In situ LA-ICP-MS U–Pb titanite dating of Na–Ca metasomatism in orogenic belts: the North...

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Int J Earth Sci (Geol Rundsch) (2014) 103:667–682DOI 10.1007/s00531-013-0978-1

ORIGInal PaPER

In situ LA‑ICP‑MS U–Pb titanite dating of Na–Ca metasomatism in orogenic belts: the North Pyrenean example

Sylvain Fallourd · Marc Poujol · Philippe Boulvais · Jean‑Louis Paquette · Michel de Saint Blanquat · Philippe Rémy

Received: 28 February 2013 / accepted: 29 October 2013 / Published online: 21 november 2013 © Springer-Verlag Berlin Heidelberg 2013

Introduction

Sodic metasomatism involves the incorporation of na into rocks infiltrated by externally derived fluids (Oliver et al. 2004). This process is quite common in many geo-logical settings throughout the Earth’s history (Kent et al. 2000; Dipple and Ferry 1992; Mc Caig et al. 1990; Turpin et al. 1988; Jaguin et al. 2013). It involves a dissolution–precipitation mechanism (Engvik et al. 2008; Hövelmann et al. 2010; Putnis and John 2010) where dequartzifica-tion and muscovitization are commonly linked with the precipitation of albite (Petersson and Eliasson 1997; Mc Caig et al. 1990). albitization often relates to the influx of surface-derived fluids (Mc Caig et al. 1990; Mc lelland et al. 2002), which tends to maintain a chemical equilib-rium with the country rocks. an increase in temperature or pressure causes an increase in the K/na activity ratio in the fluid phase (lagache and Weisbrod 1977) and, conse-quently, a gain in na and a loss in K in the infiltrated rock. as a consequence, feldspar is transformed into nearly pure albite plagioclase.

Sodic–calcic metasomatism on the other hand can be viewed as the high-temperature equivalent of the na-metasomatism, as seen for example within the context of porphyry systems (Carten 1986) or within the context of intra-granitoids metasomatism (Perring et al. 2000), where temperatures up to 500 °C have been reported. also, dur-ing hydrothermal experiments, Moody et al. (1985) have shown that plagioclase compositions varied from nearly pure albite at 300 °C to oligoclase at 400 °C and andesine at 550 °C. a na-metasomatic event has been recognized throughout the Eastern Pyrenees (Boulvais et al. 2007; Pou-jol et al. 2010) and is tentatively related to the hydrothermal system associated with the transtensive displacement of the Iberian plate relative to Europe during the end of the lower

Abstract In the Pyrenees, in association with the rota-tion of the Iberian plate around Europe during the Mid-Cretaceous, a na–Ca metasomatism is recognized as a complementary record of the hydrothermal activity that led to na-metasomatism (albitization) and talc–chlorite mineralization. It affected metasedimentary rocks as well as Hercynian granites. In situ laser ablation ICP-MS U–Pb analyses of titanite grains formed in albitites during meta-somatism date the na–Ca metasomatism between 110 and 92 Ma. The temperature of the na–Ca metasomatism is estimated to be approximately 550 °C. Both the time con-straints and temperature estimates suggest that the na–Ca metasomatism is related to the low-P high-T north Pyre-nean metamorphism.

Keywords na–Ca metasomatism · albitization · Pyrenees · Titanite · la-ICP-MS U–Pb dating

Electronic supplementary material The online version of this article (doi:10.1007/s00531-013-0978-1) contains supplementary material, which is available to authorized users.

S. Fallourd · M. Poujol (*) · P. Boulvais Géosciences Rennes, UMR CnRS 6118, OSUR, Université Rennes 1, 35042 Rennes Cedex, Francee-mail: marc.poujol@univ-rennes1.fr

J.-l. Paquette laboratoire Magmas et Volcans, UMR CnRS 6524, Université Blaise Pascal, 63038 Clermont-Ferrand Cedex, France

M. de Saint Blanquat Géosciences Environnement Toulouse/Observatoire Midi-Pyrénées, 14 avenue Edouard-Belin, 31400 Toulouse, France

P. Rémy Imerys, Route de Boulzane, 11140 Salvezines, France

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Cretaceous. Here, we document a na–Ca metasomatism of comparable extent for which we propose a genetic link with the HT/lP so-called north Pyrenean Metamorphism event that took place at the transition between the lower and Upper Cretaceous.

Geological setting

The Pyrenees represent a 450-km-long and 80–150-km-wide mountain belt located between Spain and France (Fig. 1). The belt is located on a plate boundary, which was active several times since the upper Paleozoic, and there-fore constitutes a good example of the effects of succes-sive orogenies on the same crustal segment, the best known being the Variscan and alpine orogenies.

The Variscan orogeny in the Pyrenees, which took place during the namuro–Westphalian (320–300 Ma), is marked by a polyphased tectogenesis accompanied by a large calc-alkaline magmatism and a HT–lP metamorphism. The Pyrenees are classically interpreted as an external zone of the Variscan orogen in southwestern Europe (Carreras and Capella 1994; Barnolas and Chiron 1995; laumonier et al. 2010).

The alpine Pyrenees results from the late Cretaceous to lower Miocene (85–20 Ma) inversion of a sinistral tran-stensional rift, which accommodated the rotation of the Iberian plate during the Cretaceous opening of the Bay of Biscay (115–85 Ma) (Choukroune and Mattauer 1978; Choukroune et al. 1989; Roure et al. 1989; Debroas 1990; Olivet 1996; Vergés et al. 2002).

