U-Pb SYSTEMATICS and the PETROGENETIC EVOLUTION of INFRACRUSTAL

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U-Pb SYSTEMATICS AND THE PETROGENETIC EVOLUTION OF INFRACRUSTAL

ROCKS IN THE PALEOZOIC BASEMENT OF WESTERN GALICIA (NW SPAIN)

R.P. KUIJPER

VERHANDELING NR. 5 ZWO LABORATORIUM VOOR

ISOTOPEN-GEOLQGIE, AMSTERDAM; 1979 i

The research described in this thesis was carried out at

the ZWO Laboratorium voor Isotopen~Geologiet Be Boelelaan

108St 1031 HV Amsterdam, supported by the Netherlands Or- ganisation for the Advancement of Pure Research (ZWO), Xt forms part of the investigations by the "Wovking Group Galiaia" of the State University at Leiden.

A separate paper is in preparation "Evidence for a Late Archean/Early Proterozoic continental crust in the Hespe- rian Massif of N.W. Spain by means of U-Pb zircon dating" by R.P. Kuijper, N.A.I.M. Boelrijk, E.H. Hebeda, H.N.A. Priem, E.A.Th. Verdurmen & R.H. Verschure. t:

U-Pb SYSTEMATICS AND THE PETROGENETIC EVOLUTION OF INFRACRUSTAL ROCKS IN THE PALEOZOIC BASEMENT OF WESTERN GALICIA (NW SPAIN)

PROEFSCHRIFT

ter verkrijging van de graad van doctor in de Wiskunde en Natuurwetenschappen aan de Rijksuniversiteit te Leiden, op gezag van de Rector Magnificus Dr. A.A.H. Kassenaar, Hoogleraar in de faculteit der Geneeskunde, volgens besluit van het College van Dekanen te verdedigen op woensdag 19 december 1979 te klokke 14.15 uur

door

RENE PAUL KUIJPER

geboren te 's-Gravenhage in 1949

- 1 - f.

PROMOTOREN : PROF. DR. H.N.A. PRIEM PROF. DR. E. DEN TEX

CO-REFERENTEN: DR. E.H. HEBEDA DR. A.C. TOBI

- 2 - STELLINGEN

1. Lokale gravitnetrische informatie is slechts van beperkte waarde voor de interpretatie van geologische structuren. Bayer* R. & Matte, Ph., 1979: Is the mafia/ ultramafia massif of Cabo-Ortegal (northwest Spain) a nappe emplaced during a Variaaan obduation? - A new gravity interpretation. Teatonophysics 57, T9-T18. Overmeeren, R.A. Van, 1975: A gravity investigation of the oatazonat rock complex at Cabo Ortegal (NW. Spain). Teatonophysias 26, 293-307.

2. De term discordia in U-Pb concordia-diagrammen dient slechts te worden gebruikt voor lijnen welke gedefinieerd worden door co-genetische monsters.

3. De veronderstelling dat de lage initiële B7Sr/8BSr verhoudins in archaeïsche en proterozoïsche gneis-complexen noodzakelijker- wijs duidt op een juveniele herkomst van de oorsprongsgesteenten, is in zijn algemeenheid onjuist. Moorbath, S., 1975: Evolution of Preaambrian arust from strontium isotopio evidence. Nature 254, 395-398.

4. Een opeenvolging van metamorfe fasen mag slechts dan polymetamorf worden genoemd wanneer het totale geologische beeld wijst op een polyorogene ontwikke1ing.

5. Ontwikkeling van een op waterstof berustende energie-cyclus kan een belangrijke bijdrage leveren tot het oplossen van het energie probleem.

6. De basalt-eklogiet transformatie in de oceanische lithosfeer is onvoldoende om plaat-bewegingen te initiëren.

7. De ogenschijnlijke afwezigheid van driekwart van de huidige aard- korst vóór 1.0 Ga geleden kan niet worden verklaard door een ver- onderstelde expansie van de aarde. Glikson, A.Y., 1979: The missing Precarribrian crust. Geology 7, 449-454.

8. Wanneer sommige ziekten als "volksvijand nummer één" worden aange- duid, is het logischer de bestrijding van deze kwalen te finan- cieren uit de defensie-begroting dan met behulp van collecte-bussen.

Stellingen behorende bij dissertatie R.P. Kuijper. L

... e ca esperanza, ca esperanza ardente! de a Galicia tornar ... (Terra a nosa, Rosalia de Castro)

Aan Vivian, Cleo en Roderick

- 3 - ""*•?_ ~--l

ACADEMISCHE STUDIE

Op verzoek van de Faculteit der Wiskunde en Natuurwetenschappen volgt hieronder een overzicht van de academische studie van de schrijver.

In september 1967 werd hij voor de eerste maal ingeschreven als student in de geologie aan de Rijksuniversiteit te Leiden. In juni 1972 legde hij het kandi- daatsexamen (letter 63) af en in september 1975 het doctoraalexamen Geologie. Hij studeerde onder leiding van de hoogleraren Dr. A. Brouwer, Dr. J.G. Hagedoorn, Dr. P, Hartman, Dr. A.J. Pannekoek, Dr. E. den Tex en Dr. H.J. Zwart.

Voor het eerste bijvak petrologie beschreef hij de geologie van een gebied in het Hesperisch Massief in westelijk Galicië (NW. Spanje). Voor het tweede bijvak kristallografie werd een studie gemaakt van de Si-Al orde-wanorde relaties in alkali-veldspaten. Van augustus 1974 tot oktober 1975 was hij als student- assistent verbonden aan de vakgroep Petrologie, Mineralogie en Kristallografie.

Van oktober 1975 tot juni 1979 was de schrijver in dienst van de Nederlandse Or- ganisatie voor Zuiver-Wetenschappelijk Onderzoek (ZWO) en uit dien hoofde als doctoraal-assistent verbonden aan het ZWO Laboratorium voor Isotopen-Geologie te Amsterdam.

- 4 - CONTENTS SUMMARY

I. INTRODUCTION 1.1. Scope of work 9 1.2. Geological setting 9 1.3. Previous geochronological research 12 1.4. Constants used 14

II, U-Pb SYSTEMATICS OF MINERALS 11.I. Introduction 16 11.2. Concordia diagrams 16 11.3. Geological setting and petrography of the investigated rock units 18 11.3.1. The Mellid Complex 18 11.3.2. The Sobrado/Teijeiro Complex 21 11.3.3. The Blastomylonitic Graben 24 11.3.4. Late-Paleozoic Granitic Rocks 25 11.4. Rb-Sr investigations 27 11.5. Results and discussion of the U-Pb mineral analyses 28 11.5.1. Orthogneisses of the Mellid Complex 28 11.5.2. Gneisses of the Sobrado/Teijeiro Complex 33 11.5.3. Gneisses of the Blastomylonitic Graben 38 11.5.4. Late-Paleozoic Granites 42 11.6. General discussion of the U-Pb mineral data 11.6.1. The zircon upper intercept ages 45 11.6.2. The zircon lower intercept ages 45 11.6.3. Relation between the degree of disturbance of the zircons and whole-rock composition 46 11.6.4. Monazites 46 II.G.5. Source rocks of the granites and orthogneisses 47 11.7. Concluding remarks 48

III. PETROGRAPHY AND WHOLE-ROCK U-Pb SYSTEMATICS OF HIGH-GRADE METAMORPHIC MAFIC ROCKS 111.1. Introduction 50 111.2. Metamorphic history of the high-grade mafic rocks south of Teijeiro 50 111.2.1. Granulite facies and younger metamorphic events 51 111.2.2. Eclogites 52 111.2.3. Eclogite-granulite relations 52 111.2.4. Timing of the clinopyroxene exsolution 53 111.2.5. P-T conditions 54 111.3. Investigated samples 111.3.1. Samples 55 111.3.2. Whole-rock compositions 57 111.3.3. Mineral compositions 58 111.4. U-Pb whole-rock data and discussion 60 111.5. Conclusions 66

- 5 - IV. THE OROGENIC EVOLUTION OF THE HESPERIAN MASSIF IV.1. Introduction 67 IV,2. An outline of the geological history of the Hesperian Massif 67 IV.2.1, Supracrustal history. 67 IV.2.2. Infracrustal history 73 IV.3. Inferred geological evolution of the Hesperian Massif 74 IV.4. Tectonic setting of the Hesperian Massif 77 IV.5. Concluding remarks 78

REFERENCES 80

ACKNOWLEDGEMENTS 87

APPENDIX EXPERIMENTAL PROCEDURES' AND UNCERTAINTY ESTIMATES I. SAMPLE PREPARATION 1.1. Mineral separation 89 1.2. Whole-rock samples used for U-Pb analysis 89 II. MAJOR AND TRACE ELEMENT ANALYSIS 89 III. Rb-Sr ANALYSIS 89 IV. U-Pb CHEMISTRY IV. 1. Minerals 90 IV.2. Whole-rock samples 90 V. MASS-SPECTROMETRY 91 VI. DATA REDUCTION VI.1. Bias factors 91 VI.2. Blanc and common lead correction 92 VII. ESTIMATION OF UNCERTAINTIES OF U-Pb MINERAL DATA 93 VII.1. Measured isotopic ratios 94 VII.2. Spike calibration factors, F' and F" 94 VII.3. Pb/U ratios 95 VII.4. Pb and U contents 95 VII.5. Uncertainty ellipses 95 VII.6. Influence of uncertainties in the common lead composition 95 VIII. DISCORDIAS AND INTERCEPT AGisS 98 IX. ESTIMATION OF UNCERTAINTIES OF WHOLE-ROCK U-Pb DATA 98

REFERENCES 99

SAMENVATTING 100

- 6 - SUMMARY

Western Galicia is situated in the northern part of the central zone of the Hesperian Massif, which is mainly constituted by supracrustal sequences of Late Precambriao to Late Paleozoic age, and Early and Late Paleozoic granites and migmatites. Fault-bounded, Early Paleozoic catazonal complexes occur ex- clusively in western Galicia and northern Portugal, This thesis presents the results of investigations on the U-Pb systematics of zircons, monazites and whole-rocks from the Paleozoic basement of western Galicia,

The U-Pb isotopic data obtained on suites of zircons and monazites from six orthogneisses, one paragneiss and two granites are presented in chapter II, In the concordia diagram the U-Pb systems of the suites of zircons reveal a large range of upper and lower intercepts. Most intercepts are regarded as geochro- nologically meaningless; they are interpreted in terms of a multi-stage model: variable lead loss from zircons with discordant U-Pb systems. The primary discordance is attributed to mixing of a new zircon generation with a small zircon component of older age. The age of these two components has only been preserved in the zircons of the two orthogneisses of the Mellid complex, which were apparently very little disturbed by later lead loss. Suites of zircons from these samples reveal upper intercepts at 2548 ±\l\ Ma and 226? ±\\ Ma, which are interpreted as approximating the age of the zircon component inherited from the metasediments from which the granitic precursors were generated. The lower intercept ages of 482 ± 12 Ma and 459 til Ha are interpreted as approximating the time of new zircon growth during the emplacement and consol- idation of the granitic magma. The suites of zircons from the other four orthogneisses and from the two granites show varying degrees of later lead loss, the extent of which was possibly influenced by the whole-rock chemistry. The age of this event of lead loss could not be established, but should be later than about 300 Ma ago; possibly, it was due to (sub)recent weathering.

The suite of zircons from the paragneiss sample reveals an upper intercept at 2272 lyg Ma, which is interpreted as the provenance age of the Masanteo group, the oldest sedimentary sequence. From the Rb-Sr data a maximum sedimentation age of some 1.5 Ga may be deduced. The lower intercept age of 476±12 Ma is regarded as approximating the onset of the first phase of granulite facies metamorphism, to which the rocks of the catazonal complexes were subjected and the generation of the granitic precursors of the orthogneisses was related.

A concordant monazite age of 471 Ma of the same paragneiss sample is likewise interpreted as reflecting the onset of the granulite facies matamorphism. Monazites from the two samples of Late Paleozoic granites display both concordant and discordant U-Pb systems; the concordant monazites reveal ages of a- bout 300 Ma.

Rb-Sr whole-rock data of four of the orthogneiss samples from which zircon U-Pb data were obtained are presented in chapter II. The data accord with the previously published Rb-Sr whole-rock ages of these rocks: two of 470-460 Ma and one of about 400 Ma. In addition to these ages, a Rb-Sr whole-rock isochron age of 450 ± 25 Ma was obtained for the orthogneisses from the Sobrado/ Teijeiro complex. The ages of about 470 Ma (Rb-Sr whole-rocks and U-Pb zircon lower intercepts) are interpreted as approximating the time of generation and emplacement of the granites. For the Mellid orthogneisses the U-Pb zircon lower intercepts correspond to an age of 480-460 Ma, versus the corresponding whole- rock isochron age of 400 Ma; this difference is interpreted as reflecting a prolonged Sr isotope migration after the intrusion of tiie rocks, which were emplaced at a lower crustal level than the granitic precursors of the other orthogneisses.

- 7 - The petrography of the high-grade metamorphic mafic rocks from the Sobrado/ Teijeiro complex is described in chapter III, along with the U-Pb whole-rock data of samples from eclogite and granulite facies mafic rocks in the Sobrado/Teijeiro and Cabo Ortegal complexes. The mineral relationships, mainly the relative timing of the clinopyroxene exsolution, provide arguments for the existence of a time interval between the conditions of the eclogite and those of the granulite facies. Taking into account the metamorphic history, the U-Pb data are interpreted in terms of a four-stage model. It is concluded that the emplacement age of the mafic rocks exceeds some 1.0 Ga, which sets a minimum age to the sedimentation - of the Masanteo group. •, -

In chapter IV the supra- and infracrustal history of the Hesperian Massif is sum-, marized. From these data an entirely intracratonic geological evolution is "inferred, which took place during a single dynamothermal cycle, starting at least 600 Ma ago and lasting until about 280 Ma ago. A tnantle-plume/aulacogen model for the geological evolution of the Hesperian Massif is shown to fit the available data. ' ' •--.:-. ;

- 8 - F— ••-

CHAPTER I

INTRODUCTION

I.I. SCOPE OF WORK

This study was initiated in the framework of the geological investigations ex- ecuted since 1956 by research teams of the State University at Leiden in the western part of the former Kingdom of Galicia (the present-day provinces of LaCoruna, Pontevedra, Lugo and Orense in northwestern Spain). A review of the previous work has recently been given by Arps et al, (1977). The region, forms the northwestern part of the Hesperian Massif (the Paleozoic basement underlying the western Iberian Peninsula) and is essentially characterized by a metasedimentary-granitic basement containing several complexes of catazonal rocks. A generalized geological sketch map, including most sample localities discussed in this study, is shown in Fig. 1.

Initially it was assumed that the catazonal complexes represent isolated basement upthrusts of Precambrian age, emplaced in Silurian to Devonian time in a terrain of Upper Precambrian to Lower Paleozoic sediments (Den Tex & Floor, 1971). Isoto- pic dating of the high-grade metamorphic rocks, however, revealed Paleozoic ages for the catazonal complexes (van Calsteren et al., 1979). On the other hand, the occurrence of Upper Precambrian and Cambrian detrital rocks (including greywackes and arkoses) implies the presence of an older granitic source terrain, so the existence of a Precambrian crust is still apparent. In order to investigate the age and the extent of this older crust, a study of the U-Pb systems in zircons from different basement rocks was undertaken, as well as a reconnaissance study of the U-Pb whole-rock systematics of the mafic catazonal rocks. Another purpose of this study was to test models of reworking of continental crust and complementary granulite-granite systems (Den Tex, 1974). In the last part of this thesis a model is presented for the supra- and infracrustal history of the Hesperian Massif.

1.2. GEOLOGICAL SETTING

In its simplest form, the basement of western Galicia can be divided into suites of Late Paleozoic granites and older elements. The latter include complexes of catazonal rocks and various units of mesozonal gneisses and inesozonal-epizonal metasedimentary sequences. The catazonal rocks form a number of subcircular, wholly to partially fault-bounded complexes (Fig. 1): those of Sobrado/Teijeiro, Mellid, Santiago de Compostela, Castriz/Bazar and Agualada, which together form the peripheral belt of the Ordenes complex, and the Cabo Ortegal complex. In northern Portu- gal similar complexes occur near Braganca and Morais. All these catazonal complexes are composed of high-grade metamorphic mafic and ultramafic rocks, paragneisses and orthogneisses in varying proportions. They were subjected to metamorphism up into the clinopyroxene-almandine subfacies of the granulite facies in Early Paleo- zoic time (M-l), followed by retrogradation under hornblende-granulite (M-2), amphibolite (M-3) and greenschist facias (M-4) conditions (Table 1). Penetrative deformation, resulting in folding on mainly NS axes, accompanied the metamorphism.

Mesozonal rocks occur in the Ordenes complex, the Lalin unit and the blastomylordt- ic graben (Fig. 1). The Ordenes complex consists mainly of intermediate to low- grade quartzo-feldspathic metasediments; they have tentatively been correlated with the Lat« Proterozoic Villalba Series in eastern Galicia and the Serie Negra in the central and southern parts of the Iberian Peninsula (Den Tex & Floor, 1971; Engels et al., 1972). The Lalin unit (Hilgen, 1971) is mainly composed of mafic rocks and paragneisses with thin layers of peralkaline orthogneiss. A narrow belt of orthogneisses and p^ragneisses, the "blastomylonitic graben" (Den Tex & Floor, sc

- 9 - r m

[ | low-grade metamorphic sediments

^^^J poiy-meiamarphic mafic and ultramafic rocks

[, /_] schists

j j calcalkalme to peralkaline granites (gneissified)

j | atoalifte granges

!;::;::i|i deformed catcalkaline granites

| undeformed calcalkalme granites

^^^| nabbro

\ (A • ' l\

Fig. 1. Simplified geologiaal map of western Galiaia (after van Calsteren et al.j 1979). The open circles and numbers refer to the locations of the samples discussed in this study. For the Sobrado/Teigeiro area: see Fig. 2.

- 10 - 1 • a;?;

Table 1: Correlation of geological phenomena in western Galioia (after Den Tex," 1980).

Upper Palaeozoic migmatites, InFracrustal rock complexes granitic and metamorphic rocks Lower Palaeozoic rock sequences t catazonal mesozonal supracrustal rocks dismembered meta- ophiolites

F 5: chevron folding normal faulting and thrusting. M4 ca. 290 greenschist lacies granite intrusions granite intrusions dolerite dykes.

amfibolite fades. F4:N-S development of foliation in L- IP metamorphism. anchi- to low- grade metamorphism Devonian to Carboniferous M3 ca. 310 isoclinal folding, vertical axial plane granites, wrench faulting, deformation. carbonates and greywackes. granite emplacement granite emplacement, migmatization melange formation 3S0 hornblende granulite lacies. gabbro stocks, serpentinites, f2. E-W. FyN-S horizontal axial plane. M2 blasiomylonitization, gneissification blastcmylonitization. gneissification. Silurian to Devonian Silurian to Devonian intrusion peralkaline granites, sedimentation of carbonates sedimentation of carbonates. intrusion of mafic plugs intrusion mafic dikes pelites and quartzites pelites and quartzites with bimodal volcanics with bimodal volcanics 400 gabbros, hybridization, granite emplacement granite emplacement Iprotolith of ortho- and augengneisses) {protolilh of ortho- and augengneisses) M, granulite facies Cambriar to Ordovician Cambrian to Ordovician 500 Ft:N-S, horizontal axial plane sedimentation of arenaceous sedimentation of arenaceous to pelitic rocks, with bimodal to pelitic rocks with bimodal >600 mantle-plume activity starts volcanics volcanics

Mo prograde eclogitization (B type). Fo eclogitization |C type) burial burial >1000 emplacement of mafic rocks, emplacement mafic rocks. sedimentation of (semi) pelites. sedimentation of (semi) pelites <1500 greywackes, sandstones, carbonates. greywackes, sandstones, carbonates cherts etc cherts etc erosion, 1.2500 granitic basement ? 1967), extends parallel to the Atlantic coast from Malpica to Tuy. In all these

mesozonal rocks the M-l phase of metamorphism has not been recognized (Table 1). '•*:, •

Metasediiuents outside the above complexes are largely confined to two areas, the central Galician schist area and the schist belts near the Atlantic coast. These sequences consist mainly of pelitic rocks with minor quartzitic and quartzo-feldspathic intercalations. In the southern part of the belts near the Atlantic coast the rocks are the continuation of a thicker metasedimentary sequence in northwestern Portu- gal, to which an Early Paleozoic age has been assigned (Buiskool-Toxopeus et al., 1978). All metasedimentary successions in western Galicia are reduced in comparison with neighbouring areas, where the stratigraphy is much better known and where hiatusses and unconformities can be distinguished. A locally important hiatus, stretching from the Llanvirnian to the Upper Silurian (van Adrichem Boogaert, 1967; Apalategui, 1977; Paris & Robardet, 1977) is followed in southeast Galicia by polymictic conglomerates of Late Wenlockian to Early Devonian age, containing pebbles of metamorphic and felsic volcanic rocks and reworked, probably Early Ordovician fossils (Martinez-Garcia, 1972, 1973; Aldaya et al., 1976). Conglom- erates of Middle to Late Devonian age, containing pebbles of metamorphic rocks, occur in northern Portugal near the complexes of Braganca and Morals (Ribeiro, 1974), Further evidence for tectonic disturbances in the Early Paleozoic is provided by the folding of Lower to Middle Ordovician strata, which did not affect the overlying Silurian (Martinez-Garcia, 1973). Along the southern and southeastern margins of the Cabo Ortegal and Ordenes complexes, metavolcanic rocks of mafic to felsic composition are associated with the metasediments; a Silurian or Devonian age has been assigned to the metavolcanics near Cabo Ortegal (van der Meer Mohr, 1975).

Three folding phases can be distinguished in the metasedimentary sequences, a first phase on NS axes with horizontal axial plane, a second phase, also on NS axes but with a more steeply inclined axial plane, and a third phase of local chevron folding. Remnants of a still older S-plane, crenulated by the first phase, indicate an older phase of low grade metamorphism (Aldaya et al., 1973; van Meer- beke et al., 1973). The first two important folding phases on NS axes can be correlated with the F-3 and F-4 phases in the catazonal complexes, respectively, while the phase of chevron folding can be correlated with the F-5 phase (Table 1).

The main phase of metamorphism (correlated with the M-3 phase in the catazonal complexes) reached low-to-intermediate-pressure amphibolite facies conditions and resulted in widespread anatectic melting, giving rise to the formation of a large variety of migmatic rocks, inhomogeneous anatectic granites and allochthonous intrusive granites: the "alkaline series" (Capdevila & Floor, 1970). Emplacement of this series took place during and after the deformation phases. Another group of granites, the "calcalkaline series", consists mainly of megacrystal-bearing biotite (-muscovite) granodiorites and granites. Their intrusion was syn- to post- kinematic, as they can be subdivided into a deformed and an undeformed series.

1.3. PREVIOUS GEOCHRONOLOGICAL RESEARCH

Most of the isotopic dating in western Galicia has been carried out in the ZWO Laboratory of Isotope Geology, Amsterdam. Data on the geochronology of basement rocks from western Galicia have been reported by Priem et al. (1970) and Van Calsteren et al. (1979). The data obtained by these and other workers are presented in Table 2. Where necessary, the data have been recalculated using the I.U.G.S. recommended constants (chapter I.4.). The most important results are the Early Paleozoic ages, which evidence that magmatic activities were not restricted to the Late Paleozoic, but started at least 500 Ma ago and lasted until about 280 Ma ago.

The high-grade metamorphic events were dated primarily in the Cabo Ortegal area. The Rb-Sr whole-rock age of 477+122 Ma of the lherzolites is interpreted as re-

- 12 - Table 2: Compilation of age data on Galioian rocks*

Method** Age Initial 87Sr/86Sr

1. Two-mica granite, La Guardia Rb-Sr WR 311 + 21 .7143 2. Orthogneisses, Mellid Rb-Sr WR 400 + 23 .7100 3. Orthogneisses, Vigo Rb-Sr WR 456 + 25 .7103 4. Orthogneisses, Malpica/Moya Rb-Sr WR 452 .708 5. Mafic granulitefacies rocks Rb-Sr WR 347 + 17 .70446 Cabo Ortegal + bi 6. Lherzolites, Cabo Ortegal Rb-Sr WR 477± 122 .7033 WR + phlogopite 386 ± 10 .70425 WR + edenite 337 ± 10 .70473 7. La Pioza eclogites, phengites Rb-Sr model 362, 370 phengites K-Ar 328, 341 paragonites K-Ar 362, 352 3. Cabo Ortegal, minerals of high grade metamorphic rocks K-Ar 378-441 9. Mineral ages (mainly bi, some ms) of granitic rocks (post- kinematic granites, pre-kinematic Rb~Sr model 278-316 granites, orthogneisses) some K-Ar 10. Puebla de Farga, pre-kinematic calcalkaline granite Rb-Sr WR 323± 10 11. Friol, synkinematic alkaline granite Rb-Sr WR 315 + 10 12. Forgoselo, post-kinematic calcalkaline granite Rb-Sr WR 314 ±6 13. Puebla de Sanabria, orthogneisses Rb-Sr WR 494 ±20 14. Viana del Bollo series, Viana del Bollo amphibole, amph-bearing quartzite K-Ar 370 + 12 idem, basic dyke in orthogneiss K-Ar 312 ±9 idem, garnet amphibolite K-Ar 310 + 11 15. Mineral ages of Late Paleozoic K-Ar granitic rocks and some ortho- bi rsnge 269-308 gneisses ms range 287-3^0

* 1-8: van Calsteren et al. (1979); 9: Priem et al. (1970); 10-12: Capdevila & Vialette (1970); 13: Ibarguchi (1978); 14: Cantagrel (1973); 15: Ries (1979). Where necessary, the ages were recalculated using the I.U.G.S. recommended set of constants (Steiger & Jager, 1977). For 13, the used constants are not given in the original publication. ** WR, whole-rock; bi, biotite; ms, muscovite.

