A unique sequential melting mechanism for the generation of anatectic granitic rocks from the Penafiel area, northern Portugal

Lithos 155 (2012) 110–124

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A unique sequential melting mechanism for the generation of anatectic granitic rocks from the Penafiel area, northern Portugal P.C.S. Carvalho a,⁎, A.M.R. Neiva a, M.M.V.G. Silva a, F. Corfu b a b

Geosciences Centre and Department of Earth Sciences, University of Coimbra, 3000‐272 Coimbra, Portugal Department of Geosciences, University of Oslo, PB1047 Blindern, N-0316, Norway

a r t i c l e

i n f o

Article history: Received 4 April 2012 Accepted 25 August 2012 Available online 1 September 2012 Keywords: Granitic rocks Geochronology Geochemistry Sequential partial melting

a b s t r a c t Granitic rocks derived by sequential partial melting are very uncommon occurrences worldwide with only three known cases in Portugal (Tourém, Guarda–Sabugal and Penafiel) and one in Argentina (Achiras complex). The suite documented in this paper comprises two of the three main phases of a batholith in the Penafiel area. These Variscan peraluminous two-mica granite and granodiorites formed at 310 and 305 Ma during and near the end of the Variscan D3 deformation and have the characteristics of S-type granites. None of the metamorphic country rocks have an adequate Sr isotopic composition to be a protolith. It is argued that progressive isotopic re-equilibration of crustal material, probably a metagreywacke for G1 and G3 with also some metapelitic contribution for G2, during granulization of the lower Variscan crust explains the difference in isotopic composition between outcropping granitic rocks and metamorphic country rocks. Biotite > muscovite granite G1 and biotite > muscovite granodiorite G3 have similar (87Sr/86Sr)i and εNdt values and were derived by sequential partial melting of the same source material. This is indicated by coherent trends of the whole rock major and trace element compositions, the compositions of feldspar, biotite, muscovite and tourmaline, the presence and type of xenocrystic mineral phases and xenoliths, and whole rock oxygen isotope compositions. This mechanism was the result of post-thickening Variscan extension and mantle upwelling. Another member of the same batholith is a biotite ≈ muscovite granodiorite (G2), with an intermediate age of 307 Ma, but with geochemical characteristics suggesting that it corresponds to a chemically distinct pulse of magma, although likely stemming from a comparable, though more metapelitic, lower crustal source. © 2012 Elsevier B.V. All rights reserved.

1. Introduction The most dramatic geological imprint on the Iberian Peninsula resulted from the Carboniferous Variscan–Appalachian–Ouachita orogeny, which caused widespread deformation and produced immense amounts of granitic crust in the core of Pangea (e.g. Gutiérrez-Alonso et al., 2008). Variscan granitic rocks from northern and central Portugal were mainly emplaced during the last ductile deformation phase D3 (Dias et al., 2002; Ferreira et al., 1987) and have been classified as syn-D3 (~ 313–319 Ma), late-D3 (~ 306–311 Ma), late- to post-D3 (ca. 300 Ma) and post-D3 (290–296 Ma) (Dias et al., 1998). Most Portuguese Variscan granites are of S-type, whereas I-type granites are rare. Hybrid granites with transitional mineralogical and chemical characteristics between S- and I-types also occur (Neiva and Gomes, 2001; Silva and Neiva, 2000). The S-type granites were mainly derived by partial melting of metasedimentary material and subsequently evolved by fractional crystallization. By contrast, granites derived from sequential melting are rare and so far have only been described in a ⁎ Corresponding author. Tel.: +351 239860500; fax: +351 239860501. E-mail address: [email protected] (P.C.S. Carvalho). 0024-4937/$ – see front matter © 2012 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.lithos.2012.08.019

few cases, in Portugal (Holtz and Barbey, 1991; Neiva et al., 2011) and Argentina (Otamendi et al., 1998). In this paper we document a new occurrence of granites derived from sequential partial melting from the Central Iberian Zone (CIZ) of the Iberian Massif. Sequential partial melting is a model proposed by Holtz and Barbey (1991) and Holtz and Johannes (1991) to explain the genesis of granitic partial melts segregated from the same melting zone, but at increasing temperature. The granitic rocks derived from these partial melts do not define single linear arrays in the variation diagrams, but they display their own differentiation trends. We describe the geology, petrography, U–Pb geochronology, whole rock geochemistry, isotopic Rb–Sr, Sm–Nd and δ 18O data and mineral geochemistry and discuss the genesis of these rocks and the response of chemical compositions of whole rocks and minerals during sequential partial melting in four areas. 2. Geological setting The Penafiel area is located about 18 km at east of Porto, in the Dúrico–Beirão region, northern Portugal (Fig. 1a). Geologically it lies within the Central Iberian Zone (CIZ) of the Iberian Massif. The CIZ

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Fig. 1. a — Location of Penafiel and Sabugal areas and Tourém complex in the Iberian tectonic zones: CZ — Cantabrian Zone, WALZ — West Asturian–Leonese Zone, GTMZ — Galicia Trás-os-Montes Zone, CIZ — Central Iberian Zone, OMZ — Ossa–Morena Zone, SPZ — South Portuguese Zone, defined by Julivert et al. (1974) and Farias et al. (1987). b — Geological map of the Penafiel area (adapted from Medeiros et al., 1964, 1980). 1 Schist–greywacke complex, 2 Arenigian quartzites with intercalated clay schists, 3 Llandeilian–Llanvirnian schists, greywackes, quartzites, 4 Silurian schists, graywackes and quartzites, 5 Devonian clay schists and quartzites, 6 medium- to coarse-grained porphyritic biotite > muscovite granite (G1), 7 medium-grained porphyritic biotite ≈ muscovite granodiorite (G2), 8 fine-grained porphyritic biotite > muscovite granodiorite (G3), 9 aplite, aplite–pegmatite, pegmatite veins, 10 quartz veins, 11 fault and overthrust, 12 probable faults, 13 towns and villages.

is the innermost region of the Iberian Massif and contains a large variety and abundance of granitic rocks, but granites dominate. The country rocks consist largely of late Precambrian to Lower Paleozoic supracrustal rocks representing the ancient margin of Gondwana, but Cambrian–Ordovician granitic rocks of 462–498 Ma also occur (e.g. Antunes et al., 2009; Bea et al., 2007; Neiva et al., 2009; Rubio-Ordóňez et al., 2012; Solá, 2007). The emplacement age range of Variscan granitic rocks from the CIZ, based on U–Pb data of zircon and monazite, is 325 to 280 Ma (e.g. Antunes et al., 2008; Carvalho, 2010; Costa, 2011; Dias et al., 1998; Fernández-Suarez et al., 2000; Martins et al., 2009; Neiva et al., 2009; Solá et al., 2009; Teixeira, 2008; Valle Aguado et al., 2005; Zeck et al., 2007). This age range is shorter than the 280–360 Ma obtained by whole rock Rb–Sr and K–Ar data of micas (e.g. Neiva and Gomes, 2001; Serrano Pinto and Gil Ibarguchi, 1987; Villaseca et al., 1995), because whole rock Rb–Sr ages can be affected by alteration and K–Ar data of micas yield cooling and or recrystallization ages. Three different types of granite occur in the Penafiel area (Fig. 1b). The most abundant variety is a medium- to coarse-grained porphyritic biotite>muscovite granite (G1), which intruded Cambrian, Ordovician and Silurian metasedimentary rocks producing micaschists and hornfels contact metamorphic aureoles. Granite G1 contains microgranular enclaves, 5–10 cm in diameter, showing sharp contacts (Fig. 2a). Close to the contact with the country rock, G1 also has metasedimentary xenoliths. Medium-grained porphyritic biotite ≈ muscovite granodiorite (G2) intruded G1 (Fig. 1b) and contains some microgranular enclaves and metasedimentary xenoliths (Fig. 2b). It shows a NW–SE alignment and faulted contacts. Fine-grained porphyritic biotite>muscovite granodiorite (G3) intruded both G1 and G2 (Fig. 1b) and contains abundant surmicaceous enclaves, rare metasedimentary xenoliths and microgranular enclaves which reach some tens of centimeters in diameter (Fig. 2c, d).

A series of NE–SW and NW–SE trending aplite, aplite–pegmatite and pegmatite veins cut Llandeilian–Llanvirnian schists and greywackes, Silurian schists and the three granitic rocks, but particularly G1. Quartz veins orientated N35°E cut the metasedimentary rocks to the west and all three granitic units. NW–SE dolerite cuts the schist–metagreywacke complex and the three granitic rocks, particularly G1 and G2. NW–SE and NE–SW-trending faults cut the three granitic rocks, but mainly G1. 3. Analytical methods Samples were crushed in a jaw crusher and pulverized in an agate mill. The major and trace elements were determined by ICP-AES and ICP-MS, respectively. Detection limits for oxides were 0.01 wt.%, except for MnO and TiO2 that were 0.001 wt.%. Detection limits for all trace elements were 5 ppm, except for Sb (0.2 ppm) and Zn (30 ppm). The precision for major elements and Rb was 1% and for trace elements 5%. Arsenic, Cr, Sb and Sc were determined by INAA (Instrumental Neutron Activation Analysis) and precision is below 5%. Lithium was determined by ICP-AES and F by selective ion electrode analysis, both with a precision of about 5%. These determinations were carried out by Actlabs, Ontario, Canada. Whole rock FeO was determined by titration with a standardized potassium permanganate solution and H2O + using a Penfield tube, both methods with a precision of about ±1% in the Department of Earth Sciences, University of Coimbra. Zircon and monazite separation was carried out by a combination of magnetic separation and heavy liquids and the grains were mechanically abraded (Krogh, 1982). U–Pb isotopic data for zircon and monazite were obtained by isotope dilution thermal ionization mass spectrometry (ID-TIMS) using a Finnigan Mat 262 mass spectrometer at the Department of Geosciences, University of Oslo, Norway (Corfu,

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Fig. 2. Granitic rocks from the Penafiel area; a) microgranular enclave enclosed in granite G1, showing sharp contacts; b) fusiform metasedimentary xenolith and microgranular enclave enclosed in granodiorite G2; c) metasedimentary xenolith with an angular shape and surmicaceous enclaves enclosed in granodiorite G3; d) granodiorite G3 showing its phenocrysts and small surmicaceous enclaves (black points).

