Drouin et al., 2009] and Scogna Rocchetta Vara share the inclusion of olivine-rich troctolites and mantle peridotite bodies.

[67] The most primitive gabbroic rocks of the Scogna-Rocchetta Vara pluton (Mg # in clinopyroxene cores ranging between 85 and 87) have high concentrations of Na2O in plagioclase (An = 60–61 mol % (Table 6)). Plagioclase with a low proportion of anorthite component was also reported for primitive olivine gabbros and troctolites from another gabbroic sequence of the Internal Ligurian ophiolites [Tiepolo et al., 1997; Rampone et al., 1998]. This chemical feature was also documented for abyssal gabbroic rocks from Mid-Cayman Rise [Elthon, 1987], where high Na2O abundances are further recognized in associated glasses [Thompson et al., 1980]. Conversely, many abyssal gabbroic sequences have primitive olivine gabbros and troctolites with plagioclase displaying 65–75 mol % of anorthite component [e.g., Suhr et al., 2008]. We thus propose that the parental melts of Scogna Rocchetta Vara gabbroic pluton had high Na2O concentrations, which may be correlated with a low degree of melting of an asthenospheric source [see also Dick et al., 1984; Meyer et al., 1989; Kempton and Casey, 1997].

[68] The mantle bodies from Scogna-Rocchetta Vara gabbroic pluton show the same origin and tectonomagmatic evolution of the mantle sequence enclosing the gabbroic pluton (Figure 1c). In particular, the structures of the mantle lenses and the enclosing mantle sequence are geometrically concordant. The occurrence of these mantle remnants together with the presence of the gabbroic sills in the mantle sequence overlying the gabbroic pluton (Figure 1d) may be reconciled with a process of pluton growth through a series of sill-like separate intrusions. The process of sill accretion is also consistent with the occurrence of troctolite layers at different depths within the gabbroic pluton.

[69] The sill accretion model is currently proposed for the building of gabbroic plutons exposed along slow and ultraslow spreading ridges, where a steady state magma chamber is considered to be absent [e.g., Kelemen et al., 2007; Godard et al., 2009]. In particular, field observations (troctolite layers up to a few tens meters in thickness) are consistent with the model of Grimes et al. [2008] for the construction of the gabbroic complex from Atlantis Massif, which imply that single sills are on the average ten meters thick. However, we cannot define whether the sills represent melt injections intruded at random depths [Schwartz et al., 2005; Grimes et al., 2008], or they formed at nearly constant depths beneath an exhuming sequence [Dick et al., 2000, 2002; Suhr et al., 2008]. Note that the composition of the gabbroic pluton (i.e., made up of clinopyroxene-rich gabbros and minor troctolites) requires the presence of primitive cumulates deeper in the mantle section.

7.2.3. Origin of the Olivine-Rich Troctolites

[70] Renna and Tribuzio [2011] proposed that the olivine-spinel matrix of the olivine-rich troctolites formed in mantle melt conduits of replacive nature. One of the arguments against a cumulate origin of the olivine-rich troctolites was furnished by the Rocchetta Vara section (Figure 1c), where the olivine-rich troctolites are exposed along the contact between the gabbroic complex and the overlying mantle section. There are no evolved gabbroic rocks overlying the olivine-rich troctolites, thus indicating the absence of the hypothetical residual melts after the cumulus process [Renna and Tribuzio, 2011].

[71] The new data reported in this work are consistent with the interpretation proposed by Renna and Tribuzio [2011]. We observed that the replacive dunites from the studied ophiolite are locally crosscut by gabbroic dikes with diffuse contacts and that the olivine-rich troctolites are locally crosscut by gabbroic sills with sharp planar contacts (Figure 1d). The geochemical signature of clinopyroxene from the gabbroic dikes and the olivine-rich troctolites similarly show a formation by MORB-type melts (Figure 8), whereas the dunite-forming melts and the melts impregnating the mantle section were most likely depleted in incompatible elements with respect to typical NMORB compositions. Therefore, the olivine-rich troctolite bodies may have formed in conjunction with the onset of the gabbroic dyking, most likely when large amounts of MORB-type melts were injected into dunitic conduits. Note that the mantle section overlying the olivine-rich troctolites in the Rocchetta Vara section contains meter-scale replacive dunite bodies, which could be originally connected to the large dunitic conduits (order of tens of meters in scale) presently represented by the olivine-rich troctolites. We suggest that these large dunitic conduits impregnated by MORB type melts were dissected by the gabbroic sills, similar to the explanation of the occurrence of mantle lenses within the gabbroic complex.