- 13 - presenting the intrusion age and the start of the M-l granulite facies raetamorphism. Ages obtained on whole-rock/mineral pairs are interpreted as reflecting the closure of the minerals to Rb-Sr and K-Ar upon cooling of the lherzolites. K-Ar ages clustering around 400 Ma should represent the end of the M-l phase. Both the Rb-Sr whole-rock/biotite age of 347 ±17 Ma of the granulites and similar Rb- Sr and K-Ar mineral ages of the eclogites from the blastomylonitic graben are interpreted as recording the end of the hornblende-granulite facies conditions (M-2).

Orthogneisses were dated in the blastomylonitic graben and the Me11id Complex. They represent a suite of calcalkaline to peralkaline granites, emplaced in Silu- rian time. An age of 494 ±20 Ma obtained by Vialette on orthogneisses near Fuebla de Sanabria, has been reported by Ibarguchi (1978); the analytical data and the decay constant used in the calculation of this age are not given, however.

Events cf Late Paleozoic metamorphism and granite intrusion have been dated at several places. In the extreme west, the age of 311 ±21 Ma of the foliated two- mica granites near La Guardia is interpreted as the age of the Late Paleozoic migmatization and folding, M-3/F-4 (van Calsteren et al., 1979). Towards the east, the age of 315 ± 10 Ma for the synkinematic alkaline granite near Friol is related to the F-4 phase of deformation (Capdevila & Vialette, 1970),

Cooling ages have been reported from all over western Galicia. Rb-Sr model ages between 278 and 316 Ma of biotites and some muscovites from various granitic rocks are interpreted as cooling ages after the latest phases of metamorphism and intrusion (Priem et al., 1970). Rb-Sr whole-rock ages obtained on the same samples (Priem et al., 1970) are not quoted here; the samples were taken from different massifs and, although varying little in initial Sr ratios, do not meet isochron requirements. K-Ar ages of 18 biotites and 8 muscovites from Late Paleozoic granitic rocks and some orthogneisses cluster between 275 and 295 Ma, and around 305 Ma, respectively; they have also been interpreted as cooling ages (Ries, 1979). The same interpretation presumably holds for three K-Ar amphibole ages reported from rocks from the Viana del Bollo series in southeast Galicia (Canta- grel, 1973).

Apart from some K-Ar mineral ages in the Cabo Ortegal complex which have been interpreted as excess-argon data (van Calsteren et al., 1979), the only Precam- brian ages of western Galicia basement rocks have been reported for granulites and charnockites dredged from the Danish Bane, north of Gijon (similar rocks have been dredged north of Cabo Ortegal). Biotites yield Rb-Sr ages of 1400 Ma (Lamboy, 1976; Vidal, 1977). Rb-Sr analysis of nine whole-rock samples from the granulites revealed a strong scatter of data-points in a 87Rb/86Sr versus 87Sr/86Sr diagram within an envelope with boundary lines corresponding to ages of 3400 and 1400 Ma (Vidal, 1977). No description of the samples and no analytical data have been given, however, so the significance of these Proterozoic ages remains uncertain.

A U-Pb analysis of a single zircon fraction from the orthogneisses in the Mellid complex revealed a discordant age pattern (T-206 = 485 Ma, T-207 = 516 Ma, T-207/ 206 = 658 Ma). This could point to a Precambrian age (R.H. Steiger, E.T.H. Zurich, 1975, pers. comm.).

1.4. CONSTANTS USED

Throughout this study the constants recommended by the I.U.G.S. subcommission on geochronology have been used (Steiger & Jager. 1977):

- 14 - 235 X U = 9.8485 s*'?1 X238U = 1.55125 x 10--1010a-a » XB7Rb =1.42 x lO"1^-1, X232Th - 4.9475 x 10"" a"1, *"U/238U „ 1/137.88.

For the isotopic composition of laboratory contaminant lead (blanc), the "Average Modern Lead" (AML) of Stacey & Kramers (1975) is taken: a = 18.700, B = 15.628 and y = 33.630. For primordial lead the composition of the Canon Diablo troilite lead is taken (Tatsumoto et al., 1973): o - 9.307, 3 = 10.294 and y = 29.476.

- 15 - t.v". CHAPTER 11

U-Pb SYSTEMATICS OF MINERALS

11.I. INTRODUCTION

Zircons in sedimentary rocks should be the most appropriate material to have retained evidence regarding the age of the older crust, but the humid climate of Galicia has caused strong weathering of most rock types, especially the metasediments. Only one sample of high-grade metasedimentary rock could therefore be collected that was sufficiently fresh to be included in this study. Of the other high-grade rocks, the mafic granulites contain locally a few zircons, but an attempt to separate these zircons failed because of the low proportion of zircon in the concentrates of high-density minerals. Since most of the granitic rocks in western Galicia are thought to have been generated by partial melting of crustal material (Capdevila et al., 1973), and since studies by other workers (e.g. Pidgeon & Johnson, 1974) have shown that zircon U-Pb systems in some cases can survive a melting event, the orthogneisses and granites in this area are suitable candidates to provide information about the older crust by zircon U-Pb analysis.

11.2. CONCORDIA DIAGRAMS

In order to facilitate the discussion of the analytical results, a concise introduction into the interpretation of concordia diagrams is given. Reviews of the U-Pb systematics and methods have been given by Doe (1970), Gale & Mussett (1973) and Faure (1977).

The U-Pb system of a zircon which has remained closed with respect to migration of U, intermediate daughters and Pb, or which has lost all previously accumulated radiogenic lead, produces equal or concordant ages for the two U-Pb decay series. Such a zircon will plot on the concordia curve in the radiogenic Z07Pb/2 U versus radiogenic Z06Pb/238U or "concordia" diagram (Wetherill, 1956). In most cases however, the two U-Pb ages are not equal; the U-Pb system is then called discordant and the U-Pb data-points do not plot on the concordia curve. Usually, several subsystems of the mineral population show a different degree of discordance, resulting in a linear array of discordant data points: the disaordia.

Apart from the concordia diagram a diagram of 238U/206Pb versus 207Pb/20GPb, originally intended to facilitate investigation of initial isotopic lead compositions in whole-rock samples (Tera & Wasserburg, 1972), is frequently used in mineral studies.

The simplest way of obtaining a discordant U-Pb system is mixing of two popula- tions of different age and different though concordant U-Pb systems. Another cause of discordance is open-system behaviour towards U and/or Pb, in which the discordance depends on the extent and direction of the element migration. The upper intercept of the discordia chord with the concordia curve is customarily regarded as the best possible estimate of the (re)crystallization age of the zircon.

The pattern of discordance in a suite of zircons can in most cases be ascribed to loss of radiogenic lead. Uranium-gain without the growth of a new zircon generation seems to be a very rare phenomenon (Grauert et al., 1974). In order to explain the amount and the tiring of lead loss and hence to interpret the lower intercept of the discordia with concordia, several models have been proposed. 1. The episodic lead-loss model postulates expulsion of part or all of the accumulated radiogenic lead from the crystal during a relatively short period

- 16 - w • (Wetherill, 1956). The lower intercept gives the age of the event of lead loss. 2, In order to explain the occurrence of lower intercept ages which could not readily be interpreted in terms of the regional geological history, an alternative model has been developed in which the loss of lead is due to continuous diffusion (Tilton, 1960). The often noticed increase of the degree of lead loss with increasing U content and decreasing grain size (Silver & Deutsch, 1963) was thought to support such a continuous diffusion. According to this model the lower intercept ages are geologically meaningless and result from the graphical extension of a linear array of data-points along the upper part of the diffusion-loss curve. The model was improved by introducing a diffusion coefficient varying with the radiation damage imposed by decaying nuclides on the crystal structure, and hence varying with time (Wasserburg, 1963), A major objection to continuous diffusion is the fact that the diffusion coefficients for lead in zircon'are several orders of magnitude lower than those needed in the continuous^ diffusion model (Shestakov, 1972). 3. A third model of lead loss is the dilatancy-model, according to which radiation damage woi^tld lead to the formation of micro-cracks in the mineral. Opening of such cracUs due to pressure-release upon .uplift would give way to solutions capable of leaching the accumulated radiogenic lead. According to this model the lower intercept would date the time of uplift (Goldicii & Mudray, 1972). More complex models, involving one or more phases of episodic loss superimposed on continuous diffusion-loss have also been developed. The resulting discordances were calculated by Wetherill (1963) and Allegre et al. (1974).

Detailed investigations of the discordance patterns in cogenetic zircon suites revealed irregularities which could not simply be interpreted in terms of one of the models mentioned above. Indications that the zircon system is bimodal and that the discordia is essentially a mixing-line between two end-members, one concordant and the other highly discordant, were presented by several authors (e.g. Grilnenfelder, 1963; Steiger & Wasserburg, 1966, 1969; Naylor et al., 1970). Nev- ertheless, the lead loss of the discordant phase was still assumed to be mainly due to continuous diffusion.

Recently, an elaborate investigation of the physicochemical properties of some zircon suites provided strong evidence for the concept of a bimodal zircon system (Sommerauer, 1976). Based on electron microscope observations, electron probe micro-analysis (EPMA), cathodoluminescence studies, X-ray data, etc., the following model has been developed: Zircon constitutes a polyphase system, independent of age ana chemical environment, Grains or sub-grains low in minor elements (< 2 mol %) are constituted by one phase, a highly stable zircon phase denominated ZK. Grains or sub-grains with a higher minor element content (> 2 mol %) are made of two phases, oriented domains of the stable Zjj-phase embedded in a gel-like, highly instable matrix in which the minor elements are heterogeneously distributed, the Zjj-phase. Further exsolution of this Zjj-phase, probably aided by the energy of decaying nuclides, results in spherical ZrO2-rich aggregates in a Si02-rich matrix. The mobility of lead is very low in the ZR-phase, where only volume diffusion is possible, but migration of lead by grain-boundary diffusion is very easy in the Zjj-phase. Discordant U-Pb systems are a mixture of the stable, highly concordant ZK subsystems and the unstable, highly discordant Zty subsystems. The degree of discordance depends mainly on the amount of ZJJ, the geometry of Zif-Zu intergrowths and the possiblity of element transport. Loss of lead takes place only during short, episodic events.

A grossly similar model of metamictization by exsolution has been proposed for the metamictization observed in complex niobium and tantalum oxides (Graham & Thornber, 1974). According to this model metamictization is an ordering process during which the disordered structure containing cations strongly differing in radius is exsolved on a microscale into an ordered domain structure. The presence of radioactive elements is probably a rate-controlling factor.

- 17 - Although the formation of a polyphase zircon system appears to be independent of the chemical environment (Sommerauer, 1976), the extent to which the system res- ponds to disturbing events is possibly controlled by the environment. Under experimental hydrothermal conditions the rate of lead loss and annealing of a metamict zircon was shown to be the highest in solutions rich in sodium chloride and hydrochloric acid (Pidgeon et al., 1966, 1973). In the case of an investigated natural system the degree of discordance and the extent to which an inherited radiogenic lead component has been retained seem to be related to the alkalinity of the whole-rock system: with increasing alkalinity the proportion of the inherited older zircon component decreases (Pidgeon & Johnson, 1974).

Contrary to zircon in which concordant U-Pb systems are relatively rare, monazite often contains concordant U-Pb systems. It also shows more frequently a reversed discordance than zircon, i.e. U-Pb data-points plotting above the concordia. The difference between the two minerals in the behaviour of the U-Pb systems is possibly related to the crystal structure. In this respect, reference can be made to the ThSiOu polymorphs thorite and huttonite, thorite being isostructural with zircon and huttonite isostructural with monazite. The main difference between the thorite and huttonite structures is the presence of large void spaces in the former mineral lattice, forming channels parallel to the c-axis (Taylor & Ewing, 1978). The fact that thorite is often totally or partially metamict, whereas metamict huttonites do not occur, is another strong indication for a structural control of the metamictization process. It can be speculated that the limited susceptibility to element migration of monazite is caused by a structure much more stable regarding ordering and/or exsolution phenomena than the zircon-type structure. Metamict monazites are indeed extremely rare and only one (possible) example has been reported (Deer, Howie & Zussman, 1962). Contrary to zircon, monazite often contains intracrystalline iron-bearing alteration products and then shows a good cleavage (Molloy, 1959).

II.3. GEOLOGICAL SETTING AND PETROGRAPHY OF THE INVESTIGATED ROCK UNITS

U-Pb investigations were made on zircons and monazites from the following rocks: 76-Gal-l augengneiss, Mellid complex, 3 km southwest of Mellid. 76-Gal-A orthogneiss, blastomylonitic graben, 5 km southeast of Vigo. 77-Gal-6 migmatic augengneiss, 12 km north of Noya. 77-Gal-12 two-mica granite, Montesalgueiro. 77-Gal-13 two-mica granite, Cabo Silleiro. 77-Gal-lA orthogneiss, Sobrado/Teijeiro complex, 5 km northeast of Curtis. 77-Gal-15 muscovite-chlorite gneiss, Sobrado/Teijeiro complex, 2 km southwest of Sobrado de los Monjes. 77-Gal-19 blastomylonitic augengneiss, Mellid complex, 7 km north of Mellid. 77-Gal-23 orthogneiss, blastomylonitic graben, 4 km south of Noya. The sampling sites are shown in Fig. 1. Table 3 gives the major and selected trace element composition of the samples. A description of the geological setting of the sampled rock units is given below. The various metamorphic and deformation phases are numbered according to Table 1.

II.3.1. The Mellid Complex

The high-grade metamorphic rocks around the township of Mellid form one of the complexes into which the peripheral belt of the Ordenes complex has been divided (Keasberry et al., 1976; Arps et al., 1977). A large unit of calcalkaline granitic and granodioritic orthogneisses forms the northern and western part of the complex, while low-grade metamorphic mafic and ultramafic rocks surround the high-grade center in the east and south. These low-grade rocks are epidote-amphibolites and serpentinites, forming part of a melange of metavolcanics, metasediments, amphibolites and serpentinites, extending as an arcuate zone along the southern rim of the Ordenes complex (Den Tex, 1980). The central part of the complex is occupied

- 18 - Table 3; Whole-rook major (Wt %) and selected traae (\ig/g) element contents

76-Gal-l 76-Gal-4 77-Gal-6 77-Gal-12 77-Gal-13 77-Gal-14 77-Gal-15 77-Gal-19 77-Gal-23

SiO2 68.96 74.10 70.61 73.55 73.03 73.31 54.46 70.80 73.28 liO2 .77 .32 .37 .14 .18 .24 1.07 .59 .30 A12O3 14.26 12.66 14.73 14.30 14.68 13.90 21.83 14.17 13.31 Fe2O3 .26 .07 .09 .01 .04 .14 .50 .35 .04 FeO 3.78 2.56 2.60 1.08 1.28 2.10 8.47 3.09 2.85 MnO .06 .04 .06 .03 .03 .06 .14 .06 .04 MgO .90 .45 .70 .31 .38 .28 2.79 .70 .48 CaO 2.47 1.22 1.07 .84 .71 1.11 1.26 2.10 1.86 Na20 3.41 3.19 3.15 4.24 3.07 4.00 1.37 3.07 3.61 K20 3.58 4.34 4.88 3.96 4.92 4.11 3.86 3.98 3.19 P2Os .25 .05 .31 .16 .27 .09 .19 .22 .05 H20 ± .72 .56 .94 .69 1.27 .50 4.42 1.00 .57 CO2 .1 • 1 • 1 .1 .1 .1 .1 .1 .1 Total 99.42 99.56 99.51 99.31 99.86 99.84 100.36 100.13 99.58

Rb 78 197 290 300 238 141 138 105 99 Sr 169 52 73 67 64 59 150 139 116 Rb/Sr .46 3.8 4.0 4.5 3.7 2.4 .92 .75 .85 Pb 15 12.4 25 40 40 20 10 20 10 Th 30 16.5 20 15 15 15 15 25 15 U 2 4,.6 15 20 10 5 2 4 6 Y 54 70 42 5 10 64 45 71 47 Zr 169 62 65 179 192 281 191

Na20+ K20 .67 .79 .71 .79 .71 .79 .30 .66 .71 A12O3 (molar ratio) Katanorm Qz 28 34 29 30 33 29 . - 13 31 33 Ab 36 32 33 41 32 39 - 27 32 36 i An 13 6 4 4 2 5 11 11 10 - Or 25 28 34 26 33 26 49 27 21 VO 1 by foliated, partially serpentinized, spinel-pargasite peridotites with subordinate amounts of clinopyroxene and green spinel.

Mafic granulite facies rocks partially surround the central ultramafics. The oldest metamorphic assemblage in the granulites was formed under conditions of the clinopyroxene-almandine subfacies (M-l). Following a phase of blastomylonitiza- tion, conditions of the hornblende-clinopyroxene-almandine subfacies prevailed (M-2). Adaption to conditions of the amphibolite (M-3) and greenschist facies (M-A) occurred during later retrogradations. Chemically, the rocks are similar to continental tholeiites, locally with alkali-olivine basaltic affinities (Hu- bregtse, 1973a).

In the southern part of the complex, fine to medium-grained kyanite-bearing paragneisses showing planar to planolinear textures occur adjacant to the metamafics. The high-pressure granulite facies parageneses of minor gabbroic intrusions within these paragneisses show that the rocks underwent the same metamorphic history as the mafic rocks (Hubregtse, 1973a).

The orthogneisses were probably derived from one large, compositionally zoned, plutonic body, which after intrusion (intrusive relationships with the Ordenes schists have been found in a few places, Hubregtse, 1973a) underwent amphibolite facies metamorphism during M-l. An assemblage of plagioclase, K-feldspar, garnet, biotite, muscovite and quartz was formed during this phase. Minor gabbroic intrusions, often displaying phenomena of contamination or hybridization, occur locally within the orthogneisses and show that the granitic and basaltic magmatism were coeval.

Vertical movements led to differences in crustal level after M-l and hence different metamorphic and structural evolutions, on the basis of which the gneisses are divided into two units. The low-grade unit is mainly composed of granitic gneisses, which were strongly (blasto)mylonitized during F-3, resulting in cataclastic textures. The intermediate-grade unit is largely composed of granodioritic gneisses in which a second generation of garnet developed at the expense of biotite during M-2, concurrent with strong recrystallization indicating amphibolite facies conditions. Penetrative deformation during F-4, accompanied by metamorphism under amphibolite facies conditions, gave rise to the characteristic augentexture of the rocks (Hubregtse, 1973a).

North of Mellid these gneisses are separated from the high-grade mafic rocks by a discontinuous zone of fine to medium-grained planar to planolinear, rather dark-coloured gneisses containing small feldspar-augen. More leucocratic rocks occur dispersed throughout this zone. Mineral relics show that these blastomylonitic gneisses were subjected to high-grade metamorphic conditions. Most of them have been interpreted as retrograded metasedimentary granulites, although the presence of blastomylonitized orthogneisses was also inferred from field relationships (Hamel, 1971). Investigation of this zone during tht sampling has led to the opinion that most or all of the rocks are strongly deformed orthogneisses, forming a blastomylonitic border zone of the large gneiss unit. Two samples from gneisses in the Mellid complex have been investigated: Sample 76-Gal-l was obtained by blasting an exposure in intermediate-grade augengneisses some 3 km southwest of Mellid (Fig. 1). The sample is an augengneiss in which two foliation planes can be recognized, the youngest being the main one. The sample consists of large K-feldspar and plagioclase porphyroclasts and polycrystalline quartz/feldspar augen and lenses, separated by rather coarse-grained biotite bands containing some muscovite and relics . of garnet; the bands define the main foliation. Both quartz and feldspar are strongly recrystallized, resulting in granoblastic aggregates. Later deformation is apparent from mortar rims around the constituent grains. Accessory minerals include allanite, apatite, epidote/clinozoisite, sphene and zircon.

- 20 - The zivoons are mostly pink and clear, hut some brownish, turbid grains also occur. In transmitted light most grains are colorless and clear. Apart from a large amount of broken crystals, most zircons are sub- to anhedral and prismrtic to spindle-shaped. They show rounded edges and pitted crystal-faces, caused by corrosion. Most of the grains contain tiny inclusions of zircon, rutile and opaque matter. A few grains show a fine zoning. Rounded cores, in some cases surrounded by only a narrow rim of new material, have only been encountered in the larger grains. Partly metamict grains are scarce. The magnetic fraction differs from the non-magnetic zircons only in containing a larger amount of opaque inclusions and displaying a slightly stronger corrosion.

Sample 77-Gal-19 was derived from a boulder in the blastomylonitic border zone, about 7 km north of Mellid (Fig, I). This strongly foliated, dark-coloured gneiss consists of plagioclase and some K-feldspar porphyroclasts and polycrystalline quartz/feldspar augen embedded in a fine-grained matrix of quartz and feldspar in which micas, mainly biotite, are concentrated in thin bands and layers. Late-stage cataclastic deformation is apparent from the occurrence of mortar rims around previously recrystallized quartz and feldspar, the mortar textures being more or less confined to narrow zones. Apart from relics of garnet, the accessory minerals are allanite, apatite, epidote/clinozoisite, rutile and zircon.

The zircons are mostly pink, but colorless and brownish grains are also present. In the finer-size fractions the proportion of colorless to pinkish, clear grains is higher. Otherwise the zircons show the same characteristics as in 76-Gal-l.

II.3.2. The Sobrado/Teijeiro Complex The fault-bounded high-grade metamorphic complex of Sobrado/Teijeiro forms the northern part of the peripheral belt of the Ordenes complex. Like the adjoining Mellid complex, it consists of ultramafic rocks surrounded by mafic granulite facies rocks, paragneisses, amphibolites and orthogneisses. The central ultramafic rocks are almost completely serpentinized, containing only a few relics of olivine, clinopyroxene and brown hornblende. Bands and lenses of talc-schists occur along the rims and locally within the serpentinites.

The mafic rocks form a heterogeneous unit of granofelses, granulites, garnet amphibolites, epidote-rich amphibolites and eclogites. Their metamorphic history is more or less the same as in the Mellid area, but retrogradation and deformation after M-l was somewhat weaker. The most important difference with the Mel- lid complex is the occurrence of (retrograded) eclogites and clinopyroxene- plagioclase-symplectites in granofelses and mafic granulites, indicating the original existence of a jadeite-rich clinopyroxene (Kuijper, 1979), These features are attributed to a pre-M-1 paragenesis, formed under eclogite facies conditions (chapter III).

The metasediments surrounding the mafic rocks are garnet-biotite gneisses consisting of clasts of K-feldspar, plagioclase and garnet in a groundmass of K- feldspar, plagioclase, quartz, biotite and muscovite. Kyanite is a minor component in a few samples. All rocks show planar to planolinear textures. Blastomyloni- tic textures appear near the contacts with the orthogneisses. The concentration of K-feldspar and quartz in small augen and streaks is interpreted as resulting from a minor degree of mobilization. The abscence of migmatic parts suggests that anatexis on a large scale did not take place. Without trace element analysis, however, the possibility cannot be ruled out that parts of the garnet-biotite gneisses constitute residues of partial melting (Drury, 1979). Rounded inclusions of retrograded mafic granofelses, wrapped around by the foliation of the gneisses, and small gabbroic intrusions with mineral parageneses indicating granulite facies metamorphism occur throughout the paragneisses. A larger gabbro intrusion without high-grade metamorphic parageneses is situated near Teijeiro (Kuijper, 1979).

- 21 - metasediments I I garnet-amphibolites

garnet- biotite gneisses E I granulite facies mafic rocks

Espenuca granite metagabbro

Coba da Serpe granite serpentinites

augengneisses km

Fig. 2. Simplified geological map of the Sobrado/Teigeiro complex (after Arps et al.j 1979).^ The open civales and numbers refer to the locations of the samples discussed in this study.