In the region close to the north Pyrenean fault (nPF), the transtensive deformation was accompanied by the for-mation of pull-apart basins filled by 4–5-km-thick flysch series (Debroas 1990), surrounded by tectono-sedimentary breccias, comprising large pieces of exhumed mantle rocks (lagabrielle and Bodinier 2008; Clerc et al. 2012). Hydro-thermal activity of regional extent was associated with the thermal anomaly generated during the lithospheric exten-sion and is recorded by the presence of numerous talc–chlorite (Moine et al. 1989; Schärer et al. 1999; Boulvais et al. 2006) and albitite bodies (Boulvais et al. 2007; Poujol et al. 2010) as well as the serpentinization and carbonation of peridotites (Clerc et al. 2013). The propagation of the thermal anomaly at shallow crustal levels induced a low-pressure high-temperature metamorphism (up to 600 °C) during the albian to Coniacian interval (110–85 Ma) (albarède and Vitrac 1978; Goldberg and leyreloup 1990).

Fig. 1 Simplified geological map of the Pyrenees showing the albi-tites, albitized rock and talc mineralization occurrences. Ant antras (Fig. 3c), Jar Jarnat (Fig. 3e), Pet Petches (Fig. 3b), Tar Tarerach

(Fig. 3a), Sar Sarrat-Blanc (Fig. 3d). Ar arize massif, As aston mas-sif, M Millas massif. Data from Schärer et al. (1999), Boulvais et al. (2007) and Poujol et al. (2010)

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Discrete magmatic activity is also known in the area (Mon-tigny et al. 1986), as demonstrated by the presence of small magmatic albititic intrusives that have been dated at 101 ± 2 Ma (Pin et al. 2001, 2006).

In the Pyrenees, metasomatic albitites grew at the expense of different protoliths including Variscan granites (e.g., the ax-les-Thermes granite in the massif of aston, 306 ± 2 Ma, Denèle, in review, or the Millas granite, 300–315 Ma, Romer and Soler 1995), ortho- and para-gneisses or even migmatites (arize massif; Poujol et al. 2010; Fig. 2). Clavières (1990) and Pascal (1979) estimated that the conditions of the albitization were bracketed between 350 and 450 °C, 2 and 1.5 kbar. The albitite occurrence at Salvezines (Fig. 1) has been dated at 117.5 ± 5 Ma (40ar/39ar method on newly formed muscovite; Boulvais

et al. 2007). The Rocher du Bari and lansac occurrences (Fig. 1) have been dated at 110 ± 8 Ma (U–Pb on titanite) and 98 ± 2 Ma (U–Th–Pb on monazite), respectively (Pou-jol et al. 2010).

Field description and petrography

In this study, we focus on five new albitized rock occur-rences both to the north and to the south of the nPF (Table 1 and Fig. 1). The samples were collected either in working quarries for feldspar and silica (Tarerach), in abandoned quarries (Petches and antras), or on natural out-crops (Sarrat-Blanc and Jarnat). The albitized zones vary in size, from approximately 5 meters in width for Jarnat

Fig. 2 Field aspects of the albitized rocks and sampling sites. a Contact between mig-matite and albitized migmatite in the abandoned quarry of antras. b Contact between gneiss and albitized gneiss in the natural outcrop of Jarnat. c Contact between granite and albitized granite in the abandoned quarry of Petches. d Contact between granite and albitized granite in the quarry of Tarerach. e albitite in the natural outcrop of Sarrat-Blanc. The dashed lines highlight the albitization metasomatic front

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to more than 50 m for Tarerach. The albitized zones form anastomozed subvertical zones, with the exception of the Sarrat-Blanc outcrop which is subhorizontal (Pascal 1979). The albitized rocks are either associated with ductile shear zones (Petches), faults (Tarerach and antras), or unde-formed variscan protolith (Sarrat-Blanc and Jarnat).

Two albitized rocks were developed at the expense of a high-grade sillimanite-bearing migmatitic gneiss (antras and Jarnat; Fig. 2a, b); these metasediments were migma-tized during the Variscan orogeny, but no age is available regarding the metamorphic recrystallization, classically attributed to the Variscan HT–lP metamorphism.

The protolith of one albitic rock (Petches) is the ax-les-Thermes granite (Fig. 2c), which was emplaced ca. 306 Ma ago. The ax-les-Thermes peraluminous leucogran-ite contains quartz, plagioclase and microcline as the main phases, while muscovite and biotite are present in minor amounts, and both zircon and apatite complete the mineral assemblage. at the sampling site, some petrographic het-erogeneities can be observed with (1) a homogeneous and fine-grained facies containing two types of enclave (meta-sedimentary xenoliths and biotitic schlieren) and (2) peg-matites associated with aplitic zones. The sampled albitized rocks were developed within a sub-vertical ductile shear zone showing evidence of post-albitization “hot” (300–400 °C) mylonitization.

Finally, two albitized rocks were sampled within the Millas composite intrusion (Tarerach, Sarrat-Blanc; Fig. 2d, e, f), which was emplaced 300 ± 5 Ma ago (U–Pb on zircon; Romer and Soler 1995). In the Millas pluton, it was not possible to collect unaltered samples either because of the pervasive albitization or the strong weathering. The unique unalbitized fresh sample was taken a few tens of meters away and corresponds to a granodiorite (sample VIn 12-10). It contains plagioclase, quartz, orthoclase, and biotite as the main constituents, although apatite and zircon are also present.

For the six different sampling locations, the mineralogy of the metasomatic rocks is relatively simple and consists of quartz, feldspar, and muscovite (accounting for 80–90 % of the whole rock) with apatite, titanite, and zircon as acces-sory minerals. The feldspars are always plagioclase (SEM analyses gave an4 to an7 in the Sarrat-Blanc sample, an13

to an16 in the Petches sample, an37 to an43 in the Jarnat sample); the Sarrat-Blanc occurrence contains numerous potassic feldspars. In addition, epidote (Jarnat, Tarerach, Petches, antras and Sarrat-Blanc), chlorite (Tarerach, Jar-nat and antras), carbonates (Tarerach), and sillimanite (Jar-nat) are also encountered.

