Early Ordovician metabasites from the Spanish ...

Gondwana Research 27 (2015) 392–409

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Early Ordovician metabasites from the Spanish Central System: A remnant of intraplate HP rocks in the Central Iberian Zone C. Villaseca a,b,⁎, P. Castiñeiras a, D. Orejana a a b

Departamento de Petrología y Geoquímica, Facultad de Geología, Universidad Complutense de Madrid, 28040 Madrid Spain Instituto de Geociencias IGEO (UCM,CSIC), 28040 Madrid Spain

a r t i c l e

i n f o

Article history: Received 16 November 2012 Received in revised form 25 September 2013 Accepted 3 October 2013 Available online 9 November 2013 Handling Editor: T. Gerya Keywords: Early Ordovician metabasites SHRIMP data Zircon geochronology High-pressure belts European Variscan orogeny

a b s t r a c t The rocks at Tenzuela represent the only known outcrop of HP metabasites in the internal sectors of the Iberian Belt, in the vast Central Iberian Zone (CIZ). These rocks appear as lenses of meta-tholeiites of continental affinity intruded within the Neoproterozoic metasedimentary sequences. The mafic types have high Th/Yb and Ta/Yb ratios that suggest, combined with initial εNd values ranging from + 4.4 to + 6.0, derivation from a slightly enriched mantle source. SHRIMP zircon data indicate an Early Ordovician age (206Pb/238U mean age of 473 ± 2 Ma) for this magmatic event. However, the U–Pb data do not yield a constrained age for the highpressure metamorphic event undergone by the metabasites, which nevertheless can be estimated at just before 335 Ma, during the Variscan collision. The HP metabasites rarely overpass peak pressure conditions of 15 kbar, and thus are remarkably lower than those of eclogites related to oceanic subduction in ophiolite- or allochthonous complexes from NW Iberia. On the other hand, these pressures are similar to those of the metabasites intercalated in orthogneisses along the Coimbra–Córdoba Shear Zone, in the boundary with the Ossa-Morena Zone (OMZ). In this regard, the new geochronological data set together with already published ages shows that the subduction-related Iberian eclogites (allochthonous complexes) are older (Lower to Middle Devonian) than the HP metabasites from more intracontinental settings, mostly recrystallized at lower HP conditions during the Visean. Thus, the studied metabasites should form part of an inner, diachronous belt of HP rocks within the European Variscan orogen, far away from the main subduction front. Nevertheless, the suggested geodynamic scenario needs future confirmation by accurate geochronological data for the HP metamorphism. © 2013 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1. Introduction Most of the high pressure rocks in the European Variscan Belt are associated with mantle-derived peridotite bodies, and usually appear as tectonically emplaced allochthonous nappes related to major suture zones (e.g., Bouchardon et al., 1989; Ballèvre et al., 2009). These suture zones include metamorphosed ultramafic, mafic and sedimentary rocks, but also contain significant volumes of felsic rocks (e.g., Bellot et al., 2010). These rocks are interpreted as Late Cambrian–Early Ordovician remnants of oceanic crust (e.g., Peucat et al., 1990; Santos Zalduegui et al., 1996; Abati et al., 1999). In this regard, ultramafic and mafic rocks with MORB or boninite chemistry have been interpreted as oceanic lithosphere from a back-arc setting (e.g., Bodinier et al., 1988; Couturier et al., 1994; Bellot et al., 2010). Early Ordovician bimodal magmatism has been interpreted as a continental breakup event in Western Europe that fragmented continental masses with the opening

⁎ Corresponding author at: Departamento de Petrología y Geoquímica, Facultad de Geología, Universidad Complutense de Madrid, c/José Antonio Novais, 2. 28040 Madrid Spain. Tel.: +34 913944910; fax: +34 915442535. E-mail addresses: [email protected] (C. Villaseca), [email protected] (P. Castiñeiras), [email protected] (D. Orejana).

of pre-oceanic to oceanic rifts (e.g., Pin and Marini, 1993; Briand et al., 2002; Murphy et al., 2008; Bellot et al., 2010; Díez-Fernández et al., 2012). In addition, the high- to ultrahigh-pressure relicts preserved in some of these rocks suggest that all the domains created during the Early Ordovician breakup were deeply buried during the Eo-Variscan continental subduction and later exhumed during the Devonian–Carboniferous continental collision by combined thrust and wrench tectonics. These processes occurred under intermediate pressure, Barroviantype metamorphism (e.g., Pin and Peucat, 1986; Martínez Catalán, 1990; Matte, 2001). The eclogites can be regarded as of the high- and low-temperature types (HT- or LT-eclogites, respectively) with a boundary value in excess of 700 °C (Medaris et al., 1995; Ballèvre et al., 2009). HT eclogites in the Iberian Variscan Belt, are the most common types, and concentrate in the uppermost units of the allochthonous nappes of NW Iberia (Galicia–Tras-os-Montes Zone) and more scarcely near to the southern suture zone (Ossa-Morena Zone) (Fig. 1). On the other hand, the LTeclogites are related to the Basal and Lower Allochthonous units of the NW Iberian Massif, below extensive slices of peridotites, granulites and eclogites (Arenas et al., 2004). An additional group of Variscan eclogites and HP metabasites crop out as lenses or boudins embedded within medium- to high-grade

1342-937X/$ – see front matter © 2013 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.gr.2013.10.007

C. Villaseca et al. / Gondwana Research 27 (2015) 392–409

393

Galicia 1 Tras-os-Montes Zone

West Asturian Leonese Zone

3 2

Cantabrian Zone 4

Study Area

Post-Variscan sediments Turégano

Granites Metasediments

5

Metabasites

Central Iberian Zone

6 7 8

Orthogneisses

Ossa-Morena Zone

D2ductile shear zones

South Portuguese Zone

Tenzuela massif Sotosalbos

Segovia

La Granja

0

24

68

km

Fig. 1. Geological map of central Spain depicting the Tenzuela metabasites (modified after Barbero and Villaseca, 2000). Inset shows its position in the Iberian Variscan Belt and the location of the Iberian HP massifs cited in text: (1) Cabo Ortegal, (2) Órdenes, (3) Malpica-Tui, (4) Bragança, (5) Coimbra–Córdoba Shear Zone: CCSZ, (6) Safira, (7) Viana do Alentejo (both last massifs are from the Évora–Aracena Belt) (8) Internal OMZ ophiolite sequences (IOMZOS).

crustal rocks usually generated at lower pressures and temperatures than the eclogites from the allochthonous complexes; these are mostly intraplate metabasic rocks unrelated to any oceanic suture. Only a small number of HP metabasite outcrops, have been described in the Central Iberian Zone (Villaseca, 1983; Barbero and Villaseca, 2000), and they are localized in a reduced area of the autochthonous series (far away from the Galicia–Tras-os-Montes Zone). These metabasite outcrops appear as scattered boudins interbedded within the Neoproterozoic metasediments, in the Sierra de Guadarrama area (central-eastern side of the Spanish Central System) (Fig. 1). We have selected for this study the complex mafic-to-felsic outcrop of the Tenzuela area (Segovia), which was metamorphosed under high- to intermediate-pressure, close to eclogite-facies conditions (Barbero and Villaseca, 2000). We have done U–Pb zircon geochronology in these rocks to constrain the age of the magmatic event and to better understand the pre-Variscan geodynamic setting and the early burial

stage of the Variscan orogeny in central Spain. We have also obtained new whole-rock geochemical data (including trace elements and Sr and Nd isotopic ratios) to enlarge the small existing dataset and to discuss the origin and significance of these tholeiitic metabasites. We suggest and discuss the existence of a paired belt of HP rocks within the European Variscan orogen: a northernmost eclogite belt close to the main oceanic suture, and a younger innermost array of HP metabasites and eclogite-facies outcrops formed during the thickening of these more intracontinental areas. This may help to constrain the Paleozoic evolution of the Variscan belt in Western Europe. 2. Geological setting The studied metabasites crop out in the Sierra de Guadarrama (Spanish Central System) within pre-Ordovician metamorphic sequences that are composed of metasedimentary and meta-igneous

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rocks, into which voluminous peraluminous late Variscan granites intruded (Fig. 1). These metasedimentary sequences are of Neoproterozoic age and consist of schists and paragneisses, with minor quartzites and discontinuous layers of marbles, amphibolites and calcsilicate rocks. The meta-igneous rocks comprise different augen, porphyritic, and leucocratic varieties that intrude the pre-Ordovician metasedimentary sequence. In the vicinity of the orthogneisses, the metapelites exhibit a hornfelsic appearance. Geochronological data for these orthogneisses cluster around Early Ordovician time (500– 470 Ma, Rb–Sr whole rock, Vialette et al., 1987; 490–462 Ma, U–Pb in zircon, Valverde-Vaquero and Dunning, 2000; Bea et al., 2007; Castiñeiras et al., 2008; Navidad and Castiñeiras, 2011; Talavera et al., 2013). The protolith of these Early Ordovician orthogneisses has a granodioritic to granitic composition. The peraluminous character of these rocks increases as the felsic composition decreases, a typical feature of S-type granitoids. The metabasites appear as discontinuous boudins or lenses within the metasedimentary sequence. They have never been found interlayered within the abundant orthogneisses (Fig. 1), thus they have been interpreted as older than the Early Ordovician orthogneisses (Barbero and Villaseca, 2000). These metabasites crop out in the Tenzuela area defining a complex massif showing local compositional layering of meta-gabbro to meta-leucotonalite types (Villaseca, 1983). The residual HP mineral assemblage is more evident in the most mafic metabasites. Peak baric conditions of the HP event have been estimated at about 14 kbar and 750 °C, followed by an intermediate-pressure granulite recrystallization at about 800 °C and 7–8 kbar (Barbero and Villaseca, 2000; Villaseca et al., 2002). Clinopyroxene relicts in garnet of these metabasites are Na-rich clinopyroxene rather than omphacite because most of the analyses plot in the quadrilateral (En + Fs + Ws) field (Fig. 4 of Barbero and Villaseca, 2000). The low jadeite mole fraction of these clinopyroxenes (b20.5%, Barbero and Villaseca, 2000) indicates that they underwent HP conditions close to or below the broad boundary between amphibolite and eclogite facies. In this regard, the very low jadeite content of the clinopyroxene analyzed in this study (see section below) confirms the estimates for the baric peak made by Barbero and Villaseca (2000). Other HP relicts are found in the host metapelitic paragneisses (kyanite + rutile + pyrope-rich garnet paragenesis, Villaseca and Barbero, 1994), indicating that the whole sequence experienced a deep burial during the Variscan collision, and that these HP metabasites do not represent allochthonous units. The metabasites were strongly amphibolitized during exhumation, being transformed from HP amphibolite/granulite into amphibolitic gneisses and garnetiferous mafic amphibolites. This probably occurred during the migmatization of the metasedimentary and meta-igneous wall-rocks at c. 330 Ma (Castiñeiras et al., 2008). Besides, recrystallization led to retrogression and the widespread growth of pale green amphibole, epidote, chlorite and sericite. Previous whole-rock geochemical data for the metabasites suggest that these rocks were continental tholeiitic sills richer in LILE and HFSE than the typical MORB (Barbero and Villaseca, 2000). The presence of felsic types in the Tenzuela massif is interpreted as derived from low-pressure crystal fractionation of the associated tholeiitic gabbros, typical of intracontinental settings. The rarity of such metabasic rocks in the Central Iberian Zone, and the preservation of HP relicts within them, makes these rocks important witnesses of the Variscan evolution in the Iberian Belt.