- 22 - The high-grade metamorphic center of the complex is almost completely surrounded by orthogneisses, which in the east and south are part of the intermediate-grade unit of the Mellid complex. The coarse-grained, biotite-rich orthogneisses in the west form a separate unit, with a blastomylonite zone forming the contact with the high-grade rocks. Intrusive relationships, although largely obliterated by faulting, are visible along the contact with the Ordenes metasediments. The gneisses contain a variable amount of feldspar porphyroclasts and are composed of K- feldspar, plagiocla.^e, quartz, biotite and minor muscovite. Deformation (F-4), accompanied by amphibolite facies metamorphism (M-3), increases from north to south, showing small-scale variations in intensity. Pre-kinematic garnet, biotite and feldspar porphyroclasts are the oldest recognizable fabric elements (Kuijper, 1975, 1979).

From the gneisses in the Sobrado/Teijeiro complex two samples have been investigated: Sample 77-Gal-14, an orthogneiss, comes from a small quarry about 5 km NE of Curtis (Fig. 2). The foliated augengneiss consists of a few relics of strongly granulated feldspar porphyroclasts set in a recrystallized matrix of fine-grained quartz and feldspar. Biotite and muscovite are concentrated in bands and layers parallel to the foliation and containing some garnet-relics. The rock was affected by polyphase deformation, as is evidenced by the occurrence of mortar teM- tures of recrystallized feldspar grains surrounding the larger feldspar relicts. Accessory minerals are allanite, apatite, rutile and zircon.

The ziraons are mostly colorless to pinkish, occasionally with brownish tinges. Metamictization is variable, but always rather low. The mostly (sub-)euhedral grains are rather long-prismatic and show a large amount of cracks. Minor corrosion is apparent from the occasional occurrence of rounded crystal edges. A fine zoning is visible in a number of grains, in some cases forming a finely zoned rim surrounding a euhedral core. The zircons of the magnetic fraction differ from the non-magnetic zircons in containing a larger amount of opaque inclusions and displaying a somewhat stronger corrosion. Moreover, the proportion of brownish and wholly or partly metamict crystals (occasionally surrounded by a clear rim) is higher.

Sample 77-Gal-15 belongs to the unit of garnet-biotite gneisses. Comparable gneiss units of the Cabo Ortegal and Mellid complexes were shown to be of sedimentary origin on the basis of their chemical composition (Vogel, 1967; Hubregtse, 1973a). Although somewhat high in AI2O3 (Table 3), the chemical composition of sample Gal-15 is within the range of compositions of common shales (Pettyjohn, 1957).

The muscovite-chlorite gneiss 77-Gal-15 was taken in a quarry some 2 km SW of Sobrado de los Monjes (Fig. 2). It is a bluish-grey, foliated and laminated, fine to very-fine grained gneiss consisting of muscovite-chlorite bands with subordinate quartz and feldspar, separated by quartz-rich bands. Relics of pre- kinematic garnet occur, strongly altered into chlorite and muscovite. Growth of large, often prismatic, carbonate-poikiloblasts is a late-stage phenomenon. Accessory minerals are allanite, apatite, monazite, rutile, tourmaline and zircon. The very fine grain size, the relatively high density of most mineral constituents and the presence of other non-magnetic high-density minerals (rutile, sulphides) made the zircon separation rather difficult.

Few zircons are present, all smaller than 60 ym. The crystals are variable in shape. Most grains are strongly rounded and oval-shaped, but rather long-prismatic (sub-)euhedral and spindle-shaped grains also occur. The zircons are pink to colorless, the pink grains being turbid (especially the rounded ones) and the colorless grains being clear. In transmitted light all crystals are colorless and clear, containing numerous tiny inclusions. Occasionally, rounded or irreg- ular cores occur.

- 23 - s-— -

I?" Monazite occurs in a larger amount and shows a wider range in grain size. They r are light yellowish-green, mostly strongly rounded and turbid. Large grains are greenish in transmitted light, but small grains are colorless and clear.

IX,3,3. The Blastomylonitic Graben A curved, narrow, fault-bounded belt of granitic to granodioritic and metasedimentary rocks extends from Malpica to Tuy (Fig. 1). The metasediments of this unit consist: of paragneisses and micaschists in which layers of para-amphibolites, me-aquartzites and graphite-schists occur, along with calc-silicate lenses. The paragneisses are mainly biotite-muscovite gneisses with variable textures, ranging from planar or linear to massive. A characteristic feature is the occurrence of plagioclase metablasts, in which quartz, micas and zoned, locally resorbed garnet are present as inclusions (Floor, 1966; Arps, 1970). In the northern part of the graben lenses of type-C eclogites occur within paragneisses containing relics of a high-pressure mineral paragenesis, i.e. garnet, clinopyroxene, phengitic mica and zoisite (Arps et al,, 1977; Van der Wegen, 1978).

A large variety of granitoid rocks is present within and along the rims of the graben. The oldest granitic rocks have been divided into two series, both intruded before F-3, Probably they were spatially separated, since no intrusive relationships have been observed. One series is an alkaline suite of coarse- grained megacrystal-bearing two-mica granites. On the basis of composition and relative age, the other series has been subdivided into two suites separated in time by the intrusion of a mafic dike swarm: an older suite of biotite granodiorites of calcalkaline character and a younger suite of ferrohastingsite (-biotite) granites and peralkaline astrophyllite-bearing aegirine-riebeckite granites. The mafic rocks are now present as numerous amphibolite lenses, but these are conspicuously absent within the subseries of peralkaline granites.

Penetrative deformation affected the granites during F-3, resulting in linear to planolinear gneisses and (blasto)oylonitic augengneisses. During the subsequent phase of metamorphism (M-3), which in these parts of Galicia reached its maximum just before F-4, large-scale anatexis outside the graben gave rise to the formation of migmatites and anatectic granites. Since only a very subordinate amount of migmatic rocks occur inside the graben, this part of the crust must have been at a slightly higher level during M-3. Alkaline anatectic granites did intrude into the graben, however, forming a series whose constituent rocks were emplaced from before F-4 until after F-5. Intrusion of granites belonging to a calcalkaline megacrystal-bearing granodiorite series covered the same time-span (Floor, 1966; Arps, 1970; Arps et al., 1977).

Three samples of gneisses from within or just outside the graben have been investigated: Sample 76-Gal-4 was obtained from a quarry in orthogneisses, about 5 km SE of Vigo (Fig. 1). It is a strongly linear to planolinear, medium- to fine-grained rock, consisting of polycrystalline quartz/feldspar lenses and bands and perthitic K-feldspar porphyroclasts, separated by biotite-bands. Accessory minerals are allanite, apatite, epidote/clinozoisite, a rather dark-green hornblende, sphene and zircon.

The zircons are pink to colorless, occasionally brownish. Most grains are turbid. (Sub-)euhedral and prismatic crystals showing many cracks prevail. A fine, euhedral zoning is often visible. Most grains contain a large amount of tiny inclusions, such as zircon, biotite, rutile and opaque matter. Euhedral to rounded brownish cores, occasionally metamict and surrounded by a more or less clear rim, occur especially in the larger grains. The zircons of the magnetic fraction contain a larger amount of inclusions. They have a brown colour and display a stronger metamictization. Corrosion is only apparent in reflected light by the occa-

- 24 - Ifiv sional occurrence of slightly rounded crystal edges and pitted crystal faces. A characteristic feature of the zircons, especially those of the magnetic fraction, is the lack or the strongly subdued occurrence of the {110} crystal faces. • (V-

Sample 77-Gal-23 is a very similar rock. It also comes from the graben, taken in a quarry about 4 km south of Noya (Fig. 1), The linear to planolinear biotite orthogneiss contains the same minerals as 76-Gal-4,

The zircons are pink to brownish and mostly turbid. The (sub-)euhedral crystals show the same characteristics as the Gal-4 zircons, including the poorly developed {110} faces. Rounded to euhedral cores occur in the larger grains, mostly brownish and surrounded by a finely zoned, more or less clear rim. The smaller grains are not or only very slightly zoned.

Sample 77-Gal-6 has been taken just outside the graben in a quarry about 12 km N of Noya (Fig. 1). The rocks in this quarry range from migmatic augengneiss through augengneiss-migmatite to inhomogeneous granitic rocks. Unfortunately, the large inclusions of biotite-rich melanosome, which were also sampled, did not contain enough zircon for analysis. The investigated sample is a migmatic augengneiss, consisting of a medium-grained, granoblastic quartz/feldspar matrix in which granulated feldspar and quartz porphyroclasts occur. Biotite and Muscovite are more or less concentrated in thin bands and layers. Accessory minerals are allanite, apatite, monazite, tourmaline and zircon.

The zircons are pink and clear, occasionally brownish and turbid. Both (long-) prismatic and spindle-shaped, facetted grains occur, the latter usually very clear and pinkish. Crystal edges are mostly somewhat rounded. Pitted crystal faces occur occasionally.. The amount of (partly) metamict grains is low in the coarser fractions, but increases with decreasing grain-size: about half of the 30-40 \m fraction consists of cracked, partly metamict and turbid grains. In transmitted light most grains are colorless. A few brownish or clear cores occur in the largest grains only.

II.3.4. Lace-Paleozoic Granitic Rocks The basement of western Galicia is characterized by large amounts of granitic rocks, varying in composition, structure and texture. Mainly on petrographical grounds and aided by field and some chemical data, the granites have been divided into two series (Capdevila & Floor, 1970; Capdevila et al., 1973). The first is a palingenetic, alkaline, often aluminous series, predominantly consisting of two-mica leucogranites. They range from migmatites through (par-)autochthonous granites to allochthonous, homogeneous two-mica granites and are related to the M-3 phase of metamorphism. The oldest members of this series were emplaced prior to F-4 and are highly tectonized, while the youngest are small, port-kinematic, oval-shaped intrusions. Granites of this series originated by processes of anatexis in the middle and higher parts of the crust, including the Lower Paleo- zoic sediments. The second series is mainly represented by calcalkaline megacrystal-bearing biotite (-muscovite) granites and granodiorites. Rocks of this series are often associated with mafic rocks, occasionally giving rise to hybridization. Limited assimilation of the country rocks resulted locally in the occurrence of alumina-silicates. As regards the time of emplacement, two groups can be recognized, a pre- to syn-kinematic group of mainly megacrystal-bearing granodiorites and a second, post-kinematic group. Rocks of the calcalkaline series probably originated by partial melting in the deeper and/or drier parts of the crust (Capdevila et al., 1973).

Two samples of the Late-Paleozoic granites have been investigated: Sample 77-Gal-l2 was taken from a large quarry in the center of a late-kinematic intrusion of two-mica granite (the Espenuca granite), north of the Sobrado/Tei- jeiro complex (Fig. 1). The granite belongs to the calcalkaline series (Kuijper,

- 25 - f—

Table 4: Rb-Sv whole-rook data

Sample Rb Sr Rb/Sr 87Sr/86Sr 87Rb/86Sr No, (yg/g) (Mg/g) (Wt/Wt)

Mellid ovthogneisses Gal-1 78.2 169.5 0.462 0.71725 1.337 Cuvtis ovthogneisses Cor 28 78.2 202.5 0.387 0.71566 1.120 Cor 29 76.8 206.0 0.373 0.71579 1.080 Cor 83/84 100.5 114.0 0.882 0.72394 2,554 Cor 85/86 80.7 180.5 0.448 0.71718 1.296 Gal-14 142,0 59.1 2.400 0.75370 6.97

Vigo or-thogneisses Gal-4 197.0 52,3 3.764 0.78366 10.97

Noya ovthogneisses Gal-23 97.9 115.5 0.848 0.72589 2.456

Calculated ages Age initial 87Sr/86Sr

Mellid ovthogneisses van Calsteren et al. (1979): 400 ± 23 Ma 0.71000 + 0.00047 Gal-1 included : 400 ± 24 Ma 0.70995 ±0 .00050

Curtis ovthogneisses 450 + 25 Ma 0.70854 + 0 .00091

Vigo ovthogneisses van Calsteren et al. (1979): 456 + 25 Ma 0.71031 ±0 .00287 Gal-4 included : 461 ± 26 Ma 0.70997 ±0 .00274 Gneisses E of Vigo only : 469+ 8 Ma 0.71001 + 0.00081

Noya ovthogneisses van Calsteren et al. (1979): 452 Ma* 0.708 Gal-23 included : 445 Ma* 0.709 Cor 23 excluded : 459 Ma* 0.708

Reference age.

- 26 - t;-

1979), It is a medium-grained inequigranular granite, containing some K-feldspar megacrysts and consisting of K-feldspar, oligoclase/andesine, quartz, muscovite and biotite, with accessory zircon, rutile and monazite. A weak deformation is reflected by the undulose extinction and sutured grain boundaries of nearly all minerals.

Zircons occur in very low amounts: only 35 mg could be separated from 20 kg of rock. The crystals are colorless to pinkish, mostly (sub-)euhedral and occasionally very long-prismatic or spindle-shaped. Only a few metamict grains have been encountered. Occasionally, brownish, euhedral cores occur surrounded by a zoned, less strongly coloured rim.

Monazite is more abundant. The crystals are short-prismatic to oval-shaped and always strongly rounded. The brownish to colorless grains are mostly turbid and contain a large amount of inclusions, mainly opaque matter.

Sample 77-Gal-13 was obtained from a quarry in inhomogeneous anatectic granites near the coast at Cabo Silleiro (Fig. 1). The foliated two-mica granites along -the coast north of La Guardia belong to the alkaline series (Buiskool-Toxopeus et al., 1978), They intruded before and were deformed by F-4. The sample shows a very weak foliation and consists of large K-feldspar and plagioclase porphyroclasts surrounded by a medium to fine-grained groundmass of unoriented feldspars, quartz, muscovite and biotite. Accessory minerals are apatite, monazite and zircon.

The zircons are (sub-)euhedral and translucent pink to colorless, varying from long-prismatic to spindle-shaped. A very low proportion of metamict crystals is present. In transmitted light the grains are colorless to brownish and mostly strongly zoned. Most of the larger grains are composed of a euhedral or corroded, occasionally metamict brownish core surrounded by a clear, zoned rim. In the clear, colorless grains no cores were encountered.

Monazites are mostly euhedral and strongly rounded, showing light yellowish-green colours. Thin overgrowths occur on a number of grains. The occasional subhedral grains are prismatic to tabular.

II.4. Rb-Sr INVESTIGATIONS

A number of gneiss samples used for U-Pb zircon study was also analysed for Rb/Sr and Sr isotopic composition. These samples come from four gneiss bodies: the Mellid orthogneisses, the Curtis orthogneisses, the orthogneisses near Vigo, and the orthogneisses in the blastomylonitic graben (chapter II.3.). For all these gneisses Rb-Sr data were already available. The analytical data are presented in Table 4.

Sample 76-Gal-l plots on the Rb-Sr whole-rock isochron of the intermediate grade Mellid orthogneisses (van Calsteren et al., 1979). When this sample is included in the isochron calculation, the reported age of 400 + 23 Ma (Table 2) remains the same. The initial Sr ratio changes slightly from 0.71000+0.00047 to 0.70995 ± 0.00050.

Of the Curtis orthogneisses four samples were collected previously, but they appeared to have only a small spread in Rb/Sr. Fortunately, sample 77-Gal-l4 has a substantially higher Rb/Sr ratio, which allows an isochron calculation together with the other samples. From the five samples of Table 4, two (Cor 83/84 and Cor 85/86) are combinations of sma'ler samples from the same outcrop. The five points define an isochron (Fig. 3) with a slope corresponding to an age of 450±25 Ma (2a) and an intercept of 0.70854 ± 0.00091 (2a). In view of the metamorphic history of these rocks (Kuijper, 1975, 1979), they were probably emplaced at a relatively high crustal level; the Rb-Sr whole-rock isochron age is therefore interpreted as closely approximating the time of intrusion of the original granites.

- 27 - CURTIS ORTHOGNEISSES

0,150

0.725 AGE 450 i 25 Ma (a7Sr/39sr)o-0.10B54 i 0,00091 errors 2o

87R

2.5 5.0 7.5

Fig. 3. Rb/Sr isoahron diagram of whole-rook samples from the Curtis ox'thogneisses.

The orthogneisses near Vigo revealed an age of 456±25 Ma (van Calsteren ec al., 1979). However, two of the five samples were taken from gneiss bodies which are not continuous with the larger occurrence just east of Vigo. These two samples come from a body of augengneisses near Pontevedra, which possibly belong to the alkaline series of megacrystal-bearing two-mica granites, whereas the linear orthogneisses near Vigo form part of the calcalkaline biotite-granodiorite series (chapter II.3.3.)- Consequently, the line through the data-points cannot be denominated an isochron. Combining sample 76-Gal-4 with the three samples of the orthogneisses east of Vigo results in an isochron age of 469 ±8 Ma (2o) and an intercept of 0.71001 ±0.00081 (2a). Although this figure is within the analytical error of the previously published age, 469 ± 8 Ma is taken here as approximating the intrusion age.

For six samples taken from different orthogneiss occurrences in the blastomylonitic graben between Malpica and Noya, van Calsteren et al. (1979) published a reference age of 452 Ma. Inclusion of sample 77-Gal-23 lowers this age slightly to 445 Ma. The original suite of samples included one sample of probably metased- inentary origin (Cor 23), however. If this sample is omitted, the reference age is 459 Ma. This latter age is interpreted as approximating the time of intrusion.

II.5. RESULTS AND DISCUSSION OF THE U-Pb MINERAL ANALYSES

The analytical data of the nine investigated suites of zircons and monazites are presented in Table 5. For each sample the data are plotted in a concordia-diagram The analytical procedures and estimates of the maximum errors are given in the Appendix.

II.5.1. Orthogneisses of the Mellid Complex Sample 76-Gal-l Six size and magnetic fractions of the zircons have been analyzed. In addition

- 28 - two aliquots of the NM 130-160 pro fraction were analyzed separately. A third aliquot of the latter size fraction was leached for two hours in cold hydrofluoric acid (48%) in an attempt to alter the core/overgrowth ratio and to investigate .(1. whether any loosely bound zircon or U-Pb components are present; both the leached zircon and the leachate have been analyzed. For all analyzed zircon fractions the amounts of common lead are small and range from 1.2 to 6.5 ng (total Pb 90 to 340 ng), which in view of the procedure blanc values of 2 and 4 ng fully can be ascribed to laboratory contamination.

The eight data-points define a discordia which intersects concordia at 2548 1123 Ma and at 482±12 Ma. One point (M>160 ym, VI) lies off the discordia outside the analytical error (Fig. 4). Both the magnetic and the non-magnetic fractions show a regular arrangement along the discordia with respect to the grain size. Only the three magnetic fractions display a regular increase of U content with decreasing grain size.

MELLID ORTHOGNEISS

013 GAL 1

0.10 I: NON-MAGNETIC, > IBOyum H: NON-MAGNETIC, 130-160 M: NON-MAGNETIC, 130-160 JX: NON-MAGNETIC, 130-160* 2: NON-MAGNETIC, 40 - 50 3tt: MAGNETIC, > 160 2H: MAGNETIC, 110-130 207, Pb/235., VJH: MAGNETIC, < 30 •* LEACHATE

0.S 10 1.5

Fig. 4. Concordia diagram of zircons from the Mellid orthogneiss 76-Gal-l. Discordia: Y = 0.03808X + 0.05450. The uncertainty ellipse shown is for fraction NM>160 um (a = 0.032, b = 0.0006). All other ellipses are smaller (the smallest for M <30 \m: a = 0.014, b = 0.0005).

The differences between the two aliquots of the NM 130-160 vim fraction (II, III) fof ra2afr small» amounting to about 3% for 207Pb/23SU and less than 1% for 6Pb/238U. Taking into account the small sample size (14 and 18 mg before splitting) and the relatively large grain size, resulting in less than 5000 grains per aliquot, the two aliquots are regarded as different samples. The data of the leached aliquot (IV) are also within the analytical uncertainty equal to those of the other two aliquots. Even the 208Pb/206Pb ratio lies within the same range, notwithstanding the fact that Th is probably not contained in the zircon structure, but occurs in local Th-rich phases (Mumpton & Roy, 1961; Sommerauer, 1976). When the data of the leachate (*) are plotted, a line connecting this point with that of the leached zircon intersects the positive x-axis, indicating that a phase with a higher 207Pb/206Pb ratio has been leached preferentially. The

- 29 - I Table 5: U-Pb data of the investigated zircons and monazitea* o I Sample/Fraction Weight in mg 20%Fb/206fb 207Pb/206Pb 2«8Pb/206Pb 207Pb/235U 206Pb/238U Pb(yg/g) U(yg/g) (unspiked/spiked) (x 101*)

76-Gal-l NM >160 ym 3.0 2.4 4.8 0.1040 0.1364 1.500 0.1117 29. 247. NM 130-160 7.1 6.8 4.0 0.08482 0.1330 1.041 0.09540 22. 223. NM 130-160 8.3 9.5 0.91 0.08308 0.1302 1.073 0.09516 23.0 231. NM 130-160* 14.1 9.6 2.6 0.08322 0.1302 1.074 0.09794 24 235, NM 40-50 7.2 8.8 3.7 0.06482 0.1299 0.6579 0.08021 23 286. M > 160 5.4 5.4 5. 8 0.08460 0.2539 0.8891 0.08442 21 215. M 110-130 7.6 7.8 7. 0.07243 0.2192 0.7053 0.08255 20 218. M <3O 7.3 5.7 4. 0.06509 0.1444 0.6371 0.07811 26.8 327. NM leachate 3 2 69. 0.2007 0.3715 0.8532 0.05809 1.9 21.2 76-Gal-4 NM 90-110 ym 9.0 9.4 47. 0.1253 0.2771 0.5237 0.06628 112. 1280. NM <30 8.7 9.2 69. 0.1569 0.3644 0.5065 0.06464 155. 1599. M >90 9.2 11.1 79. 0.1708 0.3809 0.4288 0.05538 234. 2714. M 30-50 9.7 11.6 63. 0.1481 0.3411 0.4672 0,05973 235. 2716. corrected with a 17.631. 6 15.498, 37.371 NM 90-110 0.5227 0.06617 NM <30 0.5050 0.06448 M >90 0.4273 0.05522 M 30-50 0.4660 0.05960

K-feldspar 546. 0.8503 2.085 548. 0.8494 2.084 77-Gal-6 NM 130-160 ym 5.5 /' 2.2 4.4 0.07614 0.1036 0.7082 0.07355 58.5 777 NM 90-110 7.6 /' 10.1 4.5 0.06873 0.09274 0.5857 0.06821 57.0 829 NM 50-60 8.2 1' 8.1 3.8 0.06402 0.06957 0.5191 0.06428 55.3 877 NM 30-40 6.3 1' 8.9 4.0 0.06237 0.06079 0.4763 0.06113 59.3 997 77-Gal-12 NM >60 ym 0.7 /' 0.5 21. 0.08531 0.1823 0.2946 0,03891 109 2459 NM 40-60 8.Q /' 5.8 13. 0.07293 0.1916 0.2765 0.03717 111 2671 Mon > 90 5.1 / 3.1 0.82 0.05366 1.937 0.3285 0.04542 1384 11769 Mon 60-90 7.8 1 8.1 0.25 0.05330 1.584 0.2696 0.03695 1185 m 14070 Mon 40-50 9.9 / 5.9 0.46 0.05322 1.527 0.3652 0.05040 1663 14783 r r

Table S - continuation

77-Gal-l3 NM > 130 ym 3 .2 ,1 1 .1 9 .1 0.07249 0.08884 0.3885 0.04752 116. 2393 NM 90-110 7 .0 ,f 2 .8 12 0.07607 0.08939 0.3890 0.04753 117. 2398 NM 40-50 10.9 /f 5 .8 10 0.07335 0.08421 0.3910 0.04825 119. 2421 NM <30 10 .0 t f 3 .7 11 0.07395 0.09769 0.3641 0.04587 128. 2708 Mon > 90 11 .2 j' 9.8 3 .4 0.05724 0.7702 0.2634 0.03656 1230. 21203 77-Gal-l4 NM 60-110 ym 12 .3 /' 5,.9 5 .0 0.06695 0.06967 0.5938 0.07219 86.9 1219 NM <30 9,.8 /' 13,.5 6,.5 0.06687 0.07620 0.5122 0.06478 99.6 1544 M >60 12 .0 /' 7,.0 9,.0 0.07156 0.08505 0.5326 0.06601 108. 1621 M <30 10,.4 /' 10,.2 12, 0.07517 0.09959 0.4602 0.05840 118. 1957 77-Gal-I5 NM 30-60 ym 2,.9 / 2,,5 5,.8 0.08933 0.09121 1.111 0,09934 72.0 706 NM 20-30 2,.9 / 3,.0 2..1 0.07480 0.06774 0.8923 0.09011 57.3 645 NM <20 6. 3 / 7. 8 0.,68 0.06807 0.06228 0.7822 0.08456 63.0 766 MODL >60 2. 2 / 2. 0 0.,75 0.05798 1.751 0.5938 0.07571 922. 5007 77-Gal-19 NM 130-160 ym 10. 0 /' 9. 3 21. 0.1156 0.1581 1.245 0.1040 36.9 311 NM 40-50 12. 8 / 8. 1 5. 8 0.06896 0.1022 0.6463 0.07748 31.6 400 NM <30 13. 2 / 11. 7 8. 5 0.07147 0.1156 0.6199 0.07611 34.7 439 M <30 13. 5 / 9. 9 8. 2 0.07038 0.1384 0.6023 0.07475 29.3 372 77-Gal-23 NM > 160 ym 0. 2 / 0. 2 14. 0.09166 0.1782 0.7755 0.07866 132. 1498 NM 130-160 2. 1 / 1. 4 3. 7 0.06529 0.1084 0.6368 0.07708 124. 1580 NM 110-130 6. 9 / 7. 2 7. 2 0.06865 0.1241 0.6105 0.07617 129. 1624 NM 90-110 8. 9 / 8. 5 3. 7 0.06314 0.1119 0.5929 0.07439 116. 1535 NM 40-50 12. 0 / 8. 7 3. 2 0.06225 0.1146 0.5576 0.07020 91.9 1285 M > no 3. 5 / 2. 5 6. 0 0.06657 0.1241 0.5681 0.07121 135. 1836

Column 1: NM, non-magnetic; M, magnetic; Mon, monazite;*, leached zircon fraction. Column 2: sample weight in milligrams for unspiked and spiked aliquots. Columns 3 through 5: measured 20l>Pb/206Pb, 207Pb/206Pb and 208Pb/206Pb ratios. Columns 6 and 7: ratios radiogenic 207Pb/235O and 206Pb/23eU, after correction for the common lead. Columns 8 and 9: total lead and uranium content.