2004; Corfu and Andersen, 2002; Krogh, 1973). The decay constants used are those of Jaffey et al. (1971). Compositions calculated with the Stacey and Kramers model (1975) were used to do the initial Pb correction. The Isoplot programme (Ludwig, 2003) was used for plots and regression. All uncertainties of analyses and ages are given at the 2σ level. Whole rock Sr and Nd isotope analyses were obtained at the Laboratório do Serviço Geral de Investigação de Geocronologia e Geoquímica Isotópica of the University of País Basco, Espanha. Samples were digested with an ultrapure acid mixture, following Pin et al. (1994) and Pin and Santos Zalduegui (1997), and analyzed by thermal ionization mass spectrometry in a Finnigan Mat 262 mass spectrometer (Menéndez Martínez, 2001). The external precision was better than ± 0.006% for 87Sr/ 86Sr (2σ) and 0.000009% for 143 Nd/ 144Nd (2σ) and measurements of the standards were 87Sr/ 86 Sr = 0.710276 ± 0.000019 (2σ) for NBS 987 and 143Nd/ 144Nd = 0.511854 ± 0.000006 (2σ) for la Jolla and 143Nd/ 144Nd = 0.512105 ± 0.000005 (2σ) for Jndi-1. The Rb, Sr, Sm and Nd concentrations used are those obtained at Actlabs, Canada, and 87Rb/ 86Sr and 147Sm/ 146Nd were calculated using the methods of Faure and Mensing (2005). Whole rock oxygen isotope analyses were determined by gas mass spectrometry and the oxygen extraction was carried out at the University of Western Ontario, Canada according to the methodology of Clayton and Mayeda (1963). A quartz standard was used. The precision was ±0.2‰. Mineral analyses were performed using a Cameca SX-50 electron microprobe at the Laboratório de Geologia e Geocronologia dos Serviços Comuns de Investigação of the University of Oviedo, Spain, operating at 15 kV accelerating voltage, 20 nA beam current and a beam diameter between 2 and 5 μm. The ZAF corrections were applied. Detection limits (3σ above mean background) for oxides were b0.03 wt.%, except for F (0.1 wt.%), Cl (0.01%) and BaO (0.06 wt.%) with counting times of 80 s for F, Cl and BaO. The contents of Fe2O3 and FeO in biotite

were calculated by multiple linear regression using the equation of Bruyin et al. (1983). The Li2O contents of micas from granites were calculated from the equation Li2O = [(2.1 / (0.356 + MgO)] − 0.088 for biotite (Tischendorf et al., 1999) and Li2O = 0.393 F 1.326 for muscovite (Tischendorf et al., 1997). 4. Petrography The most widespread unit of the Penafiel area, G1, is a monzogranite whereas G2 and G3 are granodiorites (Le Maitre et al., 2003). These three rocks have a subhedral granular texture and contain microcline and plagioclase phenocrysts (Table 1), which are more abundant in G1. The rocks contain quartz, plagioclase, microperthitic microcline, biotite, muscovite, tourmaline, zircon, monazite, apatite, ilmenite and rutile (Table 1); G1 and G2 also contain rare andalusite and sillimanite and G1 and G3 also have rare cordierite. Small amounts of secondary minerals such as muscovite, chlorite, titanite, epidote and calcite also occur locally. All the rocks show the effects of stress, decreasing from G1 to G3. Quartz has undulatory extinction, but is fractured in G1 and G2 and shows recrystallization in G1. K-feldspar also shows undulatory extinction and micas are bent in G1 and G2. Plagioclase is deformed in G1 and G2. Quartz is anhedral and contains inclusions of muscovite, apatite and rutile. Plagioclase is subhedral, polysynthetically twinned and concentrically zoned. The phenocryst plagioclase compositions are albite– andesine in G1, oligoclase in G2 and oligoclase–andesine in G3 (Table 1) whereas matrix plagioclase compositions are albite–oligoclase in G1 and G3 and albite in G2. Myrmekite occurs locally. Plagioclase contains inclusions of biotite, muscovite, apatite and rutile. Plagioclase presents irregular grain boundaries with embayment now occupied by quartz and it is corroded by microcline and muscovite and its fractures are filled by muscovite. Microperthitic microcline is subhedral,

P.C.S. Carvalho et al. / Lithos 155 (2012) 110–124 Table 1 Petrographic information on the granitic rocks from the Penafiel area, northern Portugal. G1 Quartz Plagioclase Microcline Biotite Muscovite Apatite Accessory minerals Chlorite Secondary muscovite

23.7 37.8 25.5 7.2 2.3 0.3 0.9 0.8 1.5

G2 47.1 27.2 15.9 6.9 2.1 0.2 0.2 – 0.4

37.7 33.2 11.6 7.3 7.9 – 0.15 0.6 1.5

G3 33.6 34.9 15.8 6.2 6.6 0.1 0.1 2.0 2

39.9 31.6 8.3 6.9 10.2 0.1 0.1 0.4 2.5

34.3 34.7 11.7 13.9 1.9 0.4 0.7 0.8 1.6

Phenocrysts

Microcline, plagioclase

Microcline, plagioclase

Microcline, plagioclase

Average dimensions (mm) An of plagioclase (mol%)

Phen.

40 × 12 up to 100 × 35 8–38

15 × 5 up to 50 × 25 10–17

10 × 5 up to 50 × 25 11–43

Matrix Phen.

1–14 83–93

1–7 81–97

2–16 85–89

Matrix Phen.

87–97 ≤0.06–0.27

86–94 ≤0.06–0.31

80–96 ≤0.06–0.38

Matrix

≤0.06–0.08

≤0.06–0.21

0.12–0.36

Or of K-feldspar (mol%) Wt.% BaO of microcline

(–) Absent, G1 — medium- to coarse-grained porphyritic biotite>muscovite granite; G2 — medium-grained porphyritic biotite≈muscovite granodiorite; G3 — fine-grained porphyritic biotite>muscovite granodiorite.

cross-hatch twinned and contains inclusions of plagioclase, quartz, biotite and muscovite. Microcline phenocrysts are less abundant in G1 than in G2 and G3. Plagioclase phenocrysts are more abundant than microcline phenocrysts in the three granitic rocks. Biotite and muscovite are subhedral and locally intergrown, but anhedral muscovite also occurs. Biotite is strongly pleochroic from γ- and β-reddish brown to α-pale yellow. Both micas have inclusions of zircon, monazite, apatite and ilmenite. Muscovite also has inclusions of quartz and plagioclase and is locally intergrown with quartz. Andalusite is subhedral, slightly pleochroic from α-light pink to β= γ-colourless, and is more abundant in G1 than in G2. It occurs included in muscovite. Sillimanite occurs as needles hosted in muscovite and quartz. Cordierite is euhedral, but altered to a mixture of chlorite and muscovite in G1 and G3. Tourmaline is rare, anhedral, slightly pleochroic from ω-brown to ε-light brown and contains inclusions of quartz, feldspars and muscovite. It replaces micas. Zircon is euhedral to subhedral, zoned and included in plagioclase, biotite, muscovite, cordierite, apatite and ilmenite. Monazite is euhedral and included in micas. Apatite is euhedral to subhedral, included in micas, plagioclase and quartz. Rutile is euhedral to subhedral and included in quartz and plagioclase. Ilmenite is subhedral to anhedral, included in biotite and muscovite. Some chlorite formed by alteration of biotite shows rutile exsolutions. Secondary muscovite replaces mainly plagioclase, but also microcline. Rare secondary titanite is associated with the borders of altered biotite. Rare epidote and calcite result from the alteration of plagioclase. 5. ID-TIMS U–Pb results on zircon and monazite One representative sample of each of the three granitic rocks from the Penafiel area was selected for U–Pb work. The isotopic analyses are given in Table 2. Granite G1 (sample 123‐3) has a large population of zircons dominated by long prisms, but also with short to equant crystals which, at least in part, contain visible cores. The long-prismatic crystals without cores, and tips of such crystals, were chosen for abrasion. Five zircon fractions and two monazite fractions were analysed for granite G1 (Table 2 and Fig. 3a). Four zircon analyses define a discordia line with