Barbero and Villaseca (2000) as the result of the successive metamorphic recrystallization stages. Firstly, HP relict phases comprising minerals grown during the M1 stage. This is the main thickening episode in this region, when rocks reached their maximun burial depths (e.g., Rubio Pascual et al., 2013). Minerals of the metabasites related to this stage include garnet, Na-rich diopside, rutile, and apatite, usually as inclusions in garnet domains. Second, a granulitic paragenesis overlaps the previous assemblage during the M2 stage and it is composed of plagioclase, clinopyroxene (low-Na), orthopyroxene, and amphibole. Minor biotite and K-feldspar are present in the more felsic varieties. Finally, the retrogression of the previous assemblages during the M3 stage produces minerals such as chlorite, actinolite, quartz, and epidote. Three new samples were selected for whole-rock geochemical analyses, U–Pb zircon dating and trace element analyses in zircon. Two are garnet-bearing mafic amphibolites of eclogite appearance (110406 and 110407), and the third one is a felsic amphibolite (110410). Samples were taken both from outcrops (felsic types) and as boulders (mafic amphibolites). The mafic amphibolites have abundant poikiloblastic garnet porphyroblasts, with a granoblastic matrix composed of clinopyroxene, amphibole, plagioclase, orthopyroxene, quartz and accessory phases. Garnet has minor inclusions of clinopyroxene, quartz, plagioclase, amphibole, biotite, apatite, rutile and ilmenite. Garnet (of an averaged Prp9 Alm60 Sps3 Grs28 composition, Suppl. Table 1) also shows a plagioclase corona, which is attributed to the M2 recrystallization. Clinopyroxene appears either as small inclusion in garnet (relict M1 paragenesis) or as idioblastic crystals in the matrix. All the new analyzed clinopyroxenes plot close to the WEF corner of the QUAD field for having a very low jadeite component (b8.6 mol%) (Supplementary Table 2). No chemical differences between matrix and included clinopyroxene have been found. Clinopyroxene inclusions are usually surrounded by a thin plagioclase coating that isolates them from the garnet. Orthopyroxene has been found only in one sample showing an XMg ranging from 0.31 to 0.36 (Supplementary Table 2). The presence of plagioclase coronas around the clinopyroxene inclusions suggests that plagioclase recrystallized during the second metamorphic stage, as could be also the case for amphibole, biotite, orthopyroxene, ilmenite and quartz. The most calcic plagioclases (An26 to An35) are those included in zircon or garnet (Supplementary Table 4). Amphibole has a hastingsite–ferropargasite to edenite composition (Supplementary Table 3), is quite abundant and its growth is clearly late with respect to garnet, clinopyroxene and part of the plagioclase. Some amphibole appears in intergrowth with plagioclase (chemical data in Supplementary Table 4) as an incomplete kelyphitic corona around garnet. Amphibole also appears as an apparent inclusion within garnet, where it probably represents retrograded transformation of previous clinopyroxene inclusions. Biotite, orthopyroxene and quartz are minor phases that appear mainly in the matrix, except for some biotite inclusions in garnet. Accessory phases are ilmenite, rutile (only as inclusions in garnet), titanite, apatite and zircon. The felsic amphibolite has accessory allotrioblastic garnet (with an averaged composition of Prp4 Alm63 Sps1 Grs32, Supplementary Table 1) as high-pressure mineral phase. This rock is composed of plagioclase, quartz and poikilitic amphibole, thus defining a granoblastic medium-grained texture. Amphibole is mostly hastingsite (Supplementary Table 3), whereas accessory phases are apatite, zircon, titanite and ilmenite. 4. Analytical techniques

3. Sample description Three rock types can be distinguished in the Tenzuela metabasites (Segovia area): mafic amphibolites interlayered with granoblastic felsic amphibolites, and medium- to fine-grained garnet-bearing mafic amphibolites with a massive structure and plagioclase coronas around garnet that resembles kelyphitic textures typical of decompressed eclogites. Different mineral assemblages were distinguished by

Whole-rock major and trace elements were analyzed at Actlabs (Canada). The powdered samples were melted using LiBO2 and dissolved in HNO3. The solutions were analyzed by inductively coupled plasma atomic emission spectrometry (ICP-AES) for major elements, whereas trace elements were determined by ICP mass spectrometry (ICP-MS). Uncertainties in major elements are bracketed between 1 and 3%, except for MnO (5–10%). The precision of ICP-MS analyses at

C. Villaseca et al. / Gondwana Research 27 (2015) 392–409

low concentration levels has been evaluated from repeated analyses of the international standards BR, DR-N, UB-N, AN-G and GH. The precision for Rb, Sr, Zr, Y, V, Hf and most of the REE is in the range from 1 to 5%, whereas it ranges from 5 to 10% for other trace elements, including Tm. More information on the procedure, precision and accuracy of ICP-MS analyses will be provided by Actlabs upon request. In addition to the new three samples, another three samples from Barbero and Villaseca (2000) have been analyzed to complete their trace element dataset (see Table 1). Sr–Nd isotopic analyses were carried out at the “CAI de Geocronología y Geoquímica Isotópica” (Universidad Complutense, Madrid), using an automated VG Sector 54 multicollector thermal ionization mass spectrometer. Analytical data were acquired in multidynamic mode. The analytical procedures used in this laboratory have been described elsewhere (Reyes et al., 1997). Repeated analyses of NBS 987 gave 87Sr/ 86 Sr = 0.710234 ± 30 (2σ, n = 12) and for the JM Nd standard the values of 143Nd/144Nd = 0.511854 ± 3 (2σ, n = 63) were obtained. The 2σ error on ε(Nd) calculation is ± 0.3. The measured Sr and Nd isotopic ratios for the metabasites are given in Table 2. Zircons were isolated following standard separation techniques, including crushing, grinding, sieving, Wilfley table, magnetic separator and heavy liquid (methylene iodide). The zircon grains were handpicked under a binocular microscope at the Universidad Complutense. The three samples were mounted on a double stick tape on glass slides in

395

1 × 6 mm parallel rows together with some chips of zircon standard R33 (Black et al., 2004). After setting them in epoxy resin, the zircons were ground down to expose their central portions and imaged with transmitted and reflected light on a petrographic microscope, and with cathodoluminescence on a JEOL 5600LV scanning electron microscope (housed at the University of Stanford) to identify internal structure, inclusions, fractures and physical defects. Following the analysis, secondary electron images were taken to determine the exact location of the spots. Zircon U–Th–Pb analyses were conducted on the sensitive highresolution ion microprobe-reverse geometry (SHRIMP-RG) operated by the SUMAC facility (Stanford-USGS micro analysis center) at the University of Stanford. Secondary ions are generated from the target spot with an O2− primary ion beam varying from 4 to 6 nA, which typically produces a spot with a diameter of ~20 μm and a depth of 1–2 μm for an analysis time of approximately 12 min. Measurements were made at mass resolution of M/ΔM = ~ 8500 (10% peak height), which eliminates all interfering atomic species. The reverse geometry of the USGS-Stanford SHRIMP provides very clean backgrounds and, combined with the high mass resolution, the acid washing of the mount, and rastering the primary beam for 90–120 s over the area to be analyzed before data is collected, assures that any counts found at mass 204 Pb are actually from the zircon. The accuracy using ion microprobe techniques is achieved at the expense of the precision of the results, as a quite small volume of material is sampled. Nevertheless, the precision

Table 1 Major (wt.%) and trace element (ppm) composition of the Tenzuela metabasites. Sample

96895

110406

96896

110407

Mafic amphibolites SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total Ba Rb Sr Pb Th U Zr Nb Y Co V Cr Ga Ta Hf Cs La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

46.34 3.81 12.82 18.66 0.31 4.77 9.54 3.33 0.15 0.45 0.00 99.92 25 4 61.6 0.92 1.32 0.37 191 12.3 50.3 79 562 31 22.1 1.41 4.73 0.52 9.06 23.01 3.26 16.33 5.25 1.80 7.52 1.36 8.32 1.71 4.80 0.80 4.51 0.74

67056

110410

Felsic amphibolites 47.07 3.81 12.73 18.26 0.29 4.68 9.70 3.49 0.14 0.35 0.00 100.40 9 6 93 1.03 1.76 0.58 217 11.5 50.8 44 617 19 22.0 0.91 5.50 0.60 13.60 32.60 4.90 20.70 6.05 2.07 7.45 1.57 9.29 1.88 5.45 0.79 5.04 0.74

48.23 3.46 13.12 17.75 0.28 4.19 8.95 3.39 0.36 0.52 0.01 99.96 67 15 75.9 0.98 1.84 0.58 241 14.7 63.0 62 458 19 23.8 1.68 6.04 0.90 13.04 36.66 5.17 24.61 7.90 2.52 9.59 1.69 10.96 2.15 6.14 0.94 5.88 0.93

50.55 3.20 13.24 15.76 0.29 4.03 8.22 2.88 0.40 0.53 0.01 98.94 53 6 107 2.02 2.83 0.77 365 19.5 62.8 39 398 50 19.0 1.39 7.30 0.90 20.60 48.40 6.84 32.70 9.00 2.71 10.70 1.91 11.90 2.37 6.92 1.01 6.53 1.04

69.44 0.58 13.68 5.28 0.04 0.66 3.30 5.08 1.46 0.20 0.22 99.94 224 30 122 5.23 14.83 3.17 665 20.3 74.1 5 18 19 26.0 1.57 15.20 0.52 41.72 98.07 10.91 45.78 11.37 2.22 11.13 1.83 11.46 2.32 6.86 1.07 6.66 1.05

72.73 0.38 13.47 3.23 0.03 0.37 2.37 5.68 0.42 0.08 0.00 99.31 146 9 240 6.00 13.4 4.52 609 19.2 82.7 13 9 15 23.0 1.55 15.40 0.70 43.20 113.01 12.90 44.92 10.70 1.87 10.61 2.22 14.10 2.95 8.56 1.30 8.56 1.22

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Table 2 Sr and Nd isotope data and concentrations (ppm) of the Tenzuela metabasites. Sample