1 I I -. TI ' -i

amount of lead leached is extremely small, which indicates that the hydrofluoric acid attack had only a very minor effect. This may be interpreted as indicating I that the zircon is composed largely of the highly resistant ZR-phase according to Sommerauer (1976), which is also in accordance with the rather low U contents (between 215 and 330 yg/g). The discordance of this alleged Z^-phase is ascribed to the mixing of two zircon phases (see below). Sample 77-Gal-W Four size and magnetic zircon fractions have been analyzed. When the common lead amounts of 9 to 40 ng (total Pb 370 to 455 ng) are compared with the procedure blanc of 1.6 ng, it is evident that a small initial lead component is present. Therefore, the common lead correction was made by subtracting first an amount of 1.6 ng with AML composition, and then the remaining non-radiogenic lead with a calculated composition (see the Appendix), The four data-points define a discordia which intersects concordia at 2267 i and at 459 ±11 Ma (Fig. 5). They display a regular arrangement along the discordia with respect to the grain size. Except for the magnetic fraction, the U content increases with decreasing grain size.

MELLID ORTHOGNEISS

0.15 BORDER ZONE

=> GAL 19

UNCERTAiNTY ELLIPSE To 2267 Ma

0.10

I: NON-MAGNETIC,130-160^m H: NON-MAGNETIC, 40-50 11: NON-MAGNETIC, < 30 SE: MAGNETIC, <30

0.5 1.5

Fig. S. Conaordta diagram of zircons from the border zone of the Mellid orthogneiss 77-Gal-19. Diaaordia: Y = 0.04484X + 0.04818. The uncertainty ellipse shown is for fraction M 130-160 ]sm (a = 0.029, b = 0.0010). All other ellipses are smaller (the smallest for M < 30 \m: a = 0.014, b = O.0O0S). Interpretation of the intercept ages Taking into account the tectonometamorphic history it is highly improbable that the upper intercept ages would reflect the crystallization age of all zircons, i.e. the intrusion age of the granites from which the gneisses were derived during the Paleozoic orogenesis. The presence of rounded zircon cores, especially in the coarse size fractions (chapter II.3.1.), suggests that the discordia relationships within each suite of zircons result from a mixture of a small amount (<10%) of an old zircon component and a large amount of zircon formed in

- 32 - 1p Early Paleozoic time. Accordingly, the upper intercepts are interpreted as reflecting the age of the crustal components from which the granitic precursors of the gneisses have been generated. A similar explanation, involving some inherited lead, has been invoked for other areas with strongly (>90-95%) discordant zircons (Gulson & Krogh, 1973; Krogh & Davis, 1973; Pidgeon & John- son, 1974).

Whether the difference between the two upper intercept ages is real, remains a matter of speculation. The crustal material from which the granites have been generated could very well have contained components of different age. 3ecause the discordia of sample 77-Gal-19 is based upon one coarse fraction only, the influence of the finer size fractions is stronger. The isotopic systems of the finer fractions are more easily disturbed (see below), which would result in a counter-clockwise rotation of the discordia, In this respect it may be of importance that in both suites of zircons the lowermost data-points appear to show a slight tendency to display a downward bending array.

Anyhow, the presence of minor amounts of an older zircon component with an age of the order of 2500 Ma is evident. Whether this age represents the primary age of the protolith of the original granites or the provenance age of a < 2500 Ma old sediment, is as yet unclear.

The lower intercept ages of about 480 Ma and 460 Ma can be connected with the generation and emplacement of the granitic precursors of the gneisses. These ages differ substantially, however, from the Rb-Sr whole-rock isochron age of 400±23 Ma (van Calsteren et al., 1979 and chapter II.4.). This age difference is puzzling. Possibly it might be explained in terms of a slow cooling after the granite emplacement. The high-grade mineral parageneses of the blastomylonitic border rocks indicate an initial emplacement of the granites at a relatively low crustal level, from which a low cooling rate may be inferred. From the leaching experiment and the rather low U contents of the zircons, it is probable that all zircons consist largely of the stable Zjj-phase according to Sommerauer (1976). This stable zircon phase is characterized by a closed-system behaviour towards uranium and lead migration up to high temperatures (800-900°C, Sommerauer, 1976). It could thus be imagined that the Zg-phase closed to U-Pb migration at a higher temperature than the whole-rocks closed to Sr isotope migration. Such an explanation is supported by the higher Rb-Sr whole-rock isochron ages of about 450-460 Ma revealed by samples of orthogneisses in the blastomylonitic graben (van Cal- steren et al., 1979). These rocks were derived from granites emplaced at a higher crustal level (Arps, 1970) and consequently with a higher cooling rate after emplacement. In this respect the possible slight downward bending arrays of the finer zircon fractions could be explained in terms of a higher amount of the phase in these fractions, which lost part of its accumulating radiogenic lead during the slow cooling. Alternatively, such downward bending arrays could reflect a minar later lead loss, as in the case of the other orthogneisses. Both explana- tions would account for the lower ages of the 77-Gal-19 intercepts.

II.5.2. Gneisses of the Sobrado/Teijeiro Complex II.5.2.1. Sample 77-00.1-14, Curtis orthogneies Zircons from four different size and magnetic fractions have been analyzed. The amount of common lead in the zircons ranges from 16 to 93 ng (total Pb 970 to 1290 ng). In view of the procedure blanc of 1.6 ng, this points to minor initial lead component.

The four3data-points define a line which intersects concordia at 775 ^'^ at 286 ±5iMa (Fig. 6). The data-points display a regular arrangement with respect to grain size along this line. Although the U content increases with decreasing grain size and increasing magnetic susceptibility, the variation with discordance is not regular.

- 33 - CURTIS ORTHOGNEISS

GAL

UNCERTAINTY ELLIPSE

X : NON-MAGNETIC, 60-HO^n I: NON-MAGNETIC, < 30 HI: MAGNETIC, > 80 H: MAGNETIC, < 30

0.05

Too

Fig. 6. Conaordia diagram of zircons from the Curtis orthogneiss 77-Gal-14. Disaordia: X = 0.10051X + 0.01266. The uncertainty ellipse shorn, is for fraction NM 60-110 \im (a = 0.013, b = 0.0004). All other ellipses are smaller (the smallest for NM < 30 ]m: a = 0.011, b = 0.0004).

Interpretation of the intercept ages In view of the Rb-Sr whole-rock isochron age of 450+24 Ma (chapter II.4.) and the tectonometamorphic history of the rocks, the intercept ages probably do not reflect the true ages of crystallization and lead loss of the zircons. The intercept ages of other orthogneiss samples (next paragraphs) support this view. If the upper intercepts are interpreted as reflecting the intrusion ages of the granitic precursors, this would indicate a large number of events of granite intrusion throughout the Late Proterozoic and the Early Paleozoic. Such an interpretation would also involve the metamorphic resetting of the Rb-Sr whole-rock systems, without any indication for mobilization or high-grade metamorphism.

On the other hand, the intercepts can be interpreted in terms of a multi-stage model (Fig. 7). In such a model a discordant zircon suite (crosses) with ages Ti and To, can move to a new, more or less linearly correlated discordant array due to a (second) event of lead loss at T3 (open circles). The resulting intercept ages T| and T£ will have no geological meaning. The original age T, can either be older or younger than Tj, depending on the range of the extent of lead loss. T2 will always be situated between Tj and T', while T3 will be younger than Ti. When the real data-points along a discordia are envisaged, however, the analytical uncertainties may lead to arrangements of data-points producing Ti intercepts younger than the T^ age.

Following this model, the data of the Curtis gneisses are interpreted as follows: Immediately after the generation of the granites by anatexis of older crustal rocks, the suite of zircons constituted a mixing line, similar to the discordia of the Mellid gneisses (Fig. 8). From the higher U contents and the larger degree of metamictization, it may be inferred that the proportion of the Zfj-phase is larger in the Curtis zircons than in the Mellid zircons. Such a higher proportion of the unstable phase, possibly aided by the slightly higher alkalinity of the

- 34 - Fig. ?. Conaordia diagram of multi-stage model. For explanation see text.

0.1

• CURTIS GNEISS DATA-POINTS A MELLID GNEISS DATA-POINTS

0.5 0.75 1.00

Fig. 8. Conaordia diagram of the preferred model for the evolution of the U-Pb systems in the Curtis gneiss ziraons. The full line is the Mellid orthogneiss disaordia. The dashed lines give the range of pre-loss discordance for the Curtis zircons.

- 35 - rocks (as exemplified by the molar (Na20 + K2CO/AI2O3 ratio of 0.79 vs. 0.66 for the Mellid gneisses, chapter II.6. and Table 3) can have favoured lead loss during a later event.

The timing of this later event remains a matter of speculation. For lead loss in the Late Paleozoic (300-320 Ma ago), the arrangement along the initial mixing line, assuming that this was subparallel to the Mellid discordia, must have been such that the finest size fractions had the highest Pb/U ratios and thus contained the highest proportion of old component. This seems highly unlikely and points to later lead loss.

Lead loss can have been caused by processes related to the opening of the Atlantic Ocean. The increase in temperature during the first (Triassic-Early Liassic) and the second (Late Jurassic-Early Cretaceous) rifting events along the continental margin of NW Spain (Maufret et al., 1978; Sibuet et al., 1978) may have resulted in lead loss of the investigated zircons. This temperature increase did not result in metamorphisra in the area, however, even not of very low grade.

During uplift in the Tertiary (e.g. Floor, 1966; Vogel, 1967), possibly related to compressional tectonics along the continental margin (Maufret et al., 1978; Sibuet et al., 1978), pressure-relaxation may have resulted in opening of micro- cracks in the zircon, through which lead can have been leached. Microscopical examination shows that microcracks are abundant on the zircon surface.

A third possibility could be recent lead loss due to (sub)recent weathering. If we assume again an original mixing line similar to the Mellid discordia, a line connecting the zero point of the concordia with the lowermost data-point (the finest fraction) intersects this mixing line approximately at the lower intercept. This would indicate an originally concordant U-Pb system for the finest fraction, implying that no older component was present. Both in the case of lead loss in Jurassic to Tertiary time and in the case of (sub)recent lead loss, the original arrangement of the data-points along the mixing line would have been regularly with respect to the grain size, i.e. an increasing proportion of the old component with increasing grain size.

II.5.2.2. Sample 77-Gal-153 Sdbrado musoovite-chlorite gneiss Three different zircon size fractions and one monazite fraction have been analyzed. The amounts of non-radiogenic lead are low and range from 2.2 to 6.4 ng (total Pb 166 to 395 ng for the zircons and 2030 ng for the monazite). In view of the procedure blanc of 1.6 ng, only the largest zircons and the monazites contain possibly a small amount of initial common lead.

The three zircon data-points define a discordia which intersects concordia at 2272 t\\ Ma and at 476 ±12 Ma (Fig. 9). The monazite data-point is concordant at 471 Ma (T-206 = 470 Ma, T-207 = 473 Ma and T-207/206 = 487 Ma); within the analytical errors this point lies at the lower intercept of the zircon-discordia. A line through all four data-points intersects concordia at 2231 ijjjj Ma and at 472±12Ma. The U content of the zircons shows no regular variation with the grain size, but the degree of discordance increases with decreasing grain-size.

Interpretation of the intercept ages Contrary to the orthogneisses, this discordia is not regarded as a mixing line between two zircon generations. The data are interpreted as representing a detrital zircon suite subjected to an event of lead loss. This interpretation is supported by the sedimentary origin, the fact that the rocks did probably not go through a stage of partial melting, and the absence of overgrowths on the zircons.

Following this interpretation, the upper intercept age of 2270 Ma is taken as reflecting the time of (re)crystallization of the zircons. Although the possibility

- 36 - SOBRADO MUSCOVITE-CHLORITE GNEISS

0.15 GAL 15

I I UNCERTAINTY ELLIPSE

0.10 • ZIRCON: I 30-60,um n 20-30 m <2o ft MONAZITE >60/um

0.5 1.0 1.S

Fig, 9. Concordia diagram of zircons and the monazite from the Sobrado muscovite- chlorite gneiss 77-Gal-15. Discordia: Y = 0.044S0X + 0.05002. The uncertainty ellipse shown is for zircon fraction 30-60 ]im (a = 0.024, b = 0.0006). All other ellipses are smaller (the smallest for fraction < 20 ym: a - 0.017, b - 0.0004).

cannot be ruled out that the age of 2270 Ma represents a complete metamorphic resetting of an older zircon, this Early Proterozoic age is regarded here as the provenance age of the sediment. The sedimentation should than have taken place between about 2250 and 500 Ma ago.

The lower intercept age of 470 Ma would accordingly represent the time of major lead loss (up to 98%). This age approxiiaates the inferred intrusion ages of the orthogneisses containing the least-disturbed zircon suites (Mellid and Noya) and the Cabo Ortegal lherzolites (Table 2). Following the interpretation of the geochronological data of the latter intrusions (van Calsteren et al., 1979), this age may also be interpreted as reflecting the beginning of the M-l phase of granulite facies metamorphism.

The U-Pb system of the monazite, concordant at 471 Ma, closed then also during M-l. A metamorphic growth of monazite under conditions of the higher amphibolite and granulite facies has been inferred from the relation between the amount of monazite and the metamorphic facies of the host rock (Overstreet, 1967) and from concordant monazite ages in the Lepontine Alps (Koppel & Griinenfelder, 1978).

Because of the low mechanical strength, metamict, unstable zircon phases will largely be destroyed during erosion, and transport; detrital zircons should therefore be composed mostly of the well-crystallized, stable ZK-phase (Sommerauer, 1976). The low U and Th contents (the latter estimated from the radiogenic Z08Pb at 120-160 ug/g) of the Gal-15 zircons are in agreement with this view. Never- theless, in spite of the inferred absence of substantial amounts of unstable zircon phases (no metamict or partially metamict grains were encountered), a very large proportion of the lead was lost. An explanation for this high degree of lead loss can be found in the instability of zircon during granulite facies metamorphism (Marshall, 1969).

- 37 - ,-*;.;•

II.5.3. Gneisses of the Blastomylonitic Graben

II.5.3.1. Sample 76-Gal-43 Vigo orthogneiss Four different size and magnetic zircon fractions have been analyzed. The amounts of common lead in these zircons range from 260 to 933 ng, constituting 26 to 44% of the total lead (1000 to 2280 ng). Although the procedure blanc is rather high (5.4 ng), the very large proportion of common lead and the proportionallity of the common lead content to the amount of inclusions and turbidity indicate that a substantial initial common lead component is present. The usual correction procedure can therefore not be applied in this case.

In order to estimate the isotopic composition of the initial lead, two K-feldspar concentrates were analyzed (Table 5). Since the plurifacial metamorphism and polyphase deformation has led to several phases of recrystallization of the feldspar, a higher proportion of radiogenic lead can be expected for the feldspar-lead than for the initial lead of the zircons, which did not recrystallize during these events. In order to derive the composition of the initial lead, it is assumed that (1) the last recrystallization of the feldspar took place around 300 Ma ago, (2) the lead now contained in the feldspar evolved before 300 Ma ago in an environment with U/Pb and Th/Pb ratios similar to the present whole-rock ratios (Table 3), and (3) the zircon crystallized and incorporated common lead approximately at the time calculated when supposing that all common lead has the AML-composition (about 540 Ma). Using 238U/2oVPb = y = 23.4 and 23ZTh/2OlfPb = W « 86,7, the isotopic composition of the initial lead becomes:a= 17.317, $ = 15.476 and Y = 37.019. After sub- traction of a laboratory contaminant of 5.4 ng, this composition is taken for the remaining lead.

After this correction, the data-points of the four zircon fractions define a line intersecting concordia at 551 ±\\M a and 17° -35M a (Fi8- 10)- The data-points show no regular arrangement with respect to the grain size, but the U content increases with increasing discordance and decreasing grain size.

^-""600

VIGO ORTHOGNEISS ^,,^51 Ma)

GAL 4 3 (1075 "7

|S UNCERTAINTY ELLIPSE "^

aoso I: NON-MAGNETIC,90-1W^im II: NON- MAGNETIC, < 30 ffl: MAGNETIC, > 90 12: MAGNETIC, 30-50

(UBS yMnOMa)

lOOy'

azs a'50 0.75

Fig. 10. Conaordia diagram of zircons from the Vigo orthogneiss 76-Gal~4. Disaordia: Y = 0.116091 + 0.00SS7. The uncertainty ellipse shown is for fraction NM 90-110 \m (a = 0.022, b = 0.0012); all other fractions have about the same ellipse.

- 38 - rv>' Due to the large common lead correction, the errors of the Pb/U ratios are less strongly correlated (correlation coefficients ranging from 0.33 to 0.50), resulting in error ellipses with a lower angle of tilt and a rather low aspect ratio. In the Appendix the influence of an error of 5% in the initial lead composition is estimated. This results in a systematic error of±12% for the 207Pb/235U ratio of the NM < 30 ym fraction (II), while the systematic errors for the 20GPb/238U ratios are less than 1%. Consequently, the uncertainty ellipses can have extreme positions centered about the end-points of the±12% bars for these systematic errors. In view of the position of the data-points in the concordia diagram, the errors must be smaller, however. A parallel upward displacement of the discordia, arising from the uncertainty in the initial lead composition, is limited by the concordia. A downward displacement is constrained by the parallel line through the origin of the diagram.

Interpretation of the intercept ages The upper intercept age is somewhat higher than the corresponding Rb-Sr whole-rock isochron age of 469+8 Ma (chapter II.4.). However, the zircon data probably do not reflect the true ages of crystallization and lead loss. The zircon systems of this suite are again interpreted in terms of a multi-stage model. The granitic precursors of the blastomylonitic gneisses in the graben were probably emplaced at a relatively high crustal level (Arps, 1970). From this a rather high cooling rate may be inferred, suggesting that the intrusion time of the rocks is most nearly approached by the Rb-Sr whole-rock isochron age. If the zircons contain some inherited radiogenic lead, the U-Pb systems should have constituted a mixing line prior to the lead loss.

Again, the timing of this lead loss cannot be estimated with any precision. It can only be stated that lead loss around 300 Ma ago is unlikely, since in that case the arrangement of the zircon data-points along the supposed original mixing line would indicate an increase of the old component with decreasing grain size.

Following this interpretation, the common lead correction should be carried out using a 470 Ma old initial lead, instead of the correction with 540 Ma old lead described above. In Table 5, a set of atomic ratios is therefore also given on the basis of the correction with the younger lead. The position of the discordia remains virtually the same, however.

II.5.3.2. Sample 77-Gal-233 Noya orthogneiss Zircons of six different size and magnetic fractions have been analyzed. The amounts of non-radiogenic lead range from 1.6 to 41.1 ng (total Pb 24 to 1100 ng). In view of the procedure blancs of 1.8 and 2.4 ng, the common lead contents indicate the presence of a minor initial lead component in most of the zircons.

In the conuordia diagram (Fig. 11), the data-points do not define a discordia, but they constitute a curved array, which is tentatively interpreted as consisting of two sub-linear parts intersecting at the data-point of the NM 110-130 ym fraction (III). The four lowest data-points (III-VI) define a line which intersects concordia at 559 ±\* Ma and 170 1" Ma. A line through the data-points of the three largest non-magnetic fractions (I-III) intersects concordia at 4170 t\H Ma and 473+12 Ma. This upper intercept is evidently much too high, even if the large error is taken into consideration. The slope of this line depends heavily on the position of the NM>160 ym data-point.(I), but this fraction was very small, containing only 69 zircons; consequently, only low amounts of U and Pb were loaded and measured. In view of the intensities of the measured peaks, the maximum errors are taken at 25% for the 2OI*Pb/206Pb ratio and 1.5% for all other ratios. The uncertainty ellipse for this fraction is therefore very large (Fig. 11). If this point is discarded, a line through the remaining two points gives upper and lower intercepts of 2700 and 468 Ma, respectively.

- 39 - foj.Y

NCYA ORTHOGNEISS M 1 I GAL 23 i y

UNCERTAINTY f ^X x ^ 1 ELLIPSE jS m

X NON-MAGNETIC >160/um S NON-MAGNETIC 130-160 X HI NON-MAGNETIC 110-130 x SE NON-MAGNETIC 90-110 x Z NON-MAGNETIC 40-50 SE MAGNETIC > 110

0.5 0.7S 1.00

Fig. 11. Conoordia diagram of zircons from the Noya orthogneisa 77-Gal-23. The fraction NM >160 ym has a large uncertainty ellipse (a = 0.065, b = 0.00201 see text). The other uncertainty ellipse shown is for fraction NM 130- 160 ym (a = 0.014* b = 0.0004); all other fractions have about the same ellipse.

Except for the magnetic fraction, the data-points show a regular arrangement with respect to the grain size. The U content increases with decreasing grain size for the three largest non-magnetic fractions, but decreases then with grain size for the non-magnetic fractions smaller than 110 \m.

Interpretation of the intercept ages The rather peculiar discordia picture is again interpreted in terms of a multi- stage model. As is the case with the Vigo gneisses, a relatively high crustal emplacement of the granitic precursors is probable (Arps, 1970). The Rb-Sr whole- rock reference age of 459 Ma (chapter II.4.) may therefore be regarded as approximating the intrusion time. In that case the coarsest non-magnetic zircon fractions would contain a small amount of an inherited older zircon component, and the zircons along the upper trajectory would represent a mixing line (although probably disturbed) between the two zircon generations.

The intercepts of the lower trajectory, resulting from a younger episode of lead loss, should then have no geological meaning. From the position of the data- points it is unlikely that lead loss took place in Paleozoic time, since this would imply that the smallest zircons contain the highest proportion of the old component; any time later than about 250 Ma ago is possible, however.

X-ray and optical investigations indicate that all zircons constitute a single suite. The discordia pattern can therefore only be explained by an episode of isotopic disturbance, which had its greatest impact on the finest grains or on the grains containing magnetic inclusions. An interesting phenomenon is that the break in the discordant array coincides with the maximum U content. The decrease in U content with increasing discordance along the lower trajectory is apparently at variance with models of lead loss. A possible explanation may be found in the geometry of zircon-phase intergrowths. The larger grains are mostly zoned, often containing brownish cores, in which the outermost zone is usually clear. The

- 40 - smaller grains are not or only very little zoned and do not contain the clear rim (chapter II.3.3.). It could be imagined that the clear rim around the larger zircons is composed of the resistant Zj(—phase, which shielded the crystals during a later disturbance. In spite of the lower U contents and, consequently, the lower proportion of Z^-phase, the smaller zircons may thus have suffered more of the later lead loss.

The zircons of samples 76-Gal-4 and 77-Gal-23 show virtually the same characteristics. Nevertheless, the discordia picture of 76-Gal-4 differs from that of 77-Gal- 23. The U contents and the geometry of zircon-phases therefore cannot be the only factors influencing the degree of lead loss. Possibly, the composition of the environment may also be of importance. In this respect it may be noted that the alkalinity of sample 77-Gal-23 is lower than that ol 76-Gal-4 (Table 3).