Various pervasive mineralogical transformations linked to the albitization are observed in the thin sections (Fig. 3): dequartzification (Jarnat; Fig. 3a), secondary silicification (Fig. 3b, c) and muscovitization (Fig. 3a), which are ubiquitous, as well as carbonation (Tarerach and antras) and chloritization (Tarerach, antras and Jar-nat). In the cases of Petches and antras, silicification is strong, affecting 70–80 % of the whole rock. Muscoviti-zation developed before silicification. Carbonation devel-oped late in the alteration history and post-dates the silici-fication and muscovitization alteration processes. Epidote developed locally as an alteration product of plagioclase, either in the protolith or in the albitized rock itself, where it appears as a late phase.

In the Jarnat albitite, silimanite is associated with the newly formed quartz (silicification) and presents a fibrous habit (Fig. 3d). We interpret this feature as representing a free growth facilitated by the porosity created during albitization. This also applies to the magnesium chlorite observed at antras, which presents a radial habit.

From a geochronological point of view, zircon and titan-ite were identified in four of the five samples (Fig. 4). They represent the best candidates in order to date the albitization process. However, one has to make sure that their growth can be indisputably attributed to the metasomatic event, as both minerals are also found in the unaltered protoliths, when available. a careful examination of the thin sections by either petrographic microscopy or backscattered imag-ing reveals that the titanite grains (Petches) present in the protolith are rather big (5 mm) and display magmatic zona-tions, while the titanite in the albitized rocks are smaller (max 2 mm), zoning-free and are systematically associated with the neoformation of quartz, albite and/or muscovite. In the case of the zircon grains, however, they appear to be zoned in both the unaltered protolith and their albitized equivalent and have comparable sizes (50–200 μm), even when they are spatially associated with alteration minerals.

Table 1 Sample location coordinates and the age and nature of the protoliths

Sampling site GPS coordinates Protolith age of the protolith (in Ma)

Tarerach n42°40.678 E2°30.873 Millas granite 300 ± 5

Sarrat-Blanc n42°40.836 E2°32.047 Millas granite 300 ± 5

VIn.12.10 n42°39.285 E2°31.521 Millas granite 300 ± 5

Petches n42°42.433 E1°50.749 ax-les-Thermes leucogranite 306 ± 2

Jarnat n42°52.860 E1°37.987 Migmatized metasediments Variscan ?

antras n42°53.270 E1°38.942 Migmatite Variscan ?

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Whole-rock geochemistry

Major elements

The chemical analyses were performed by ICP-aES (major elements) and ICP-MS (trace elements) at the SaRM labo-ratory (CRPG, nancy, France). The whole-rock composi-tions were determined for the metasomatic rocks that devel-oped after the granites but not for the ones that developed after the high-grade metasediments (Table 2), because, in the latter case, the observed petrographic heterogeneities were too strong to get reliable information on the element mobility.

The first important result is that none of the studied samples is a pure albitite, as the maximum na2O content is only 9.2 wt% in samples PET 15.15a3 and PET 12.16, well below the theoretical value of 11.5 wt%. On the other hand, the CaO content is rather high in samples PET and TaR (between 1.8 and 6.6 wt%) as is the K2O content in samples TaR (between 1.1 and 5.0 wt%) and SaR (6.6 and 6.7 wt%). Poujol et al. (2010) also reported this alkali metasomatism in the Rocher du Bari site where the pro-tolith is a migmatitic rock comparable to the one sampled at the Jarnat and antras sites. The SiO2 content is also highly variable, reaching 79 wt% in sample PET 12.15b,

where silicification has been observed in thin section. The chemical–mineralogical evolutions associated with metaso-matism can be deciphered in the millicationic diagram of Debon and Fort (1983). This diagram (Fig. 5) first confirms that (1) the main characteristics of sample PET 12.15b can be linked to albitization and silicification, (2) many sam-ples have undergone both dequartzification and albitization and even a slight carbonation as evidenced by the negative value of the Q parameter, and (3) the TaR samples are also overprinted by muscovitization. The position of the SaR samples to the right of the TaR samples in Fig. 5 reflects both muscovitization and the presence of large K-feldspar grains. Finally, note that whereas sample PET 12.17 cor-responds to an unaltered granite, which compares well with the other Pyrenean granites, sample VIn 12.10 closely matches the composition of a granodiorite.

Trace elements

The REE patterns are poorly affected by metasomatism (Fig. 6). In the ax-les-Thermes granite, the large dif-ference between the REE-poor and REE-rich spectra is likely due to a magmatic history during which the frac-tionation of REE-rich phases likely induced the lower-ing of the REE content in the evolved liquid. The role of

Fig. 3 Cross-polarized thin sections images of albitized rocks. a Dissolution of quartz replaced by muscovite. b Silicification of pla-gioclase. The dashed line highlights the limit of the replacement. c neoformation of euhedral titanite in a silicification zone. Quartz 1:

magmatic quartz. Quartz 2: silicification. d Fibrous sillimanite in a zone of newly formed feldspar. Qtz quartz, Pl plagioclase, Ms musco-vite, Ttn titanite, Sil sillimanite

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monazite (controlling light REE) and zircon (controlling heavy REE) is evidenced by the extremely low Th and Zr contents in the REE-poor samples. Independently, the fact that the unaltered granite PET 12.17 is indeed a dif-ferentiated melt can be deduced from its rather high SiO2 content and its high content in hygromagmatophile ele-ments such as W and Cs. also, the observation of the vari-ous petrographic facies in the field is consistent with the idea of a coexistence between differently evolved melts in the sampling site. Between the granite PET 12.17 and the silicified albitite PET 12.15b, the strong correspond-ence between the REE patterns and the similar contents in Zr, Th, Ti and al strongly supports the idea that these elements were immobile during the fluid–rock interac-tions associated with both albitization and silicification. This can be confirmed in an isocon diagram (Fig. 7; Grant 1986) where these elements plot along a straight line pass-ing through the origin. The slope of the isocon directly

relates to the mass change associated with the metaso-matism and, if the density is measured, the changes in volume can be calculated. In this case, the slope of 0.82 points to a dilution of the immobile element in the prod-uct (silicified albitite) relative to the protolith (unaltered granite) and thus to a mass gain of 1/slope = 1.21, i.e., 21 %. Clearly, most of this mass gain can be attributed to the influx of Si and na (plus Ca), as these elements plot above the Isocon (in the field of “gains”) during silicifica-tion and albitization.