Rb

Sr

87

Rb/86Sr

87

Sr/86Sr

96896 110406 110407 110410

15.1 6.0 6.0 9.0

75.9 93 107 240

0.199 0.187 0.162 0.109

0.708951 0.706212 0.707349 0.708410

± ± ± ±

06 05 05 06

(87Sr/86Sr)0

Sm

Nd

147

Sm/144Nd

143

Nd/144Nd

0.704850 0.704882 0.706193 0.707637

7.90 6.05 9.00 10.70

24.61 20.70 32.70 44.90

0.1941 0.1767 0.1664 0.1441

0.512938 0.512796 0.512779 0.512591

± ± ± ±

03 03 03 03

E(Nd)0

TDM

6.03 4.37 4.69 2.45

961 1070 918 1023

Uncertainties for the 87Sr/86Sr and 143Nd/144Nd ratios are 2 sigma errors in the last two digits.

can be statistically improved by pooling a set of equivalent data together. In addition, zircons younger than 1000 Ma yield poor 207Pb/206Pb ages, thus reducing the efficacy of the U–Pb dating method. In order to minimize this drawback, the counting time for 206Pb was increased to improve counting statistics and precision of the 206Pb/238U age. Concentration data were normalized against zircon standard CZ3 (550 ppm U, Pidgeon et al., 1995), and isotope ratios were calibrated against R33 (419 Ma, Black et al., 2004), which was analyzed repeatedly throughout the duration of the analytical session. Data reduction followed the methods described by Williams (1997), Ireland and Williams (2003), and Squid 1.08 and Isoplot 3.0 software (Ludwig, 2002, 2003) were used. The U–Pb zircon data are shown in Table 3. Trace element analysis in zircon was also carried out on the SHRIMPRG. After the isotopic analysis, the zircon mounts were lightly polished to remove the original gold coating and sputtered pits, and recoated with gold. Methods follow those discussed by Mazdab (2009). The high mass resolution of the SHRIMP-RG permits the measurement of a large set of trace elements, including Li, Sc, Nb, Y, all the REE, Hf, Pb, Th and U (Mazdab, 2009; Barth and Wooden, 2010). The smaller spot diameter (about 15 μm) and less energetic O2 beam (between 1 and 2 nA) needed for trace elements permitted that the analyses were conducted in a volume adjacent to that analyzed for isotopic compositions, as well as in additional areas to better describe trace elements variations. The primary standard is a gem quality crystal from Madagascar that has been extensively characterized in-house and found to be very chemically homogeneous (Mazdab and Wooden, 2006). The secondary zircon is CZ3, a Sri Lankan zircon megacryst described by Ireland and Williams (2003). These standards were analyzed every ten unknowns over multiple analytical sessions to establish precision of the trace element analyses. The procedure to obtain concentrations from raw counts is described in Schwartz et al. (2010). Precision for Y at 2σ is ±6%; for the measured REE (excluding La), Hf, Th, and U, 2σ precision ranges from ±8 to 18%; the precision for La is ±30%.

5. Whole-rock geochemistry and Sr–Nd isotopes The analyzed samples from the Tenzuela massif show a wide compositional range from mafic (metagabbros) to felsic rocks (leucotonalites), and from silica saturated to oversaturated (Table 1) (see also Barbero and Villaseca, 2000). The subalkaline character and their low K2O content (from 0.14 to 1.46 wt.%, Table 1), combined with their high Fe/Mg ratio (Fig. 2), define the tholeiitic nature of these metabasites, as was also stated in previous works (Villaseca, 1983; Barbero and Villaseca, 2000). On a MORB-normalized trace element diagram the mafic samples are more enriched in most traces than a N-MORB, but they show conspicuous negative anomalies in some LILE, such as Ba, K and Sr (Fig. 3a). In a similar primordial mantle-normalized diagram, the mafic samples plot mostly in the range from 2 to 50 times normalized values for most of the trace elements (Fig. 3b). LILE negative anomalies appear in most samples. The studied metabasites lack a negative Nb–Ta anomaly, thus suggesting the absence of subduction or crustal components in their source. The felsic varieties show a general increase in most of the trace element contents except for Ti–P, and display a minor negative Nb–Ta anomaly in the multi-trace element diagram,

thus suggesting some minor crustal contamination during fractional crystallization processes (Fig. 3c). The REEt contents in the studied metabasites range from 88.5 to 162.6 ppm in the mafic types, and from 252 to 276 ppm in the felsic amphibolites (Table 1). They have chondrite-normalized REE patterns from flat (mafic types) to slightly LREE-enriched (felsic types), whereas a small negative Eu anomaly appear in the more felsic metabasites (Fig. 4). The consistent flat HREE patterns suggest that neither garnet nor amphibole were involved in significant amounts in the origin of the tholeiites (Barbero and Villaseca, 2000). Moreover, the studied metabasites do not have characteristics of cumulate rocks. Accordingly, they can be regarded as intrusive magmas that evolved by low-pressure crystal fractionation towards intermediate and felsic types. Mafic metabasites yield εNd (t = 475 Ma) values from + 4.37 to + 6.03 and TDM values from 918 to 1070 Ma (Table 2). The felsic variety shows a lower εNd ratio (+ 2.45) but a similar T DM value (1023 Ma) than the mafic metabasites. This sample also shows a higher (87Sr/86Sr)t ratio of 0.7076, plotting outside the compositional field of the associated mafic types (Fig. 5a). The lack of evidence of crustal contamination in the trace element composition of the metabasites suggests their derivation from a slightly LREE-enriched mantle source at about 1.0 Ga. On the other hand, the whole-rock geochemistry of the felsic amphibolites points to a minor crustal assimilation process during fractional crystallization. 6. U–Pb and trace element SHRIMP analysis 6.1. Zircon description Zircon crystals in mafic amphibolite 110406 are scarce. They are usually broken and show an allotrioblastic habit and irregular surfaces, and have a brownish translucent color. In mafic amphibolite 110407, zircon crystals are abundant. Grains are mostly equant and allotrioblastic, but some of them look like zircon aggregates. Color varies from light to dark yellow. In contrast, felsic amphibolite 110410 produced the highest zircon yield. Grains are variable in size even though their aspect is metamorphic, that is, mostly allotriomorphic, with a brown translucent color. Under cathodoluminescence, zircon grains from the mafic amphibolites were non-luminescent, preventing distinction of any textural feature in their interiors. This low luminescence is usually related to high U and rare earth elements contents (Rubatto and Gebauer, 2000). BSE signals vary depending on the average atomic number of the elements present in a region of a zircon crystal; high signals largely correspond with U, Th, Hf and REE. Backscattered images of zircons from sample 110406 reveal two main textural types (Fig. 6a). A few grains show low BSE intensity cores surrounded by high BSE intensity rims, all of them apparently homogeneous, although a broad oscillatory zoning can be discerned in a couple of rims (Fig. 6a, #2 and #6). However, most zircon crystals in this sample display veined or patchy textures, which are probably related to recrystallization during metamorphism (Fig. 6a, #3 and #5). These patchy textures consist of low BSE intensity domains replacing moderate to high BSE intensity areas, the latter interpreted as the original magmatic zircon. We detected a myriad of inclusions in these recrystallized areas, which make the selection of the spots quite difficult. Most of the inclusions are plagioclase, amphibole

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Table 3 SHRIMP-RG U–Th–Pb analytical data for zircon of the Tenzuela metabasites. 207

Uncorrected ratios

206

238

Pb corrected

Spot name and description

Common 206Pb (%)

U (ppm)

Th (ppm)

Th/U

Pb/238U age

U/206Pb

204

Pb corrected ratios

207

Pb/206Pb

238

U/206Pb*

207

Pb*/206Pb*

110406 (TZ-1, UTM: 416495, 4546692, 1086) and 110407 (TZ-2, UTM: 416427, 4546702, 1088): Mafic amphibolites 1-1 l,g,p 0.09 1281 2715 2.19 473.5 ± 1.0 13.11 1-2 l,g,p 0.60 772 649 0.87 459.7 ± 1.4 13.45 1-3 rd −0.79 2778 10466 3.89 618.1 ± 1.0 10.02 1-4 l,g,p 0.42 633 510 0.83 408.1 ± 1.5 15.24 1-5 d,g,p 0.43 319 289 0.94 417.1 ± 2.1 14.90 1-6 l,g,p 1.91 762 221 0.30 441.3 ± 1.6 13.84 1-7 d,g,p 0.39 253 138 0.57 448.2 ± 2.1 13.83 1-8 d,g,p 0.09 395 337 0.88 415.4 ± 1.7 15.01 1-9 d,g,p 0.46 182 182 1.03 403.3 ± 2.6 15.42 1-10 hCP 9.08 541 711 1.36 458.0 ± 16.1 12.35 1-11 he 0.94 381 454 1.23 376.4 ± 18.5 16.47 1-12 d,g,p 0.20 356 251 0.73 453.2 ± 2.2 13.70 1-13 d,g,c 0.64 60 87 1.50 391.5 ± 4.0 15.87 2-1 h,c 0.18 472 408 0.89 458.1 ± 1.5 13.55 2-2 h,c 0.14 545 456 0.86 462.4 ± 1.6 13.43 2-3 h,c 0.11 557 486 0.90 476.1 ± 1.5 13.03 2-4 r 0.72 850 645 0.78 436.3 ± 1.4 14.18 2-5 h,c 0.13 1963 3324 1.75 464.8 ± 0.8 13.36 2-6 h,c,d 0.12 147 74 0.52 467.0 ± 2.8 13.29 2-7 h,c 0.06 522 431 0.85 472.5 ± 1.6 13.14 2-8 l,r 0.32 1216 1567 1.33 449.1 ± 1.0 13.81 2-9 o,r 0.12 555 467 0.87 462.6 ± 1.4 13.43 2-10 h,c 0.08 527 501 0.98 468.5 ± 1.5 13.25 2-11 h,c 0.20 742 731 1.02 465.0 ± 1.2 13.34 2-12 h,c 0.09 700 713 1.05 461.3 ± 1.3 13.47 2-13 h,c,d 1.41 238 83 0.36 444.3 ± 2.3 13.82 2-14 h,c 0.03 660 646 1.01 469.5 ± 1.3 13.23 2-15 h,c 0.08 455 319 0.73 470.8 ± 1.5 13.19 2-16 h,c 0.31 575 540 0.97 463.6 ± 1.4 13.37 2-17 h,c 0.07 969 1026 1.09 474.2 ± 1.1 13.09 2-18 h,c 0.14 1121 1308 1.21 475.0 ± 1.1 13.06 2-19 h,c,d 0.03 609 503 0.85 462.0 ± 1.4 13.45 2-20 h,c 0.15 1032 1038 1.04 464.9 ± 1.0 13.35

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

0.2 0.3 0.2 0.4 0.5 0.3 0.5 0.4 0.6 3.1 5.0 0.4 1.0 0.3 0.4 0.3 0.3 0.2 0.6 0.3 0.2 0.3 0.3 0.3 0.3 0.5 0.3 0.3 0.3 0.2 0.2 0.3 0.2