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an upper intercept at 309.6 ±1.0 Ma. Another fraction (5) plots to the left of this line giving a technically concordant age of 304 Ma. Using all five zircon fractions the upper intercept age is 309.3 ± 2.3 Ma, but the MSWD is 2.7, thus resulting in a lower precision than the age calculated using the four fractions. The two monazite fractions are reversely discordant giving a 207Pb/235U age of 305.4± 0.5 Ma, which is lower than the zircon age. This could be explained by later monazite recrystallization from fluids derived from another granitic intrusion. The more probable age of this granite should be 309.6 ± 1.0 Ma. Granodiorite G2 (sample 124‐8) has a small population of zircons, most of them with inclusions and fractures and some with inherited cores. Clear, transparent and core-free zircons are rare. Two groups of long-prismatic crystals and tips were selected for abrasion. The eight zircon fractions analyzed for sample 124‐8 represent single grains or small groups of abraded prisms and tips (Table 2 and Fig. 3b). Seven fractions define a discordia line whose upper intercept with the concordia curve is at 307.0 ± 3.2 Ma (MSWD = 2.0, anchoring at 0 Ma). One fraction (7) is discordant suggesting that it contains an inherited component (Table 2 and Fig. 3d). Two analyses of monazite (fractions 8 and 9, 1 and 2 grains, respectively) plot slightly reversely discordant (Fig. 2b), a fact commonly linked to initial 230Th excess, which eventually results in an excess of 206Pb and reverse discordance. The two fractions of monazite give an average 207Pb/ 235U age of 304.9 ± 0.7 Ma. The age of 307.0 ± 3.2 Ma is considered as the crystallization age of this granite. Granodiorite G3 (sample 124‐19) has a large population of zircons and it was possible to select multi-grain fractions with very clear and transparent core-free long prismatic zircons. In this sample six fractions of zircons and two fractions of monazite were analyzed (Table 2 and Fig. 3c). The five zircon fractions plot in a cluster very close to the concordia curve (Fig. 3c). One additional zircon fraction (fraction 6) is discordant suggesting the presence of an inherited component (Table 2 and Fig. 3d). If the decay constant uncertainty is considered, the five analyses define a concordia age of 305.1±0.4 Ma (Fig. 3c). The monazite fractions are reversely discordant and have a 207Pb/235U age of 306.0± 0.6 Ma. The more probable age for this granite is 305.1±0.4 Ma. The two inherited zircons of sample 124‐8 (fraction 7) of G2 and 124‐19 (fraction 6) of G3 plot on a common discordia line forced through a lower intercept at 307.6 ± 5 Ma (Fig. 3d). The upper intercept of about 587 ± 100 Ma (Fig. 3d) indicates the age of the inherited cores present in the two fractions.

6. Whole rock geochemistry The major and trace element contents of the three granitic units are given in Tables 3 and 4. The molecular ratio Al2O3/(CaO + Na2O + K2O) is 1.11–1.23 in G1, 1.18–1.27 in G2 and 1.06–1.14 in G3, so all the granitic rocks are peraluminous. The SiO2 content ranges from 65.50 to 74.47%, K2O is always higher than Na2O and the average CaO/Na2O ratio is 0.45 in G1, 0.30 in G2 and 0.75 in G3. The average normative corundum content is 2.70% in G1, 3.45% in G2 and 2.29% in G3. These characteristics are consistent with those of the S-type granitic rocks. Plotted in the diagrams of Frost et al. (2001), they mainly belong to the alkali-calcic series; granite G1 is ferroan, whereas granodiorites G2 and G3 are magnesian. G1 and G2 plot in the syn-collision granite field and G3 plots in the post-collision to late-orogenic fields of La Roche et al. (1980) and they all plot in the field of syn-collision granites of Pearce et al. (1984). In the variation diagrams, the samples in each unit define a curved trend, with a decrease in P2O5, Al2O3, MgO, Ba, and Sr, and an increase in Cs, Rb/Sr and Rb/Ba with decreasing FeOt (Fig. 4), suggesting fractionation of feldspars, biotite and apatite. The trend defined by granodiorite G3 continues into that of granite G1, suggesting that they are related, but each one showing specific degrees of differentiation. By contrast, granodiorite G2 defines a separate trend, parallel to the trend of G1

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Table 2 U–Pb data of granitic rocks from the Penafiel area, northern Portugal. Mineral characteristics

Fraction

Weight (μg)

Ua (ppm)

Th/Ub

Pbcc (pg)

206 204

Pb/ Pbd

207 235

Pb/ Ue

206

2σ (abs)

238

Pb/ Ue

2σ (abs)

Rho

D (%)f

Ages (M.y.) 206 238

Pb/ Ue

2σ (abs)

207 235

Pb/ Ue

2σ (abs)

207 206

Pb/ Pbe

2σ (abs)

Sample 123‐3: G1 Z, LP (21) Z, LP (3) Z, tips (6) Z, LP (6) Z, LP, “s-br” (10) Mz (3) Mz (2)

1 2 4 5 6 7 8

33 25 17 9 7 15 19

628 314 665 775 503 1375 1777

0.14 0.31 0.19 0.14 0.3 13.31 19.59

9.3 2 2.89 2.38 2.27 6.4 8.76

6722 11981 11868 8913 4708 9858 11768

0.34954 0.35299 0.35067 0.35041 0.34857 0.35084 0.35102

0.00084 0.00083 0.00082 0.00083 0.00099 0.00088 0.00085

0.04825 0.04868 0.0484 0.04836 0.04822 0.04874 0.0487

0.0001 0.0001 0.0001 0.0001 0.00011 0.00011 0.0001

0.93 0.92 0.94 0.9 0.8 0.95 0.96

303.8 306.4 304.7 304.4 303.6 306.8 306.5

0.6 0.6 0.6 0.6 0.7 0.7 0.6

304.4 307 305.2 305 303.6 305.4 305.5

0.6 0.6 0.6 0.6 0.7 0.7 0.6

309 311.2 309.1 309.6 303.9 294.3 297.6

2 2.1 1.9 2.3 3.9 1.9 1.7

1.7 1.6 1.4 1.7 0.1 −4.4 −3.1

Sample 124‐8: G2 Z, Tips (1) Z, LP (1) Z, LP (1) Z, Tips (1) Z, LP (3) Z, LP, “s-br” (5) Z, LP, Py (4) Z, Tips (9) Mz (1) Mz (2)

1 2 3 4 5 6 7 8 9 10

9 1 1 1 1 1 1 13 1 1

300 844 1199 1302 2201 1635 422 868 2630 1654

0.34 0.09 0.21 0.45 0.13 0.23 0.21 0.19 20.12 11.07

7.93 1.2 1.01 1.96 1.12 3.82 1.13 3.77 3.51 3.63

1041 2013 3600 1976 5568 1302 1291 8965 2298 1408

0.348 0.32827 0.35005 0.34038 0.32463 0.34713 0.40389 0.34569 0.34962 0.35115

0.00144 0.00186 0.00192 0.00208 0.00151 0.00153 0.003 0.00094 0.00112 0.0015

0.04802 0.04525 0.04822 0.04714 0.04492 0.0479 0.05432 0.04778 0.04861 0.04872

0.00012 0.00019 0.00023 0.00021 0.0002 0.00014 0.00025 0.00012 0.00011 0.00014

0.62 0.69 0.88 0.77 0.89 0.65 0.65 0.91 0.75 0.61

302.3 285.3 303.6 296.9 283.3 301.6 341 300.9 306 306.7

0.8 1.2 1.4 1.3 1.2 0.9 1.5 0.7 0.7 0.9

303.2 288.2 304.8 297.5 285.5 302.6 344.5 301.5 304.4 305.6

1.1 1.4 1.4 1.6 1.2 1.2 2.2 0.7 0.8 1.1

309.9 312.4 314 301.6 303.5 310 367.8 306.1 292.6 297.3

7.4 9.2 5.9 8.9 4.8 7.6 12.6 2.5 4.8 7.7

2.5 8.9 3.4 1.6 6.8 2.8 7.5 1.7 −4.7 −3.2

Sample 124‐19: G3 Z, LP (10) Z, LP small (15) Z, tips, (9) Z, LP (3) Z, LP, small (14) Z, tips, “s-br” (12) Mz (3) Mz (3)

3 4 1 5 6 2 7 8

34 10 29 13 12 7 11 3

567 778 534 666 528 695 247 960

0.19 0.26 0.19 0.2 0.22 0.29 21.49 26.68

4.47 2.93 4.38 2.97 2.36 1.9 5.86 3.42

13099 8072 10735 8878 8184 8113 1432 2598

0.35073 0.35043 0.35067 0.35054 0.35112 0.37005 0.3512 0.35212

0.00082 0.00085 0.00082 0.00086 0.00097 0.00096 0.00124 0.00106

0.04841 0.04841 0.04846 0.04845 0.04857 0.05061 0.04872 0.04889

0.0001 0.0001 0.0001 0.0001 0.00012 0.00012 0.00011 0.00011

0.95 0.92 0.95 0.91 0.85 0.87 0.66 0.75

304.8 304.8 305 305 305.7 318.3 306.6 307.7

0.6 0.6 0.6 0.6 0.8 0.7 0.7 0.7

305.3 305 305.2 305.1 305.6 319.7 305.6 306.3

0.6 0.6 0.6 0.6 0.7 0.7 0.9 0.8

309 307.3 306.6 306 304.2 329.9 297.9 295.6

1.6 2.2 1.8 2.3 3.3 2.9 6 4.6

1.4 0.8 0.5 0.3 −0.5 3.6 −3 −4.2

G1, G2 and G3 are as in Table 1; Z — zircon; Mz — monazite; LP — long prismatic; “s-br” — small and broken”; (N) — number of grains in fraction. a Weight and concentrations are known to better than 10%, except for those near and below the c. 1 μg limit of resolution of the balance. b Th/U model ratio inferred from 208/206 ratio and age of sample. c Pbc is total common Pb in sample (initial + blank). d Raw data corrected for fractionation. e Corrected for fractionation, spike, blank and initial common Pb; error calculated by propagating the main sources of uncertainty; initial common Pb corrected using Stacey and Kramers (1975) model Pb. f Degree of discordancy.