For the Millas Massif, there are also some heteroge-neities in the REE patterns. The patterns for the albitized samples plot above the one for the unaltered sample VIn 12.10. This difference cannot reasonably be attributed to a strong REE enrichment during metasomatism. Rather, we argue that the albitized rocks were formed at the expense of a more evolved granitic rock than the granodiorite VIn 12.10, a hypothesis that is consistent with the high contents

Fig. 4 BSE imaging of some grains of titanite (a, b, c) and zircon grains (d, e, f) found in samples TaR12.04 and PET12.15a3. The scale meas-ure is 50 μm. Ttn titanite, Zr zircon, Ms muscovite, Qtz quartz, Olg oligoclase

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Table 2 Chemical and isotopic (δ18O) compositions of the albitized rocks from the ax-les-Thermes and Millas massifs

location ax-les-Thermes Millas

Sample PET 12.15a3

PET 12.15b

PET 12.16

PET 12.17

TaR 12.02

TaR 12.04

TaR 12.6b

TaR 12.6c

TaR 12.6d

SaR 12.07a

SaR 12.07b

VIn 12.10

Rock type albitized gneiss

albitized mylonite

albitized granite

Granite albitized granite

albitized granite

albitized granite

albitized granite

albitized granite

albitized granite

albitized granite

Grano-diorite

SiO2 62.00 79.00 62.55 73.45 58.00 57.05 57.03 57.94 58.13 63.28 63.98 65.46

al2O3 21.46 12.18 21.85 14.42 23.70 24.47 23.70 24.37 24.41 20.57 19.93 16.54

Fe2O3 0.10 0.03 0.12 0.89 0.81 1.42 0.74 0.41 0.51 0.47 0.29 4.59

MnO 0.01 0.00 0.01 0.02 0.04 0.00 0.01 0.00 0.00 0.00 0.00 0.06

MgO 0.10 0.02 0.11 0.22 0.29 0.62 0.16 0.17 0.13 0.25 0.15 1.61

CaO 4.29 1.78 3.96 0.85 3.96 2.75 6.57 5.81 6.29 0.85 0.84 3.73

na2O 9.18 5.72 9.17 3.06 5.84 5.53 7.26 7.22 7.29 6.18 6.27 3.20

K2O 0.41 0.12 0.45 5.30 4.61 5.03 1.27 1.52 1.15 6.63 6.73 3.32

TiO2 0.39 0.15 0.39 0.14 0.71 0.79 0.90 1.09 0.92 0.42 0.42 0.52

P2O5 0.92 0.23 0.53 0.30 0.22 0.26 0.24 0.38 0.32 0.20 0.19 0.17

PF 0.83 0.20 0.70 0.73 1.30 1.82 1.85 0.80 0.96 0.84 0.74 0.70

Total 99.69 99.43 99.82 99.38 99.47 99.74 99.74 99.70 100.12 99.70 99.52 99.89

na2O/CaO 2.14 3.22 2.32 3.59 1.48 2.01 1.10 1.24 1.16 7.24 7.48 0.86

δ18O (‰) 7.7 7.2 7.5 10.0 7.4 8.1 8.3 8.3 8.4 9.4 9.6 8.8

la 29.79 5.575 38.09 6.143 84.81 24.11 121.5 80.81 55.35 83.74 65.23 30.57

Ce 59.45 10.85 77.66 12.04 155.5 44.15 225.6 157.3 96.14 155.2 115.3 55.1

Pr 6.842 1.182 8.588 1.32 17.86 4.659 24.72 18.6 9.91 13.8 11.37 5.701

nb 19.7 4.99 18.57 6.91 14.97 16.65 18.6 18.76 18.92 13.22 15.41 11.69

nd 30.31 5.4 37.39 6.054 67.51 22.72 96 65.19 46.62 64.04 52.28 25.78

Sm 6.481 1.296 7.359 1.525 15.61 6.249 23.01 14.23 11.39 11.69 12.52 4.964

Eu 0.572 0.163 0.619 0.39 3.507 1.39 8.256 4.692 4.007 1.504 1.88 1.28

Gd 5.414 1.279 5.292 1.574 16.66 8.276 27.08 14.85 14.5 7.927 11.57 4.296

Tb 0.942 0.236 0.834 0.295 2.365 1.359 3.883 2.338 2.409 1.074 1.602 0.592

Dy 5.845 1.525 4.748 1.905 12.51 8.193 20.2 13.54 14.85 5.712 8.395 3.13

Ho 1.101 0.294 0.863 0.351 2.233 1.591 3.503 2.557 2.918 0.989 1.394 0.554

Er 3.235 0.861 2.49 1.01 5.788 4.357 8.29 6.736 7.97 2.637 3.654 1.465

Tm 0.493 0.129 0.378 0.159 0.768 0.596 1.033 0.885 1.103 0.374 0.504 0.195

Yb 3.352 0.87 2.69 1.045 5.015 3.855 6.291 5.341 6.964 2.407 3.22 1.343

lu 0.511 0.126 0.411 0.143 0.731 0.575 0.939 0.772 1.029 0.344 0.456 0.201

Cs 0.63 0.21 0.55 3.78 1.61 3.16 0.76 0.95 0.70 1.94 1.64 4.94

Rb 17.6 2.38 21.0 187 178 289 70.5 85.0 62.4 251 218 108

Ba 41.6 17.8 41.2 242 1,104 736 156 226 149 799 863 947

Sr 1,037 537 1,043 73.4 384 525 767 712 785 175 190 204

ni <l.D. 5.24 <l.D. <l.D. <l.D. 6.21 <l.D. <l.D. <l.D. <l.D. <l.D. 13.5

Co 0.21 <l.D. <l.D. 0.93 0.20 0.34 <l.D. <l.D. <l.D. <l.D. 0.22 9.19

Cr 15.3 33.1 17.1 20.3 41.3 45.1 54.6 69.8 41.3 25.0 19.2 52.9

V 13.9 3.91 15.9 4.28 82.9 85.6 65.8 43.8 56.6 40.8 24.4 53.5

Zn <l.D. <l.D. <l.D. 14.2 24.6 <l.D. <l.D. <l.D. <l.D. <l.D. <l.D. 66.0

Pb 6.44 2.47 7.11 28.8 18.6 4.61 7.57 7.04 7.80 17.3 20.9 18.2

Ga 24.9 13.1 25.1 18.5 30.9 37.6 26.8 23.9 26.0 22.2 19.7 21.4

Sn 5.62 1.09 5.13 6.69 13.7 20.4 15.9 13.9 16.3 5.82 5.36 4.42

W 0.30 <l.D. <l.D. 2.68 0.87 1.23 0.48 0.48 0.45 1.12 0.95 0.45

Ta 2.30 0.81 2.26 0.93 2.53 1.44 1.76 1.45 1.94 1.61 1.79 1.01

Th 31.9 2.96 40.6 3.26 23.8 19.5 33.1 21.0 22.4 29.4 29.2 10.4

U 6.15 1.31 5.38 4.56 6.36 2.56 15.7 5.30 5.82 7.34 5.64 1.72

Zr 185 33.2 252 33.9 247 278 404 276 398 220 240 177

Hf 5.58 1.06 7.53 1.13 6.91 7.73 9.86 7.01 9.34 6.60 6.98 4.86

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of the incompatible elements such as nb, Ta, Th, U, Zr and Hf in these samples.

Stable isotopes

The oxygen isotope composition of the whole rock has been measured at the stable isotope laboratory of Univer-sity of Rennes 1, using the conventional fluorination tech-nique (see details in Boulvais et al. 2007).