7.3. The Tectonomagmatic Evolution Leading to Exposure of the Gabbro-Peridotite Association at the Seafloor

[72] This section considers the tectonomagmatic evolution recorded by the gabbroic pluton after its solidification, which comprises ductile to brittle shearing. The brittle deformation was associated with injections of basalt dikes and led to exposure of the gabbro-peridotite association at the seafloor, as shown by the calcite- and hematite-bearing structures in both gabbroic and mantle rocks at the contact with the sedimentary cover.

7.3.1. Ductile Deformation in the Gabbroic Pluton

[73] The gabbroic pluton was affected by ductile deformation along localized shear zones. The shearing foliation crosscuts at a low angle the modal/grain size layering of the gabbros and the magmatic foliation of the troctolites. The shear zones show a retrograde evolution characterized by early recrystallization of clinopyroxene + plagioclase (±accessory Ti pargasite) at ∼850°C, followed by a hornblende + plagioclase amphibolite facies event at ∼710°C. Neoblastic hornblende occurs in high modal proportions and contains significant amounts of Cl (0.2 wt % (Table 7)), thus suggesting that downward infiltration of seawater-derived fluids occurred along the amphibolite facies shear zones. Sheared gabbros with similar microstructural and compositional features were reported for other gabbroic bodies from the Internal Ligurian ophiolites [Molli, 1995, 1996; Tribuzio et al., 1995; 2000; Menna, 2009]. In a few cases, the high-temperature shearing affecting the gabbros was shown to also involve the associated mantle peridotites, with development of plagioclase-bearing mylonitic peridotites [see also Cortesogno et al., 1987].

[74] In gabbroic sections from modern (ultra)slow spreading ridges, sheared gabbros are commonly localized in discrete zones that overprint the magmatic foliation [e.g., Mével, 1987; Cortesogno et al., 2000; Dick et al., 2000; Escartín et al., 2003; Ildefonse et al., 2007]. Most of these gabbros are characterized by neoblastic aggregates of pyroxene (±accessory Ti pargasite) and plagioclase, that are considered to have been formed under high-temperature conditions (T > 800°C) before seawater penetration (just after magmatic crystallization). At the 15°N section of Mid Atlantic Ridge, the ductile shearing locally affected both the gabbroic and mantle sequences, and formed plagioclase-bearing mylonitic peridotites in the mantle section [Kelemen et al., 2007; Achenbach et al., 2011]. Amphibolite facies shearing was observed in some of these oceanic sections, e.g., Hole 735B at Atlantis II Bank at Southwest Indian Ridge [MacLeod et al., 1999; Dick et al., 2000], Mid Cayman Rise [Ito and Anderson, 1983] and Vema fracture zone of Mid Atlantic Ridge [Honnorez et al., 1984] and interpreted to act as pathways for seawater-derived fluids.

[75] The foliated sheared gabbros from the Scogna-Rocchetta Vara ophiolites therefore show striking structural and compositional similarities with their counterparts from modern (ultra)slow spreading settings. We recognized a retrograde evolution from near solidus to amphibolite facies conditions, which is consistent with the idea proposed by Manatschal et al. [2011] that the ductile shear zones act as a decoupling horizon at the ductile-brittle transition. The sheared gabbros thus probably formed at the interface between a magma-rich ductile layer and a fluid-rich brittle layer.

7.3.2. Intrusion of Basalt Dikes

[76] The Scogna-Rocchetta Vara gabbroic pluton is locally crosscut by chilled basalt dikes forming a high angle with respect to the intrusive fabric of enclosing gabbros. Dikes were therefore injected into a solidified gabbroic pluton, i.e., when the gabbro-peridotite association was subjected to a brittle deformation regime. Similar relationships were reported for another gabbroic pluton of the Internal Ligurian ophiolites [Cortesogno et al., 1987; Menna, 2009]. Taken as a whole, the basalt dikes from the Internal Ligurian ophiolites share a common incompatible element signature (see also Figure 9).

[77] Clinopyroxene phenocrysts from the basalts dikes and clinopyroxene cores from gabbroic rocks show the same trace element signature, which implies a formation by MORB-type melts with a similar geochemical fingerprint (Figures 8c and 8d). Clinopyroxene from the basalt groundmass shows an incompatible element pattern that is nearly parallel to that of the clinopyroxene phenocrysts, but at higher absolute concentrations (Figure 8d). We attribute these variations to fractional crystallization controlled by plagioclase, olivine and clinopyroxene, as indicated by the occurrence of these minerals as phenocrysts. In particular, the involvement of plagioclase in the crystallization process is confirmed by the development of negative Eu and Sr anomalies in the clinopyroxene from the groundmass.