0.0573 0.0611 0.0540 0.0583 0.0586 0.0712 0.0591 0.0558 0.0585 0.1297 0.0617 0.0576 0.0597 0.0576 0.0574 0.0575 0.0614 0.0574 0.0574 0.0570 0.0585 0.0572 0.0571 0.0579 0.0570 0.0672 0.0567 0.0571 0.0588 0.0571 0.0577 0.0565 0.0575

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

0.7 0.8 0.4 1.0 1.5 1.8 1.5 1.5 2.2 10.4 1.1 2.9 3.3 1.2 1.0 1.0 1.9 0.5 1.9 1.1 1.0 1.0 1.0 0.9 0.9 1.3 0.9 1.0 1.0 0.8 0.7 1.0 0.7

13.12 13.52 10.02 15.27 14.99 14.10 13.96 15.07 15.67 13.62 16.59 13.76 16.11 13.56 13.45 13.05 14.29 13.36 13.30 13.16 13.85 13.44 13.30 13.35 13.48 13.98 13.25 13.20 13.40 13.10 13.06 13.46 13.35

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

0.2 0.4 0.2 0.4 0.6 0.5 0.5 0.4 0.8 3.5 5.0 0.5 1.3 0.3 0.4 0.3 0.4 0.2 0.6 0.3 0.2 0.3 0.3 0.3 0.3 0.6 0.3 0.3 0.3 0.2 0.2 0.3 0.2

0.0568 0.0570 0.0538 0.0565 0.0532 0.0565 0.0519 0.0527 0.0452 0.0538 0.0559 0.0545 0.0474 0.0570 0.0563 0.0562 0.0551 0.0569 0.0568 0.0561 0.0565 0.0563 0.0545 0.0577 0.0562 0.0577 0.0558 0.0563 0.0570 0.0564 0.0578 0.0559 0.0573

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

0.7 2.4 0.4 1.5 5.3 5.5 3.6 2.5 7.9 37.3 2.2 3.7 15.5 1.3 1.2 1.3 4.8 0.6 2.1 1.4 1.4 1.2 1.7 0.9 1.2 3.5 1.2 1.3 1.5 0.9 0.7 1.2 0.8

110410 (TZ-3, UTM: 416561, 4546799, 1092): Felsic amphibolite 1 r,c 0.14 109 2 c 0.26 405 3.1 r,c 0.24 88 3.2 c 0.13 340 4 c −0.01 200 5 c −0.09 197 6.1 r,c 0.23 126 6.2 c 0.24 253 7 c −0.01 432 8.1 r,c 1.55 125 8.2 c −0.07 195 9 r,c 0.56 104 10 c 0.04 178 11 c 0.13 217 12.1 i,c 0.09 143 12.2 c 0.19 239 13 r,c 0.49 65 14 c 0.09 347 15 c 0.05 140 16 r,c −0.05 277 17 r,c 0.44 129 18 r,c 0.13 256 19 r,c 0.34 81 20 r,c 0.08 529 21 r,c 0.18 634 22 r,c −0.02 512

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

0.8 0.4 1.0 0.5 0.6 0.6 0.9 0.6 0.4 0.8 0.7 0.6 0.6 0.6 0.7 0.6 0.8 0.4 0.6 0.5 0.6 0.5 0.8 0.3 0.3 0.3

0.0564 0.0587 0.0568 0.0577 0.0568 0.0560 0.0547 0.0582 0.0565 0.0676 0.0559 0.0595 0.0569 0.0577 0.0576 0.0581 0.0577 0.0571 0.0567 0.0526 0.0579 0.0542 0.0548 0.0536 0.0541 0.0526

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

2.6 1.3 2.9 1.5 1.9 2.0 3.0 1.8 1.4 2.9 2.2 2.1 2.0 1.8 2.2 1.8 2.8 1.4 1.9 2.5 2.2 1.7 2.8 1.5 1.3 1.3

14.75 13.00 15.59 13.02 12.82 13.03 19.84 13.43 13.06 15.13 13.21 15.23 13.09 12.96 12.83 13.07 17.44 13.27 13.41 19.30 16.23 18.83 21.72 19.19 19.99 19.58

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

0.8 0.4 1.1 0.5 0.6 0.6 1.0 0.6 0.4 1.2 0.7 0.7 0.6 0.6 0.7 0.6 0.9 0.4 0.6 0.5 0.7 0.5 0.9 0.3 0.4 0.3

0.0506 0.0581 0.0464 0.0558 0.0539 0.0525 0.0467 0.0574 0.0562 0.0477 0.0548 0.0527 0.0564 0.0557 0.0558 0.0576 0.0497 0.0556 0.0537 0.0527 0.0499 0.0509 0.0481 0.0534 0.0500 0.0516

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

5.3 1.6 8.1 2.1 3.0 4.2 8.5 2.0 1.4 16.0 2.5 5.2 2.3 2.5 2.8 1.9 7.4 1.8 3.1 2.5 6.0 2.9 6.6 1.5 2.9 1.7

33 394 26 251 108 120 20 148 295 12 96 29 89 130 58 146 24 275 66 1 33 25 0 15 12 2

0.32 1.01 0.31 0.76 0.56 0.63 0.16 0.60 0.71 0.10 0.51 0.29 0.52 0.62 0.42 0.63 0.37 0.82 0.48 0.00 0.27 0.10 0.00 0.03 0.02 0.00

425.2 476.8 405.0 477.4 485.8 478.9 319.4 462.3 475.7 416.4 471.2 411.0 474.8 479.8 484.6 474.8 361.1 468.9 465.2 325.8 387.4 334.5 291.6 327.2 315.7 321.6

± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ± ±

3.4 1.9 4.2 2.2 2.9 2.9 2.8 2.6 2.0 3.2 3.2 2.6 3.0 2.7 3.4 2.7 3.0 2.1 2.8 1.6 2.5 1.6 2.4 1.1 1.1 1.1

14.65 12.99 15.39 12.99 12.78 12.98 19.64 13.42 13.06 14.75 13.20 15.11 13.08 12.92 12.80 13.06 17.27 13.24 13.36 19.30 16.07 18.75 21.54 19.19 19.89 19.55

he, high error; rd, reversely discordant; hCP, high common Pb. d, dark; l, light; g, gray; p, patchy; r, recrystallized; h, homogeneous; o, oscillatorry; c, core; i, inner; r, rim.

and quartz, and more rarely ilmenite, clinopyroxene (diopside) and garnet. Compared to the previous case we found in sample 110407 that homogeneous and oscillatory zonings (Fig. 6b, #3, #4 and #5) are more common than the recrystallized zircons. Yet, limited recrystallization occurs in some of the crystals (Fig. 6b, #1 and #2), where inclusions are common. The homogeneous zones have moderate to low BSE

intensity, whereas oscillatory zones consist of broad cyclic variations from high to low BSE intensity. The zircons from both mafic amphibolite samples record different episodes of the same evolving history. In this regard, the homogeneous and oscillatory zones represent the igneous protolith, whereas the patchy areas evidence recrystallization during the subsequent metamorphism.

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1000

8

Sample/N-MORB

FeOT/MgO (wt%)

Felsic amphibolites

6 5 4 3

Mafic amphibolites

a

7

L TO C L CA

Mafic amphibolites

100

10

1

2 1 50

60

70

80 1000

SiO2 (wt%)

7. U–Pb results To minimize the analytical bias of the SHRIMP technique in samples with high uranium content (Butera et al., 2001), the analyses were performed in the areas with the lowest BSE response (dark areas in the image), which in turn correspond to domains with the lowest U content (e.g., the mafic amphibolite zircons: samples 110406 and 110407). As stated in the previous section, zircon textures from both mafic amphibolites are the result of a common geological history. For this reason, we consider as a whole the analyses performed on both samples. After an initial evaluation of the thirty-three analyses, three of them were discarded for the age discussion because of their reverse discordance (110406, spot 1-3), high common Pb content (110406, spot 1-10) or high error (110406, spot 1-11) (Table 3). The ages of the remaining 30 analyses are distributed between 475 and ~ 400 Ma (Fig. 7a), but attending to the zircon texture, two age groups can be distinguished (Fig. 7b). The first group comprises analyses obtained from homogeneous and oscillatory-zoned domains, most of them from sample 110407, where magmatic zircon is best preserved, and one of the analyses in sample 110406. The second group comprises most of the analyses performed in the recrystallized areas (veined or patchy zoning textures). The seventeen analyses obtained from the first group plot closer to concordia and yield the older ages, spreading from 475 to 460 Ma. The mean age of all the analyses in this group is 468 ± 3 Ma, but the high mean square of weighted deviation (MSWD = 16) suggests that more than one age population is included in this concordant dataset. For this reason, we used the Sambridge and Compston (1994) statistical approach that has been designed to deconvolute an assemblage of dates that overlap within their assigned errors but have more than one age-component. Following this approach, two classes can be established, namely an older age of ~ 473 Ma and a younger age of ~ 464 Ma (Fig. 7c). The mean age of 473 ± 2 Ma is very common in igneous protoliths from the Spanish

Sample/P. Mantle

Cathodoluminescence (CL) images show that most of the zircons from the felsic amphibolite (sample 110410) display oscillatory zoning, sometimes combined with sector zoning (Fig. 6c, #1, #2, #5 and #8). Occasionally, they exhibit broad homogeneous inner zones (Fig. 6c, #3, #7 and #11) and subrounded distinct cores (Fig. 6c, #6). The latter can be interpreted either as xenocrysts or as evidence of a complex magma evolution (Corfu et al., 2003). Additionally, irregular domains of homogeneous, low- to moderate-U zircon cut across growth zoned areas. These transgressive domains occur dominantly as lobes at crystal terminations, but irregular patches and sealed fractures are also present within the interior of a crystal (Fig. 6c, #4, #9 and #10).

Mafic amphibolites

b 100

Zr

10

K Sr 1

Ba

1000

c Sample/P. Mantle

Fig. 2. FeOt/MgO vs SiO2 diagram for the Tenzuela metabasites. Data are from this work (Table 1) and from Barbero and Villaseca (2000) (gray circles). The line dividing tholeiitic (TOL) and calc-alkaline (CALC) fields is taken from Winter (2010).

Felsic amphibolites

100

10

1

Ba Th K Ta Ce Sr Nd Zr Eu Tb Tm Lu Cs Rb

U Nb La Pb

P Sm Hf

Ti

Y Yb

Fig. 3. Multi-element normalized diagrams for the Tenzuela metabasites: mafic amphibolites (a and b), and felsic amphibolites, including data from sample 66726 of Barbero and Villaseca (2000) (c). N-MORB values from Klein (2004), and primitive mantle values from Taylor and McLennan (1985).