but at lower FeO contents. These relationships suggest that G2 is not related to G1 and G3. The chondrite-normalized REE patterns (Fig. 5a) are similar for the three rocks, but with differences in the total abundance: 182–257 ppm in G1, 153–167 ppm in G2 and 386–404 ppm in G3. They show similar enrichment in LREE relatively to HREE (La/YbN =25–31) and a negative Eu anomaly (Eu/Eu* =0.34–0.42) (Table 4). The primitive mantle-normalized diagram (Fig. 5b) shows a general negative slope and Rb, Th, U, K, and La positive anomalies, Ce and Ba, Nb and Ti negative anomalies, which are characteristics of a crust dominant source. The negative Ba and Ti anomalies suggest fractional crystallization of mainly K-feldspar, biotite and ilmenite.

partial melting from the Guarda–Sabugal area, central Portugal (Fig. 1a; Neiva et al., 2011) also plot in the same isotopic field, but with somewhat less negative εNdt values (Fig. 6). Several other peraluminous granitic rocks from the CIZ also plot in the field of lower crustal felsic granulites (e.g. Dias et al., 2002; Villaseca et al., 1998, 1999) and some authors suggest mantle–crust interaction (e.g. Castro et al., 1999; Costa, 2011; Dias et al., 2002; Pinarelli and Rottura, 1995; Silva et al., 2000). The mean oxygen isotopic compositions of nine representative samples of G1 to G3 range from 10.90 to 11.36‰ (Table 5). Such high δ18O values are typical of Variscan granitic rocks in Europe (Hoefs and Emmermann, 1983), and have been explained as a result of heterogenous anatexis from sources of metasedimentary origin (Hoefs, 2009).

7. Whole rock Rb–Sr, Sm–Nd and oxygen data

8. Mineral chemistry

Fifteen whole rock samples were analyzed to obtain Rb and Sr isotopic compositions and sixteen samples for the Sm and Nd isotopes (Table 5). The ratios of (87Sr/86Sr)310 =0.7085±0.0005 for G1, (87Sr/ 86 Sr)307 =0.7080±0.0006 for G2 and (87Sr/86Sr)305 =0.7085±0.0006 for G3 indicate that the three units are mainly derived from partial melting of crustal material. The mean εNdt values are −6.6 (G1), −7.1 (G2) and −6.9 (G3) (Table 5). The similar (143Nd/144Nd)i values suggest that these rocks were derived by partial melting of similar crustal material. The mean TDM ages range from 1.3 Ga to 1.5 Ga (Table 5), which are typical values for Variscan granites (e.g. Dias, 2001; Liew and Hofmann, 1988). Granitic rocks from the Penafiel area match the isotopic composition of lower crust felsic granulites (Fig. 6). Granitic rocks formed by

8.1. Feldspars The compositions of K-feldspar and plagioclase are given in Table 6. The orthoclase contents in phenocryst and matrix K-feldspar of the three granitic rocks are similar (80 to 97 mol%; Table 1). The Ba content of K-feldspar decreases from phenocryst to matrix in G1, G2 and G3 and is higher in phenocryst and matrix of G3 than in phenocryst and matrix of G1 (Table 1). Plagioclases from G1, G2 and G3 are zoned, with the anorthite content decreasing from core to rim and from phenocryst to matrix in all three rocks (Fig. 7). The anorthite content of phenocryst and matrix plagioclase from G3 is higher than the anorthite content of phenocryst and matrix plagioclase from G1 (Table 1), whereas G2 has the

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115

Fig. 3. Concordia diagrams displaying the U–Pb data for zircon and monazite of three granitic rocks from the Penafiel area, northern Portugal. Error ellipses are drawn at 2σ. a) Medium- to coarse-grained porphyritic biotite > muscovite granite (G1); b) medium-grained porphyritic biotite ≈ muscovite granodiorite (G2); c) fine-grained porphyritic biotite > muscovite granodiorite (G3); d) two zircon fractions with some inheritance of sample 124‐8 (G2) and 124‐19 (G3). See Table 2 and Section 5 for discussion of the results.

lowest anorthite content for phenocryst and matrix plagioclase. The P2O5 content of both feldspars is ≤0.39 wt.%, showing a low degree of fractionation of G1, G2 and G3 (Table 6). 8.2. Micas The average compositions of biotite and muscovite are given in Table 7. Biotites have Mg/(Mg+ Fe2+ + Fe3+) ranging from 0.29 to 0.40 (Rieder et al., 1999) and compositions similar to those found in biotites from aluminium-potassic rock series of the biotite ± cordierite and biotite± muscovite fields (Nachit et al., 1985). Biotite from granodiorite G3 has more Mg and less Li and F than biotite from granite G1 (Table 7). Muscovites from G1, G2 and G3 have high TiO2 and Al2O3 and low MgO (Table 7) and are therefore magmatic (Miller et al., 1981; Monier et al., 1984). Muscovite from G3 has more Mg than muscovite from G1. 8.3. Accessory minerals Andalusite and sillimanite are accessory minerals in G1 and G2. Andalusite contains 0.33 and 0.31 wt.% FeO in G1 and G2, respectively (Table 7). It is subdehral, without inclusions and is mantled by single crystals of muscovite. It corresponds to the magmatic type of Clarke et al. (2005). Restitic andalusite is rare because it is a low pressure

mineral and thus unlikely to occur in a truly restitic assemblage (Clarke et al., 2005). Cordierite is present in G1 and G3, but could not be analyzed due to its alteration. Tourmaline plots in the alkali group and is of schorl composition (Table 8; Fig. 8a, b, c). Schorl from G1 is the richest in Fe, whereas schorl from G3 is the richest in Mg (Fig. 8d), suggesting a higher melt temperature in G3 than in G1 (Benard et al., 1985). The (□Al)(NaFe)−1 substitution is important in schorl from G1 and G2, whereas the FeMg−1 substitution occurs in schorl from G3 (Fig. 8d). The (□Al)(NaR)−1 substitution dominates in schorl from G1, G2 and G3 (Fig. 8e). Ilmenite has an MnO content ranging from 1.74 to 3.57 wt.% in G1 and 2.04–2.12 wt.% in G3. 9. Discussion 9.1. Anatectic granitic rocks and their protoliths The granitic rocks from northern and central Portugal have been classified according to their relationship to the last Variscan ductile deformation phase (D3) because the majority were emplaced during that phase (Ferreira et al., 1987). The 309.6±1.0 Ma granite (G1) and the 307.0± 3.2 Ma granodiorite (G2) are classified as late D3 (~306–311 Ma) (Dias et al., 1998). They project in the field of syn-collision granites in the R1– R2 diagram (La Roche et al., 1980) and in the tectonic discrimination

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Table 3 Chemical analyses (in wt.%) and trace elements (in ppm) of granitic rocks from the Penafiel area, northern Portugal. G1 SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 LOI Total F Nb Zn Sn Li Zr Cu Y Sr Pb Ba Rb Th U Ga Cr V Sc Co Cs W Be Hf Sb As Bi Be

69.44 0.58 14.97 0.30 2.75 0.04 0.75 1.42 2.87 5.41 0.40 0.89 99.82 0.16 18 90 8 150 238 10 21 137 40 578 320 29 16 27 21 30 6.4 6 16 2 7 7 0.30 13.8 3 7

G2 70.42 0.51 14.64 0.14 2.31 0.03 0.67 1.57 2.80 5.34 0.34 0.99 99.76 0.15 14 440 11 135 202 10 13 120 47 500 336 25 10 25 17 28 4.0 5 16 2 6 6 a

14.6 1 6

70.78 0.47 14.70 0.35 2.19 0.04 0.63 1.27 2.91 5.57 0.34 0.84 100.09 0.14 15 90 11 162 194 a

16 107 41 369 344 24 12 25 12 27 4.0 4 22 2 8 5 0.80 17.3 1 8

71.61 0.51 14.39 0.32 2.40 0.04 0.65 1.10 2.68 4.86 0.36 1.11 100.02 0.13 16 70 10 167 212 10 16 101 34 341 329 27 7 25 19 25 4.9 5 22 2 5 6 0.50 15.9 3 5

72.53 0.42 13.83 0.12 2.27 0.04 0.60 1.12 2.80 4.85 0.33 0.80 99.71 0.16 15 80 10 203 172

74.47 0.32 13.40 0.12 1.76 0.03 0.43 1.04 2.75 4.69 0.26 0.78 100.04 0.11 11 60 11 189 137