The overall range of δ18O for the albitized rocks is rather restricted, between 7.2 and 9.6 ‰ (Table 2). However, when looked at more closely, each sampling site is somewhat dis-tinct from the others. The PET albitized rocks have δ18O values between 7.2 and 7.7 ‰, which are lower than the TaR albitites (between 7.4 and 8.4), and which are them-selves lower than the SaR albitized rocks (9.4 and 9.6 ‰). as a consequence, we infer that, even if all of the albitized sites share a common history, which we will propose using geochronological arguments, each of them developed under a peculiar regime of fluid–rock interaction, during which either the temperature or the fluid–rock ratio may have been different. at the PET site, the unaltered protolith has a significantly higher δ18O value than its albitized counter-part. The absolute value of 10.0 ‰ is lower than the one expected for such an evolved magma (δ18O near 12 ‰; see for example Tartese and Boulvais 2010) and may indicate some interaction between this sample and the fluids respon-sible for the albitization. Regardless of the exact cause for this specific value, it is clear that the albitized rocks have much lower values, which is a clear sign of an interaction with low δ18O fluids. Considering that the original signature

of the rock is lost during alteration, one can estimate the δ18O of the original fluid because, under these elevated fluid/rock ratios, the fluid keeps at least a significant part of its original signature. When considering the δ18O value for the albitized samples but not for the silicified samples PET 12.15a3 and PET 12.16 (δ18O = 7.6 ‰), the δ18O value of an aqueous fluid in equilibrium would be near 4.5–5.5 ‰ at 350–400 °C (using the fractionation coefficient between water and feldspar given by Zheng 1993). These values correspond to the lower limit for metamorphic fluids (Shep-pard 1986) and also for surface-derived waters that under-went isotopic exchange with crustal rocks during their cir-culation in the crust. More interesting is the low value for the silicified albitite PET 12.15b (7.2 ‰). The fluid in equi-librium with this quartz-rich sample has a lower δ18O value because quartz–water equilibrium fractionation is higher than feldspar-water fractionation. Considering a mixture of Qz–Fsp in the proportion 30–70, the δ18O value of water in equilibrium with PET 12.15b at 350–400 °C would be in the range of 3.3–4.5 ‰, reinforcing the surface-derived hypothesis for the origin of the fluid.

In the Millas intrusion, the value of the unaltered pro-tolith (δ18O = 8.8 ‰) is close to the one expected for a granodioritic magma. In the absence of an isotopic contrast between the unaltered and the albitized rocks, very little can be concluded from the δ18O values of the whole rocks.

Geochronological results and interpretation

Unfortunately, no minerals suitable for U–Pb dating were found in the sample from Sarrat-Blanc. For the four

Fig. 5 Chemical mineralogi-cal Q–P diagram of Debon and Fort (1983). The Q and P parameters are in molar propor-tions multiplied by 1,000. The albitized rocks (black squares) are compared with the Hercyn-ian granites from the Pyrenees (crosses; from Debon and Fort 1983). Grey lozenges indicate the location of the common igneous rocks in the diagram: gr granite, ad adamellite, gd granodiorite, to tonalite, sq quartzo-syenite, mzq quartzo-monzonite, mzdq quartzo-mon-zodiorite, dq quartzo-diorite, s syenite, mz monzonite, mzgo: monzogabbro, go gabbro

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Fig. 6 Chondrite-normalized rare earth elements (REE) pat-terns of granite and albitized rocks from the Millas and ax-les-Thermes massifs (chondrites values from Evensen et al. 1978). The Millas granite is also reported (Fourcade 1981)

Fig. 7 Isocon diagram (Grant 1986) showing element mobility during albitization/silicifica-tion in the case of Petches. The gained elements (black circles) lie above the Isocon, while the lost elements (empty circles) lie below. The Isocon is defined by the immobile elements (black triangles) and is almost entirely controlled by al. Its slope (a = 0.83) allows us to calculate the net mass variation fm = a−1 (here the mass gain is 21 %). Scaling factors have been added to the elements

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remaining samples, in situ U–Pb ages for the zircon and titanite grains in the thin sections were obtained by la-ICP-MS dating performed at the laboratoire Magmas et Volcans in Clermont-Ferrand (France) using a Resonetics Excimer laser coupled to a quadripole 7500 agilent ICP-MS. Suitable minerals were identified using a petrological microscope and imaged using SEM backscattered imaging.