Central System (Castiñeiras et al., 2008). For this reason, we interpret this age as the best estimate for the crystallization of the mafic amphibolite protolith. On the other hand, the fourteen analyses from the second group, carried out mainly in the recrystallized areas, are variably discordant and are scattered between 460 and 392 Ma. These data are here interpreted as the result of a lead loss event younger than 392 Ma. In the felsic amphibolite (sample 110410), twenty-six analyses were performed in twenty-two zircon grains (Fig. 8a). The target of thirteen analyses was the area of zircon with magmatic zoning. For these analyses, the best estimate was obtained by putting together eight analyses, which yielded a mean age of 476 ± 1.7 Ma, with a MSWD of 0.86, here interpreted as the age of the magmatic protolith (Fig. 8b). The remaining thirteen analyses were aimed to the discordant domains. The youngest analysis (~ 290 Ma) is clearly affected by lead loss. Only six rim analyses yield Variscan ages, distributed from 335 to 316 Ma (Fig. 8c), whereas the ages scattered between 425 and 360 Ma are

C. Villaseca et al. / Gondwana Research 27 (2015) 392–409

1000

Sample/Chondrite

Felsic amphibolites

100

Mafic amphibolites

10

La 1

Pr Ce

Eu Nd

Sm

Gd

Tb

Ho Tm Lu Dy

Er

Yb

Fig. 4. REE patterns for the Tenzuela metabasites. The field for mafic types represents data from Table 1. Chondrite values from McDonough and Sun (1995).

399

tendency towards the metamorphic field of the data from the recrystallized zircon in the mafic amphibolites indicates that the zircon composition from them was probably modified during recrystallization. In the felsic amphibolite sample (110410), the Th/U ratio varies between 0.35 and 1.00 in the magmatic cores and the data define a narrow array, whereas in the metamorphic rims Th/U is lower than 0.45 and the data are much more scattered, with a weighted mean of 0.04 (Fig. 9c). Yb/Gd values for the magmatic domains cluster together around a weighted mean of 11, whereas this ratio is much more scattered in the metamorphic rims, with a weighted mean close to 40. In the U/Ce versus Th plot (Fig. 9d), all of the core analyses are placed in the magmatic field, whereas most of the analyses from the rims plot in the metamorphic field. As in the geochronology section, where some analyses yielded intermediate ages, the rim analyses that plot in the magmatic field can be interpreted as the result of a mixed composition by partially overlapping magmatic domains. 8. Discussion

mixed ages as a result of hitting older domains owing to the irregularity and the thinness of the rims. The dispersion of ages hinders the obtaining of a valid Variscan age from these six analyses. However, metamorphic zircon ages can be further adjusted considering published ages for the metamorphism in the Spanish Central System (see Section 8 below).

8.1. Significance of the Ordovician mafic magmatism in Iberia Most of the Ordovician mafic magmatic rocks present in the internal areas of the Variscan Iberian Belt appear within the allochthonous 12

8

Nd475Ma

4

0 Bragança eclogite

CO Ultramafics

-4

-8

Metasediments

-12

0.700

0.705

0.710

0.715

0.720

87Sr/86Sr 475 Ma

12

b 8

Nd475Ma

We performed seventy-six trace element analyses in zircon grains in the mafic amphibolites (Supplementary Table 5). Twenty-eight analyses were discarded on account of their high Fe content. Thus, only twenty-five analyses from sample 110406 and twenty-three analyses from sample 110407 are considered. We can define two main compositional categories in the mafic amphibolite zircons (Supplementary Table 5 and Fig. 9). The first one represents the magmatic zircon (including both low BSE intensity cores and high BSE intensity rims) as well as undisturbed areas of the recrystallized zircon. The second one stands for the patchy domains of the recrystallized zircons. In sample 110410 (felsic amphibolite), we carried out a total of sixtysix analyses in twenty-seven zircon grains (Supplementary Table 5) and we only discarded four analyses owing to their high Fe content. We organized the remaining sixty-two spot analyses in two broad sets depending on their textural characteristics and chemical composition. One includes all of the magmatic analyses, regardless of the zoning type (oscillatory, homogeneous or sector), whereas the other corresponds to the discordant zircon domains, irrespective of their luminescence. In the mafic amphibolite samples (110406 and 110407), we observed a negative exponential correlation between Th/U and the fractionation of the heavy rare earth elements (denoted by the Yb/Gd ratio), with magmatic zircon having the lowest variation in Yb/Gd and patchy areas showing a significantly higher variation (Fig. 9a). Furthermore, there is a slight compositional variation between core and rims in the magmatic grains, and a trend towards higher Th/U from core to rim is evident. Compared to these data, the recrystallized zircon has lower Th/U ratios and somewhat higher Yb/Gd ratios than in the magmatic domains, whereas the patchy domains have the lowest Th/U contents and a variable content in Yb/Gd. The same relationships are evident in Fig. 9b, where data in the magmatic grains show a trend of increasing Th from core to rim, whereas an opposite tendency is obvious in the recrystallized zircon, with the data of the patchy domains smearing in the direction of the 1:2 line. This can be used to discriminate between metamorphic or fluid-precipitated zircons and common igneous zircon using a U/Ce = 2×Th line (Bacon et al., 2012). The consistent trends between Th/U and Yb/Gd in the magmatic grains suggest that fractional crystallization processes are involved in the evolution of the mafic magma during zircon growth (Fig. 9a). The

Tenzuela Cabo Ortegal (CO) ultramafics Cabo Ortegal eclogites& granulites Metabasites CCSZ

a

7.1. Trace element results

MORB-type NW ophiolites to 0.403

4 0 -4

Bragançae clogite

Tenzuela Cabo Ortegal ultramaf. Cabo Ortegal eclog+granul NW ophiolites CCSZ Ossa-Morena

-8 -12

Metasediments

0.050

0.100

0.150

0.200

0.250

0.300

0.350

147Sm/144Nd

Fig. 5. Nd and Sr isotopic data for Ordovician metabasites from the Iberian Belt. (a) εNd vs 87 Sr/86Sr at 475 Ma plot. The fields for eclogites and ultramafics rocks from Cabo Ortegal are taken from Peucat et al. (1990) and Santos Zalduegui et al. (2002), respectively. The eclogite sample from Bragança Complex is taken from Peucat et al. (1990). The metabasite samples from the Coimbra–Córdoba Shear Zone (CCSZ) are from Gómez-Pugnaire et al. (2003). (b) εNd vs 147Sm/144Nd at 475 Ma. Data for ophiolites from NW Iberia are taken from Sánchez-Martínez (2009) and Murphy et al. (2008). Data for Cabo Ortegal are taken from Santos Zalduegui et al. (2002) and two samples from the basal units of Murphy et al. (2008). Data from the CCSZ are from Gómez-Pugnaire et al. (2003). Data from OssaMorena metabasites are from Chichorro et al. (2008). Isotopic field of Neoproterozoic metasediments from the Spanish Central System is taken from Villaseca et al. (1998).

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Fig. 6. BSE images of zircons from the mafic amphibolites (110406 and 110407) and CL images of zircons from the felsic amphibolite (110410) of Tenzuela. Ellipses represent U–Pb spots and are labeled with the spot number and the age in Table 3. Circles delineate trace element spots and are tagged with the spot number in Supplementary Table 5.

complexes of the Galicia–Tras-os-Montes Zone (Fig. 1). They are mostly mafic and ultramafic rocks with an oceanic signature related to the suture zone (ophiolitic unit), to the upper units above the suture, or to the basal units below the suture (e.g., Arenas et al., 2007a). These rocks have been dated in the Órdenes Complex at 497 ± 4 Ma (Arenas et al., 2007b), and from 520 to 490 Ma in the Cabo Ortegal Complex (Fernández-Suárez et al., 2007). The mafic rocks from the allochthonous complexes of NW Iberia have different components varying from a MORB type tholeiite affinity (e.g., ophiolitic rocks: SánchezMartínez, 2009), to more enriched island-arc tholeiites with subduction-related components (e.g., Vila de Cruces ophiolite unit in the Órdenes Complex, Arenas et al., 2007b; or Cabo Ortegal upper unit eclogites and pyroxenites: Peucat et al., 1990; Santos Zalduegui et al., 1996) (Fig. 5). Murphy et al. (2008) found crustal components in the MORB-type ophiolites of Cabo Ortegal (Moeche unit with protolith

ages of 480 Ma) suggesting either contamination during subduction or derivation from an enriched subcontinental lithospheric mantle. This mixed signature has been interpreted as a transitional oceaniccontinental setting proximal to the northern Gondwanan margin. Moreover, eclogitized metabasites of calc-alkaline affinity in the Malpica-Tui basal unit have yielded an age of 494 ± 3 Ma (Abati et al., 2010). An important feature of the basal units of the allochthonous complexes of NW Iberia is the abundant presence of coeval felsic igneous rocks defining an extended stage of magmatism ranging from 495 to 475 Ma (e.g., Abati et al., 2010 and references therein), which evolved from Late Cambrian calc-alkaline and peraluminous to Early Ordovician alkaline and peralkaline types (e.g., Díez-Fernández et al., 2012) (Fig. 10). This Late Cambrian to Early Ordovician, mainly peraluminous magmatism, is similar to the voluminous metagranitic rocks that crop out in the Central Iberian Zone, thus forming an almost continuous belt along its northern

C. Villaseca et al. / Gondwana Research 27 (2015) 392–409

401

0.078

a

a

0.074

207

Pb/206Pb

0.070 0.066 0.062 0.058

500 480

440

0.054 0.050 12

13

420

14

400 15

380 16

17

238

U/206Pb

b

b

Recrystallized zircon Magmatic zircon

490

470

450

430

410

390

370

Age (Ma)

c

4 Age ±2 fraction ±2 464.04 0.92 0.56 0.37 473.14 1.1 0.44 --relative misfit = 0.339

Number of analyses

c 3

2

1

0 456

460

464

468

472

476

480

484

Age (Ma) Fig. 7. Geochronological results obtained in zircon from the mafic amphibolite samples (110406 and 110407): (a) U–Pb Tera–Wasseburg diagram where dark and light gray ellipses represent two distinct magmatic age groups, white ellipses stand for recrystallized zircon, and the arrow depicts the lead loss evolution, (b) weighted average diagram with the three groups separated by dashed lines, and (c) graphical result of the Sambridge and Compston algorithm applies to the magmatic zircon from the mafic amphibolite samples.

part including the studied Spanish Central System (e.g., Bea et al., 2007; Montero et al., 2009). Other Ordovician metabasites have been reported to the South to the Central Iberian Zone. A bimodal (felsic–mafic) igneous complex of Cambrian to Ordovician age has been described in the Évora–Aracena Belt, close to the boundary with the South Portuguese Zone, and is characterized by the dominance of late MOR basalts and gabbros (Chichorro et al., 2008). The presence of a Cambro–Ordovician bimodal magmatism in the Coimbra–Córdoba Shear Zone, close to the limit with the Central

Fig. 8. (a) U–Pb Tera–Wasserburg diagram for zircon from the felsic amphibolite (110410), where gray and white ellipses represent analyses carried out in magmatic and metamorphic areas, respectively, (b) weigthed average diagram showing data from the magmatic zircons, and (c) diagram showing the dispersion of data from the metamorphic areas in zircon. Black bars in (b) and (c) are analyses considered to calculate the age.