II.5.3.3. Sample 7?-Gal-6> Noya migmatitia augengneiaa Four different zircon size fractions have been analyzed. The amounts of non-radiogenic lead range from 3.5 to 16.2 ng (total Pb 320 to 450 ng). In view of the procedure blanc of 3.1 ng this indicates the presence of a minor amount of initial lead; around 1 yg/g for the three smallest fractions.

The four data-points do not show a good linear correlation in the concordia diagram; they rather form a curved array (Fig. 12). A line through all four data- points intersects concordia at 2027 t\H Ma and 379±10 Ma. Both the discordance and the 0 contents shov a regular variation with respect to the grain size. Unlike

NOYA MIGMATIC AUGENGNEISS o.H GAL 6

I 130-160^111 I 90-110 IDC 50- 60

0.05 3Z 30- 40

207Pb/235u

0.5 0.15 1.00

Fig. 12. Concordia diagram of zircons from the Noya migmatia augengneiss 77-Gal- 6. Disoordia: Y = 0.0S706X + 0.03685. The uncertainty ellipse shown is for fraction NM 130-160 m (a = 0.015, b = 0.0004). All other ellipses are smaller (the smallest for NM 30-40 ym.* a = 0.010s b = 0.0004).

the zircons of the other orthogneisses, in which the 208Pb content varies erra- tically or increases along with an increase in U content and decrease in grain size, the radiogenic 208Pb content of the Gal-6 zircons slowly decreases with the grain

- 41 - size. In comparison with the increase in U content, however, the decrease of the estimated Th contents is rather small (215-140 ug/g). F

Interpretation of the intercept ages t •: Although no Rb-Sr data have been obtained from this sample, a comparison with other orthogneisses renders an intrusion age of about 450 Ma likely, Migmatiza- tion of this gneiss took place during M-3, to which phase an age of around 310 Ma has been assigned (chapter 1.3.)- The discordia pattern reflects thus again the presence of an older zircon component.

In view of the isotopic disturbance suggested by the apparent curved array, the intercepts give minimum estimated for the ages of both the old and the new zircon components constituting the original mixing line. The later isotopic disturbance event leading to the curvature of the trajectory may be explained by either one of two alternative processes: growth of a third zircon generation during the migmatization, or lead-loss during the migmatization or at some other time.

The gneiss 77-Gal-6 shows only minor effects of the migmatization (chapter II.3.3.). This, along with the absence of overgrowths and the very slight corrosion of the zircons make the growth of a third zircon generation during M-3 unlikely. Moreover, such a new component must amount to 30-50% of the finest zircon fraction in order to explain the position of the data-points.

A better explanation for the discordia pattern is lead-loss from zircons with an already discordant U-Pb system. The maximum age for the episode of lead-loss is indicated by the intersection with concordia of a line through the two lowest data-points: 359 Ma. An obvious time of lead-loss would be the M-3 migmatization. Any time later than 310 Ma is also possible, perhaps even more likely taking into account the probable post-300 Ma lead loss of some other orthogneisses.

II.5.4. Late-Paleozoic Granites U.S.4.1. Sr:nple 77-Gal-lZ, Cabo Sillei.ro granite Four different zircon size fractions and one monazite fraction have been analyzed. The amounts of common lead in the zircons range from 7 to 44 ng (total Pb 370 to 1290 ng), the procedure blanc was 1.7 ng; this indicates the presence of a small initial lead component. For the monazite the amount of common lead is high (165 ng), also indicating the presence of an initial common lead component, but it still constitutes only 1.5% of the total lead (13780 ng).

The four zircon data-points show little spread in Pb/U: the three coarsest fractions cluster, while the finest fraction has a somewhat lower ratio (Fig. 13). A line through all four data-points has upper and lower intercepts of 1502 t!\\\ Ma and 254 Z\B Ma, respectively. The U content of the zircons increases with decreasing grain size. The monazite data-point is nearly concordant: T-206 = 231 Ma, T-207 = 237 Ma and T-207/206 = 296 Ma. Due to the very high U content (21000 Ug/g), the sample/spike ratio was unfavourable, giving rise to a large error enlargement factor for single isotope dilution. This resulted in an uncertainty ellipse nearly twice as long in the x-direction as the ellipses of the zircon data-points (Fig. 13).

Interpretation of the intercept ages Both the Rb-Sr whole-rock isochron age of 311 ±21 Ma and its relation to the Late Paleozoic migmatization, metamorphism and deformation, define the age of this palingenetic granite rather accurately. The discordia pattern of the zircons can therefore only be interpreted in terms of a multi-stage model. The anatectic nature of the granites of the alkaline series renders an inheritance of zircons from the protolith likely. This finds support in Arps' (1970) observation that

- 42 - there is an inverse relationship between the degree of homogenization of autochthonous granites and the proportion of relict sedimentary zircon cores.

3 CABO SILLEIRO GRANITE 8 0.06 I" GAL 13 §

UNCERTAINTY ^^ ^^^--'^(To 1502 Ma) aos

UNCERTAINTY^ ^^^^'^ • ZIRCON: ao4 ElllPSE^^^ ^^"^ I >130/um MONAZITE ^^ >^ IE 90-110 IE 40-50 S<30

* MONAZITE >90/um

0.2 0.3 0.4 0.5

Fig. 13. Concordia diagram of zircons and the monazite from the Cabo Silleiro granite ?7-Gal-13. Discordia: Y = 0.071SSX + 0.01988. The uncertainty ellipse shown for zircon is for the fraation 40-50 \m (a = 0.009, b = 0.0004); all other zircon fractions have about the same ellipse. The uncertainty ellipse for the monazite is also shown (a = 0.016t b = 0.0002).

The proportion of euhedral or resorbed zircon cores is quite large, especially in the coarser fractions. Nevertheless, there can have been no substantial contribution to the U-Pb systems of the total zircon suite unless a post-1000 Ma age of the old component and a high degree of post-310 Ma lead loss is supposed. This apparent discrepancy can be explained when it is assumed that the older zircon component lost part of its radiogenic lead during anatexis. The presence of metamict cores could support such an explanation.

If it is assumed that the monazite did not inherit any pre-310 Ma component, a line connecting the 310 Ma point on concordia with the monazite data-point should approximate the time of the post-310 Ma lead loss. This line intersects concor- I dia at 46 til Ma- Tne error limits are obtained by connecting the maximum points of the 310±21 Ma interval on concordia with the end-points of the uncertainty ellipse axes of the monazite.

II.5.4.2. Sample 77-Gal-123 Espenuaa granite Two zircon and three monazite size *.-'actions have been analyzed. The common lead contents of the zircons are around 9 Ug/g, amounting to 6.4 and 48.1 ng for the > 60 ym and 40-60 ym fractions, respectively (total Pb is 76 and 890 ng). The amounts of common lead in the monazites range from 7 to 13 ng (total Pb 7010 to 16530 ng). In view of the procedure-blancs of 1.6 ng and 3.5 ng, the presence of an initial non-radiogenic lead component can be inferred.

- 43 - 0.06

ESPENUCA GRANITE

0.05 GAL 12 (To939Ma)

UNCERTAINTY 0.04 ELLIPSE MONAZITE

0.03 UNCERTAINTY ^ • ZIRCON: ELLIPSE ZIRCON I : >«0/uni H: 40-60 A MONAZITE:

IE: >90/um 12:80-90 1: 40 - SO

0.01

0.1 0.2 0.3 0.4

Fig. 14. Concordia diagram of zircons and monazites from the Espenuca granite 7?-Gal-12. Discordia through the zircon data-points: Y = 0.09610X + 0.01060. Discordia through the monazites: Y = 0.14112X - 0.00106. The uncertainty ellipse shown for zircon is for the fraation 40-60 \im (a = 0.007, b = 0.0004); the other zircon fraction has about the same ellipse. The uncertainty ellipse shown for monazite is for the fraction 40-50 ]im (a = 0.035, b = 0.0002); the fractions 60-90 ]im and > 90 ]m have smaller ellipses (a = 0.029, b = 0.0002, and a = 0.016, b = 0.0002, respectively).

In the concordia diagram, the two zircon data-points are discordant (Fig. 14): A line through these two points intersects concordia at 940 Ma and at 200 Ma. The three monazite data-points show a rather peculiar pattern in the concordia diagram. The coarsest size fraction (> 90 Jim) is concordant at 287 Ma (T-206 = 286 Ma, T-207 = 288 Ma, T-207/206 = 306 Ma), the intermediate fraction is discordant, while the finest fraction (40-50 ym) is again concordant, but with an age of 316 Ma (T-206 = 317 Ma, T-207 = 316 Ma, T-207/206 = 309 Ma). A line through all three data-points has an upper intercept at 308 +8 Ma and intersects the positive x- axis. As is the case for the Cabo Silleiro monazite, the U and Pb contents of the monazites are very high; the uncertainty ellipses are therefore two to five times those of the zircons. The U contents of the zircon and monazite fractions increase with decreasing grain size.

Interpretation of the intercept ages In view of the late-kinematic setting, the age of this granite can be taken at around 290 to 300 Ma (Kuijper, 1975). The two zircon data-points reflect thus the presence of an inherited zircon component and a post-290 Ma disturbance of the U- Pb systems. A maximum age of 200 Ma can be inferred for this disturbance.

The difference between the two concordant monazite size fractions is not significant when the error limits are taken into consideration: as is evident from Fig. 14, the error ellipses of the>90 pm and the 40-50 ym data-points overlap each other. The lower intersection to the right of the origin may also be explained by the large uncertainty limits, as the zero-point of concordia lies within the error limits of the line through the three monazite data-points.

- 44 - In spite of the large error limits, the concordant-discordant-concordant pattern of three successive size fractions cannot be regarded as insignificant. From the brownish colour, the often rather good cleavage and the occurrence of tiny opaque inclusions, an incipient alteration of the raonazites can be inferred (Molloy, 1959). Although no variation in the degree of alteration with grain size could be detected, the alteration may have played a role in determining the susceptibility to later lead loss.

The monazite analyses confirm thus the inferred age of the Espenuca granite. The zircon data again indicate the presence of an older, inherited zircon component.

II.6. GENERAL DISCUSSION OF THE U-Pb MINERAL DATA

II.6.1. The Zircon Upper Intercept Ages All of the investigated suites of zircons contain a minor older zircon component. This older component is interpreted as being inherited from a protolith from which the rocks were generated by processes of anatexis. This indicates that all granitic magmas were derived from an older crust. The same conclusion has been drawn for the Late Paleozoic calcalkaline and alkaline granites on the basis of REE analyses (Albuquerque, 1978). The initial Sr ratios of the orthogneisses also suggest the involvement of older crustal components.

The presence of the inherited zircon component gives rise to discordant arrays in the concordia diagrams. For most of the investigated suites of zircons this pattern is severely disturbed by a later event of lead loss. Samples with lower intercept ages comparable to the corresponding Rb-Sr whole-rock ages (Mellid orthogneisses) are the least disturbed and have the highest Pb/U ratios. For such samples the upper intercepts are interpreted as approximating the (re)crystallization age of the older zircon component. Some samples were subjected to a higher degree of lead loss, but they still contain clear evidence of an inherited component in their coarsest zircon fractions. In such cases the upper intercept ages are only minimum estimates of the age of the older zircon component, depending on the extent of lead loss. The zircons of the remaining samples have lost still higher proportions of their radiogenic lead and the upper intercept ages are geologically meaningless.

The zircon data do not allow any conclusion about the geological significance of the Early Proterozoic to Late Archean ages. If the protolith consisted of metasediments, two interpretations are feasible: the upper intercepts could represent the time of metamorphic resetting of detrital zircons, the sedimentation being older, or they could represent the provenance age of the sediment. As large-scale zircon growth during sedimentation or diagenesis is highly improbable, the upper intercepts cannot be interpreted as sedimentation ages.

II.6.2. The Zircon Lower Intercept Ages As is the case with the upper intercept ages, the geological meaning of the lower intercepts depends on the degree of lead loss. The lower intercept ages of the least disturbed samples are interpreted as reflecting the time of intrusion of the granitic rocks. The lower intercepts of the strongly disturbed samples can be used to estimate the time of lead loss. On the basis of the position of the most discordant data-points, it has been shown that lead loss in Paleozoic time is unlikely. In the Mesozoic a rise in temperature during the opening of the Atlantic Ocean could have facilitated lead loss, but no evidence whatsoever has been found for such an increase in the regional temperature. Another possibility could be the uplift in the Tertiary in relation to compressional tectonics along the continental margin; this could have provided a pressure-release resulting in leaching of lead through micro-cracks in the zircon crystals according to the dilatancy-model (Goldich & Mudrey, 1972). The uplift in the Tertiary has been relatively small, however. A third, perhaps more feasible process is (sub-)recent lead loss during weathering. A, 11,6.3. Relation between the degree of disturbance of the zircons and whole-rock composition Regarding the degree of lead loss the investigated samples can tentatively be divided into three groups. Sample 77-Gal-15 did not loose significant amounts of lead after 470 Ma ago; the proportion lost from samples Gal-1 and Gal-19, if any, must also have been very small. These three samples constitute the least disturbed group. The proportion of lead loss from the zircons of samples Gal-6 and Gal- 23 were evidently larger, but not large enough to obliterate the presence of an inherited component. Sample Gal-13 can also be classed under the moderately disturbed samples. The remaining samples, Gal-4, Gal-!2 and Gal-14, contain the strongest disturbed suites of zircons.

In table 6 the samples are arranged according to the order of increasing degree of disturbance explained above, along with the U and Th contents of the zircons and the alkalinity of the host rocks. The data show a rather poor correlation between the U and Th contents and the degree of lead loss. The absence of a clear correlation between the U content and the degree of discordance was already noted in the paragraphs discussing the different investigated suites of zircons. This implies that the U content, although important, is not the only factor governing the degree of lead loss. From Table 6 an increase of the degree of disturbance with increasing alkalinity of the host rocks can be inferred. This correlation is still rather speculative and based on a small number of samples from a restricted area only. The study of the zircon U-Pb systems in three phases of the Cam Chuinneag granite in Scotland, however, lends support to the hypothesis that the chemical composition of the environment influences the degree of lead loss (Pidgeon & Johnson, 1974).

Table 6: Investigated suites of zircons arranged according to increasing degree of lead loss* U Th alkalinity of (ug/g) (ug/g) host rock

w Gal-15 645-766 122-156 .30 aw w Gal-1 215-330 83-160 .67 o o Gal-19 311-440 83-130 .66 S Gal-23 1284-1836 414-597 .71 S Gal-6 780-997 142-214 .71 Gal-13 2390-2710 346-470 SIN G

STU R .71

II.6.4. Monazites The few investigated monazite fractions show the varying behaviour of the U-Pb systems of this mineral. The degree of lead loss is possibly related to the extent of alteration. Monazite Gal-15 is clear and without any cleavage and has a concordant U-Pb system. Monazite Gal-13, although clear, is slightly altered in view of the occurrence of cleavage (Molloy, 1959) and has a discordant U-Pb system. Finally, the monazites Gal-12 are brownish and show a good cleavage, indicating a stronger degree of alteration; they have both concordant and discordant U-Pb systems.

- 46 - 11.6,5. Source rocks of the granites and orthogneisses Field relationships show that the Early Paleozoic sediments acted as source rocks for the Late Paleozoic alkaline granite series. The Late Paleozoic calcalkaline granites are thought to have been generated by partial melting of deeper and/or drier parts of the crust (Capdevila et al., 1973), The REE patterns of rocks from both series are consistent with a derivation by 30-60% partial melting of metasediments with a pelitic to greywackish composition (Albuquerque, 1978),

Both the zircons of the sample of high-grade metasediments and the inherited zircon components in the orthogneisses display Early Proterozoic ages. It may be supposed that some genetic relationship exists between both rock types. The normative Qz-Qr-Ab-An values of the orthogneiss samples are presented in Table 3. In the Qz-Ab-Or and An-Ab-Or projections (Winkler et al., 1976), they plot on or close to the cotectic surfaces (Fig, 15), Trtis indicates that the gneisses represent relatively low-temperature granitic melts (Winkler et al., 1975, 1977). The presumably metasedimentary rocks of the high-grade complexes are mostly garnet-

J9» * ™ 33 U&7&

31 "" 29,, 31 <,

Fig. 15. Normative Qz-Or-Ab-An values of orthogneisses (o) and garnet-biotite gneisses (%). An and Qz values (Qz + Or + Ab + An - 100) are given for each point in the upper and lower diagram, respectively. Temperature-contours3 An values and Qz values on the ooteatio surfaces after Winkler et al. (1975); these Values relate to Vt=PRz0~7 ^* For the pressures of about 10 kb probably prevailing during M-l (Chapter III.3.) the temperatures should be lowered by about 15°C and the An-contents raised by 1-2% (winkler et al., 1975).

- 47 - [fe, biotite gneisses (Vogel, 1967; Engels, 1972; Hubregtse, 1973a; Kuijper, 1979), in which the proportion of alkali-feldspar is variable (0-30 vol %). The garnet- biotite gneisses of the Mellid complex have a greywackish to semi-pelitic composition (Hubregtse, 1973a); normative Qz-Or-Ab-An values, not corrected for modal biotite, are given in Table 7. In the Qz-Or-Ab-An tetrahedron, most of the Mellid paragneisses fall in the plagioclase space, close to the plagioclase-quartz cotectic surface (Fig. 15). For the mafic rocks the P-T conditions during M-l are estimated at around 800-850°C and around 10 kb (see Chapter III). For the meta- sediraents Pt must have been the same, while the temperatures can have been somewhat lower. Under these conditions rocks like ^he Mellid garnet-biotite gneisses can easily yield large volumes of granitic melt with compositions comparable to those of the orthogneisses, provided that enough water is available (Winkler, t 1974; Winkler et al,, 1975, 1977). In view of the restricted occurrence of partial melting phenomena in the outcropping paragneisses, it is assumed that during M-l PHaO was *ess tnan pt (Engels, 1972). If the granitic precursors of the orthogneisses were derived from the metasediments, either local variations in PH2O must, have played a role, or much larger volumes of metasediments yielded only low proportions of granitic melt. In the latter case large amounts of partial melt residues should be present. Trace element analysis of garnet-hornblende gneisses from the Cabo Ortegal complex revealed that these gneisses are indeed residues' of partial melting (Drury, 1979), which strongly supports the above hypothesis. In other garnet- biotit" occurrences the proportion of K-feldspar is probably too large for a melt residue. For those units it must be assumed that local variations in PH,0 Save rise to melt production in some parts and only minor anatexis in other parts. In any case, a trace element survey should be needed in order to detect the residual and still fertile parts of the metasedimentary units.

Table ?: Normative Qz-Or-Ab-An values of garnet-biotite gneisses of the Mellid complex, not corrected for modal biotite. Sample numbers taken from Hubregtse (1973a).

Mel Mel Mel Mel Mel Mel Mel Mel Mel Mel Mel 27 28 29 30 31 32 33 34 35 36 37

Qz 39 28 30 25 31 28 44 49 29 47 41 Ab 33 35 37 32 30 28 25 29 34 24 28 An 10 12 13 15 12 10 11 8 13 11 9 Or 18 24 21 28 28 34 20 15 24 18 22

If the high-grade metasediments are accepted as possible source-rocks of the granitic precursors of the orthogneisses, an attempt can be made to estimate the sedimentation age, as follows: The Rb/Sr ratios (by weight) determined by XRF analysis on 20 samples of high-grade metasediments collected all over the Teijeiro area range from 0.08 to 0.65, averaging 0.31. It is difficult to estimate the 87Sr/86Sr ratio of the sediments at the time of deposition, as strontium in non-carbonate sediments does not always equilibrate with the strontium of ocean water and ratios between 0.704 and 0.743 have been reported (Faure & Powell, 1972). A ratio of 0.704 can be taken as a minimum value, however. The initial 87Sr/86Sr ratios of the investigated orthogneisses average at 0.709 with an average Rb-Sr whole-rock age of about 460 Ma (chapter H.4.). If the range in Rb/Sr ratios of the metasediments and the minimum initial 87Sr/86Sr ratio of 0.704 are related to the initial 87Sr/ 8ESr ratio of the granitic magmas generated from them about 460 Ma ago, assuming that the Rb-Sr evolution in the metasediments has proceeded in an approximately closed system after the deposition, a maximum sedimentation age can be calculated of about 800 Ma for Rb/Sr - 0.3 and about 1500 Ma for Rb/Sr =0.1.

II.7. CONCLUDING REMARKS The discordia intercepts at about 480 Ma may be relat.ed to the beginning of the conditions of granulite facies metamorphism (M-l). All investigated orthogneisses

- 48 - and granites from western Galicia contain an older zircon component, which for some samples revealed an Early Proterozoic age. This older zircon generation was presumably inherited from the protolith from which the granitic rocks were generated by anatexis around 460-480 Ma and 300 Ma ago, respectively. Events of major crustal reworking during the Paleozoic are therefore apparent. In view of the occurrence of Early Proterozoic zircons in the high-grade metasediments of the Sobrado area, it can be assumed that the metasedimentary units of the catazonal complexes constitute the source rocks from which the granitic precursors of the orthogneisses were derived. The major element chemistry of one metasedimentary unit is consistent with this view, provided that enough H2O should have been available in order to produce significant amounts of melt. In one particular case it has been shown that under lower-crustal conditions Rb-Sr whole-rock systems can yield ages younger than the corresponding zircon U-Pb age, presumably due to prolonged Sr isotopic equilibration at elevated temperature. A comparison of the major element chemistry of the host rocks and the degree of (sub-)recent lead loss of the suites of zircons indicates that the alkalinity of the environment may be of importance in determining the extent of lead loss from zircons.

- 49 - _. I

CHAPTER 111 PETROGRAPHY AND WHOLE-ROCK U-Pb SYSTEMATICS OF HIGH-GRADE METAMORPHIC MAFIC ROCKS

III.l. INTRODUCTION

Detailed petrological and geochemical studies of the eclogite and granulite facies mafic rocks of western Galicia (Vogel, 1967; Engels, 1972; Hubregtse, 1973a, 1973b; van Calsteren, 1978) have provided much information about the metamorphic history and origin. Nevertheless, several questions are still unanswer- ed. Geochemical data show that the rocks are comparable to continental plateau basalts (van Calsteren, 1978), but the time of intrusion or extrusion is unknown. After emplacement, the rocks were subjected to metamorphism in the eclogite and granulite facies. The time relationship between both catazonal facies is uncertain, but the preferred orientations of the clinopyroxenes in eclogites and granulites at Cabo Ortegal (Engels, 1972) could point to different tectonic regimes during the development of the two mineral assemblages. The mantle plume model proposed for the geological history of Galicia (van Calsteren, 1977) also implies an older age for the event of eclogite facies metamorphism. No isotope geochronological data are available regarding this alleged age difference, however, as a Rb-Sr whole-rock investigation of Cabo Ortegal eclogites did not reveal an isochron relationship (van Calsteren et al., 1979). The conditions of the granulite facies metamorphism (M-l/M-2) are thought to have prevailed from about 480 Ma ago (chapter II.5.2.2.) until 347+17 Ma ago (Rb-Sr whole-rock isochron age determined for the Cabo Ortegal migmatic granulites), although the K-Ar mineral ages cluster around 400 Ma (van Calsteren et al., 1979).

The occurrence of eclogites, granofelses and transitional rock types in the northern part of the Sobrado/Teijeiro complex makes this area very suitable for investigation of the relationship between the eclogite and granulite facies parageneses. Since both the Rb-Sr whole-rock and the U-Pb zircon investigations have shown to be unsuitable for dating these metamorphic events in the mafic rocks, a reconnaissance investigation of the U-Pb whole-rock systems was undertaken. In such an investigation constraints on the number of possible stages in the lead evolution have to be provided by petrological data, hence evidence for a time difference between the two events will be presented first. In order to establish a link between the Teijeiro area and the much better known Cabo Ortegal complex, samples were investigated from both complexes. Although the rocks have identical chemical and modal compositions and underwent the same metamorphic history, the requirement of an identical initial lead composition is not necessarily fullfilled.

III.2. METAMORPHIC HISTORY OF THE HIGH-GRADE MAFIC ROCKS SOUTH OF TEIJEIRO

The geological setting of the high-grade metamorphic Sobrado/Teijeiro complex has been described in chapter II.3. The catazonal mafic rocks occur as a 0.5-1 km wide zone around the central serpentinites and along the western border of the high-grade center (Fig. 2). No differences have been observed between the rocks on both sides of the serpentinites south of Teijeiro, except for the larger occurrence of (retrograded) eclogites in the eastern band. The rocks form a heterogeneous unit of granofelses*, mafic granulites, garnet amphibolites, epidote-rich amphibolites, eclogites and metagabbros. Similar rocks are found as rounded inclusions with diameters of about 0.1-1 m in the bordering garnet-biotite gneisses, while two outcrops of paragneiss and a serpentinite inclusion occur in the mafic unit. Most of the heterogeneity is caused by different degrees of deformation and retrogradation, which both vary on the scale of an outcrop (Kuijper, 1975, 1979).

