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The pre-Cenozoic tectonic evolution of the FMC
The evolution of the Variscan belt can be regarded as a typical Wilson cycle, exhibiting subduction and extinction of oceanic basins and collision of continental plates (Fig. 2; e.g., Behr et al., 1984; Matte, 1986, 1991; Pin, 1990; Faure et al., 2009). Prior to the onset of the Variscan orogeny, the Rheic Ocean firstly emerged in the early Ordovician and separated the two ancient continents Gonwana and Laurussia (Faure et al., 2009; Nance et al., 2010).
It continued expanding at the expense of the Iapetus Ocean, which subducted beneath Armorica (Matt, 1886, 1991; Faure et al., 2009). The Rheic Ocean achieved its largest width (~4000 km) in the Silurian and then commenced closing and subducting southward in the Devonian (Nance et al., 2010). The extinction of the Rheic Ocean marked the onset of the Variscan orogeny, i.e., the collision between Gondwana and Laurussia during Devonian-Carboniferous times (Lardeaux et al., 2001; Nance et al., 2010), followed by two large-scale extension events and the gravitational collapse inside the continent (Faure et al., 2009). Many plutons distributing in the FMC (e.g., mostly granites) developed during this stage. After that, the sedimentation processes gradually predominated and filled the extension-induced basins through the Mesozoic. Then intense volcanic activities related to the West-European Rift system occurred during Cenozoic times (Ziegler, 1992; Michon and Merle, 2001).
The Cenozoic volcanism in the FMC
The Cenozoic volcanism within the FMC made it the largest magmatic province in the Western Europe (Michon and Merle, 2001). The volcanic activitiy started in the Eocene and lasted until the most recent volcanism in the Chaîne des Puys ~6,900 yr ago (Juvigné, 1992). The spatial distribution of volcanic activities spread over several regions (Fig. 2-2), such as Cantal, Devès, Velay Oriental, Aubrac, Mont-Dore, and Chaîne des Puys. The entire volcanism has been classified as three magmatic phases: pre-rift, rift-related, and major magmatic events (Michon and Merle, 2001). They have outcrops of different volumes in different localities. The pre-rift phase started in the late Cretaceous and was active in the Paleocene and Eocene. It has minor volumes and mainly scatters in the northern part of the FMC. The lavas erupting in this stage are under-saturated (alkali basalt to melilitite; Wilson and Downes, 1991; Michon and Merle, 2001). The rift-related volcanism is spatially restricted in the northern part of the FMC and could be ascribed to low degrees partial melting resulting from decompression during extension. In the Oligocene, the crust in the northern part of the FMC underwent significant thinning associated with graben fromation, especially the Limagne graben (Merle et al., 1998; Michon and Merle, 2001, whereas the crust thinning was very subtle in the southern part of the FMC and no related volcanic activity has been reported there. The lavas associated with this phase are mainly alkaline basalts and nephelinite. The major magmatic phase commenced at different times in the northern and southern parts of the FMC (5.5 Ma and 15 Ma ago, respectively; Michon and Merle, 2001). It lasted from Miocene till recently and produced the few large magmatic provinces presently observed in the FMC (Fig. 2-2).
The lavas have alkali basaltic compositions and sometimes exhibit two series. Based on geochemical and Sr-Nd-Pb isotopic compositions, Wilson and Downes (1991) proposed that these lavas could be regarded as binary mixtures of two components of different origins: the primary magmas and the lithospheric mantle metasomatized during Variscan orogeny. The primary magmas probably had signatures between DM and HIMU, which has been explained by partial melting of subducted oceanic lithosphere during Variscan orogeny after it entered the convecting asthenosphere. Then they were contaminated by phlogopite or amphibole-bearing lithospheric mantle during ascent towards the surface. Meantime, due to the harmony of major magmatic localities and the thermal anomalies underneath the FMC, the extensive volcanism during this stage has been considered to be prompted by an asthenospheric upwelling (Michon and Merle, 2001). Recently, it has been also suggested that the presence of fertile lithologies (pyroxenites and/or amphibolites) in the FMC mantle had also enhanced the volcanism (Pilet et al., 2008; France et al., 2015).