Iberian Zone (Gómez-Pugnaire et al., 2003; Pereira et al., 2010), and in other parts of the Ossa-Morena Zone (Sánchez-García et al., 2008, 2010), suggests an important Ordovician tholeiitic magmatism also in SW Iberia (Fig. 10). The studied metabasites of central Spain have important differences with the other metabasic rocks from NW Iberia. On the one hand, their age is slightly younger (Early Ordovician, ~475 Ma, compared to Cambrian ages in NW Iberia, Fig. 10) and they are coeval with the abundant peraluminous granitoids of the region, although the absence of mixing structures at the outcrop scale suggests a slightly younger age of emplacement of these tholeiitic rocks. On the other hand, the absence of geochemical fingerprints of direct crustal contamination for the origin of their parental basic magmas is a remarkable feature. In a Th/Yb versus Ta/Yb diagram (Fig. 11), the Tenzuela metabasites show a more

402

C. Villaseca et al. / Gondwana Research 27 (2015) 392–409

Mafic amphibolite

Felsic amphibolite 1000

a

Yb/Gd

200

Magmatic domain Patchy domain

Recrystallized zircon

150

c

Rim Core

Magmatic zircon

100

Yb/Gd

250

100

10

Magmatic Metamorphic

50

0 0.0

0.5

1.0

1.5

2.0

2.5

1 0.001

3.0

0.1

1

10

10000

=T

U/

1000

Ce

U/

Ce

=2

d

b

h

xT

h

1000

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Fig. 9. Plots of an assortment of zircon trace elements concentrations and ratios plots for the mafic amphibolites (a)–(b), and the felsic amphibolite (c)–(d) from Tenzuela. Lines in (c) and (d) give reference U/Ce:Th ratios. Non-igneous zircon plots at U/Ce N 2×Th.

enriched composition than many of the Ordovician mafic rocks from Iberia, especially those from ophiolite complexes that display NMORB-like geochemical signatures. In the same diagram, the metabasic rocks from the upper units of the allochthonous complexes also exhibit a tendency towards higher Th/Yb values, due to minor crustal contamination (Fig. 11) (see also Murphy et al., 2008). However, other Iberian tholeiitic metabasites, usually in the allochthonous basal units and not related to ophiolitic sequences, show similar geochemical features to those of central Spain; for example, the metabasites intercalated within orthogneisses in the Coimbra–Córdoba Shear Zone (Gómez-Pugnaire et al., 2003), the E-MORB-like group-II of anorogenic tholeiites from the Évora–Aracena Belt (Chichorro et al., 2008) and some mafic granulites from the Cabo Ortegal Complex (e.g., Galán and Marcos, 1997) (Fig. 11). Thus, in the εNd – 147Sm/144Nd plot (Fig. 5b) the Tenzuela metabasites plot in an intermediate field between typical MORB-like ophiolites from NW Iberia (Peucat et al., 1990; Sánchez-Martínez, 2009), and those mafic–ultramafic rocks with high negative εNd values, but mostly overlapping the field of metabasites from Cabo Ortegal and Ossa-Morena (Fig. 5a–b). There is still a debate regarding the geodynamic setting of the Iberian Cambro–Ordovician magmatism. Its location along the northern margin of Gondwana, has led to the proposal of arc or back-arc settings

related to an active subduction zone during the Ediacaran (ValverdeVaquero and Dunning, 2000). Some other models suggest an evolution towards rifting conditions in the Middle Ordovician, leading to the opening of the Rheic Ocean (Abati et al., 2010; Díez-Fernández et al., 2012). Another proposal favors the development of intracontinental rifts situated away from the volcanic arc and associated back-arc basin (i.e., the Rheic Ocean) (Montero et al., 2009; Díez Montes et al., 2010; Ballèvre et al., 2012). The long distance of the SCS with respect to recognizable suture zones in the Central Iberian Zone favors a geodynamic context of intracontinental extension. However, the minor presence of mafic metaigneous rocks in this area suggests that rifting in central Spain was aborted or not significant at that time. In NW Iberia, the local association of peraluminous orthogneisses with mantle-derived magmas (alkaline or tholeiitic) has been explained in terms of crustal melting related to the underplating of mafic magmas in an extensional setting (e.g., Malpica-Tui Complex, Galicia, Montero et al., 2009). The coeval intrusion of peraluminous felsic and mafic magmas in central Spain also supports an extensional geodynamic context, with an apparent dominance of crustal melting over a mantle input. In this regard, the fact that Tenzuela metabasites are derived from an isotopically slightly enriched mantle with no apparent influence of arc-like signatures suggests that the region was not affected by subduction related processes

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Fig. 10. Summary of U–Pb ages for Cambrian–Ordovician metabasites (and related felsic magmatism) and HP metamorphic events in the Variscan Iberian Belt, including those from this work (t.w.). 1 = Eclogites from the Malpica-Tui unit (Abati et al., 2010). 2 = Felsic gneisses from the Malpica-Tui unit (Santos Zalduegui et al., 1995; Abati et al., 2010). 3 = Alkaline orthogneisses from the Malpica Tui unit (Rodriguez et al., 2007). 4 = Felsic orthogneisses from the Malpica Tui unit (Díez-Fernández et al., 2012). 5 = Eclogites from the Cabo Ortegal Complex (Ordóñez Casado, 1998). 6 = Mafic high-P granulite from Cabo Ortegal (Fernández-Suárez et al., 2007). 7 = Felsic gneiss intruding the Vila de Cruces ophiolite (Arenas et al., 2007b). 8 = Corredoiras orthogneiss, Órdenes Complex (Abati et al., 1999). 9 = Eclogite from the Bragança Complex (Roger and Matte, 2005). 10 = Orthogneisses from the Spanish Central System (Valverde-Vaquero and Dunning, 2000; Bea et al., 2007; Castiñeiras et al., 2008; Navidad and Castiñeiras, 2011; Talavera et al., 2013). 11 = Orthogneisses from the Schistose– Greywacke Domain (Rubio-Ordóñez et al., 2012). 12 = Orthogneisses from the Urra Formation (Solá et al., 2008). 13 = Gneiss from the Coimbra–Córdoba Shear Zone (Pereira et al., 2010). 14 = Felsic magmatism in the Coimbra–Cordoba Shear Zone (Pereira et al., 2012, and references therein). 15 = Azuaga gneiss formation (CCSZ) (Ordóñez Casado, 1998). 16 = Felsic orthogneisses from the Evora–Aracena Belt (OMZ) (Chichorro et al., 2008). 17 = Felsic gneisses from the OMZ and dated 517 Ma plagiogranite related to mafic bodies (Sánchez-García et al., 2008). 18 = Metagabbro of a dismembered ophiolite sequence within the Ossa-Morena Zone, Internal Ossa-Morena Zone Ophiolite Sequences (IOMZOS) (Ribeiro et al., 2010).

such as those of NW Iberia and the southern part of the Ossa-Morena Zone. 8.2. Age of the metamorphic events The REE profile of a zircon that has grown during an eclogitic event is characterized by the absence of an Eu anomaly and a flat HREE pattern (e.g., Liu et al., 2012; Peters et al., 2013) in response to the lack of plagioclase and the presence of garnet, respectively (Rubatto, 2002). The presence of a negative Eu anomaly (Eu/Eu*bb1) and the strongly fractionated HREE contents (Yb/GdNN1) in the recrystallized zircon

Fig. 11. Th/Yb vs Ta/Yb discrimination plot after Pearce (1983) for Cambro–Ordovician metabasites from the Iberian Belt. Samples from Cabo Ortegal are those of Santos Zalduegui et al. (2002) and Galán and Marcos (1997). Metabasites from Coimbra–Córdoba Shear Zone (CCSZ) are from Gómez-Pugnaire et al. (2003). The averaged values from OssaMorena metabasites (groups 1, 2 and 3) are from Chichorro et al. (2008). Ophiolites from NW Iberia are taken from Sánchez-Martínez (2009). Note that from Tenzuela outcrop only mafic terms are plotted.

from the mafic amphibolite samples (110406 and 110407) (Supplementary Table 5) indicate that zircon did not achieve complete reequilibration during the HP event. This re-equilibration was not attained in the U–Th–Pb isotopic system either, resulting in a group of discordant ages distributed between 460 and 392 Ma (Fig. 7) which precludes precise dating of such HP event in the mafic amphibolites (of maximum P-T conditions of 14 kbar and 750 °C, after Barbero and Villaseca, 2000). Moreover, the chemistry of the metamorphic zircon domains of the felsic amphibolite does not correspond to the expected composition in eclogitic zircon either, complicating the dating of the HP event. In addition, there are no inclusions in zircon that directly link its growth to HP conditions. With the data presented here, we can only constrain the age of this HP metamorphic event between the youngest datum obtained in the recrystallized areas of samples 110406 and 110407 (~ 390 Ma), and the oldest datum from the rims of sample 110410 (~335 Ma). Scattering in metamorphic zircon ages in eclogites is a common feature (e.g., Tichomirowa and Köhler, 2013). This could be a reflection of continous stages of zircon growth or partial inheritance rather than prolonged periods of episodic metamorphic events (Tichomirowa and Köhler, 2013), thus making necessary future works combining different accurate geochronological methods to precise the age of HP metamorphism in central Spain. Other published geochronological estimates for the metamorphism of this area can help us to limit the HP metamorphic event. A K–Ar hornblende age of 341 ± 7 Ma was obtained for the Tenzuela metabasites (Barbero et al., 1995), whereas in a recent work, the early stage of crustal thickening has been estimated at 347 ± 4 Ma (Ar–Ar plateau age of Ordovician slates, Rubio Pascual et al., 2013). The low “closing-temperatures” of some minerals used in these Ar-geochronological studies adds uncertainty in the age estimation for peak conditions in this area. Nevertheless, the subsequent low-P, high-T stage (4.5 kbar and 725 °C, M2 in Fig. 12) has been dated by zircon U–Pb geochronology in the studied area at around 330 Ma (Castiñeiras et al., 2008), in accordance with ages for the metamorphic thermal peak in the eastern part of the SCS (in the range of 337 to 326 Ma) (Escuder Viruete et al., 1998). We argue that decompression from high-pressure conditions in these rocks should have occurred in a short time interval due to the isothermal (or with a slight heating) decompressive path underwent by