Titanite dating by la-ICP-MS remains a difficult pro-cess mostly because of the presence of a non-negligible common Pb content even in the standards. When dated by la-MC-ICP-MS (see e.g., Poujol et al. 2010), a common Pb correction can be successfully applied. However, in the case of the quadripole ICP-MS analysis, and because of the non-negligible mercury content of the ar gas creating an isobaric interference between 204Pb and 204Hg, a common Pb correction can be quite hazardous.

Prior to the analyses of the titanite grains found in the albitized rocks, two “standard” titanites have been ana-lyzed (the Khan titanite (c. 522 Ma; Heaman 2009) and the laC (c. 520 Ma; Pedersen et al. 1989)) using ablation spot diameters of 33 μm with repetition rates of 3 Hz. These analyses show that both standards contain a non-negligible common Pb content (up to 6 %). If one standard is used to correct the other for U–Pb fractionation and for mass bias, the resulting corrected ratios do not yield the expected ages. Therefore, as titanite is a silicate, we decided to use the GJ-1 zircon standard (Jackson et al. 2004) as the exter-nal standard. Sun et al. (2012) claimed that titanite can be up to 12 % younger than their known ages using either spot or raster analyses when a zircon standard is used. although this problem was not encountered with our approach (e.g., the cogenetic titanite and monazite grains associated with the talc mineralization yield comparable ages of 101 ± 4 and 106 ± 3 Ma, respectively, Poujol, unpublished data), we cannot completely rule out the fact that the ages meas-ured in this study may be slightly younger than their “true” ages. But we will demonstrate that this possibility does not significantly affect our conclusions.

For the zircon dating, we used an ablation spot diameter of 20 μm with repetition rates of 3 Hz, and the data were corrected for U–Pb and Th–Pb fractionation and for the mass bias by standard bracketing with repeated measure-ments of the GJ-1 zircon (Jackson et al. 2004).

For both titanite and zircon, no common Pb correction was applied prior to data reduction using the GlITTER® software package developed by the Macquarie Research ltd. (Jackson et al. 2004). ages and diagrams were gener-ated using Isoplot/Ex (ludwig 2001). all errors given in the supplementary data tables 3 and 4 are listed at one sigma, but where the data are combined for the regression analy-sis, the final results are provided with 95 % confidence lim-its. Further information on the instrumentation and the ana-lytical technique is detailed in Hurai et al. (2010).

Sample ANT 12.21

no zircon grains were analyzed for this sample. Eight-een spots on eighteen titanite grains were analyzed (sup-plementary data Table 4). In a Tera–Wasserburg diagram (Fig. 8d), they plot in a discordant to very discordant posi-tion with 12 analyses plotting very close to the concordia curve. a regression through the 18 analyses anchored by the composition of common Pb at 110 Ma (with a 50 % error) following Stacey and Kramers (1975) model for Pb evolution yields a lower intercept date at 109.9 ± 4.4 Ma (MSWD = 1.5) that we interpret as the age of crystalliza-tion for these titanites.

Sample PET 12.15a3

nine zircon grains were analyzed (supplementary data Table 3). They display fairly consistent U and Pb con-tents, but their Th/U ratios are very scattered with val-ues ranging from 5.15 down to 0.6. This scattering is also observed in a concordia diagram (Fig. 8f) where they plot in a sub-concordant to a very discordant posi-tion. The best way to interpret these data is to consider that most of the grains have incorporated some common Pb followed by at least one episode of Pb loss at some unknown time. This history prevents us from calculating a precise age for this sample, but a relatively poorly con-strained date of 335 ± 27 Ma (MSWD = 0.48) can be calculated using the least discordant analyses. This date is within error with the known age for the protolith (ax-les-Thermes peraluminous granite) dated at 301 ± 15 Ma (Folkert and Harry 1987). Consequently, the zircon U–Pb system, although perturbed, did not record the age of the metasomatic event.

Eighteen titanite grains associated with the neoforma-tion of quartz, albite and/or muscovite were analyzed (sup-plementary data Table 4). They plot in a discordant posi-tion in a Tera–Wasserburg diagram (Fig. 8b). a regression forced through the Stacey and Kramers (1975) common Pb composition calculated at 110 Ma (with a 50 % error) yields a date of 105.9 ± 4.2 Ma (MSWD = 3.1) interpreted as the age of their crystallization.

Fig. 8 Concordia diagrams for samples anT 12.21, JaR 12.19, PET 12.15a3 and TaR 12.04. all errors ellipses in the concordia diagrams are plotted at 1σ and the age uncertainties are given at the 2σ level. a Tera–Wasserburg diagram for titanite grains from sample JaR 12.19. b Tera–Wasserburg diagram for titanite grains from sample PET 12.15a3. c Tera–Wasserburg diagram for titanite grains from sample TaR 12.04. d Tera–Wasserburg diagram for titanite grains from sam-ple anT 12.21. e Tera–Wasserburg diagram for zircon grains from sample JaR 12.19. f Wetherill concordia diagram for zircon grains from sample PET 12.15a3. g Tera–Wasserburg diagram for zircon grains from sample TaR 12-04

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Sample JAR 12.19

Eight zircon grains were analyzed with the U content rang-ing from 424 down to 94 ppm, while the Pb contents vary from 32 down to 8 ppm. The Th/U ratios are also very variable with values ranging from 4.52 down to 0.65 (sup-plementary data Table 3). Once again, they plot in a sub-concordant to discordant position in a Tera–Wasserburg diagram (Fig. 8e). This can be interpreted as a sign of com-mon Pb incorporated in the zircon lattice followed by a Pb loss. The best estimate for the zircon crystallization age is given by the 206Pb/238U date of 304 ± 3 Ma found for the least discordant grain (Zr4, supplementary data Table 3). The signification of this date can be tentatively attributed to the age of the migmatization that affected the metasedi-mentary protolith.