* Named after Behr et al. (1971).

- 50 - The various metamorphic and deformation phases to which these rocks were subjected are presented in Table 1. The granulite facies metamorphism and younger events are described in chronological order. For a detailed description reference can be made to Vogel (1967) and Hubregtse (1973a).

111.2.1. Granulite facies and younger metamorphic events M-l Granulite faoies/F-1 phase of deformation During the first phase of granulite faeies metamorphism a paragenesis of garnet + diopsidic clinopyroxene + quartz + plagioclase (25-45% An) + rutile ± zoisite ± pargasitic hornblende was formed in the mafic rocks. The conditions were those of the (hornblende-)clinopyroxerce-almandine subfacies of the granulite facies (De Waard, 1965). The M-l/F-1 fabric is an inequigranular polygonal granoblastic assemblage of garnet, clinopyroxene and plagioclase. These rocks are designated as (hornblende-)garnet-clinopyroxene granofelses. Although preferred orientations seem to be absent, a folding phase (F-l) accompanying the M-l metamorphism is postulated, as is the case for the Cabo Ortegal (Engels, 1972) and Mellid (Hubregtse, 1973a) areas.

A pre-M-1 fabric is indicated by aji inclusion pattern within a garnet porphyro- clast in a banded mafic granulite^ This M-l garnet contains parallel bands of quartz grains at an angle of 37° with the foliation (Kuijper, 1975).

M-2 Hornblende-granulite faoies/F-2 and F-3 phases of deformation In addition to a partial recrystallization of the clinopyroxene during M-2, also pargasitic hornblende was formed at the expense of the pyroxene, either as porphyroblasts or as rims around clinopyroxene. A second generation of garnet was formed as rims around M-l garnet, or as new grains. The M-2 paragenesis garnet + clinopyroxene + pargasitic hornblende + rutile + zoisite + plagioclase (20-45% An) + quartz is indicative for conditions of the hornblende granulite facies (hornblende-clinopyroxene-almandine subfacies) (De Waard, 1965). Evidently, T and Ptotal did not change significantly between M-l and M-2. The major difference between the conditions of the two phases was an increase in Pfiu£d during M-2, as can be inferred from the growth of pargasitic hornblende and local scapolite (Engels, 1972; Kuijper, 1979).

Penetrative deformation (F-3) accompanied the M-2 metamorphism and gave rise to blastoraylonitic textures and banding on the scale of millimeters to centimeters. The preferred orientation of pargasite porphyroblasts and M-2 garnets defines the F-3 foliation. Locally, pargasite blasts in parallel orientation overgrow narrow (blasto)mylonitic zones, which developed during a pre-M-2 deformation (F-2). The typical rock resulting from the M-2 metamorphism and the F-2/3 deformation is a banded mafic granulite with blastomylonitic textures. Granulite textures occur locally in bands rich in quartz and feldspar.

Near the central serpentinites (Fig. 2, sample 77-Gal-18), very localized migmatization occurred during M-2. The resulting migmatic mafic granulite consists of quartz-feldspar bands containing large greenish-brown hornblende porphyroblasts with relict garnet inclusions, and bands of hornblende clasts and relics of garnet and clinopyroxene embedded in a fine-grained groundmass of hornblende and interstitial plagioclase.

Minor gabbroic intrusions were emplaced prior to M-2. They show either relict ophitic textures with garnet coronas around polycrystalline amphibole-clinopyroxene aggregates, or flaser textures.

M-3 Amphibolite faaies/F-4 phase of deformation During M-3 a paragenesis of green hornblende + sphene + epidote + plagioclase (20- 40% An) + quartz + garnet developed. This retrogradation resulted locally in the

- 51 - formation of (garnet-)amphibolites containing relics of the older high-grade parageneses,

The M-3 metamorphism was accompanied by a penetrative deformation (F-4) responsible for the subvertical NW to NE trending main foliation, expressed by a preferred orientation of the M-3 hornblendes. Blastomylonitic cextures, related to F-4, occur in most of the samples, especially in the domains rich in quartz and feldspar,

M~4 Gveenschtst faaiee Replacement of the older mineral parageneses took place during M-4 either along narrow veins (<1 mm) C* along broader zones (at least up to several centimeters wide). In strongly retrograded samples nearly all older minerals have (partly) been replaced by a new paragenesis of actinolitic hornblende + blue-green hornblende + chlorite + (Fe-poor) epidote/clinozoisite + plagioclase + quartz.

F-S Phase of deformation A late phase of deformation is apparent from the undulatory extinction of M-4 actinolites and chlorites. This phase was not accompanied by any recrystallization or development of a new foliation.

111.2.2. Eclogites East of the central serpentinites a relatively large amount of eclogites occurs scattered throughout the band of high-grade mafic rocks. The eclogites contain clinopyroxene and garnet as main constituents, occurring as large porphyroblasts or -clasts. Accessory minerals are quartz, a-zoisite, rutile and kyanite. Clino- pyroxene-plagioclase symplectites replace most of the original clinopyroxene. These symplectites are occassionally surrounded and partly replaced by symplectites of brown hornblende + plagioclase. Retrograde effects include the development of brown and green hornblende, epidote/clinozoisite, chlorite, sphene and sericite (around kyanite). In most of the rocks foliation is absent or weakly developed. Blastomylonitic textures occur locally, often only along narrow zones.

111.2.3. Eclogite-granulite relations 111.2.3.1. Transitional rooks Many samples of granofels and granulite show features which can be interpreted as indicating a transition between the eclogite and granofels (M—1) parageneses. In the granofelses of the Teijeiro area, rutile-needles in garnet, although without any preferred orientation, are a common feature. Most of the garnets are surrounded by kelyphitic rims of hornblende + plagioclase. The clinopyroxenes are nearly colourless and show hardly any pleochroism, while the garnets are colourless to light-pink. Kyanite has been found in a few samples; textural relationships with the M-l minerals have not been preserved, but in one sample kyanite occurs enclosed in M-2 pargasitic hornblende. All the above features have been described as characteristic for the (retrograded) eclogites of the Cabo Ortegal area (Vogel, 1967). Both a-zoisite and 3~zoisite occur in the Teijeiro granofelses. According to Vogel (1967) and Den Tex (1971), a-zoisite should be attributed to the eclogite facies, while g-zoisite should belong to the granulite facies paragenesis.

111.2.3.2. Clinopyroxene-plagioclase symplectites The occurrence of clinopyroxene-plagioclase symplectites (designated as cpx-plag symplectites), is a much stronger indication for a transition between eclogite and granofels parageneses. In the Teijeiro area nearly all granofelses and granulites contain intergrowths of clinopyroxene and plagioclase secondary after

- 52 - clinopyroxetie. The cpx-plag symplectites show the same stages in their development as described from exsolved omphacites in the Cabo Ortegal eclogites (Vogel, 1967): a first stage of narrow dusty zones along cracks in the crystal, intermediate stages of increasingly coarser symplectitic •intergrowths, and finally clinopyroxenes containing (vermicular) plagioclase inclusions with identical optical orientation. In most of the rocks the development of symplectites has proceeded until the final stages.

The textural relationships of the cpx-plag symplectites with other minerals are difficult to unravel. In a few samples recrystallization or replacement phenomena have been found:

Sample 7672Q-A (77-Gal-i8, Fig. 2) is a weakly foliated mafic granulite consisting of large porphyroc^asts of garnet, clinopyroxene and amphibole, and amphibole-aggregates with clinopyroxene relics, all embedded in a fine-grained groundmass of amphibole, plagioclase and quartz showing blastomylonitic textures. Most of the clinopyroxene contains vermicular plagioclase inclusions or forms symplectitic intergrowths with plagioclase. Locally, the rims of the clasts and the cpx- plag symplectites have recrystallized into inequigranular, polygonal granoblastic aggregates of very-fine grained clinopyroxene and some hornblende with interstitial plagioclase. From the relationships with the adjoining exsolved clinopyroxene it is apparent that this recrystallization took place after the development of the symplectites. Since M-2 was the last phase of stable clinopyroxene, exsolution of the pyroxenes must have taken place prior to M-2.

Samples 76721 and 76721-A (77-Gal-16/17, Fig. 2) are massive retrograded eclogites, with blastomylonitic textures along narrow zones. Both samples consist of porphyroclasts of garnet, clinopyroxene and amphibole with some interstitial quartz. The clinopyroxene is exsolved both into cpx-plag symplectites and into clinopyroxene with vermicular plagioclase inclusions. Pargasitic hornblende partially replaces the clinopyroxene. Small pargasite grains along the rims of symplectites have formed after the exsolution, since vermicular plagioclase is continuous from the symplectite into the amphibole.

Sample C2-143 (77-Gal-16/17, Fig. 2) is also a massive retrograded eclogite. Most of the original clincpyroxene is replaced by cpx-plag symplectites and, to a lesser extent, by symplectites of brown hornblende and plagioclase. The latter have also replaced locally the cpx-plag symplectites.

In view of the growth of pargasitic hornblende during M-2, it can be inferred from the last three samples that the cpx-plag symplectites developed before M-2.

III.2.4. Timing of the clinopyroxene exsolution If the cpx-plag symplectites developed between M-1 and M-2, both the eclogite and the granofels parageneses could still have formed during M-1. Since no difference in T or P can be postulated for the scattered occurrences south of Teijeiro, it has to be assumed then that small differences in chemical composition led to the different parageneses during M-1. The eclogites and granofelses of the Cabo Ortegal area, however, show only slight differences in MgO/FeO (van Calste- ren, 1978). The major element compositions of the eclogites and granofelses from the Teijeiro area are also indistinguishable (chapter III.3.).

Contrary to the Mellid area (Hubregtse, 1973a), only cataclastic deformation without major recrystallization occurred during F-2 in the mafic rocks near Teijeiro and in the Cabo Ortegal complex (Engels, 1972). Therefore, exsolution of pyroxene due to an increase in T between M-1 and M-2 is not a likely process.

Exsolution due to a decrease in P is also unlikely, as can be inferred from the stability curves for omphacite coexisting with quartz (Fig. 16, Kushiro, 1969).

- 53 - Fig. 16. Stability curves for jadeite in omphaoite coexisting with quartz (Kushiro, 1969) and curves for the Fe-Mg partition coefficients for coexisting garnet and clinopyroxene (Ratieim & Green, 1974). Vertical shading, estimated PT-field for the Cabo Ortegal eclogites. Horizontal shading, estimated PT- field for eclogites from the blastomylonitic graben (van der Wegen, 1978). Arrow, proposed PT-puth prior to M-l.

Unexsolved omphacites from the Cabo Ortegal eclogites contain 30-33% jadeite, while the jadeite content of the symplectitic omphacites is 18-23% (Vogel, 1967). The jadeite contents of the clinopyroxenes from the Teijeiro area (chapter III.3.) and the Mellid complex (Hubregtse, 1973b) are generally lower. In view of the association with plagioclase, the jadeite contents of the clinopyroxenes from the latter two complexes cannot be used for P estimates, however (Kushiros 1969). In order to lower the jadeite content at constant T from about 30% to about 20%, as is the case for the exsolved omphacites of Cabo Ortegal, a decrease in P of about 8 kb is necessary (Fig. 16). Such a large decrease in P is not possible in view of the similarity between the mineral parageneses developed during M-l and M-2 (chapter III.2.1.). Any decrease in Ptotal was much t0° small to cause development of low—pressure granulite facies assemblages (orthopyroxene-plagioclase subfacies).

If no major exsolution of omphacitic clinopyroxene between M-l and M-2 did occur, it must be inferred that the cpx-plag symplectites developed during M-l. The eclogite facies paragenesis then represents an older event (M-0). In view of the preferred orientation of the Cabo Ortegal omphacites (Engels, 1972), this phase must have been accompanied by a deformation phase (F-0).

III.2.5. P-T conditions For the P-T conditions during the formation of the eclogites, estimates of 630- 900°C and 14-19 kb (Vogel, 1967) and 700-75O°C and 11-13 kb (Engels, 1972) have been made. On the basis of the curves for the Fe-Mg partition coefficients for coexisting garnet and clinopyroxene (Raheim & Green, 1974) and the curves for the jadeite content of omphacite (Kushiro, 1969) (Fig. 16), however, the conditions of the eclogite facies metamorphism can be estimated at 580-640°C and 10-11 kb.

- 54 - Maaskant (1970) estimated the PT conditions for subsequent phases of recrystallization of the ultramafic rocks in the catazonal complexes. He arrived at a temperature of 800-900°C and a pressure of 10-15 kb for the phase of catazonal recrystallization, which is taken to represent the conditions of emplacement of the ultramafics in the lower crust where eclogite facies conditions prevailed. When it is postulated that a T-increase due to cae intrusion of the ultramafic bodies caused the development of the granulite facies assemblages (van Calsteren, 1977), a temperature of about 850°C can be estimated for this "contact metamorphism". As can be seen from Fig. 16, an increase of 200-250°C under a constant pressure of 10-11 kb will result in omphacites containing about 25% jadeite. This jadeite content is consistent with the composition of the exsolved omphacites of the Cabo Or- tegal eclogites. However, in addition to the symplectitization of the omphacites, also the granulite facies paragenesis (clinopyroxene + garnet + plagioclase) must have formed from the eclogites under these conditions. This involves exsolution of a jadeite component, leading to the formation of interstitial plagioclase without development of symplectitic textures. A similar exsolution without symplectite formation has been recorded from meta-anorthosites, leading to corona structures with diameters of 3-5 cm (Griffin, 1972; Griffin & Heier, 1973; Mysen & Griffin, 1973).

Conditions of the high-pressure granulite facies were maintained during M-2, although the prevailing T and F became somewhat lower towards the end of this phase (700-750°C and 8-12 kb, Maaskant, 1970, or 6OO-7OO°C and 8-10 kb, Hubregtse, 1973b). A slight increase in Pfiuid resulted in the growth of hornblende. The origin of these fluid phases remains a matter of speculation. A derivation from the country rocks is not very likely in view of the dry nature of the paragneisses. Taking into account the late-stage pargasite-edenite (± phlogopite) veins and lenses in the Cabo Ortegal lherzolites, the volatiles could have been derived from the mantle (Bailey, 1978).

III.3. INVESTIGATED SAMPLES

III.3.1. Samples All samples from the Sobrado/Teijeiro complex which were used for U-Pb whole-rock investigation come from the northern part of the complex. The sample sites are shown in Fig. 2.

Sample 77-Gal-9 is a foliated, banded mafic granulite with a rather high proportion of B-zoisite. The mineral composition and texture are typical for the M-2 metamorphism. Samples 77-Gal-10 and Cl-13 are massive clinopyroxene-garnet granofelses (M-l), the former containing minor (3-zoisite. Sample C2-29 is a massive metagabbro with relict ophitic texture, showing garnet coronas around polycrystalline amphibole-clinopyroxene aggregates. The rock is typical for the pre-M-2 emplaced metagabbros (chapter III.2.1.). The first three samples come from the same outcrop, some 1.5 km south of Teijeiro, while sample C2-29 is from an outcrop some 100 m to the east.

Samples 77~Gal-16 and 17 are massive eclogites (M-0) with minor quartz and a-zoisite, respectively. Sample C2-4 is a massive clinopyroxene-garnet granofels (M-l) with accessory ot-zoisite. Sample C2-143 is a massive eclogite (M-0) with minor a- zoisite. All four samples come from the same outcrop, some 4 km south of Teijeiro.

Sample ?7-Gal-18, a migmatic hornblende granulite (M-2, chapter III.2.1.), was derived from a small quarry about 8 km south of Teijeiro.

The three samples of Cabo Ortegal eclogite (73-Heb-38, 39 and 40) were taken from the collection of Dr. P.W.C. van Calsteren; the samples were already analyzed for Rb-Sr (van Calsteren et al., 1979). They come from a large quarry on the Monte

- 55 - I Ln

Table 8: Whols-vock major element compositions

77-Gal-9 77-Gal-10 77-Gal-16 77-Gal-17 77-Gal-l8 C2-4 Cl-13 C2-29 C2-143 44.96 45.95 44.29 47.89 61.18 46.20 47.39 46.41 45.30 TiOz 0.31 0.36 1.17 0.70 0.69 1.75 0.45 0.60 0.25 A12O3 17.41 15.86 16.42 18.05 15.09 14.56 17.65 18.41 18.23 Fe2O3 1.04 1.54 2.99 1.00 1.67 1.82 0.62 1.26 1.87 FeO 6.63 8.19 14.90 10.63 • 6.85 10.66 9.98 6.22 8.95 MnO 0.14 0.18 0.27 0.20 0.17 0.20 0.19 0.13 0.18 MgO 10.67 10.06 6.96 7.81 3.35 9.38 8.79 12.44 9.82 CaO 14.78 15.64 11.93 12.36 7.73 13.68 12.93 9.41 13,94 Naz0 0.65 0.95 0.94 0.42 2.69 1.50 1.10 1.88 0.59 K2O 0.01 0.0 0.01 0.03 0.04 0.06 0.00 0.20 0.07 P2OS 0.06 0.0 0.04 0.01 0.14 0.13 0.03 0.10 0.0 + H20 ~ 2.82 1.64 0.83 1.18 0.56 0.56 0.69 2.80 1.17 CO2 0.44 0.37 0.10 0.10 0.10 0.10 0.10 0.10 0.05 Total 99.92 100.74 100.05 100.38 100.26 100.60 99.92 99.96 100.42

Niggli values si 91.6 91.6 91.0 104.7 191.6 94.0 100. 97.3 90.7 al 20.9 18.6 59.9 23.3 27.9 17.5 21.9 22.8 21.5 fm 45.5 46.1 52.0 46.9 38.0 49.7 46.6 52.0 47.4 c 32.3 33.4 26.3 28.9 25.9 29.8 29.2 21.1 29.9 alk 1.3 1.8 1.9 0.9 8.3 3.0 2.3 4.1 1.1 mg 0.7 0.7 0.4 0.5 0.4 0.6 0.6 0.8 0.6 Catanormative mine- xl composition plag 51.8 48.1 50.7 52.6 54.8 47.1 53.8 59.6 53.0 qz 1.5 19 .7 cpx 22.5 30.3 11.8 9.5 5 .1 22.7 15.5 2.3 17.1 opx 8.4 1.9 18.3 33.9 16 .9 5.5 19.0 14.4 14.8 ol 15.5 17.4 13.3 18.8 10.0 21.0 12.8 ace 1.9 2.4 5.8 2.6 3 .5 5.9 1.7 2.8 2.5 Castrillon, near Carino. The geological setting and the petrography have been described by Vogel (1967).

III.3,2, Whole-rock compositions Major element compositions are presented in Table 8. In Fig. 17 the Niggli values al, fm, c and alk are plotted against si; the fields in which Cabo Ortegal mafic rocks plot are also given (van Calsteren, 1978). The distinction between eclogites and granofelses on the one hand and amphibolites and metagabbros on the other hand (van Calsteren, 1978) on the basis of their si values and catanormative mineral compositions, is not supported by the rocks of the Teijeiro area.

' -- w

Fig. 17. Niggli values al, fm, o and alk versus si. Symbols: «3 granofels/grcmu- lite; O, eologite; m3 metagabbro. Dashed contours define the areas for the Cabo Ortegal eclogites/granofelses (right) and amphibolites/metagabbros V fc'0J"t/ •

The si values of the eclogites and granofelses and the metagabbro are within the range of si values of Cabo Ortegal amphibolites and metagabbros. The alk values are lower and the c values are higher than those of the Cabo Ortegal rocks. With the exception of one eclogite and one (migmatic) granulite, the samples of the Teijeiro area are olivine-normative, as are the Cabo Ortegal amphibolites and metagabbros (van Calsteren, 1978).

In Fig. 18 the weight percentages (FeO + Fe2O3 as FeO), MgO and (Na2O + K20) are plotted in the alkali-iron-magnesium diagram, along with the same data of the Cabo Ortegal mafic rocks (van Calsteren, 1978) and of the high-grade mafic rocks of the Mellid complex (Hubregtse, 1973a). In comparison with the Cabo Ortegal

- 57 - eclogites and granofelses, the same rocks of the Teijeiro area are lower in alkalies and somewhat higher in MgO. A subdivision on the basis of their MgO/FeO ratio, as has been inferred for the Cabo Ortega! eclogite and granofelses (van Cal- steren, 1978), cannot be made. The data of the granulites and the granofelses of the Mellid complex also do not support such a subdivision. The data of mafic rocks of the Cabo Qrtegal complex reported by Vogel (1967; not plotted here) fall in the same fields as the data of Fig. 18.

18. AFM diagram for mafic rocks from the Teigeiro area. Symbols as in Fig. 17. Bashed contours give the areas for the Cabo Ortegal eclogites (E)3 granulites (G) and metagabbros (short dashes). Full contours define the area for the Mellid granulites. Open triangle is the average of the Mellid metagabbros.

The above data indicate that the eclogites and granofelses of the Teijeiro area have major element compositions comparable to those of the high-grade mafic rocks of the Cabo Ortegal and Mellid complexes, although they are somewhat higher in MgO and lower in alkalies. The similarity in composition suggests that the eclogite-granulite transformation was an isochemical process, at least as far as the major elemens are concerned.

III.3.3. Mineral compositions A limited number of garnets, clinopyroxenes and amphiboles from granofelses (C2-4 and Cl-13), a metagabbro (C2-29) and an eclogite (C2-'.43) have been analyzed on their major element composition.

Clinopyroxene analyses are presented in Table 9. Point analyses using the energy dispersive system (not given here) did not reveal differences in composition between cores and rims of the clinopyroxenes. All clinopyroxenes are sodic augites and are comparable to the clinopyroxenes from the Mellid complex (Hubregtse, 1973b). Although the augite contents are higher, the clinopyroxenes of the Teijeiro area have a jadeite/acmite ratio larger than one, similar to the omphacites of the Cabo Ortegal eclogites (Vogel, 1967).

- 58 - ,^,1/.*T

Table 9. Clinopyroxene compositions* Table 10. Brown hornblende compositions* Table 11. Garnet eompositions*

C2-4 CI-13 C2-29 C2-143 C274 CI-13 C2-29 C2-143 C2-4 Cl-13 C2-29 C2-143

SiO2 52.85 51 .13 54.23 52.12 SiO2 42 .33 43 .80 45 .68 44 .18 SiO2 39 .92 40. 70 41.40 40.95 AUOa 7 .55 7.87 4.97 5.40 AI2O3 15.45 14.60 14.07 13.98 A12O3 22 .20 22. 83 23.38 23.03 TiO2 0.40 0.60 0 .25 0.27 T1O2 1.25 I.55 2.20 0.27 T1O2 0.05 0.05 FeO 5.50 5.15 2.75 4.83 FeO 10.68 11.52 5.47 10,.95 FeO 21,.63 16.98 14.07 16.10 MnO 0.02 0.02 0.03 MnO 0.0G 0.10 0.07 0.13 MnO 0..40 0.35 0.27 0.25 MgO 1 1.25 12.18 14.48 13.40 MgO 12.53 12.20 16.13 13.27 MgO 8..12 10. 10 14.27 10.45 CaO 19.70 21 .70 21 .62 23 .17 CaO 11 .47 11 .85 1).17 12.78 CaO 9..08 10. 80 7.93 10.83

NaZO 2.70 1.52 I.78 0.98 Na20 2.47 1.85 1.68 1.67 K20 K2O 0.55 0.05 0.75 0 .17 Tot. 99.,95 100,.17 100,.09 100.2 Tot. 96 .81 97 .52 97 .22 47.40 Tot. 101,.35 101. 81 101.37 101.61 Cations per 6 0 's Cations per 22 0';3 Cations per 24 O's! Si 1,.925 1,.867 1,.955 1,.910 Si 6,.220 6.372 6..479 6..417 Si 6..014 5.996 5.9B7 6.013 Allv 0..075 0,.133 0..045 0..090 Al 2 .675 2.500 2,.352 2..403 Al 3..938 3.961 3.986 3.984 A1VI 0.,249 0..206 0..166 0,.136 Ti 0., 139 0 .170 0,.231 0..031 Ti 0.004 0.003 Ti 0.Oil 0.,017 0..008 0..008 Fe I..314 I.398 0,.646 1..333 Fe 2..725 2.093 1.700 1.977 Fe3+ 0.016 0..034 0,.007 0,.023 Mn 0..009 0 .009 0,.006 0,.014 Mn 0.,053 0.044 0.032 0.031 + Fe 0. 153 0.,123 0.,076 0..125 Mg 2,.741 2 .648 3,.408 2..879 Mg 1.,822 2.217 3.075 2.288 Mg 0.611 0.,663 0.,778 0..732 Ca 1 .80, 4 1..844 1,.693 1..995 Ca 1,,467 1.704 1.137 1.703 Ca 0.768 0..849 0.,835 0.,909 Na 0,.700 0..524 0..463 0.All Na 0. 190 0.,107 0. 125 0.,069 K 0..103 0.,009 0..136 0.,029 Tot.- 3.998 3.999 3.995 4..002 Tot. 15.,705 15.,474 15..414 15. 573 Tot. 16.019 16.019 15.920 15.996 Endmembevs** Site occupancies** Endmembers Ti-Di 1.1 1.7 0.8 0.8 T Si 6.22 6..37 6..48 6.,42 Aim 44.92 34. 55 28.15 32.95 Ts 5.3 9.9 3.0 7.3 Al 1.78 1,.63 1.,52 1.58 Pyr 30.03 36.60 50.92 38.14 Ac 1.6 3.4 0.7 2.3 C Al 0.90 0.,87 0.83 0.82 Spess 0.87 0.73 0.53 0.52 Jd 17.4 7.3 1 1.8 4.6 Ti 0. 14 0. 17 0.23 0.03 Gross 24. 18 28. 13 20.41 28.40 Di+Hd 70.4 73. 3 79.8 82. 8 Mg 2.74 2.,65 3.41 2.88 En+Fs 3.0 2.7 2.9 1.5 Fe 1.22 1.31 0.52 26 * The figures represent the average B Fe 0. 09 0.09 0.12 0.07 values of El'MA analyses on four Mu 0. 01 0.01 0.01 0.01 * The figures represent the average grains. Ca 80 84 j p 69 2. 00 values of EPMA analyses of three Na 0. 10 0.06 0. 17 0.02 grains. A Na 0.60 0.46 0.29 0.45 ** Endmembers calculated according; to K 0. 10 0.01 0. 14 0.03 Mysen & Griffin (1973). Ti-Di: Ti- Diopside; Ts: Ca-Tschermak's mole- The figures represent the average cule; Ac: Acmite; Jd: Jadeite; Di: values of EPMA analyses of three Diopside; Hd: Hedenbergite; En: grains. Enstatite; Fs: Ferrosilite. in Site occupancies calculated ac- ID cording to Leake (1978). C2-4: Ferroan Pargasite; Cl-13, C2-29, C2-U3: Tschermakitic hornblende.

c- Bram hornblende analyses are presented in Table 10. Like the clinopyroxenes, the amphiboles do not show zoning. According to the amphibole nomenclature (Leake, 1978), the investigated amphiboles range from ferroan pargasite through ferroan pargasitic hornblende and edenitic hornblende to magnesio and tschermakitic hornblende. In comparison with the ferroan pargasites to edenitic hornblendes from the Mellid granulites (Hubregtse, 1973b), the brown hornblendes of the Teijeiro area are slightly lower in alkalies.