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Fig. 12. (a) Summary of estimated P-T conditions of Iberian HP (eclogite- and blueschist-facies) metabasites. Estimated peak pressure conditions of SCS metabasites and the general P-T path for the studied region, taken from Barbero and Villaseca (2000) and Villaseca et al. (2002). HP peak conditions of samples from the Centro Iberian Zone or Ossa-Morena Zone are shown in gray fields, whereas those from the allocthonous Galicia–Tras-os Montes Zone are shown in white. HP conditions for eclogites from the Santiago unit in the Órdenes Complex from Rubio Pascual et al. (2002). Peak conditions in the Malpica-Tui Unit in the lower (Malpica-LA) and upper (Malpica-UA) assemblages are from Rodríguez Aller (2005) and López Carmona et al. (2010), respectively. Conditions for eclogites of Cabo Ortegal D1 and D2 are taken from Gil Ibarguchi et al. (2004). HP conditions for ultramafic rocks from the Herbeira massif at Cabo Ortegal are taken from Girardeau and Gil Ibarguchi (1991). Maximum peak conditions for Sobrado eclogites are taken from Arenas and Martínez Catalán (2002). Peak conditions for eclogites from the Bragança Complex are taken from Roger and Matte (2005). Eclogite facies conditions for metabasites from the Coimbra–Córdoba Shear Zone (CCSZ) are from López Sánchez-Vizcaíno et al. (2003). Peak conditions in the Azuaga gneiss formation (CCSZ) are taken from Ábalos et al. (1991). Conditions from Campo Maior (CCSZ) are those of Pereira et al. (2010), and from Ossa-Morena-Zone (OMZ) are those of Fonseca et al. (1999). Facies boundaries and mineral reactions are based on Brown (2007). Al2SiO5 stability diagram from Holdaway and Mukhopadhyay (1993). Facies and facies boundaries (gray lines) and position of the coesite–quartz reaction are based on Brown (2007). Low-temperature (LT) and high-temperature (HT) eclogites have 750 °C as the limit between both types (as in Ballèvre et al., 2009). (b) P-T paths observed in high-pressure terranes of the European Variscan Belt: (1) HP–LT (P-T path of glaucophane schists from the Malpica-Tui unit, López Carmona et al., 2010), (2) UHP (P-T path of coesite-bearing eclogites from the French Massif Central, Lardeaux et al., 2001) (3) HP–HT (P-T path of eclogites from Cabo Ortegal, Gil Ibarguchi et al., 2004), (4a) IP–HT fast exhumation rate (P-T path of HP metabasites described in this work, Barbero and Villaseca, 2000), (4b) IP–HT low exhumation rate (exhumation rates paths based in Rey et al., 2009).

the rocks (Fig. 12), typical of fast exhumation rates (e.g., Rey et al., 2009). Therefore, the high-pressure metamorphic peak should have occurred not more than 10–20 m.y. before 330 ± 5 Ma, the age of the Variscan low-P event in the area. Moreover, HP events in the Coimbra–Córdoba Shear Zone at the southern boundary of the CIZ yielded an average age of ~ 340 Ma (Ordóñez Casado, 1998; Pereira et al., 2010) (Fig. 10). The metabasites with HP relicts were exhumed from c. 50–60 km (Barbero and Villaseca, 2000) to c. 14–18 km of depth during the subsequent partial melting event related to the exhumation of this highgrade terrane (Castiñeiras et al., 2008). If this exhumation lasted for 10 to 20 m.y. during the Variscan orogeny, the estimated exhumation rate would be in the order of 1.6 to 3.2 mm/yr, which is similar to other rates for exhumation of high-pressure rocks in the Iberian Belt (e.g., Ordóñez Casado, 1998; Pereira et al., 2010). Conversely, this rate range is remarkably lower than the rates of unroofing in UHP terranes, required for the preservation of such metamorphic relicts (e.g., Rubatto and Hermann, 2001). 8.3. Comparison with Variscan HP rocks from Iberia Most of the Ordovician metabasites from the Iberian Massif have been described as restricted to suture zones. Thus, HP rocks are

common in the allochthonous complexes of the Galicia–Tras-os-Montes Zone in NW Iberia, where the eclogite stage shows a wide range of P-T conditions. High-temperature eclogites appear in the La Capelada unit of Cabo Ortegal Complex (780 ± 25 °C and 22 ± 1 kbar) followed by a granulite stage at 800 ± 50 °C and 15 ± 1 kbar (Mendía, 2000; Ábalos et al., 2003). Ultramafic layers record similar high-pressure equilibrium conditions at c. 800 °C and 16.5 kbar (Girardeau and Gil Ibarguchi, 1991) (Fig. 12a). U–Pb dating on these eclogites indicates metamorphic zircon growth at 392–383 Ma, as the best estimate for the eclogite event (Ordóñez Casado et al., 2001), which agrees with ages of the high-pressure assemblages from the ultramafic rocks (390 Ma, Santos Zalduegui et al., 2002). However, the high-P peak probably occurred before 410 Ma in some gabbroic protoliths of Cabo Ortegal, which is the estimated age of the subsequent HT metamorphism (Fernández-Suárez et al., 2007). Other HT-eclogites appear in the Órdenes Complex, for example those from Sobrado Unit, up to 770 °C and 17 kbar conditions (Arenas and Martínez Catalán, 2002) (Fig. 12a). Most of the other eclogites in NW Iberia show lower temperature conditions and could be classified as LT-eclogites such as those from the Agualada (695 ± 45 °C and 12–15 kbar, Arenas et al., 1997) and Santiago (480–510 °C and 11–16 kbar, Rubio Pascual et al., 2002) units, both in the Órdenes Complex. In the Malpica-Tui Complex,

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high-pressure assemblages vary from extreme peak conditions of 610 °C and 25 kbar in the lower sequence (Rodríguez Aller, 2005) towards lower grade conditions in the upper sequence: 600 °C and 12 ± 2 kbar (Gil Ibarguchi and Ortega, 1985), or 475 ± 25 °C and 17 kbar in pelite-derived glaucophane schists (López Carmona et al., 2010). In the Bragança Complex (Portugal), the high-pressure metamorphism has been estimated at 600 °C and 17 kbar (Roger and Matte, 2005) (Fig. 12a). The geochronological data of high-pressure metamorphism of these basal units yield a wide range of ages: 390 Ma (Bragança, Roger and Matte, 2005), 372 Ma (Agualada, Órdenes, Abati et al., 2010), and 365 Ma (Malpica-Tui, Rodríguez et al., 2003) (Fig. 10), although when using only U–Pb data the age range is reduced towards older ones (390–372 Ma). Other HP rocks crop out south of the Central Iberian Zone; for example, in the limit with the Ossa-Morena Zone (Coimbra–Córdoba Shear Zone: CCSZ), boudins of HP metabasites within migmatitic gneisses (Campo Maior Unit) record maximum P-T values of 850–880 °C and 15.5 ± 1 kbar (Pereira et al., 2010). The presence of HP relicts in the wall-rock gneisses suggest a continental subduction, which differs from the accreted terranes of NW Iberia, mostly derived from oceanic realms. Relict jadeite-rich clinopyroxene in some of the CCSZ metabasites yields even higher peak pressure conditions at about 19 kbar and 550 °C (López Sánchez-Vizcaíno et al., 2003) (Fig. 12a). The age of this HP metamorphism is estimated at 341 ± 9 Ma (Pereira et al., 2010), similar to the age of other HP relicts (340 ± 13 Ma) in the Coimbra–Córdoba Shear Zone (Ordóñez Casado, 1998) (Fig. 10). These were the youngest ages in the Variscan European Belt for a HT-eclogite event until our geochronological estimate for the HP event in the Tenzuela metabasites. In the Ossa-Morena Zone some allochthonous complexes, close to the southern suture zone with the South-Portuguese terrane, show also remnants of HP assemblages, such as those from Viana do Alentejo (650 °C and 15 kbar) and Safira (570 °C and 12 kbar) (Leal et al., 1997; Fonseca et al., 1999). The lack of reliable geochronological data for these eclogites prevents a correlation with the HP metabasites of the internal zones of the Variscan Iberian Belt. An examination of Figs. 10 and 12 reveals two important features of the Iberian HP metabasites. The first one reveals that eclogites from oceanic sutures (mostly those from ophiolite complexes from NW Iberia) experienced high pressure conditions, thus indicating a higher degree of burial and subduction. Two types of P-T paths can be described: HP–HT eclogites as those from Cabo Ortegal and lower sequences of the Malpica Unit (path 3 in Fig. 12b), and HP–LT eclogites (and blueschists) locally associated (path 1 in Fig. 12b). On the contrary, HP metabasites interlayered within crustal rocks rarely overpass peak pressure conditions of 15 kbar, correlatively with burials of less than 55 km, defining an IP–HT type (path 4 in Fig. 12b). This path is observed for both the studied metabasites from the Central Iberian Zone and those from the nearby Coimbra–Córdoba Shear Zone (Fig. 12). This similitude in HP conditions is suggestive of the minor oceanic subduction occurred within the boundary between the CIZ and the OMZ, in contrast to models suggesting the presence of a suture between both zones (e.g., Simancas et al., 2001). A second important feature shown by the Iberian HP metabasites is that those related to oceanic subduction are older than those from more intracontinental settings (410 to 365 Ma, and an unconstrained range of 360 to 335 Ma, respectively) (Fig. 10). The Coimbra–Córdoba Shear Zone has been interpreted as an old Precambrian boundary between an autochthonous Gondwanan terrane (CIZ) and a Neoproterozoic arc overlain by a Gondwanan passive margin rocks (OMZ) (e.g., Pereira et al., 2010). The presence of some boudins of eclogitic rocks (e.g., López Sánchez-Vizcaíno et al., 2003) and the oceanic affinity of some of these metabasites in the CCSZ (Gómez-Pugnaire et al., 2003) suggest that small remains of oceanic-affinity mafic rocks were subducted during the Variscan collision, but the question of a suture contact between the OMZ and the CIZ still remains unsettled.