The existence of thermal anomalies underneath the FMC has been continuously verified by results of seismic tomography (Fig. 2-3; Hoernle et al., 1995; Granet et al., 1995; Sobolev et al., 1997; Zeyen et al., 1997; Goes et al., 1999; Fichtner and Villaseñor, 2015), but different interpretations of this observation have been given.
On the one hand, diapiric mantle upwelling deriving from an older mantle plume, whose head spread at a depth of 400 km, have been proposed to account for these thermal anomalies (Granet et al., 1995). Goes et al. (1999) have also observed a low-velocity anomaly located at the depth of 660 to 2000 km beneath central Europe and suggest that it had played the role of the root for mantle upwellings interspersing under western and central Europe. On the other hand, these thermal anomalies have been ascribed as hotspots that originated from the upper mantle only to the bottom of the transition zone, and not from the lower mantle (Hoernle et al., 1995). A recent full-waveform tomographic model from Fichtner and Villaseñor (2015) has not shown the presence of mantle plumes beneath the FMC and they discussed the previous plume signatures observed in tomographic images as an artifact brought by vertical smearing in the upper mantle (Fichtner and Villaseñor, 2015). Alternatively, the mantle upwellings were temptatively associated with slab subduction or extension-induced lithosphere thinning during Alpine orogeny (Sobolev et al., 1997; Goes et al., 1999). In spite of the debate, upwellings of deep materials hotter than ambient lithosphere had brought heat and resulted in thinning and partial melting of lithosphere to generate a few magmatic provinces that carried large quantities of xenoliths from various depths to the surface.
The lithospheric mantle beneath the FMC
In the past decades, peridotite or pyroxenite xenoliths have been used to decipher the composition and evolution of the lithospheric mantle beneath the FMC (e.g., Mercier and Nicolas, 1975; Downes, 1987; Downes and Dupuy, 1987; Nicolas et al., 1987; Werling and Altherr, 1997; Zangana et al., 1997, 1999; Xu et al., 1998; Lenoir et al., 2000; Lorand and Alard, 2001; Lorand et al., 2003; Downes et al., 2003, 2015; Féménias et al., 2003, 2004, 2005; Berger et al., 2005, 2007; Wittig et al., 2006, 2007; Touron et al., 2008; Yoshikawa et al., 2010; Harvey et al., 2010; Uenver-Thiele et al., 2014; France et al., 2015).
Among these, Lenoir et al. (2000) integrated textural observations with bulk rock major and trace element compositions of peridotite xenoliths widely distributed in the FMC (25 volcanic centers) and identified two distinct lithospheric mantle domains to the North and South of a boundary at ~45°30’N (Fig. 2-2). They suggest that the northern domain could be considerd as a cratonic sub-continental lithospheric mantle. The southern domain lies surrounding the northern one and they merged together during the Variscan orogeny. Peridotites for the northern domain are mainly protogranular in texture, more refractory in chemical compositions with Sr-Nd isotopic compositions largely overlapping the compositional field of the European Asthenospheric Reservoir, whereas peridotites from the southern domain are coarse-granular, more fertile with MORB-like Sr-Nd isotopic compositions (Lenoir et al., 2000; Downes et al., 2003). Both domains have been metasomatized to different degrees, with highly incompatible element and LREE enrichment more pronounced in the northern domain (Lenoir et al., 2000; Downes et al., 2003). Additionally, remarkable HFSE negative anomalies as an indicator of carbonatitic metasomatism have often been observed in xenoliths from the northern domain, but not in xenoliths from the southern domain (Lenoir et al., 2000; Downes et al., 2003; Witting et al., 2007). Thus, it has been proposed that the northern and southern domain of the FMC have been metasomatized by different types of agents, fluid/carbonatitic melt and silicate melt, respectively (Wittig et al., 2007). However, the metasomatic agents recorded in peridotites and pyroxenites from the FMC have sometimes been attributed to a deep, recycled, and enriched component, possibly related to Variscan subduction (Downes and Dupuy, 1987; Touron et al., 2008; Yoshikawa et al., 2010). In-situ measurements of amphibole D/H ratios in peridotites from the FMC have presented high values, which were associated with recycled surficial materials through Variscan subduction (Deloule et al., 1991).