Garnet analyses are presented in Table 11. Differences between cores and rims, if any, are smaller than the differences between the various grains in one sample. Garnets of granofelses of the Teijeiro area are richer in pyrope than the granulite- garnets from Cabo Ortegal and Mellid (Vogel, 1967; Hubregtse, 1973b).

The compositions of the clinopyroxenes and garnets in the eclogites are thus similar to those in the granulites. This supports the inferred prograde eclogite- granulite transformation.

HI.4. U-Pb WHOLE-ROCK DATA AND DISCUSSION

Whole-rock samples have been analyzed of two eclogites, two granofelses and one (migmatic) granulite from the Teijeiro area and of three eclogites from the Cabo Ortegal complex. The U-Pb data are presented in Table 12 and plotted in Figs. 19 to 22. The analytical procedures are given in the Appendix.

In comparison with MOR basalts (Church & Tatsumoto, 1975; McDougall, 1977), the U and Pb contents of the Galician mafic rocks are rather high. The Pb isotopic ratios are more radiogenic and the y values are lower. The same holds when the data are compared with those of basalts from Iceland and the Reykjanes Ridge (Sun et al., 1975). In comparison with the scarce data of continental plateau basalts, the U and Pb contents are rather low (Upuluri, 1974; Wedepohl, 1969). Both are comparable to those of B-type eclogites (Wedepohl, 1969).

In the diagrams of 20BPb/20"Pb versus 238U/20*Pb (Fig. 19) and 207Pb/2("*Pb versus 23SU/Z01*Pb (Fig. 20), the data-points do not show a good linear correlation. Ages calculated from the slopes of best-fit lines range from 2.0 Ga when the two points highest in radiogenic Pb are included, to negative when these points are omitted.

In the lead-lead diagram (Z06Pb/201tPb versus 207Pb/20')Pb, Fig. 21), the slope of the best-fit line depends strongly on the position of the two highest data-points (Gal-10 and Gal-16). The line through all data-points intersects the mantle growth curve of Stacey & Kramers (1975) at around 3600 Ma. Although this growth curve is actually a two-stage curve, it will be regarded here as a single-stage model in view of the relatively small change in ]i value (from 7.19 to 9.74) between the two stages.

In the modified concordia diagram((Z07Pb/Z35U)* versus (206Pb/238U)*, Fig. 22) the data-points show a large spread in Pb/U and are fairly good linearly correlated. A line through all data-points intersects concordia at 4300 ± 30 Ma and 450 i 200 Ma. (It should be noted that the error limits are derived from the scatter of the data-points and not from the analytical uncertainties, see the Appendix). When the two points above the line (77-Gal-10 and 16) are omitted, the upper intercept is slightly raised to 4360+ 10 Ma and the lower intercept becomes 382 + 45 Ma. The Cabo Ortegal eclogites define a line which intersects concordia at 4366 ± 20 Ma and 464±130 Ma.

A comparison between the mafic rocks from the Teijeiro area and the Cabo Ortegal eclogites reveals a close resemblance of the U-Pb data of both suites of samples. The slightly lower U and Pb contents of the Teijeiro samples can be regarded as reflecting the slight differences in chemical composition. The lead isotopic ratios

- 60 - Table 12: Whole-rook U-Pb data of eologite and granulite faoiea rooks*

20l7 238 207p pb 2C">pb 206 Pb 2 0 7pb U 20 207pb # 206pb b Sample Pb U 2os 20t6 20 201) 20 20,p 23 23 2 , 20 No. pb pb «pb Pb *Pb b «u 5u (yg/g) (yg/g) 0 pb *Pb

77-Gal-9 1.998 0.8138 0.0515 19.44 15.81 7.93 1.28 95.8 1.2 0.15 18.8 15.8 77-Gal-10 1.933 0.7845 0.0479 21.10 16.47 5.50 2.14 159. 0.57 0.05 20.7 16.4 77-Gal-16 1.737 0.7025 0.0395 26.77 18.25 27.7 0.632 39.7 0.21 0.07 24.6 18.1 77-Gal-17 2.034 0.8271 0.0522 19.16 15.84 4.69 2.10 163 1.1 0.08 18.8 15.8 77-Gal-18 1.926 0.7837 0.0498 20.12 15.75 25.5 0.425 29.5 2.0 0.77 18.2 15.6 73-Heb-38 2.191 0.7965 0.0501 20.01 15.91 13.9 0.770 55.7 1.6 0.33 18.9 15.9 73-Heb-39 2.100 . o793. 1 0.0497 20.17 15.97 18.8 0.577 41.5 1.3 0.34 18.7 15.9 73 Heb-40 2.071 0.8315 0.0525 19.05 15.84 4.98 1.96 154. 3.2 0.24 18.7 15.8

* Columns 2 through 4: measured 208Pb/206Pb, Z07Pb/206Pb and 201(Pb/206Pb ratios. Columns 5 through 9: 206Pb/201*Pb, 207Pb/20*Pb, 239U/2OltPb, (206Pb/238U)* and (207Pb/23SU)* ratios corrected for procedure blanc, see the Appendix. Columns 10 and li: total lead and uranium contents. Columns 12 and 13: age-corrected 206Pb/2D4Pb and 207Pb/201tPPbb ratios for T = 480 Ma, using the present-day 238U/2OltPb ratios. (zo6pb/238D)* and (Z"7pb/23su)*j ratios corrected for primordial lead.

I I

CATAZONAL MAFIC ROCKS, GALICIA •'.fl 238,j-206pb ISOCHRON DIAGRAM

30-1

• TEMEIRO --- • CABO ORTEGAL

10 20 30

Fig. 19. Data of catazonal mafia rooks from western Galiaia plotted in a Z3BU- Fb iaoahron diagram. The line is the 500 Ma reference isoahron for mantle-derived lead (Staaey & Kramers* 197S).

CATAZONAL MAFIC ROCKS, GALICIA -& 23Su-207Pb ISOCHRON DIAGRAM 30 J

• TEUEIRO • CABO ORTEGAL

0.1 0.2 0.3

Fig. 20. Data of aatazonal mafia rocks from western Galiaia plotted in a 23SU-2 7Pb isoahron diagram. The line is the 500 Ma reference isoahron for mantle-derived lead (Staaey & Kramers, 1975).

- 62 - —I

CATAZONAL MAFIC ROCKS, GALICIA LEAD-LEAD DIAGRAM

• TEUEIRO • CABO ORTEGAL * RECALCULATED TO 480 Ma

10 20 25

Fig. 21. Lead-Lead diagram of oatazonal mafic rooks from western Galicia. Growth curve taken from Staoey & Kramers (19?S). Best-fit line: y = 0.32X + 9.51. 77-Gal-16 and 10 are the data-points with the highest Z06Pb/Z0*Pb ratios.

* CATAZONAL MAFIC ROCKS, GALICIA i MODIFIED CONCORDIA 2.0 4

UNCERTAINTY ELLIPSES

• TEUEIRO • CABO ORTEGAL

SO 100 150

Fig. 22. Modified oonaordia diagram of aatazonal mafia rooks from western Gali- cia. Best-fit line: y = 0.013X + 0.065. The large and small uncertainty ellipses shown are for the samples with the highest and the lowest (Z07Pb/ 2*SU)* ratio, respectively. 77-Gal-lO and 16 are the data-points above the line.

- 63 - fall within the same range. In view of these similarities, a significant difference between the initial lead compositions of the rocks appears to be unlikely. Both suites of samples will therefore be treated together.

Interpretation of the U-Pb systematias The whole-rock U-Pb evolution has been discussed by, e.g., Gale & Mussett (1973). In the lead-lead diagram the linear array of the data-points of a two-stage system passes through the upper intersection of the yi growth curve with the "zero isochron" and the point on the \i\ growth curve corresponding to the age of the beginning of the second stage. In the modified concordia diagram the linear array of a two-stage system passes through the point on concordia representing the age of the Earth (4.57 Ga) and the point corresponding to the age of the beginning of the second stage. For a lead evolution consisting of more than two stages, the intersections of the best-fit line with the m growth curve have no geochronological meaning in the lead-lead diagram; in the modified concordia diagram the upper intercept has likewise no geochronological meaning, but the lower intercept should correspond with the age of the beginning of the last stage in the U-Pb evolution, provided that the U-Pb fractionation between the first stages was constant for all subsystems. For the Galician samples, the lower intercepts of the best-fit line in both the lead-lead and the modified concordia diagrams do not correspond to the same age (3600 Ma versus 450 Ma), so we are dealing here with a U-Pb evolution consisting of at least three stages. Although the lower intercept in the modified concordia diagram (Fig. 22) has a large uncertainty, an Early Paleozoic age is in agreement with the ages of the high-grade metamorphism. When the lead isotopic ratios are recalculated to 480 Ma, using the present-day y-values (Table 12), the age-corrected data-points still reflect an evolution of at least three stages (Fig. 21). The present-day U-Pb systems possess thus at least a four-stage history.

The absence of a linear correlation in the isochron diagrams of Figs. 19 and 20 could be due to (sub)recent uranium migrations. In the. lead-lead and modified concordia diagrams (sub)recent uranium migrations would not give rise to scatter of the data-points. When the three most deviating points in the 238u-206Pb isochron diagram (Fig. 19) are omitted (Gal-10, 16 and 18), the remaining five points produce a line with a slope corresponding to an age of around 500 Ma. If it is assumed that the three omitted points fall off the line due to (sub)recent uranium migration, losses of 81% and 71% are needed for Gal-16 and Gal-10, respectively, while the U content of Gal-18 should have increased by 50%. In order to give a linear correlation in the Z35U-207Pb isochron diagram (Fig. 20), the uranium losses need to be as high as about 90%. A recent oxygen and hydrogen isotopic study, however, revealed a closed-system behaviour for these isotopes in alternating bands of eclogite and paragneiss during and after the high-grade metamorphism, indicating that element migration between these rock units was very limited Forester & Javoy, 1979).

Another possible cause of disturbance of the U-Pb systematics is the apparent influx of volatiles during the M-2 phase of granulite facies metamorphism. This influx may have given rise to addition of uranium and lead, and to a mixing of the lead either with less radiogenic mantle lead or more radiogenic lead from the en- vironmental paragneisses. The oxygen isotopic data are inconclusive regarding element migration between the mafic and ultramafic rock units, but seem to preclude an influx of volatiles derived from the paragneisses (Forester & Javoy, 1979). As the most radiogenic lead is found in rocks with the lowest U and Pb contents, however, a mixing with less radiogenic mantle lead can be excluded. Mixing with more radiogenic lead from the paragneisses can also be excluded, as the lowest •Z06Pb/20''Pb ratio is found in the sample with the strongest M-2 influence (Gal-18> and vica versa (Gal-16).

The scatter in the isochron diagrams should therefore mainly be explained in terms of varying U-Pb fractionation factors prior to the last stage of the Pb evolution

- 64 - (Gale & Mussett, 1973), The subsystems should thus have possessed different Pb isotopic compositions at the beginning of the last stage.

If the scatter of the data is meaningful and not results from recent uranium migration or mixing with exotic lead, a model can be constructed for an evolution with changes in the U/Pb ratios at discrete times. As has been shown, the U-Pb systems underwent at least a four stage evolution. Taking into account the geological history of the rocks (chapter III.3.), a number of events can be envisaged that could be responsible for U-Pb fractionation. A first event of fractionation may be related to the generation of the basaltic liquids from the mantle. The fractionation factor (fn = yn/yn+1) will be constant if the source area"was small with respect to mantle inhomogeneities, and if the mafic (in)extrusives were derived from one or several closely related magmas. The prograde raetamorphism up into the eclogite facies may have given rise to a second event of U-Pb fractionation, due to uranium loss during dehydration prior to the peak of the eclogite facies raetamorphism. U/Pb ratios are indeed reported to have been lowered in the deep continental crust either by events of partial melting, or by dehydration during prograde metamorphism (Lambert & Heier, 1968; Heier, 1973; Gray. 1977; Sighinolfi & Gorgoni, 1978). The fractionation factor for such an event is not necessarily constant for all subsystems. Finally, a last event of U-Pb fractionation may be related to the development of metamorphic banding during the granulite facies metamorphism (M-l/ M-2).

The equations governing the U-Pb evolution for multi-stage models can be derived from Gale & Mussett ,(.1973). In the modified concordia diagram the equation of a four-stage model is:

(2Q7pb/23Su) in which e ^P"*-; e ») t f (e AX?-. , f< 2 - e 2) + (e f2(eA'Ti - (eVT2 -e X'T3) and '. . Zz-fc W2/U3-

To maintain ELnear relationships for the several subsystems during such a multi- stage evolution, both i\ and f2 have tQ,,be constant for all subsystems; in that case A g-:fcves the slope of the linear array (Gale & Mussett, 1973). Any departure from this condition will result in a scatter p£ the data-points.

In view of the large number of unknown parameters, speculations on the ages of the fractionating events can only be made by assuming boundary-values. The following assumptions are made: 1. A first stage in the mantle beginning at T = 4570 Ma. The Pi value may be estimated at about 8.5, intermediate between the values of the two stages of Stacey & Kramers (1975). 2. A second (crustal) stage, beginning at Ti with the generation and (in)extrusion of ^basaltic magmas. As the basaltic rocks are intercalated with the metasedi- raents (Vogel, 1967), it may be inferred that the maximum for Ti equals the maximum sedimentation age, i.e. 1.5 Ga (chapter II.6.). For p2 a value around 18-20 can be taken (Doe, 1970).

3. A third (lower crustal) stage, beginning at T2 with the uranium loss in relation to the eclogite facies metamorphism (M-0). For the pa value it is assumed that the present-day subsystems were derived from uranium redistribution within the overall system. The average present-day y value of the eight investigated samples is 13.6. 4. A fourth stage, starting at T3 with uranium redistribution during the granulite facies metamorphism (M-l/M-2). T3 can be taken between 480 Ma and 350 Ma.

- 65 - On the basis of the estimated values for \i\, y2 and \i$ mentioned above, several T)~T2 pairs can be calculated from the equation for the slope of the line in the modified concordia diagram. When an age of 1,5 Ga is taken for Ti, T2 will be between 1.0 and 0.5 Ga, assuming that U2 is between 15 and 23 and ]i3 about 15. A Ti value of less than 1.0 Ga would imply unrealistically high U2 values: at least 35. It can thus be concluded that the emplacement of the basaltic magmas must have taken place at least 1.0 Ga ago.

The following conclusions can thus be drawn from the U-Pb whole-rock systematics: 1, An explanation of the scatter of the data-points in the isochron diagrams by migration of large proportions of the uranium is not feasible. The scatter is better explained by a variable isotopic composition of the lead at the beginning of the last stage of the U-Pb evolution (T3), 2, Linear arrays in the lead-lead and modified concordia diagrams are interpreted in terms of a prolonged crustal history of the high-grade mafic rocks. Mixing with exotic lead is unlikely. 3, The U-Pb systematics are best explained by successive events of U-Pb fractionation, probably a four stage evolution. The U-Pb fractionation factor for the first fractionating event, fi, has probably been more or less constant through the rocks, but there may have been a small spread in the £2 value. A. Regarding the timing of the various fractionating events according to a four stage model, there are several possibilities. The emplacement of the mafic rocks (Ti) must anyhow have taken place before 1.0 Ga ago, while the maximum age is 1.5 Ga. Assuming an age of 1.5 Ga for the emplacement, the eclogite facies metamorphism (T2) should have taken place between 1.0 and 0.5 Ga ago. Several other T1-T2 pairs can also be calculated, however.

III.5. CONCLUSIONS

Based on petrographical, geochemical and isotopic data, the following geological history is inferred for the catazonal mafic rocks of western Galicia (Table 1): 1. The catazonal mafic rocks were emplaced as high-level intrusions or as extrusive rocks. In view of their resemblance to continental plateau basalts, an emplacement during a phase of continental rifting has been proposed (van Calsteren, 1978). The whole-rock U-Pb systematics and the geological relationships suggest that this event has taken place between 1.5 and 1.0 Ga ago. 2. During subsidence, the rocks were subjected to prograde metamorphism up into the eclogite facies. For this metamorphism (M-0) P-T conditions of around 600°C and 10-11 kb can be estimated. The whole-rock U-Pb systematics suggest that this metamorphism has probably taken place between 1.0 and 0.5 Ga ago. 3. Around 480 Ma ago the eclogite facies rocks were intruded by lherzolite diapirs (van Calsteren et al., 1979). As a result the temperature in the country rocks increased to the P-T conditions of the high-P granulite facies (M-l), estimated at about 850°C and 10-11 kb. The eclogites were isochemi- cally metamorphosed to clinopyroxene-garnet granofelses. Locally (Cabo Orte- gal, Teijeiro), the transformation was incomplete and eclogites with exsolved omphacites or rocks with transitional mineral assemblages were preserved. K-Ar mineral dates of about 400 Ma reflect the end of M-l (van Cal- steren et al., 1979). 4. An increase in pfluid resulted in a development of hornblende-granulite facies parageneses and locally in migmatization (M-2). P-T conditions at the end of this phase are estimated at 600-700°C and 8-10 kb (Hubregtse, 1973b) or 700-750°C and 8-12 kb (Maaskant, 1970). The Rb-Sr whole-rock isochron age of 347 +17 Ma of migmatic granulites from the Cabo Ortegal complex is interpreted as recording the end of this phase of metamorphism (van Calsteren et al., 1979). 5. Finally, the rocks were metamorphosed under amphibolite (M-3) and greenschist facies conditions (M-4). These phases have been dated at about 310 to 280 Ma ago (Priem et al., 1970; van Calsteren et al., 1979).

- 66 - "-•/

CHAPTER IV

THE OROGENIC EVOLUTION OF THE HESPERIAN MASSIF

IV.1. INTRODUCTION

The geological evolution of the Hesperian Massif, which takes up the western half of the Iberian Peninsula, is frequently explained by plate collision in Late Pale- ozoic time, i.e. the "Variscan Orogeny" (e.g. Bernard & Soler, 1974; Martinez- Garcia et al., 1975; Vegas, 1978). Older folding phases affecting Lower Paleozoic rocks, with or without accompanying metamorphism (Martinez-Garcia, 1973; Aldaya et al., 1973, 1976), are not considered in these models. Extensional tectonic regimes are either ignored (Bernard & Soler, 1974) or placed in a back-arc setting (Vegas & Munoz, 1976). It has been pointed out by Van Calsteren (1977) and Den Tex (1979) that collision models do not adequately explain the orogenic evolution of western Galicia. Van Calsteren (1977) proposed a mantle-plume model for the Early Paleo- zoic history. An extended mantle-plume model, including also the Late Paleozoic events, will be discussed by Den Tex (1980). In this thesis an attempt will be made to test the model against the available geological data of the entire Hes- perian Massif.

IV.2. AN OUTLINE OF THE GEOLOGICAL HISTORY OF THE HESPERIAN MASSIF

The Hesperian Massif is characterized by an abundance of Upper Precambrian to Up- per Paleozoic supracrustal rocks, Paleozoic granites and Paleozoic migmatites. The Massif is essentially composed of a central zone displaying-the highest grade of metamorphism, flanked by zones in which the metamorphic grade gradually dimin- ishes outwards (Fig. 23). In the northern part of the central zone the occurrence of catazonal complexes, juxtaposed to non- or slightly metamorphic supracrustal rocks, reflects the importance of differential vertical movements. From geochemical and geochronological data it is evident that most or all of the granitic rocks were derived from older crustal rocks. The amount of pristine materials is small in comparison with the amount of reworks.' crustal materials. In the following, the geological history of the Hesperian Massif will be summarized.

IV.2.1. Supracrustal history On the basis of local stratigraphic successions (Fig. 24) a tentative stratigraphic column has been made for the pre-Lower Devonian sediments (Fig. 25). The names of the formations and series referred to in the text are given in Figs. 24 and 25. The stratigraphic positions and the ages of the sequences below the Lower Cambrian limestone formations (e.g. the Candana Formation) and the Cambrian (?) volcanodetrital formations (e.g. the Olio de Sapo Formation) are not known with certainty. A discussion of the current controversies about the ages of the formations is beyond the scope of this study; at this stage the ages are not essen- tial for the development of the model, however.

In the following the major units distinguished in the Hesperian Massif are brief- ly described. They are informally named after locations where complete sequences occur. The units are discussed in order from old to young.

Masccnteo group The oldest sedimentary rocks of the Hesperian Massif are the paragneisses of the catazonal complexes in the northwestern part of the peninsula. This sequence is largely composed of greywackes and semipelites with minor calcareous intercalations. Mafic sills and/or mafic extrusive rocks occur within the succession at

- 67 - [ | CANTABRIAN ZONE

|-/.;•]".':\ Narcea Series

WEST ASTURIAN - LEONESE ZONE a) Villalba Series

Olio de sapo

CENTRAL • IBERIAN ZONE

CATAZONAL complexes in Galicia and Notan Portugal

_,.+] Pedroches Batholit

OSSA-MORENAZONE

SOUTH PORTUGUESE ZONE

100 km

Fig. 23. Outline of the Hesperian Massif (after Fontbote & Julivert, 1972). The numbers refer to the aolwms of Fig. 24. an uncertain stratigraphic level. This sequence is denominated here the "Masanteo group", after the peninsula near Cabo Ortegal where banded gneisses and intercalated eclogites occur (Vogel, 1967). The basement upon which the sediments of the Masanteo group were deposited is unknown. The sedimentation age did probably not exceed 1.5 Ga, while an age higher than 1.0 Ga is inferred for the emplacement of the mafic rocks. At least part of the sediments were derived from a 2.5 Ga old source terrain (chapters II and III).

Valdelaaasa group This group is a rather monotonous sequence of greywackes and pelites with some intercalated quartzites, conglomerates, and occasionally calcareous rocks. It constitutes the lowermost part of the Upper Precambrian-Paleozoic sedimentary succession. Mafic volcanics occur locally oiar the top of this sequence, while the occurrence of fragments of plagioclase crystals has been interpreted as indicating felsic volcanism. The group locally attains a considerable thickness, e.g. 6000-7000 m for the Alcudia Series near Almaden (Saupe et al., 1977; Fig. 24). The sediments were probably deposited in a shallow marine environment (Bouyx, 1970). Tentatively, the metasediments of the Ordenes complex are also included in the Valdelacasa group. The basement underlying the Valdelacasa group is unknown. In most places the transition to the overlying Montes de Toledo group is gradual, although angular unconformities occur (Parga, 1971).