70.39 0.39 14.96 0.17 1.85 0.03 0.72 1.12 3.07 4.70 0.36 1.29 99.06 0.15 12 80 9 133 138

G3 71.03 0.37 14.40 0.23 1.68 0.04 0.65 1.19 3.03 4.64 0.34 1.47 99.07 0.16 12 80 18 124 139

71.45 0.37 14.66 0.26 1.66 0.03 0.68 1.04 3.05 4.78 0.34 1.40 99.72 0.15 12 80 9 140 138

71.93 0.39 14.68 0.29 1.92 0.10 0.75 0.89 2.96 4.69 0.34 1.41 100.36 0.17 13 90 17 117 145

72.59 0.25 14.36 0.14 1.35 0.02 0.44 0.71 3.15 4.95 0.33 1.06 99.35 0.15 13 120 16 148 101

72.70 0.27 14.01 0.10 1.50 0.03 0.46 0.89 2.99 4.72 0.33 1.00 99.00 0.15 13 70 11 148 109

72.79 0.37 14.53 0.26 1.59 0.03 0.64 0.79 3.05 4.72 0.33 1.22 100.33 0.19 12 80 21 89 136

73.40 0.26 14.41 0.18 1.34 0.02 0.44 0.82 3.19 4.79 0.35 1.04 100.25 0.16 13 80 14 157 100

65.50 0.97 15.72 0.33 3.90 0.06 1.43 2.26 2.87 5.20 0.45 0.99 99.68 0.23 16 100 7 138 333

a

a

a

a

a

a

a

a

a

a

a

17 102 39 385 370 21 16 25 18 22 4.8 4 27 3 6 5 0.60 15.1 7 6

14 78 37 229 370 19 15 22 11 14 2.9 2 26 2 8 4 0.40 24.9 7 8

11 198 40 472 324 17 13 25 16 23 3.7 3 22 2 3 4 0.30 11.7 3 3

12 167 36 432 349 19 12 25 18 20 3.3 3 20 3 5 4 0.20 14.6 3 5

10 173 39 423 347 19 13 25 12 20 3.7 3 23 2 4 4 0.50 12.8 2 4

11 171 33 459 325 20 16 26 14 23 3.6 3 12 6 4 4 0.30 32.2 1 4

7 106 34 332 386 13 18 25 5 14 2.4 2 23 5 3 3 0.40 36.6 1 3

10 118 27 341 398 14 9 26

11 155 35 384 335 19 7 25 9 22 3.3 2 16 4 4 4 0.40 16.3 1 4

9 118 31 303 398 13 27 26 11 12 2.6 2 22 3 3 3 0.50 4.9 1 3

25 198 46 671 335 60 13 28 20 57 9.0 8 17 5 4 9 0.50 16.2 2 4

a

14 2.5 2 22 3 2 3 0.20 4.6 1 2

65.94 0.84 15.19 0.45 3.31 0.06 1.15 1.83 2.78 5.36 0.40 1.70 99.00 0.23 16 100 8 116 307 20 24 169 52 698 344 59 5 26 15 45 8.0 7 15 3 5 9 a

10.1 1 5

66.12 0.98 14.81 0.38 3.96 0.06 1.41 2.47 2.73 4.62 0.42 1.34 99.30 0.23 15 110 7 115 321 a

26 191 40 642 301 57 13 26 18 56 10.0 9 14 2 5 9 0.50 11.2 1 5

66.23 0.97 15.36 0.59 3.69 0.06 1.44 2.28 2.82 4.92 0.44 1.14 99.94 0.24 16 120 6 115 347 10 26 192 43 662 326 63 8 28 27 57 9.0 9 13 2 5 9

67.38 0.91 14.85 0.73 3.26 0.06 1.30 2.15 2.78 4.81 0.42 0.95 99.59 0.23 15 90 7 118 313

69.10 0.70 14.66 0.17 3.10 0.06 0.97 1.52 2.69 5.28 0.38 1.14 99.76 0.26 16 100 15 274 284

a

a

24 184 40 603 306 58 12 25 23 50 9.0 7 14 2 4 9

23 133 38 510 369 55 14 26 13 38 7.0 6 17 3 5 8 0.30 45.4 1 5

a

a

11.4 2 5

9.9 2 4

G1, G2 and G3 are as in Table 1. a Below the detection limit.

Although they have compositional features of syncollisional granites, the trace element composition of granitoids is determined mainly by the source composition and crystallization history of the melt rather than the tectonic environment (Frost et al., 2001). The three granitic

diagrams of Pearce et al. (1984). The 305.1±0.4 Ma granodiorite (G3) is classified as late- to post‐D3 (ca. 300 Ma), although its age is closer to the age of late D3 granitoids, as it plots in the field of late- to post‐orogenic rocks in the R1–R2 diagram (La Roche et al., 1980). Table 4 Rare earth elements in ppm of granitic rocks from the Penafiel area, northern Portugal. G1 La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu ∑ REE ∑ LREE ∑ HREE (La/Lu)N Eu/Eu*

52.5 112.0 13.5 50.3 9.6 1.1 7.2 1.0 4.9 0.8 2.1 0.3 1.7 0.2 257.2 228.3 4.3 25 0.39

G2 48.0 103.0 12.4 45.7 8.6 0.9 6.5 0.9 4.1 0.6 1.6 0.2 1.2 0.2 233.8 209.1 3.2 31 0.34

G1, G2 and G3 are as in Table 1.

41.4 89.0 10.6 39.0 7.4 0.8 5.5 0.8 3.6 0.6 1.4 0.2 1.1 0.2 201.6 180.0 2.8 29 0.39

37.1 79.4 9.6 35.4 6.8 0.8 5.0 0.8 3.5 0.6 1.6 0.2 1.2 0.2 182.2 161.5 3.2 21 0.40

31.1 66.8 8.1 30.5 5.8 0.7 4.3 0.6 2.7 0.4 1.1 0.1 0.9 0.1 153.3 136.5 2.3 27 0.42

G3 32.1 69.4 8.4 32.1 6.2 0.7 4.2 0.6 2.6 0.4 1.0 0.1 0.9 0.1 158.8 142.0 2.2 30 0.39

32.0 69.2 8.5 31.7 6.1 0.7 4.3 0.6 2.6 0.4 1.1 0.1 0.8 0.1 158.2 141.4 2.2 28 0.38

30.5 67.4 8.2 30.6 6.0 0.7 4.3 0.6 2.6 0.4 1.0 0.1 0.8 0.1 153.3 136.7 2.1 29 0.38

33.8 73.4 8.9 33.6 6.3 0.7 4.3 0.6 2.7 0.4 1.1 0.1 0.9 0.1 167.0 149.7 2.3 29 0.40

77.7 174.0 22.1 86.1 14.3 1.4 9.0 1.1 5.5 0.9 2.6 0.4 2.1 0.3 397.4 359.9 5.3 28 0.35

75.6 170.0 21.5 82.8 14.0 1.3 8.8 1.1 5.4 0.9 2.5 0.4 2.1 0.3 386.6 349.9 5.2 28 0.33

76.5 168.0 21.3 82.6 13.8 1.4 8.8 1.2 5.7 1.0 2.7 0.4 2.2 0.3 385.9 348.4 5.6 26 0.36

77.4 172.0 21.6 83.8 14.2 1.3 8.8 1.2 5.6 0.9 2.5 0.3 2.0 0.3 391.9 354.8 5.1 28 0.33

79.2 177.0 22.5 87.5 14.6 1.3 9.0 1.2 5.6 0.9 2.5 0.4 2.1 0.3 404.0 366.2 5.2 29 0.33

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Fig. 4. Variation diagrams for selected major and trace elements and ratios of the granitic rocks from the Penafiel area. G1, G2 and G3 are as in Table 1. Granitic rocks G1 and G3 seem to be related, but granite G2 is independent.

Fig. 5. a) Subparallel (average) chondrite-normalized REE patterns; b) primitive mantle normalized major, trace elements and REE of the granitic rocks from the Penafiel area. Chondrite normalization values from Taylor and McLennan (1985); primitive mantle normalization values from Sun and McDonough (1989).