[78] Whole-rock compositions of the Internal Ligurian basalts are slightly LREE- and Zr-enriched relative to typical NMORB (Figure 9). Similar compositions [Kempton and Casey, 1997] were reported for basalt dikes crosscutting mantle peridotites from Site 920 at Kane Fracture Zone (Mid Atlantic Ridge). This incompatible element signature was attributed to low degree partial melting of a spinel peridotite mantle source [Kempton and Casey, 1997]. The composition of basalt the dikes and the host gabbros (section 7.2.2) from the Scogna-Rocchetta Vara mantle ophiolite therefore concordantly indicate a formation by a low degree of melting of an asthenospheric mantle source. Note that the slight depletion of HREE relative to MREE of the Internal Ligurian basalts may indicate the involvement of a minor garnet-bearing component in their mantle sources [see Hirschmann and Stolper, 1996].

8. Conclusions

[79] The gabbro-peridotite association from Scogna-Rocchetta Vara Jurassic ophiolite share many structural and compositional similarities with melt-poor sections from modern (ultra)slow spreading settings, which are characterized by an elevated lithospheric thickness [e.g., Cannat et al., 1997, 2006; Kelemen et al., 2007; Schroeder et al., 2007]. The series of magmatic and tectonic events recorded by the studied gabbro-peridotite association have allowed us to propose a conceptual model (Figure 10) for its formation and evolution, which is summarized as follows:

Figure 10.

Conceptual model for the tectonomagmatic evolution of Scogna-Rocchetta Vara ophiolite. Websterites, dunites, gabbroic dikes, and sills are exaggerated in scale. The transition between spinel and plagioclase peridotite is assumed to be 0.7 GPa [Gasparik, 1984]. See text for further details.

[80] 1. The mantle sequence was made up of moderately depleted peridotites and scattered spinel websterite layers formed by infiltrations of MORB-type melts close to the lithosphere-asthenosphere boundary.

[81] 2. A deformation event affected the mantle section after the formation of the websterite layers. This deformation led to uplift of the mantle sequence and is probably related to the emplacement of asthenospheric material at the base of the lithosphere.

[82] 3. At shallower levels (in the plagioclase stability field), the mantle sequence was subjected to further infiltrations of melts, which were most likely depleted with respect to typical MORB compositions. Melt transport occurred in form of grain-scale porous flow, either channeled or diffuse, and was favored by the spinel facies foliation planes. Reactive channeling of primitive olivine-saturated melts formed replacive dunitic conduits, whereas residual orthopyroxene-saturated melts led to melt impregnation of the mantle section.

[83] 4. New injections of melts displaying a MORB-type geochemical signature formed troctolite to olivine gabbro dikes, documenting that melt migration started to occur through fracture-controlled mechanisms. This diking event is presumably correlated with the formation of the olivine-rich troctolite bodies, by infiltration of MORB-type melts within large dunitic conduits.

[84] 5. As the mantle section cooled significantly, the dip of the melt migration structures evolved from subvertical to subhorizontal. Injected melts thus developed sill-like gabbroic bodies.

[85] 6. The growth of the gabbroic pluton (up to ∼400 m thick) is attributed to a process of accretion of gabbroic sills. Hence, the mantle section was dissected by the gabbroic sills and partly incorporated by the gabbroic intrusions.

[86] 7. The gabbroic pluton records ductile shearing at a low angle with respect to the intrusive fabric. It shows a retrograde evolution from near solidus to amphibolite facies conditions; the latter stage was most likely associated with penetration of seawater-derived fluids. This deformation most likely occurred close to the ductile-brittle transition.

[87] 8. The basalt diking represents the last event of melt injection and occurred during exhumation of the gabbro-peridotite association to the seafloor, when it was under brittle regime conditions.

[88] The conceptual model proposed shows a “hot” lithospheric evolution in which melt migration is associated with reactive crystallization (steps 1 to 3). This “hot” evolution implies that uprising melts may be trapped within the lithospheric mantle. The following “cold” lithospheric evolution is characterized by melt transport through fractures and is associated with fractional crystallization (steps 4 to 8). Taken as a whole, the melt transport evolution recognized for the Scogna-Rocchetta Vara ophiolite is consistent with the petrogenetic model recently reported by Collier and Kelemen [2010] for the melt migration mechanisms under oceanic spreading ridges.

Acknowledgments

[89] We would like to acknowledge R. Vannucci for constructive discussions and for the support in this study. We are grateful to M. Tiepolo for suggestions and assistance during LA-ICP-MS analyses and to M. R. Renna for the discussions on the origin of the dunites and troctolites. Conversations with G. B. Piccardo enhanced our understanding of the petrology of the mantle sequences from the Alpine-Apennine ophiolites. The paper was improved by insightful reviews by K. Achenbach, E. Hellebrand, and T. Morishita. We also wish to thank the Associate Editor M. Cheadle for his recommendations that enhanced the revision of the manuscript. This work was financially supported by Programma di Ricerca di Interesse Nazionale of Italian Ministero dell'Università e della Ricerca and Fondi di Ateneo per la Ricerca of Università di Pavia.

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