Twenty-two titanite grains were analyzed (supplemen-tary data Table 4). Here again they plot in a discordant position (Fig. 8a). a date of 97.5 ± 2.9 Ma (MSWD = 2.6) can be calculated for these grains when the regression is forced to the common Pb composition at 100 Ma (with a 50 % error; Stacey and Kramers 1975). Consequently, these titanite grains crystallized 97.5 ± 3 Ma ago.

Sample TAR 12.04

Eighteen zircon grains have been analyzed and contain 174–721 ppm U and 8–48 ppm for Pb. The Th/U ratios are inhomogeneous with values ranging between 9.57 and 1.57 (supplementary data Table 3).

all the data are very scattered (Fig. 8g) and rather dis-cordant. The zircon grains probably incorporated some common Pb and were affected by Pb loss, thereby prevent-ing the calculation of a reliable age for their crystallization. The protolith of this sample is the Millas granite that has been dated at 300 ± 5 Ma (Romer and Soler 1995).

Twenty-two titanite grains were analyzed (supplemen-tary data Table 4). They plot in a discordant to very discord-ant position with a group of nineteen grains plotting very close to the concordia curve (Fig. 8c). If a regression is cal-culated, anchored by the Stacey and Kramers (1975) com-position of common Pb at ca 95 Ma (with a 50 % error), the data yield a date of 91.8 ± 3.6 Ma (MSWD = 3), which is interpreted as the crystallization age for the titanite.

Discussion

Dating the na–Ca metasomatism

The four dated samples in this study are all characterized by a na–Ca metasomatic signature while the albitites pre-viously dated in the Pyrenees (Boulvais et al. 2007; Poujol

et al. 2010), with the exception of Rocher du Bari (see later) were characterized by a purely sodic metasomatism.

The first striking result of this study is that, although the protoliths are different both in age and in nature, all of the dated zircon grains contain rather high common Pb contents as well as a variable degree of radiogenic Pb loss, although when available, the ages found for the zircon grains are comparable with the known ages for the proto-liths. Therefore, none of the zircon U–Pb data can be used to date the metasomatic event that affected the protoliths. However, it can be argued that the fluids responsible for the albitization perturbed the U–Pb signatures of the zircon (incorporation of common Pb and/or Pb loss) without reset-ting them.

On the other hand, the titanite grains yield ages that are younger than the known age of their respective protoliths. although enriched in common Pb, their calculated ages range from ca 110 Ma down to 92 Ma. Because titanite growth is linked to na–Ca metasomatism, we can con-clude that this event occurred between ca 110 and 92 Ma. These ages are slightly younger than the ar–ar age pub-lished by Boulvais et al. (2007) for the Salvezine albi-tite (117.5 ± 0.4 Ma) and are comparable with the U–Pb ages published by Poujol et al. (2010) for the lansac (98 ± 2 Ma) and Rocher du Bari (110 ± 8 Ma) albitite occurrences. Furthermore, as previously noted by Poujol et al. (2010), these ages are also comparable with the ages found for the talcification in the world-class Trimouns talc deposit (between 112 ± 1 Ma and 97 ± 1 Ma, Schärer et al. 1999). If we consider the possibility that the calculated ages are biased due to the zircon standardization (as of Sun et al. 2012), we end up with ages of ca 123 Ma down to ca 103 Ma. Yet these ages are still comparable with the previ-ously published ages for both the albitization and talc min-eralization. although we consider that our calculated ages are accurate, taking both calculations into account, we can indubitably conclude that the na–Ca metasomatism took place during the Mid-Cretaceous.

Temporal and geographical repartitions of the na- and na–Ca metasomatism

In this study, we sampled sodic–calcic metasomatic rocks throughout the eastern half of the Pyrenees (Fig. 1). The first interesting result of our study is that the metasomatic event that was so far only documented to the north of the nPF is also recorded to the south of the nPF within the axial Zone.

The overall period of hydrothermal activity and alkali metasomatism spreads from 117.5 down to 91.8 Ma, i.e., it corresponds to a period of ca 25 Ma of fluid circulation in the upper crust (this 25 Ma period remains the same if we consider 123 Ma as the oldest age (this study) and 97 Ma

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as the youngest (Schärer et al. 1999). Clearly, the driving energy behind this long-lasting hydrothermal activity must be found at the scale of the plate tectonic processes. In the Pyrenees, there are two geodynamic processes that could be responsible for such thermal anomalies (Fig. 9). They are (1) the opening of the Bay of Biscay and the associ-ated rotation of the Iberian plate around Europe (125 to 112 Ma, Gong et al. 2008, or 115–85, Olivet 1996) and (2) the ensuing northward motion of Iberia, and the associated north–south shortening in the Pyrenean domain (85–20 Ma, Fidalgo Gonzalès 2001; Vergés et al. 2002).

Recent works have shown that the first stage—the Bay of Biscay opening—was characterized by a phase of extreme continental crust thinning associated with the exhumation of the subcontinental mantle (lagabrielle and Bodinier 2008; Jammes et al. 2009; lagabrielle et al. 2010). This type of setting could have enhanced both the thermal gradi-ent in the extended zone and the downward infiltration of surficial fluids into the upper crust.

We also document for the first time a na–Ca metaso-matic event that was not recognized in the previous studies by Boulvais et al. (2007) and Poujol et al. (2010), which described enrichment solely in na. There is, however, one exception with the Rocher du Bari metasomatic rock, which presents a CaO content of 4.09 % (Poujol et al. 2010). Both the na- and na–Ca metasomatic events share a common geographical distribution (i.e., the northern Pyr-enees), but when examined more closely, it is seen that the na-metasomatism is located to the north of the nPF, while the na–Ca metasomatism is located both to the north and south.