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Table 5 Whole rock Rb–Sr, Sm–Nd and O isotopic values of granitic rocks from the Penafiel area, northern Portugal. Sr (ppm)

Rb (ppm)

Rb/Sr

87

G1 123-3 123-4 123-8 123-9 135-1 135-2

107 101 137 – 102 –

344 329 364 – 370 –

3.215 3.257 2.657 – 3.627 –

9.3386 9.4631 7.7125 – 10.5419 –

G2 124-1 124-6 124-7 124-9 124-10 124-12

198 118 106 – 155 171

324 398 386 – 335 325

1.636 3.373 3.642 – 2.161 1.901

G3 124-17 124-18 124-19 124-21 124-22 135-3

184 133 191 198 169 192

306 369 301 335 344 326

1.663 2.774 1.576 1.692 2.036 1.698

Sample

Rb/86Sr

87

Sr/86Sr

σ (%)

(87Sr/86Sr)i Sm (ppm)

Nd (ppm)

147

Sm/144Nd

0.74953 0.75072 0.74267 – 0.7544 –

0.0006 0.0006 0.0005 – 0.0006 –

0.7084 0.7090 0.7087 – 0.7080 –

7.68 8.19 9.66 7.88 6.86 –

41.3 44.1 52.4 43.5 36.5 –

0.1126 0.1124 0.1115 0.1096 0.1137 –

4.7438 9.7977 10.5826 – 6.2693 5.5115

0.72931 0.74987 0.75426 – 0.73535 0.73254

0.0006 0.0006 0.0006 – 0.0006 0.0006

0.7086 0.7071 0.7080 – 0.7080 0.7085

6.38 4.36 4.41 – 6.49 6.77

33.0 23.5 21.9 – 33.0 35.5

4.8213 8.0551 4.568 4.905 5.9037 4.9222

0.7296 0.74451 0.72815 0.72958 0.73421 0.72917

0.0006 0.0006 0.0006 0.0006 0.0006 0.0006

0.7087 0.7095 0.7083 0.7083 0.7086 0.7078

14.4 13.7 14.4 14.6 14.1 14.8

85.3 78.6 85.4 87.6 82.4 86.9

143

Nd/144Nd

2σ

(143 Nd/144Nd)i εNdt

TDM δ18O

0.512131 0.512118 0.512135 0.512131 0.512130 –

0.000005 0.000005 0.000005 0.000004 0.000004 –

0.51190 0.51189 0.51191 0.51191 0.51190 –

−6.6 −6.9 −6.5 −6.5 −6.7 –

1.4 1.4 1.4 1.3 1.4 –

– 10.92 – – 11.62 11.54

0.117 0.1251 0.1215 – 0.119 0.1154

0.512106 0.512130 0.512123 – 0.512115 0.512137

0.000004 0.000005 0.000004 – 0.000005 0.000002

0.51187 0.51188 0.51188 – 0.51188 0.51191

−7.3 −7.1 −7.1 – −7.2 −6.6

1.5 1.6 1.5 – 1.5 1.4

– 11.27 – 11.18 – 11.47

0.1021 0.1054 0.1018 0.1010 0.1032 0.1033

0.512095 0.512098 0.512106 0.512103 0.512080 0.512090

0.000002 0.000003 0.000001 0.000003 0.000002 0.000006

0.51189 0.51189 0.51190 0.51190 0.51187 0.51188

−6.9 −7.0 −6.7 −6.7 −7.2 −7.0

1.3 1.3 1.3 1.3 1.3 1.3

– 11.33 10.82 – 10.54 –

G1, G2 and G3 are as in Table 1. (–) Not determined. TDM age is in Ga, calculated based on values of De Paolo (1981). δ18O values are in ‰.

rocks contain aluminum-rich minerals such as biotite, muscovite, cordierite, andalusite and sillimanite. They also have ilmenite, normative corundum >2.29%, K2O > Na2O, low CaO/Na2O, A/CNK > 1.06, (87Sr/ 86 Sr)i = 0.7071 to 0.7095, εNdt = −7.3 to −6.5 and δ18O = 10.54 to 11.62‰ (Table 5). All these characteristics point to an affinity with S-type magmas. These considerations imply that these granitic rocks came from similar sources. According to Jung and Pfänder (2007), the CaO/Na2O ratio can be used to infer the source composition of peraluminous granites, as metapelitic rocks are poor in CaO with CaO/Na2Ob 0.5, in contrast to metagreywacke or meta‐igneous rocks which have CaO/Na2O=0.3– 1.5. These ratios are 0.45 and 0.75 for G1 and G3, respectively, suggesting a metagreywacke/meta‐igneous source, whereas G2 has CaO/Na2O=0.3 and therefore its magma source is not well constrained. The ratios of (87Sr/86Sr)i are in the same range in G1 (0.7080– 0.7090) and G3 (0.7078–0.7095) (Table 5). The same is true for εNdt G1 = −6.9 to −6.5 and G3 = −7.2 to −6.7, and δ18O with 10.92– 11.62‰ in G1 and 10.54–11.32‰ in G3 (Table 5). The similarity

Fig. 6. Plot of εNdinitial vs. εSrinitial at 307 Ma for the granitic rocks from the Penafiel and Guarda–Sabugal areas. Results of field projections for Ordovician and Silurian metasediments and orthogneiss (Beetsma, 1995), schist graywacke complex (SGC), southern sector (Beetsma, 1995; Tassinari et al., 1996) and northern sector (NE Trás-os-Montes) (Teixeira, 2008), lower crust felsic granulites (Villaseca et al., 1999). Data of two granites from the Guarda–Sabugal (Neiva et al., 2011) and those from the Penafiel areas all plot in the field of felsic granulites.

supports a provenance of G1 and G3 from the same metagreywacke/ meta‐igneous source. The source of G2 granodiorite is isotopically slightly different with less radiogenic (87Sr/86Sr)i (0.7071–0.7086), but broadly comparable Nd composition (−6.6 to −7.3; Table 5) and δ18O (11.18–11.47‰). These isotopic compositions imply sources with similar crustal histories, whereas the chemical compositions indicate a stronger contribution from metapelitic components for G2 than G1 and G3. G2 has the highest A/CNK (1.27) and the highest contents of micas, particularly of muscovite (6.6–10.2%), whereas G1 and G3 have muscovite contents of 2.1–2.3% and 1.9%, respectively. Andalusite and cordierite are rare in these granites. Inherited zircons from G2 and G3 of the Penafiel area are ~587 Ma (Fig. 3d), indicating that detritus of Neoproterozoic age was involved in the origin of these granitic rocks. The isotopic compositions of the three granitic rocks are similar to those of lower crustal felsic granulite xenoliths presented by Villaseca et al. (1999) (Fig. 6), and also compares with the composition of other Portuguese Variscan granites (e.g. Azevedo and Valle Aguado, 2006; Costa, 2011; Martins et al., 2009). Major and trace element modelling (Villaseca et al., 1999) supports the idea that late-Variscan peraluminous granites from the Spanish Central System represent liquids in equilibrium with lower crustal granulites having isotopic compositions similar to those of the peraluminous granites and that the granites were derived from partial fusion of meta‐igneous sources (Villaseca et al., 1998, 1999). Two Variscan biotite>muscovite granites from the Guarda–Sabugal area also plot in the field for lower crustal felsic granulites (Fig. 6), suggesting a similar origin. Although Nd is comparable, the Sr isotopic compositions of granitic rocks from Penafiel and Guarda–Sabugal are lower than those of the surrounding schist–greywacke complex and orthogneisses (Fig. 6). This mismatch could imply that there is an apparent absence of isotopically appropriate crustal protoliths, as has also been discussed for other areas (e. g. Bernard-Griffiths et al., 1985; Bickle et al., 1988; Clark et al., 1988; Liew et al., 1989; Peucat et al., 1988; Villaseca et al., 1998). In distinct regions, however, the Sr isotopic compositions of lower continental crust tend to be less radiogenic than at upper crustal levels (e.g. Eberz et al., 1991; Villaseca et al., 1999). The 87Sr/86Sr ratios are lowered with increasing metamorphic grade until anatectic conditions are reached, because the aqueous fluids expelled during prograde metamorphism homogenize these ratios (Bickle et al., 1988). Isotopic disequilibrium during partial melting, for example the early dissolution and loss of

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Table 6 Selected chemical compositions (in wt.%) of K-feldspar and plagioclase of the granitic rocks from the Penafiel area, northern Portugal. K-feldspar

Plagioclase

G1

SiO2 TiO2 Al2O3 FeO MnO MgO BaO CaO Na2O K2O P2O5 Total An (mol%) Or (mol%)

G2

G3

G1

G2

G3

Phen

matrix

Phen

matrix

Phen

matrix

Phen

matrix

Phen

matrix

Phen

matrix

64.52 – 18.08 – – – 0.12 0.04 0.75 15.87 0.15 99.53

64.67 – 18.14 – – – 0.07 – 0.39 16.86 0.26 100.39

63.74 0.03 19.13 – – – 0.31 0.06 1.05 15.15 0.22 99.69

64.21 0.04 19.11 – – – 0.09 – 0.74 15.68 0.26 100.13

64.57 0.09 19.08 0.00 0.06 0.00 0.38 0.07 1.41 14.60 0.07 100.35

65.34 – 17.72 0.03 – – 0.12 – 0.97 15.91 0.13 100.22

66.43 – 20.74 – 0.03 – – 1.67 10.11 0.27 0.39 99.64 8

67.60 – 20.01 – – – – 0.98 10.66 0.07 0.20 99.52 5

63.65 – 22.86 – – – 0.07 3.10 9.46 0.24 0.23 114.70 15

68.04 0.03 20.46 – – – 0.03 0.47 11.25 0.15 – 102.69 2

61.83 – 24.02 0.21 – 0.10 – 4.70 8.63 0.10 0.12 99.71 23

64.89 – 22.44 – – – – 2.69 9.82 0.08 – 99.93 13

93

97

90

93

86

91

Phen — phenocryst, – not detected, analyst: P. C. S. Carvalho.