The sampling localities
Peridotite xenoliths studied herein were sampled in two localities, both belonging to the southern FMC domain (Fig. 2-2). One is the Ringue quarry in Allègre (Fig. 2-4a), situated in a solidified lava lake that filled a maar ~3 Ma ago. In the field, xenoliths from different levels were found, including granite, gneisses, and spinel peridotites.
The other is the Razas Grand quarry in Mont Coupet (Fig. 2-4b). The volcano erupted in strombolian type 2 Ma years ago. Many wonderful volcanic bombs of different sizes containing peridotites xenoliths have been collected in this quarry.
Laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS)
Cpx and Amp trace element compositions were determined using LA-ICP-MS at GeoRessources Laboratory, University of Lorraine (France). Double-polished thick sections (~0.15mm thick) or thin sections were ablated in-situ using a nanosecond excimer laser (GEOLAS Pro; 193 nm wavelength) with a spot size of 44 µm at 5 Hz with an energy density of 10 J/cm2 per pulse. The ablation products were transported in a helium flow, mixed with argon gas, and then analyzed with an Agilent 7500 ICP-MS. A complete analysis comprises 30s for background acquisition and 50s for sample acquisition. The raw intensities of ions were recorded as a function of time. SiO2 contents obtained from EPMA were used as an internal standard; NIST 612 and 614, analyzed at the beginning and end of the analytical session, were used as external standards. Some minor elements in Ol from Mont Coupet samples, such as Al, Ca, Ti, Cr and Ni, were also analyzed following the same procedure.
Secondary ion mass spectrometry (SIMS)
The secondary ion mass spectrometry is an analytical technique with the ability to quantitatively detect in situ all the elements (especially for light elements, e.g., H, Li, Be, B) in the chemical periodic table with detection limits reaching part per million or below. The instrument mainly consists of a primary ion source, a mass spectrometer and a secondary ion detection system (Fig. 3-2). In present, the obstacle to the extensive application of SIMS is the dependence on standards to quantify the matrix effects strongly dependent on the chemical compositions of the analyzed materials, and the high cost for this equipment.
The primary ions, generally O- or Cs+ ions for geological materials (Fig. 3-3), are extracted from the sources and focused by a series of electrostatic lenses (L1, L2, L3 and L4 in Fig. 3-4a) onto the sample surface. These lenses can be adjusted to get the appropriate intensity and focusing for measuring. During an entire analytical session, the intensity of the primary ion beam usually fluctuates. If it deviates too much from the initial value, it should be adjusted again because its variation has a significant influence on the instrumental matrix fractionation (Fitzsimons et al., 2000; Hauri et al., 2006a). When the primary ion beam bombards the sample surface, ions will partly transfer their high energy (4 to 20 keV) to the matrix atoms by collision. The atoms that obtain energy in excess of their binding energy will break away from the matrix and be sputtered (Fig. 3-3). Many of the released atoms are neutral while parts of them are charged as ions (Fig. 3-3) that can be accelerated by an electric field and transferred to the mass spectrometer. These secondary ions subsequently can be selectively positive or negative depending on the polarity of the accelerating voltage. Here, it’s worth noting that not only are the target ions sputtered by the primary beam, but also ions of the other elements froming the matrix and even complex ions formed by combination of two or more atoms of different elements. Commonly, complex ions with mass close to that of target ions can be produced and go through the entrance slit, which requires a mass spectrometer at high mass resolution to distinguish the ions concerned from the interference complex ions. It is achieved by combination of an electrostatic filter and a magnet sector, forming a double-focusing mass spectrometer (Fig. 3-2). The mass resolution can be controlled by adjusting the width of the entrance slit and the exit slit, which will significantly influence the intensities of secondary ions ultimately reaching the collectors. Some technical ways can be adopted to raise the secondary ion transmission (details referring to Hilton (1995)).