Furthermore, although these two types of metasoma-tism share a common time frame, to a first order, there is a slight difference in age between the two (Fig. 9). Indeed, the na-metasomatism started and ended earlier than the na–Ca one. This is for example well documented by the two extreme ages found: i.e., 117.5 Ma for the na-metas-omatism at Salvezines and 91.8 Ma for the na–Ca metaso-matism in Tarerach.

To summarize, both types of metasomatism share a common spatial and temporal distribution within the north-ern Pyrenees, suggesting that they are very probably two complementary records of the same long-lasting hydrother-mal event. In further detail, the na-metasomatism is only located to the north of the nPF and is slightly older than the na–Ca metasomatism located both to the north and to the south of the nPF.

Temperature of the na–Ca metasomatism

The experimental study of Moody et al. (1985), as well as the field studies of Carten (1986) and Perring et al. (2000), demonstrated that na–Ca metasomatism necessitates a

temperature from 400 to 550 °C, which is higher than the temperature of 350–450 °C usually required for na-meta-somatism in the Pyrenees (Pascal 1979). In the present study, two lines of evidence are in a good agreement with the higher temperature required to trigger a na–Ca-type metasomatism.

First, sillimanite is present in the Jarnat albitized rock, where it developed in a zone of newly formed feldspar (Fig. 3d). Major and trace elements analyses are not avail-able for these samples so we do not know whether they are sodic–calcic in nature or not. However, the SEM analy-ses demonstrated that the feldspar in the albitized rock is andesine (an37 to an43). according to Moody et al. (1985), a temperature above 525 °C is required in order to produce andesine during a metasomatic event. This together with the fact that the formation of sillimanite requires a temperature of at least 550 °C is a good indicator that the overall temperature during this na–Ca metasomatic event was relatively high, above 550 °C.

In the other samples where the na2O/CaO ratios are lower than five, all of the feldspars are oligoclase (an13 to an16 in the samples Petches samples, for example) to andesine, a composition which, according to Moody et al. (1985), requires a temperature above 400 °C.

These arguments demonstrate that the temperature dur-ing the na–Ca metasomatism was well above 400 °C (more probably around 550 °C) and consequently higher than the temperature recorded by the albitite resulting from the na-metasomatism (350–450 °C).

For the Petches occurrence in the ax-les-Thermes leu-cogranite, this conclusion may be at variance with the involvement of surface-derived fluids, as suggested by the O isotope characteristics. It must be recalled here that the involvement of low-δ18O fluids is especially proposed for the silicification stage, which could have occurred indepen-dently from the albitization stage. It is therefore possible that some altered samples have recorded a continuum of

Fig. 9 Diagram showing the main geological events in the Pyrenees during the middle Cretaceous. The magnetic anomalies came from Tarduno (2002)

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alteration between high-temperature metamorphic fluids and moderate-temperature surface-derived fluids.

na–Ca metasomatism within the framework of the evolution of the Pyrenees

In the diagram reporting all of the main Cretaceous events in the Pyrenees versus time (Fig. 9), it is obvious that the talcification and na-metasomatism share a common time range between ca 117 and 97 Ma. On the other hand, the na–Ca metasomatism (including the age for the Rocher du Bari) seems to start slightly later (close to 110 Ma) and remained active until ca 92 Ma. When these ages are com-pared with the known ages for the north Pyrenean meta-morphism (110–85 Ma), one can see that their respective time ranges overlap, especially toward the younger ages. In addition, in Jarnat, the presence of sillimanite implies a temperature that was greater than 550 °C, in a good agree-ment with the maximum temperature of ca 650 °C recorded during the north Pyrenean metamorphism (Goldberg and leyreloup 1990).

Furthermore, as demonstrated by the study of Moody et al. (1985), a temperature greater than 550 °C is required in order to transform plagioclase into Ca-rich feldspars. all of these arguments seem to converge toward the conclusion that the na–Ca metasomatism in the Pyrenees, as seen in this study and in Poujol et al. (2010), is related to the north Pyrenean metamorphism. This conclusion remains valid even if we consider that our calculated ages are up to 12 % too young. It is, however, difficult to determine whether the metamorphic event or the regional-scale hydrothermal event was primarily responsible for the talcification and probably the na-metasomatism, as their age ranges over-lap, at least in part. Finally, some albitized samples that show several stages of alteration, such as the samples from the PET occurrence, may have recorded both events as sug-gested above.

Conclusion

na–Ca metasomatism is reported in the Pyrenees. It occurred during Mid-Cretaceous events, in association with the rotation of the Iberian plate around Europe and the beginning of the subsequent south–north shortening. High-heat fluxes associated with the extensive tectonic regime responsible for the thinning of the crust and the mantle uprising described by lagabrielle et al. (2010) allowed flu-ids to circulate at a crustal scale. When compared to the age of the na-metasomatism (albitization) already reported in the region (117–98 Ma, limited to the north of the nPF), it seems that the na–Ca metasomatism is slightly younger (110–92 Ma) and more widespread (extending the alkali

metasomatism to the south of the nPF). also, the na–Ca metasomatism occurred at higher temperatures (ca. 100 °C) than the albitization.

at this stage, a fluid inclusion investigation is neces-sary to confirm and specify the difference in temperature between the two events (na vs. na–Ca metasomatism), but also to identify the nature of the fluids involved in each pro-cess. as suggested here, the na–Ca metasomatism might be related to the north Pyrenean metamorphism and could result from the involvement of metamorphic fluids, whereas the na-metasomatism might be related to the talc–chlorite mineralization and could result from the involvement of surface-derived fluids. The general tectonic framework that led to such differences in the fluid and temperature regimes could then be addressed.

Acknowledgments We gratefully acknowledge research funding through the CnRS (action Incitative InSU grant to MP) and the anR Pyramid (M. Ford). We would also like to thank IMERYS for facili-tating the fieldwork carried out by SF and Sara Mullin for improving the English content. Paul Sylvester and an anonymous reviewer are thanked for their constructive comments and Ingo Braun for his edito-rial handling of this manuscript.

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