biotite and K-feldspar, can also cause progressive isotopic changes with increasing metamorphic grade (Villaseca et al., 1998). Therefore, granitic magmas derived from a source at depth will not necessarily have the same isotopic composition, particularly for Sr, as the equivalent metamorphic rocks at the level of granitic emplacement. In orogenic areas, granite sources are not the outcropping metamorphic rocks, but are located at deeper crustal levels (e. g. Hanchar et al., 1994; Miller et al., 1992; Villaseca et al., 1999). In the middle and lower crust of the Spanish Central System the rocks are mainly meta-igneous and metapelitic and have the same Nd isotopes, suggesting that tectonic breaks do not occur between the middle- and lower-crust (Villaseca et al., 1999) as in crustal segments from Nova Scotia (Eberz et al., 1991). All three units G1, G2 and G3 of the Penafiel area contain metasedimentary enclaves, but these enclaves are, most probably, xenoliths from the country rocks as they are metapelitic. A metagreywacke or meta‐igneous source for these granites is thus more probable. As Clemens (2003) points out, metasedimentary enclaves represent mid-crustal xenoliths rather than restites as they do not have a melt-depleted composition, biotite is a stable phase, and they do not contain a refractory mineralogy. The Penafiel granites

also contain some microgranular enclaves, but more mafic igneous rocks do not crop out. 9.2. Sequential partial melting of G1 and G3 Granodiorite G3 has a higher biotite/muscovite proportion (7.3) than granite G1 (3.1) (Table 1). It also has more TiO2, total FeO, MgO, CaO, F, V, Zr, Y, Sr, Ba, and REE, less SiO2, and lower Rb/Sr and Rb/Ba values than G1 (Table 3, Figs. 4 and 5). Each trend defined by granite G1 in the variation diagrams is independent, but seems to continue into that of granodiorite G3, suggesting that G1 and G3 are related. But, a fractional crystallization process is not adequate to explain the genesis of these two granitoids, because G3 is the less evolved and 4 M.y. younger. The Al2O3/TiO2 ratios can be used to infer melting temperatures of unfractionated magmas. Peraluminous granites with low Al2O3/TiO2 ratios are generated at higher temperatures than those with high Al2O3/ TiO2 ratios, largely independent of source composition (Jung and Pfänder, 2007; Sylvester, 1998). As the variations diagrams suggest (Fig. 4), to evaluate fractionation within each granitic suite (G1 and

Fig. 7. Plot of plagioclase compositions of granitic rocks from the Penafiel area, showing that anorthite content decreases from phenocryst to matrix plagioclase and anorthite content in phenocryst and matrix plagioclase from G3 is higher than in those from G1.

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Table 7 Average chemical analyses (in wt.%) of biotite, muscovite and andalusite of granitic rocks from the Penafiel area, northern Portugal. Biotite

SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Li2O Na2O K2O F Cl H2O* O≡F O ≡ Cl Total

Muscovite

G1

G2

G3

34.94 2.80 18.85 1.30 21.86 0.20 5.86 – 0.25 0.12 9.72 0.62 0.08 3.58 100.21 0.26 0.02 99.93

35.24 2.91 19.07 0.84 20.56 0.25 6.87 – 0.20 0.12 9.85 0.51 0.05 3.67 100.16 0.21 0.01 99.94

34.86 2.77 18.35 1.26 20.87 0.21 6.95 – 0.20 0.08 9.86 0.33 0.12 3.70 99.59 0.14 0.03 99.42

G1 SiO2 TiO2 Al2O3 FeOt MnO MgO CaO Li2O Na2O K2O F H2O*

47.15 0.99 34.77 1.51 – 0.82 – – 0.47 10.25 – 4.54 100.5 O≡F – O ≡ Cl – Total 100.5

Andalusite

G2

G3

G1

G2

46.13 0.98 34.44 1.48 – 0.99 – 0.07 0.66 10.52 0.26 4.37 99.96 0.11 – 99.85

47.62 0.56 34.22 1.42 – Total 1.08 – – 0.43 10.55 – 4.54 100.42 – – 100.42

36.83 36.77 – 0.05 62.31 62.51 0.33 0.31 99.47 99.59

FeOt of muscovite and andalusite is total FeO. – not detected. H2O* values of micas were calculated by stoichiometry. Analyst: P. C. S. Carvalho.

G3), the least evolved sample composition was used. Since G3 has a lower Al2O3/TiO2 ratio (16.20) than G1 (25.99), the latter originated at a lower temperature. Estimates of the conditions of formation of granitic magmas may be obtained from the zircon saturation equation (Watson and Harrison, 1983), assuming equilibrium conditions. The average zircon saturation temperature (Tzr) is of 847 °C for G3 and 813 °C for G1. These Tzr values are underestimated, because there are rare inherited zircons. Anyway, the Tzr is higher for G3 than G1, indicating a higher degree of partial melting for G3 (Miller et al., 2003). Phenocryst and matrix microclines from G3 have higher Ba contents than the corresponding microclines in G1 (Table 1). Anorthite contents of phenocryst and matrix plagioclase from G3 are higher than those of phenocryst and matrix plagioclase from G1 (Fig. 7). Microcline and plagioclase from G3 have less P2O5 than the feldspars in G1 (Table 6). Biotite from G3 has more Mg and less Li and F than biotite from G1 (Table 7) and muscovite from G3 has a higher Mg content than muscovite from G1 (Table 7). Schorl from G3 has more Mg than schorl from G1 (Table 8, Fig. 8d). All these features support that G3 was formed at a higher temperature than G1. The geochemistry of minerals from these two granitic rocks also confirms that they are not related by a fractional crystallization mechanism. Table 8 Representative chemical analyses in wt.% of tourmaline of granitic rocks from the Penafiel area, northern Portugal. G1 SiO2 TiO2 B2O3* Al2O3 FeO MnO MgO CaO Li2O* Na2O K2O H2O* F O≡F Total Name

35.85 1.08 10.43 32.29 10.20 0.05 3.36 0.35 0.42 1.80 0.03 3.51 0.18 99.54 0.08 99.47 Schorl

G2 35.65 0.70 10.47 33.39 9.94 0.07 3.34 0.31 0.37 1.80 0.05 3.50 0.24 99.83 0.10 99.73 Schorl

36.10 1.18 10.50 32.59 9.23 0.05 3.50 0.37 0.55 1.91 0.04 3.53 0.20 99.75 0.08 99.67 Schorl

G3 35.65 0.74 10.48 33.29 9.00 0.07 3.86 0.32 0.40 1.92 0.05 3.48 0.28 99.53 0.12 99.41 Schorl

36.53 0.51 10.50 31.83 8.66 0.03 4.73 0.46 0.53 2.03 0.04 3.49 0.28 99.62 0.12 99.50 Schorl

36.42 0.28 10.54 32.97 8.60 0.03 4.53 0.30 0.39 2.06 0.03 3.61 0.06 99.82 0.03 99.79 Schorl

Values inferred from considerations of stoichiometry. Analyst: P. C. S. Carvalho.

The restite unmixing model of White and Chappell (1977) requires significant amounts of restitic minerals and also linear trends in the variation diagrams. By contrast, G1 and G3 do not contain restitic minerals (except for some zircon cores). Restitic minerals identification can be blurred if they were retrogressed into biotite (Stevens et al., 2007), but such an evidence was not found. Cordierite is rare and totally altered and its origin cannot be defined. Anyway, the variation diagrams do not show linear trends (Fig. 4) and hence magma mixing cannot explain the variations between G1 and G3. Other authors (e.g. Villaseca et al., 1998) point out that in the Iberian Massif the basic and intermediate rocks, associated with peraluminous granites, do not seem to have played a major role in controlling the chemical variability of the peraluminous granites. As argued above, the geochemical and isotopic data indicate that G1 and G3 are derived by partial melting from the same crustal source, but they cannot be related through a fractional crystallization process, as granodiorite G3 is the less evolved, was formed at a higher temperature and emplaced up to ~ 4 M.y. later than granite G1. Therefore, granitic rocks G1 and G3 were derived from sequential partial melting from the same lower crustal material, as other peraluminous granitic rocks in CIZ (Neiva et al., 2011) and that crustal source was probably metagreywacke (Fig. 6) (Neiva et al., 2011). Partial melting of metasedimentary sources is the response to dehydration of hydrous minerals (muscovite and biotite) (Clemens and Vielzeuf, 1987). The process begins at less than 850 °C with the breakdown of muscovite. This stage of melting formed G1. At higher temperatures of 850–900 °C biotite breaks down producing abundant peraluminous melt (Clemens and Vielzeuf, 1987; Vielzeuf and Holloway, 1988). Granite G3 is the result of this stage of melting. The granitic magmas G1 and G3 evolved separately by fractional crystallization of quartz, plagioclase, K-feldspar, biotite, ilmenite and apatite, as suggested by decreasing Al2O3, MgO, total FeO, Sr, Ba, and P2O5, and increasing SiO2, Rb/Sr and Rb/Ba (Fig. 4).