405

8.4. Position within the Variscan European Belt and geological implications for eclogite formation The presence of oceanic ophiolites within the European Variscan Belt is well known, being distinguished up to five major Paleozoic sutures (e.g., Matte, 2001; Keppie et al., 2010) dispersed in the so-called Iberian–Bohemian Belt (Keppie et al., 2010) which extends from the southern Ossa-Morena Zone to Galicia in Iberia, it continues from Armorica and the Massif Central in France to the Black Forest and the Bohemian Massif in Germany and the Czech Republic (Fig. 13). However, this scenario is complicated by the existence of different HP rocks that appear within continental crust domains, such as the metabasites studied here, that make problematic the definition of a single Variscan eclogitic belt. The discussion on the different models established for interpreting the architecture of the collision of the Gondwanan microterranes during Variscan times and the origin of the eclogite belts is beyond the scope of this work. However, the main points summarized above on the Iberian eclogites impinge on the main geodynamical framework of the European Variscan orogen, and some of these aspects are discussed below. The direct correlation of eclogites with major suture zones is misleading as shown in many Variscan sectors such as the studied CIZ. The complex distribution of eclogitic belts within the European Variscan Belt could be better explained by recent models of extrusion related to a complex scenario of continental collision (e.g., Keppie et al., 2010). In this kind of models it is possible to find different P-T-t paths depending of the position with respect to the initial configuration of the collisional plates and other parameters such as convergence rate, slab dip, or plate composition (e.g., Stöckhert and Gerya, 2005; Faccenda et al., 2008; Li et al., 2010). In the European Variscan Belt, HP rocks from the allochthonous slices represent oceanic mélanges that have been thrust over margin units, thus reflecting an oceanic suture combined with an extrusion channel that could give rise to UHP or HP–HT/LT eclogites (Fig. 13). These massifs are mainly concentrated along the OMZ and GTMZ in Iberia, and then from Armorica and French Massif Central to the Bohemian Massif of Germany and the Czech Republic, thus defining the main HP belt between two suture zones described in the literature (main suture zone is represented as line 1 in Fig. 13). This array of UHP-to-HP eclogites, mostly related to complex mafic–ultramafic associations, must represent extrusion channels close to the old collision front (the Rheic Ocean closure) or exhumation channels close to underthrust blocks of continental lithospheric units (e.g., the Bohemian Massif, Babuska and Plomerová, 2013). Most geochronological data on these eclogite massifs yielded the oldest ages for a HP event of the Variscan collision: from 420 to 380 Ma (e.g., Lardeaux et al., 2001; Santos Zalduegui et al., 2002; Schulmann et al., 2005). This suggests that UHP and HP metamorphism (paths 1 to 3 in Fig. 12b) takes place prior to collision or in the early stages of continental dragging (less than 20 Ma after plate subduction, Gerya et al., 2008). Thermomechanical models of subduction or at the onset of continental collision generate clockwise-hairpin P-T paths that are in most respects compatible with those determined from this northern Variscan HP/UHP belt (e.g., models 15 and 16 from Gerya et al., 2008; reference model of Faccenda et al., 2008; standard oceanic subduction model of Burov and Yamato, 2008). On the contrary, many other HP metabasites and eclogite outcrops within the European belt correspond to intracontinental areas such as those from the CIZ, defining a southern HP belt (Fig. 13). These HP metabasites are mostly boudins within metasedimentary or felsic meta-igneous rocks, but some HP metabasites and eclogites are occasionally related to mélanges with mafic or ultramafic rocks, thus suggesting some kind of minor suture zones (e.g., central Maures massif, Rolland et al., 2009; and Golfo Aranci outcrop in Sardinia, Franceschelli et al., 2007). Thus, they are mostly crustal terranes or marginal oceanic basins far away from the main subduction front. Consequently, they

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Fig. 13. (a) Geological sketch of the Variscan orogen in Europe with the location of HP metabasite outcrops. Paleozoic sutures and Variscan zones are taken from Martínez-Catalán et al. (2007). The probable Rheic suture (1) is taken from Keppie et al. (2010). Abbreviations in the Iberian Massif: CIZ = Central Iberian Zone, CZ = Cantabrian Zone, CCSZ = Coimbra–Córdoba Shear Zone, GTMZ = Galicia–Tras-os-Montes Zone, OMZ = Ossa-Morena Zone, SPZ = South Portuguese Zone. Other abbreviations are: BF = Black Forest, CO = Corsica, ECM = External crystalline massifs of the Alps, MN = Montagne Noir, MM = Maures Massif, PY = Pyrenees, S = Sardinia, VM = Vosges Massif. Location of HP rocks in the French Variscan Belt is mainly based in Bouchardon et al. (1989) and Ballèvre et al. (2009). Others HP metabasite locations are taken from Medaris et al. (1995) (Bohemian Massif), Franceschelli et al. (2007) (Sardinia), and Kalt et al. (1994) (Black Forest). UHP locations are mostly taken from Kalt and Altherr (1996), Lardeaux et al. (2001), Massonne (2001), and the recompilation of Faryad (2011, and references therein).

have experienced a minor degree of burial evidenced by their lower P-T conditions, mostly defining a group of IP-eclogites or HP-amphibolite/ granulite metabasites (path 4 in Fig. 12b). This second array of HP rocks extends from most of the OMZ and CIZ from Iberia, then from Pyrenees, Montagne Noir and Maures to the Corsican and Sardinian Massifs (Fig. 13). Probably they extend to the Southern Alps, and then to the Western Carpathians, as Franceschelli et al. (2007) suggest. This second group of HP rocks, some of them defined as metabasites with eclogitefacies relicts (e.g., Franceschelli et al., 2007), yields younger ages for their HP events, although they are less constrained, from 355 to 335 Ma, mostly overlapping with ages of their evolution to the LP-HT stage (e.g., Franceschelli et al., 2007; Pereira et al., 2010). This time delay is in agreement with thermomechanical models; after oceanic closure, most continental orogenic belts (mainly those crustal plates with high internal heat production rates) evolve to pure shear flow and large folding and thickening is favored (Burov and Yamato, 2008). Most of the continental material buried to HP or IP conditions are exhumed in P-T paths which are very delayed (30 to 50 Ma) with respect to the initiation of the collisional event (e.g., Faccenda et al., 2008).

The southern European HP belt has HP metabasites sharing similar orogenic evolution and P-T-t paths to the studied here. Thus, the peak pressure conditions are mostly below 18 kbar (18 kbar in SE Corsica, 15 kbar in Maures, 10 kbar in Montagne Noire, 15 kbar for most HP metabasites from Sardinia: Bouchardon et al., 1989; Franceschelli et al., 2007; Giacomini et al., 2008; Rolland et al., 2009, respectively), with a subsequent high-temperature decompression (path 4a in Fig. 12b), developed during the exhumation which lasted from 360 to 340 (IP–HT conditions) to 340–320 Ma (LP–HT metamorphism) (Franceschelli et al., 2007; Giacomini et al., 2007; Rolland et al., 2009). The extrusion path of the HP metabasites from the inner parts of the (upper) European continental plate could be favored by transcurrent shear zones within the overriding plate, as occurs in the CCSZ (Pereira et al., 2010) or in the Sardinia–Corsica–Maures Shear Zone (Elter et al., 2010). In the Montagne Noire or in the CIZ, the thermal reequilibration of a previously thickened crust induces gravitational instability in the upper crust. This facilitated extrusion of the lower crust in vertical root zones at the orogen center, which on reaching higher levels could be incorporated to nappes and displaced horizontally producing

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gneissic domes (Aerden, 1998; Rubio Pascual et al., 2013). Thermomechanical models of continental thickening reproduce the P-T-t paths described by those intracontinental Variscan HP terranes, although paths crossing from the kyanite to the sillimanite stability field (path 4a in Fig. 12b), specifically those following a near isothermal decompression episode, are typical of HT-terrains undergoing fast exhumation processes (e.g., Rey et al., 2009; Jamieson et al., 2010). In conclusion, the complex HP belts generated during the European Variscan collision could be explained both by the key role played by the subduction and later closure of the Rheic Ocean (the northernmost HP belt), and by the evolution of the orogenic pile after the collision in a more intracontinental plate scenario (including the closure of small oceanic basins), and their complex interaction during this microcontinental amalgamation. A similar scenario of several intracontinental subduction zones culminating at about 340 Ma, and unrelated to the Rheic suture, has recently been described by Kroner and Romer (2013) for the European Variscides. Thus, the Variscan orogeny might have implicated many subduction zones but the final collision must have been marked almost exclusively by intracontinental collision affecting the entire internal zone of the orogen (Kroner and Romer, 2013). Nevertheless, the suggested geodynamic hypothesis needs future confirmation by accurate age estimates for respective HP metamorphism. Acknowledgments We thank José Manuel Fuenlabrada and José Antonio Hernández from the CAI of Geocronología y Geoquímica (UCM) for their help in the TIMS analyses, and to Alfredo Larios from the CNME for the help when using the electron microprobe. Valuable constructive suggestions from the associate editor Dr. Taras Gerya are gratefully acknowledged. We also thank Dr. R. Oyarzun and two anonymous reviewers for their constructive comments. This work is included in the objectives of, and supported by the CGL2012-32822 project of the Ministerio de Economía y Competitividad of Spain, and the 910492 UCM project. Appendix A. Supplementary data Supplementary data associated with this article can be found in the online version, at http://dx.doi.org/10.1016/j.gr.2013.10.007. These data include Google maps showing the location of the Tenzuela outcropdescribed in this article. References Ábalos, B., Gil Ibarguchi, J.I., Eguiluz, L., 1991. Cadomian subduction/collision and Variscan transpression in the Badajoz-Córdoba shear belt (SW Spain). Tectonophysics 199, 51–72. Ábalos, B., Puelles, P., Gil Ibarguchi, J.I., 2003. Structural assemblage of high-pressure mantle and crustal rocks in a subduction channel (Cabo Ortegal, NW Spain). Tectonics 22, 1–21. Abati, J., Dunning, G.R., Arenas, R., Díaz García, F., González Cuadra, P., Martínez Catalán, J.R., Andonaegui, P., 1999. Early Ordovician orogenic event in Galicia (NW Spain): evidence from U–Pb ages in the uppermost unit in the Ordenes Complex. Earth and Planetary Science Letters 165, 213–228. Abati, J., Gerdes, A., Fernández-Suárez, J., Arenas, R., Whitehouse, M.J., Díez Fernández, R., 2010. Magmatism and early-Variscan continental subduction in the northern Gondwana margin recorded in zircons from the basal units of Galicia, NW Spain. Geological Society of America Bulletin 122, 219–235. Aerden, D.G.A.M., 1998. Tectonic evolution of the Montagne Noire and a possible orogenic model for syncollisional exhumation of deep rocks, Variscan Belt, France. Tectonics 17, 62–79. Arenas, R., Martínez Catalán, J.R., 2002. Prograde development of corona textures in metagabbros of the Sobrado window (Ordenes Complex, NW Iberian Massif). In: Martínez Catalán, J.R., Hatcher, R.D., Arenas, R., Díaz García, F. (Eds.), Variscan–Appalachian dynamics: the building of the Late Paleozoic basement. Geological Society of America, Special Paper, 364, pp. 73–88. Arenas, R., Abati, J., Martínez Catalán, J.R., Díaz García, F., Rubio Pascual, F.J., 1997. P-T evolution of eclogites from the Agualada Unit (Ordenes Complex, northwest Iberian Massif, Spain): implications for crustal subduction. Lithos 40, 221–242. Arenas, R., Martínez Catalán, J.R., Díaz García, F., 2004. Complejos alóctonos de Galicia–Trásos-Montes. In: Vera, J.A. (Ed.), Geología de España. SGE-IGME, Madrid, pp. 138–165.

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