Table of contents :
1.1 The lithospheric mantle
1.2.1 Geochemical features
1.2.2 The current research state of Li in the lithospheric mantle
1.3 water in the NAMs
1.3.1 Water in the earth mantle
1.3.2 The current research state of water in the NAMs
1.4 Objective of this thesis
2. Geological background
2.1 The pre-Cenozoic tectonic evolution of the FMC
2.2 The Cenozoic volcanism in the FMC
2.3 The lithospheric mantle beneath the FMC
2.4 The sampling localities
3. Analytical methods
3.1 Sample preparations
3.2 Electron micro-probe analysis (EPMA)
3.3 Laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS)
3.4 Secondary ion mass spectrometry (SIMS)
3.4.1 Basic working principles
3.4.2 Standards for Li isotopes measuring by SIMS
3.4.3 Measuring procedure for Li isotopes by SIMS
3.4.4 Measuring procedure for water content by SIMS
3.5 Fourier transform infrared (FTIR) spectrometer
3.5.1 The rationale
3.5.2 Measuring procedure for water content by FTIR
3.5.3 Quantitative calculation procedure
4.1 Petrological description
4.1.1 Mineral proportions and textures
4.1.2 Spongy textures in sample AL47P1
4.2 Mineral major element compositions
4.3 Trace element compositions
4.3.1 Cpx in samples from Allègre
4.3.2 Cpx in samples from Mont Coupet
4.3.3 Amp in samples MC36 and MC53
4.3.4 Al, Ti, Cr and Ni in Ol from Allègre and Mont Coupet
4.4 Li concentrations and isotopic compositions in minerals
4.4.1 Samples from Allègre
4.4.2 Samples from Mont Coupet
4.5 Water content results measured by FTIR
4.5.1 Infrared spectra
4.5.2 Calculated water content from infrared spectra
4.6 Water content results from SIMS measurements
6.1 Depletion by partial melting
6.2 Mantle metasomatism revealed by trace element compositions
6.2.1 Trace element variations in Cpx from Allègre
6.2.2 Trace element variations in Cpx from Mont Coupet
6.2.3 Trace element partition between Cpx and Amp in samples MC36 and MC53
6.2.4 Trace element compositions of metasomatic melts/fluids and melt-rock reaction kinetics
6.3 Li distribution and isotopic composition variations
6.3.1 intra- and inter-mineral Li and Li isotopes distribution in Allègre samples
126.96.36.199 Intra-granular and inter-mineral lithium concentration disequilibrium
188.8.131.52 Distinctive Li isotopic offsets among mineral phases induced by asynchronous metasomatic events
6.3.2 No obvious Li addition but large inter-sample Li isotopic composition variations in peridotite xenoliths from Mont Coupet
184.108.40.206 Nearly equilibrated Li partitioning among minerals phases and no concentration zoning in grains
220.127.116.11 Large isotopic composition variations caused by a recent metasoamtism .
6.3.3 Origin of melts with exceptionally light Li isotopic compositions
6.3.4 Precedence relationship of the carbannatitic mantle metasomatism and percolation of the negative δ7Li melt
6.4. Water content of minerals in peridotite xenoliths from the French Massif Central
6.4.1 Assignment of bands in the infrared spectra
6.4.2 Water partitioning among mineral phases
6.4.3 The low water content in minerals from Allègre samples – equilibrated with degassed host magmas
6.4.4 Mineral water concentration variations in xenoliths from Mont Coupet
18.104.22.168 Preservation of original water content during xenolith ascent
22.214.171.124 The processes controlling the water content in peridotite xenoliths from Mont Coupet
126.96.36.199 Relationship between water content and trace element concentrations in Ol .
6.4.5 Water content in the lithospheric mantle beneath the French Massif Central