9.3. Comparison with other granites derived by sequential partial melting The sequential partial melting evolution from Variscan biotite> muscovite granite G1 (309.6±1.0 Ma) to biotite>muscovite granodiorite G3 (305.1±0.4 Ma) from the Penafiel area can be compared with those of just a few other Portuguese Variscan granites: a) biotite > muscovite granite G2 (301.4 ± 2.6 Ma) and biotite > muscovite G5 (299.8±0.6 Ma) from the Guarda–Sabugal area, central Portugal (Neiva et al., 2011); b) felsic granites, heterogeneous and cordierite–biotite granites from the Tourém complex (303±6 Ma), northern Portugal (Holtz and Barbey, 1991; Holtz and Juteau, 1987; Neiva, 1994) (Fig. 1a). Granite G1 and granodiorite G3 from the Penafiel area, and granites G2 and G5 from Guarda–Sabugal are restite-free, except for restitic zircon grains. In the Tourém complex felsic granites are also restite-free, but heterogeneous and cordierite–biotite granites have high contents of residual phases (Holtz and Barbey, 1991) and were derived by partial melting of orthogneiss, as has been shown experimentally (Holtz and Johannes, 1991). The granitic rocks from Penafiel (G1, G3) and Guarda–Sabugal (G2, G5) have Sr–Nd isotopic compositions similar to those of felsic granulite xenoliths (Fig. 6) and probably were derived from partial melting of metasedimentary sources in the lower crust. In all three areas, the increase in the degree of partial melting is marked by the increase in whole rock TiO2, total FeO, MgO, CaO, Zr and Ce contents, while SiO2 decreases and the anorthite content of plagioclase and Mg contents of biotite and muscovite also increase. At Penafiel and Guarda–Sabugal there is also an increase in the biotite/muscovite ratio and whole rock Sr, Ba and REE contents, whereas the P2O5 contents of K-feldspar and plagioclase and the F content of biotite decrease with increasing degree in partial melting due to the higher melting temperature.

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121

Fig. 8. Tourmaline compositions of granitic rocks from the Penafiel area. a) Classification of the principal groups of tourmaline based on X-site (from Hawthorne and Henry, 1999), showing the location of the diagram b; b) tourmalines plot in the alkali group. c) X-vacancy/(Na + X-vacancy) versus YMg/Y(Mg + Fe) diagram of schorl from G1, G2 and G3, where schorl from G3 is the richest in Mg/(Mg + Fe); d) Mg versus Fe2+ diagram, where schorl from G3 has the highest Mg content and schorl from G1 has the highest Fe content; The FeMg−1 substitution occurs in schorl from G3 and the (□Al)(NaFe)−1 substitution is important in schorl from G1 and G2; e) X-site vacancy versus total Al; R represents the sum of divalent cations. The (□Al)(NaR)−1 substitution dominates in schorl from G1, G2 and G3.

Granites generated by sequential melting also occur in the Achiras complex, Córdoba, Argentina, and range from leucogranites to biotite granites and biotite ± amphibole granites (Otamendi et al., 1998). These leucogranites are 384 Ma old. There is also an increase in TiO2, total FeO, MgO and CaO with the decrease in SiO2 contents of rocks and the anorthite content of plagioclase and Mg content of biotite also increase with increasing degree of partial melting, as in the Portuguese granitic rocks from the three areas. Thus, although granitic rocks generated by sequential partial melting are uncommon, they can be found in Europe and South America in orogens of distinct ages, Late Carboniferous and Early Devonian, respectively. The sequential partial melting process implies that the source area is subjected to increasing temperature or decreasing pressure. The increase in temperature can be achieved by an increase of heat from upwelling mantle and crustal thinning. Granitic plutonism is commonly associated with post-orogenic or anorogenic crustal extension (Clemens, 2003) and the Variscan D3 deformation phase, in the Central Iberian Zone, is related to the post thickening extension of the orogen. After thickening of the crust during the compression phases ductile extension caused partial melting of the lower felsic crust. The components with the lower melting point will melt first, leading to the formation of voluminous peraluminous magmas of granitic composition. Some mantle fusion, restricted in space and time also occurred and the mafic magmas intrude the granite magma chambers, originating the microgranular enclaves. The more important factor is that crustal thinning was followed or accompanied

by mantle upwelling, which caused an increase in temperature in the thinned crust (Gutiérrez-Alonso et al., 2011). As extension continues, the continuous decrease in pressure of the thin hot crust will allow the extraction of subsequent partial magmas from the more refractory components of the crust, resulting in the production of later granodioritic peraluminous magmas. Granitic rocks from the Penafiel area, granites G2 and G5 from the Guarda–Sabugal area and granites from the Tourém complex contain mafic microgranular enclaves, which are unrelated to the country rocks, implying transport from deeper levels of the crust and supporting mantle upwelling. Sequential partial melting granitic rocks are rare, because they must be derived from the same source materials, having the same bulk composition and with an increasing degree of partial melting. A derivation from the same source can be demonstrated either through melting experiments performed on the presumed source rock (Holtz and Johannes, 1991) or using ( 87Sr/ 86Sr)i, εNdt and δ 18O data of granitic rocks (Neiva et al., 2011; this study). The generation of granitic rocks by sequential partial melting requires high geothermal gradients and these are mainly established by crustal thickening and subsequent collapse and extension (Matte, 1991; Ribeiro et al., 1990). The reason why sequential partial melting granitic rocks mainly occur in Portugal is due to the combination of extreme crustal thickening during the Variscan orogeny and the subsequent extension and mantle upwelling. This is also supported by the fact that granodiorite G3 from the Penafiel area, formed at higher

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temperatures, contains microgranular enclaves that are larger than in granite G1. The anatectic granites from Argentina derived from the same mechanism were generated during the Farmatinian orogenic cycle (Otamendi et al., 1998). 9.4. The other anatectic granite (G2) Granodiorite G2 defines a distinct trend in the variation diagrams for major and trace elements, subparallel to the trend of granite G1 (Fig. 4). Granodiorite G2 has the lowest total FeO, Zr and REE contents (Figs. 4 and 5; Table 3) and its phenocryst and matrix plagioclase have the lowest anorthite content (Table 1, Fig. 7). The ratio (87Sr/86Sr)i =0.7080± 0.0006 of G2 is slightly lower than the value for G1 and G3 and εNdt =−7.1 is the lowest and TDM ages are the oldest (Table 5), suggesting that it is derived from a slightly different source (Fig. 6). The bulk composition of the source rock is not the only factor that controls the isotopic composition of the melts, because the stoichiometry of the melting reaction and the kinetics of melting also play important roles, as shown experimentally (Hammouda et al., 1994). The G2 magma had a higher contribution from metapelitic components than G1 and G3, as suggested by the relatively low CaO/Na2O. The whole rock Al2O3/TiO2 value of 37.97 for G2 is higher than for G1 (25.99) and G3 (16.20), indicating that G2 was generated at a lower temperature than the others (Jung and Pfänder, 2007). Moreover, the average Zr content is lower for G2 (126 ppm) than for G1 (193 ppm) and G3 (290 ppm), showing that G2 was formed at a lower temperature than G1 and G3. This is supported by the estimated melt temperature of 781 °C obtained from the zircon saturation equation (Watson and Harrison, 1983), that is lower than those estimated for G1 (813 °C) and G3 (847 °C). The age of G2 is indistinguishable within uncertainty from those of G1 and G3 (Fig. 3a, b, c). The three granitic rocks were generated during the same Variscan tectonic event. G2 was generated at a lower melting temperature than G1, as a decompression event may be polybaric with decreasing P leading to a retrograde cooling during magma ascent (Clemens and Vielzeuf, 1987). Therefore, differences in the more metapelitic bulk composition of the source for G2 than for G1 and G3 are probably the main reason for the compositional differences between these granite units. 10. Conclusions In the Penafiel area, Variscan peraluminous granitic rocks have mineralogical, chemical and isotopic characteristics of S-type suites. The three granitic rocks are restite free and formed at 309.6±1.0 Ma (G1), 307.0±3.2 Ma (G2) and 305.1±0.4 Ma (G3) based on zircon. Monazite in all three units records the age of 305 Ma of the youngest intrusion. Granite G1 and granodiorite G3 have distinct major, trace and rare earth element contents and compositions of feldspars, micas and schorl. They have similar ( 87Sr/ 86Sr)i, εNdt and δ 18O values suggesting a similar source, probably a metagreywacke. In the εSri–εNdi diagram they plot in the field of lower crustal felsic granulites. Their Sr isotopic compositions are lower than those of country rock greywackes, suggesting that similar material re-equilibrated during granulization of the lower Variscan crust, losing radiogenic Sr. Granite G3 was generated at a higher temperature and from a higher degree of partial melting of the same source material than G1. Granodiorite G2 corresponds to a distinct pulse of magma with comparable isotopic characteristics as G1 and G3, but probably derived from a metagreywacke containing more abundant pelitic components, as suggested by the CaO/Na2O ratio. Granitic rocks derived by sequential partial melting are uncommon. The fact that they mainly occur in the Variscan of Portugal is due to the unusual circumstances related to Variscan extreme crustal thickening and subsequent extension and mantle upwelling, possibly with delamination, which provided the heat essential in the sequential melting process.

Acknowledgments This work is a part of the PhD thesis of P. C. S. Carvalho. Thanks are due to Prof. Andres Cuesta Fernandez for the facilities at Laboratório de Geologia e Geocronologia of the University of Oviedo, Spain; Prof. G. Ibarguchi for the Rb–Sr and Sm–Nd isotopic data obtained at Departamento de Mineralogía-Petrología, Universidad del País Vasco, Bilbao, Spain; Prof. F.J. Longstaffe for the oxygen isotope analyses obtained at the Department of Earth Sciences, University of Western Ontario, Canada. The comments of Prof. C. Villaseca and an anonymous referee helped to improve this paper. Funding was provided by a FCT (Portuguese Foundation for Science and Technology) PhD grant for P. C. S. Carvalho (SFRH/BD/21373/2005) and Geosciences Centre, University of Coimbra.

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