Montes de Toledo group The Montes de Toledo group is a Late Precambrian-Middle Cambrian heterogeneous succession with strong lateral differences in sedimentary facies. Conspicuous se-

- 68 - SOUTHERN SIERRA MORENA t MONTES DE SISTEMA NARCEA AIMADEM 131 SALAMANCA IGI EXTREMAOURA |I| EXTREMADUHA [2| TOLEDO (4) CENTRAL 151 ANTIFORM (10)

LOWER DEVONIAN

U OROOVlCtAN CALYMENES SCHISTS

L ORDOVICIAN

U CAMBRIAN

M CAMBRIAN VAIDE- CANAS L CAMBRIAN SERIES MORILLE i SERIES

EOCAMBRIAN VALDE- LACASA SERIE SERIE ALCUDIA SERIES NEGHA NEGRA SERIES PRECAM BRIAN

pelites V felsic volcanics Fig. 24. Pre-Lower Devonian stratigraphy of some areas in Ohe Hesperian Massif, pelites, greywackesj T turbidites For locationst see Fig. 23. Data obtained from; sandstones 1, Vegas (1968); 2, Vegas (1970), Gutierrez- M mafic volcanics Elorza et at. (1971); Z3 Saupe et al, (1977); quartzites 4. Moreno (1975), Capote et al. (1977), Martin- Fe ferruginous Escorza (1977); 5, Capote et al. (1977)j 6, o o o o conglomerates liartinez-Garaia & Nicolau (1973), Ordonez I (1974); 7, Martinez-Gareia (1973); 8, Martinez- limestones Garaia & Apalategui (1976); 9, Maraos (1973); 10, Perez-Estatin (1973), Perez-Estaun & Mar- dolomites tinez (1978).

hiatus

T SIMPLIFIED STRATIGRAPHY OF THE HESPERIAN OROGEN

OOOQOOO SAN VITERO GROUP (San Vitero S., Garganta F.. Ousixoiro F.) 400 Ma

' : oooooOoo

M J

,OI-N \ -450Wa NAVIA GROUP (Agüeira F., Luarca F., San Pedro de las Herrerias F., M ^- — ^Z Culebra F., Armorican Quartzites)

Fe — Fe — Fe

-510 Ma - BARALLA GROUP (Cabos S„ Baralla flyschoid S„ Puebla F., Porto S. (?) )

m S < MONTES DE TOLEDO GROUP (Ollo de Sapo. Narcea S. (?), Cândana S., Transitional S., Vegadeo F., Lancara F., u Valdecanas S., upper parts of Porto S. (?), Morille S.) 590 M —. — M — v V V VALDELACASA GROUP (Alcudia S.. Valdelacasa S., Villalba S., Serie Negra, lower parts of Porto S. (?), Ordenes S. (?) Narcea S. (?) )

^i_ —•—— v 7~- MASANTEO GROUP

Fig. 2S. Simplified pre-Lower Devonian stratigraphy of the Hesperian Orogen. For the Legend, see Fig. 24.

- 70 - quences included in this group are the felsic volcanosedimentary rocks of the Olio de Sapo and comparable formations, which attain their greatest thickness in the central zones of the Hesperian Massif, The Olio de Sapo formation consists of coarse- grained augengneisses with pelitic, quartzitic, amphibolitic and calc-silicate intercalations. Inmost places the volcanosedimentary origin of this formation is clear, butthe stronger metamorphic parts in the Sistema Central contain possibly some high-level intrusives related to the volcanism (Capote et al., 1977).

The age of the Olio de Sapo formation has been subject to debate (Parga & Vegas, 1971; Fontbote & Julivert, 1972; Bischoff et al., 1978). Age estimates range from Eocambrian to Late Cambrian/Early Ordovician (Fig. 24). A pre-Early Ordovician age seems certain, but the exact stratigraphic position has still to be established. In the Sanabria region (Martinez-Garcia, 1973; column 7, Fig. 24) and in the eastern Sierra de Guadarrama (Bischoff et al., 1978) a Late Cambrian age has been inferred, but this dating is questionable (Parga & Vegas, 1971). Probably, the volcanosedimentary formations are diachronous: felsic volcanism in the central zone may be slightly younger and of longer duration than that in other parts of the Hesperian Massif.

Alleged lateral equivalents of the Olio de Sapo formation are carbonate rocks, locally ferruginous, and turbidites. In the Montes de Toledo region the latter deposits contain intercalated olistostromes, slide conglomerates and slump sheets (Moreno, 1975). Mafic volcanics of alkaline character occur locally within the clastic successions (Parga, 1976). Towards the top of the Montes de Toledo group pelites, quartzites and limestones occur in a shelf facies (Moreno, 1975; Capote et al., 1977). The sediments of the Montes de Toledo group register vertical movements during the deposition. Locally, volcanodetrital sequences were developed on (partly) emerged areas, which persisted for the longest time in the central zone of the Massif. In the submerged areas carbonate sedimentation occurred on local highs, while in adjoining NW-SE trending furrows locally thick turbiditic sequences were deposited (Marcos, 1973; Capote et al., 1977). A more quiet sedimentation in shallow marine environment occurred in Early to Middle Cambrian time, although basins persisted locally. Infra- to intertidal environments are inferred in the Cantabrian Zone (van der Meer Mohr, 1969; Vegas, 1978).

The Narcea Series in the northern part of the Hesperian Massif is customarily correlated with the Villalba Series in de Valdelacasa group. However, the occurrence of volcanosedimentary rocks and distal turbidites (Perez-Estaun, 1973; Perez-Estaun & Martinez, 1978) make a correlation with similar rocks of the lower part of the Montes de Toledo group more likely.

Baralla group The Baralla group of Middle Cambrian to Early Ordovician age consists of pelites and sandstones with intercalated quartzites, probably deposited in a shallow marine environment. The Baralla group occurs in the northern and northwestern parts of the Hesperian Massif. In the southern and southeastern parts, the Precambrian and Cambrian sequences are separated from the overlying Ordovician by an angular unconformity and the Baralla group is nearly absent. The 3000-4000 m thick Cabos Series in western Asturias (Fig. 24) was deposited in a sublittoral to littoral environment in a narrow basin (Marcos, 1973). Alkali-olivine basalts and alkali- trachytes are intercalated especially in the lower part of this Series (Parga, 1976).

Navia group The Navia group starts with the "Arsnorican Quartzite" of Arenigian age. This quartzite occurs in nearly all parts of the Hesperian Massif (Vegas, 1978), locally followed by alternating quartzites and sandstones (Martin-Escorza, 1977). The thickness of the quartzitic formation is variable (Martin-Escorza, 1977). The overlying pelitic rocks are often ferruginous and indicate euxinic conditions (Marcos, 1973). In the northwestern part of the central zone an important hiatus is present

- 71 - in the upper Navia group (Martinez-Garcia, 1972, 1973), while a thick ( >3000 m) turbiditic sequence was deposited in a deeply subsiding trough in western Astu- rias (Marcos, 1973). A similar trough is possibly represented by the thick Ordo- vician successions of the Ossa Morena Zone (Vegas & Munoz, 1976). Locally, the Navia group contains intercalations of alkaline mafic volcanics (Parga, 1976). The sediments of the Navia group indicate that in Early Ordovician time subsidence occurred nearly ubiquitously, during which the Arenigian quartzites and overlying dark shales were deposited. Subsequently, differential vertical movements resulted locally in emerged areas and subsiding basins.

Saw Vitero group In the northwestern part of the central zone the Upper Silurian to Lower Devon- ian SanVitero group consists of pelites and greywackes with intercalated conglomerates, felsic and mafic volcanics, quartzites and limestones (Ribeiro, 1974; Aldaya et al,, 1976; Parga, 1976). Pebbles of metamorphic rocks occur locally in the conglomerates (Aldaya et al., 1976). The melange units around the Cabo Orte- gal complex and the southern rim of the Ordenes complex (Den Tex, 1980) also belong to this group (van der Meer Mohr, 1975). In the other parts of the Hespe- rian Massif this group generally starts with a dark shale formation of variable thickness, followed by pelites and (ferruginous) sandstones with locally felsic or mafic volcanic intercalations (Marcos, 1973; Ordonez, 1974; Saupe et al., 1977). Limestones of Early Devonian age, indicating a shelf environment, occur in the Cantabrian and the Ossa Morena Zones (van Adrichem Boogaert, 1967; Scher- merhorn, 1971). During the deposition of the San Vitero group there was again a shallow marine environment with locally emerged areas, especially in the central zone.

Younger sediments Middle Devonian to Carboniferous sediments are locally present in the central zone of the Hesperian Massif. They attain their greatest thickness in Cantabria and the South Portuguese Zone. Small NW-SE oriented basins in the southern part of Extremadura show a progressive decrease in age from Early to Late Carbon- iferous towards the southwest. In the Cantabrian Zone the Devonian and Carbon- iferous sequences consist of shales, sandstones and carbonaceous rocks deposited in shallow marine environments. The occurrence of fossil soils and of alternating trans- and regressive facies indicate repeated vertical movements (van Adri- chem Boogaert, 1967). In the southern Cantabrian Mountains, a thick Upper Carbon- iferous turbiditic sequence grades laterally into sediments deposited in a shallow marine environment (Maas, 1974). In the South Portuguese Zone the Devonian pelites and greywackes are separated from a thick Carboniferous turbiditic sequence by the volcanosedimentary rocks constituting the Iberian Pyrite Belt (Schermerhorn, 1971; Bernard & Soler, 1974; Vegas & Munoz, 1976). In the extreme southwest the (Upper) Visean limestones might represent a shelf facies bordering the trough in which the turbidites were deposited (Schermerhorn, 1971).

Summary of the Late Preccorbrian-Paleozaio history In the Late Precambrian subsidence started, affecting a rather broad area. Local highs and furrows along a central emerged area with widespread volcanic activities were produced by vertical movements in latest Precambrian to Middle Cambri- an time. Continued subsidence in the Early Ordovician resulted in broadening of the sedimentary basin and was again followed by a period of block-faulting, which gave rise to alternating highs and furrows and narrow, deeply subsiding troughs near the borders of the area. In the Silurian subsidence outside the central zone resulted in a transgression, while in the northwestern part the first phases of uplift can be recognized. This uplift continued up into the Late Carboniferous and ultimately affected the whole of the Hesperian Massif. The overall picture of the supracrustal history is a prolonged regime of subsidence, successively affecting a broader area, with superimposed differential vertical movements lead-

- 72 - fefcy" ing to the formation of local emerged areas. These acted as the source areas for the immature sediments in the neighbouring furrows. Biraodal volcanism occurred from the latest Precambrian up into the Carboniferous. The prolonged regime of subsidence came to an end in the Late Paleozoic and was followed by uplift.

IV.2,2, Infracrustal history Regarding the metamorphic and deformation phases affecting the rocks of the Hes- perian Massif in Paleozoic time, a subdivision in catazonal rocks and metamorphosed Precambrian to Paleozoic supracrustals has to be made.

The aataaonal complexes The catazonal complexes which are situated in the northwestern part of the Massif display the longest and most complete succession of metamorphic and tectonic events. The geological history of these complexes is dealt with in chapters II and III and is schematized in Table 1. Following prograde metamorphism up into the eclogite facies, the constituent rocks were subjected to an increare in temperature resulting in the development of high-pressure granulite faciep assemblages and migmatization. The subsequent events reflect largely a deer' ise in pressure and temperature, i.e. cooling during uplift, accompanied by repeated folding.

Metamorphosed Upper Precambrian to Paleozoic supraorustals The metamorphic and tectonic history of the supracrustal rocks in western Galicia outside the catazonal complexes has been schematized in Table 1. For the other parts of the Hesperian Massif the evolution is more or less comparable (e.g. Bard et al., 1973; Martinez-Garcia, 1973; Aldaya et al., 1973; Fuster et al., 1974; Vegas & Munoz, 1976; Capote et al., 1977; Aparicio Yagiie & Huertos, 1978). At least three Late Paleozoic folding phases are usually recognized, the first two being the most important (Bard et al., 1973). A phase of thrusting is locally inferred to have taken place between the first two phases. The first phase produced folds with (sub)horizontal to vertical axial planes, while during the second phase vertical axial planes were developed. These two phases can be correlated with the third and the fourth deformation phase in the catazonal complexes (Table I). Ear- ly Paleozoic deformation is attributed to a Late Cambrian phase and an Early Silur- ian phase, both of local importance only. Late Paleozoic metamorphism generally reached its climax before or during the second main phase of deformation. The metamorphism is of the low- to intermediate-pressure type and its intensity varies from very low grade to amphibolite facies. In areas subjected to the highest grade of metamorphism, large scale migmatization and anatexis took place, giving rise to the development of (par)autochthonous to allochthonous alkaline granites. Calc- alkaline granites are either derived from deeper and/or drier parts of the supracrustals or by larger proportions of partial melting. Locally, mineral relics belonging to an older phase of relatively high-pressure metamorphism have been preserved. Although an overall symmetry can be recognized regarding the intensity of the metamorphism in the Hesperian Massif (i.e. non to very low grade metamorphic in the outer zones and highest grades in the central zone), considerable variation occurs on a smaller scale. From the juxtaposition of low-grade and high- grade massifs and blocks showing metamorphic imprints with different PT-gradients, important vertical movements can be inferred.

In summary, the supra- and infracrustal history reveals the following sequence of main events: - More or less continuous sedimentation in predominantly shallow marine environments from the Late Precambrian to the Late Paleozoic. A mote or less continuous block-faulting, although with some peaks in activity, influenced the sedimentation by the development of alternating highs and furrows. Deeply subsiding basins along the margins were successively shifted outwards. - Bimodal magmatism, as indicated by the occurrence of felsic and (alkaline) mafic volcanics throughout the whole of the sedimentary record, and by the close associ-

- 73 - Ifc ation and often hybridization of granitic and mafic rocks (Capdevila et al., 1973; Parga, 1976; Den Tex, 1980), - Abundant granitic rocks and migmatites, providing evidence for important reworking of older crustal materials. - Temporal and spatial variation of PT-gradients. - More or less continuous infracrustal dynamo-thermal activity from at least 500 Ma ago to about 280 Ma ago, as reflected in the range of ages revealed by isotopic dating (Table 2 and chapter II).

IV.3. INFERRED GEOLOGICAL EVOLUTION OF THE HESPERIAN MASSIF

The continuous sedimentation and isotopic age record precludes a distinction between Variscan and pre-Variscan events. From the supra- and infracrustal history it is rather apparent that the magmatic and tectonometamorphic events which affected the Hesperian Massif belong to one dynamo-thermal cycle of long duration (see also Den Tex, 1980). Use of the terms "Variscan" and "pre-Variscan" has therefore been avoided in this study. For convenience the orogenic processes which built the Hesperian Massif from the Late Precarabrian up into the Late Paleo- zoic will be referred to here as the "Hesperian Orogenesis".

In the mantle-plume model for western Galicia (van Calsteren, 1977; Den Tex, 1980) the existence of an older continental crust is assumed. Although the basement underlying the sediments of the Valdelacasa group is unknown or hitherto unrecognized, the Rb-Sr whole-rock and the U-Pb zircon data suggest that the rocks of the Masanteo group formed part of the basement. Both the nature of the sediments of the Masanteo group and the occurrence of ^2.5 Ga old zircons indicate a derivation from an older continental source terrain. The chemical composition of the intercalated mafic rocks points to an intracontinental setting of the basin in which the rocks of the Masanteo group were deposited. The U-Pb systematics of the mafic rocks indicate that the age of the basement is at least 1.0 Ga (chapter III).

The Late Precambrian to Late Paleozoic rocks also provide arguments for an intracratonic setting of the Hesperian orogen: (1) most of the supracrustals were deposited in epicontinental seas; (2) deeper basins were either filled with sediments in (sub)littoral facies, or with (proximal) turbidites and mixtites which testify the proximity of the source areas; and (3) the initial 87Sr/86Sr ratios of granitic rocks indicate the involvement of continental crust in their generation.

For most of the Hesperian Massif the existence of a continental basement is usually acknowledged (e.g. Bard et al., 1973; Vegas, 1978). Only the southwestern- most part (the South Portuguese Zone) is frequently regarded as the oceanward extension of the sedimentary basin and as the site of a Late Paleozoic collision suture. No evidence for the existence of an oceanic crust in that zone has been found, however. Moreover, the bimodal volcanism of the Pyrite Belt is inferred to have occurred in an intra-continental basin (Munha, 1979). The thick Devono- Carboniferous sequences in the South Portuguese Zone are therefore better interpreted in terms of trough sequences like those in Cantabria than as a continental-rise prism.

Taking all the evidence together, it can be concluded that the sedimentation proceeded in a broad intracontinental basin, with its northeastern border probably in Cantabria. The site of the southwestern border has so far not been recognized with certainty, but may have been close to the extreme southwestern tip of the peninsula (Schermerhorn, 1971). The Hesperian Orogen should thus be an ensialic one and its evolution might be described in terms of the following model (Table 1).

Evidence pertaining to the geological history prior to the Hesperian Orogeny is scarce and can only be inferred from the catazonal complexes in the northwestern

- 74 - part of the Massif. The presence of "V2.5 Ga old zircons in metasediments of the Masanteo group, to which a maximum age of 1.5 Ga is tentatively assigned, indicates the existence of older source terrains. Basement rocks of comparable age crop out in the northern part of the Armorican Massif (Calvez & Vidal, 1978), while from other Paleozoic basement areas in western Europe similar occurrences of >2.0 Ga old zircons have been reported (Zwart & Dornsiepen, 1978). Before 1.0 Ga ago mafic rocks were emplaced as extrusives and/or as high-level intrusives, possibly in relation to a stage of continental rifting (van Calsteren, 1978). Subse- quently, these rocks were subjected to metamorphism up into the eclogite facies (M-0). Although this phase still has not been dated, the position in the lower parts ( > 30 km) of the continental crust in Early Paleozoic time appears to preclude any genetic relationship between the prograde eclogitization and the Hesperian orogenesis.

The geological evolution of the Hesperian Massif from the Late Precambrian up to the end of the Paleozoic is attributed to a mantle-plume which reached the base of the lithosphere in Late Precambrian time. As a consequence of this mantle plume a broad rift~zone was formed by updoming, crustal thinning and foundering of the domal crest. In view of the thickness and the spatial extent of the Valdelacasa group sediments, subsidence ar, 1 basin formation had already affected a rather broad area (of the order of 500 km) in latest Precambrian time. Since the base of the Valdelacasa group is not exposed, the time at which the subsidence has begun can only be estimated: probably at least some 600 Ma ago. Due to the limited number of data on the sedimentary environment and the spatial coherency of the Valdelacasa group outcrops, the initial stages of updoming and graben formation cannot be re- constructed. Continued crustal thinning, block faulting, graben formation and concomitant bimodal volcanism are clearly reflected by the'rocks of the Montes de To- ledo group. The bimodal volcanism, which started already at the end of the time of deposition of the Valdelacasa group, indicates that partial melting processes in the lower crust, probably induced by plume-derived mafic magmas and/or volatiles, began at an early stage in the development of the orogen.

About 500 Ma ago, second-order diapirs from the mantle plume were emplaced at the base of the continental crust. The heat of these lherzolite diapirs induced high- pressure granulite facies metamorphism (M-l) and migmatization in their immediate surroundings. The close association of granitic and gabbroic rocks (Den Tex, 1980) probably indicates that heat transfer proceeded not only by conduction, but also by material transport, i.e. the gabbroic partial melts produced from the lherzolites. Granitic magmas resulting from the migmatization were locally able to intrude higher levels of the crust along deep-reaching faults, probably bounding key-stone grabens (e.g. the "blastomylo-itic graben" in western Galicia). Contin- ued doming of the lithosphere and stretching and foundering of the crust caused subsidence and uplift, the former especially in western Asturias (Cabos Series) and the latter especially in the southeastern part of the central zone (the Late Cambrian to Early Ordovician hiatus). A relatively short period of more general subsidence (Armorican Quartzites and overlying pelites) was followed by renewed block faulting, leading to deeply subsiding troughs near the margins of the rift zone. Uplift of the northwestern parts of the central zone, accompanied by the formation of the local furrows in which possibly a limited amount of oceanic crust was formed (melange sequences in western Galicia, Den Tex, 1980), could be due to a secondary arching of the crust upon emplacement of the second-order diapirs.

The subsequent hornblende-granulite facies metamorphism (M-2) in the lower crust may have been induced by juvenile volatiles derived from the mantle plume or its off-shoots (Bailey, 1978). Heat transfer by such volatiles may also have caused the relatively high-pressure metamorphism at higher levels of the crust in the central zone. The presence of mantle-derived volatiles is evidenced by the occurrence of alkaline to peralkaline granitic rocks (Bard et al., 1973; Bailey, 1974, 1978; van Calsteren, 1977). Uplift of the central zone started in Devonian time (lower pressures at the end of M-2; pebbles of metamorphic rocks in conlgomerates),

- 75 - I »

while marginal subsidence due to continued stretching of the crust slowly shifted outwards and resulted in the deposition of Devonian to Carboniferous sediments in Cantabria and the South Portuguese Zone. A tensional tectonic regime is indicated for the sedimentary basins in the latter areas (Schemerhorn, 1971; Kullmann & Schonenberg, 1979; Munha, 1979). The uplift of the central zone during the continuous widening of the rift may be explained by a beginning of thermal expansion of the crust (Bailey, 1978); Bridwell, 1978). In view of the prolonged NE-SW dil- atation (Cambrian to Early Carboniferous NW-SE trending basins), the deformation and folding on horizontal axial planes prior to the low-pressure metamorphism may be related to uplift and gravity tectonics. The vergence of the F-3 folds, generally directed away from the uplifted central zone (Bard et al., 1973), seems to support this hypothesis.

The rising heat front of the first-order mantle plume caused widespread steepening of geothermal gradients, which resulted in the low-pressure M-3 raetamorphism. The large time lapse between the arrival of the mantle plume and its thermal effects in higher parts of the crust may be explained by assuming that heat was, transferred largely by conduction. Conduction is a slow mechanism of heat transfer (Carmichael et al., 1974), but it may have been slowed even more by the (sub)horizontal foliations arisen from the first folding phases (Talbot, 1971). Contrary to the isograds of the high-pressure metamorphism, those of the low- pressure metamorphism are generally related to intrusive granitic rocks (Fuster et al., 1974). With the beginning of widespread anatexis, material transport may therefore have become increasingly important for the transfer of heat. The subsequent retrogressive M-4 metamorphism possibly reflects local readjustments to conditions of lower pressure and temperature during uplift.

The continuous uplift, affecting an increasingly broader part of the rift zone, may partially be related to thermal expansion of the crust during M-3 and partially to compression. The post-F-3 thrusting (e.g. Bard et al., 1973) possibly marks the beginning of the compressional tectonic regime, during which the F-4 folds with vertical axial planes and the later minor folds were produced. Em- placement of the Ordenes and Laiin complexes, which occurred prior to M-3 (Den Tex, 1980), may either ba related to the post-F-3 thrusting or to the secondary doming of the crust at an earlier stage. As a result of the compression, the last stage of up-thrusting of the catazonal complexes in the northwestern part of the Massif took place after M-4 (Engels, 1972; Den Tex, 1980).

Regarding the final stages of compression and uplift, an explanation could be envisaged as follows:

The cooling and contraction of the mantle plume, followed by collapse of the domed lithosphere caused compression in the (upper) crust and concomitant (minor) thickening of the crust (Hofman et al., 1974). Both the collapse of the lithosphere-dome and the contraction upon cooling of the crust will have counteracted uplift. The final uplift may have been caused, however, by isostatic readjustment upon thickening of the crust and upon the increase in density of the cooled and depleted plume-material.

On the other hand, the uplift might have been related to external sources. Late Carboniferous to Early Permian wrench faults indicate NW-SE to NNE-SSW directed compression (Arthaud & Matte, 1975). The rotation of the compressional stress field from a direction perpendicular to the rift zone to the directions inferred from the wrench faults may be due to plate motions unrelated to the mantle plume.

The mantle-plume model for the evolution of the Hesperian Massif can be summarized as follows:

A mantle plume arrived at the base of the lithosphere at least some 600 Ma ago, and caused updoming of the lithosphere. Continued crustal thinning and downwarp

- 76 - -X *

. f-~--

Nv: of the domal crest resulted in an increasingly broader, slowly subsiding sedimentary basin, in which block faulting influenced the sedimentation. Plume-derived materials (second-order diapirs and possibly volatiles) caused high-pressure metamorphism and migmatization in the (lower) crust. At a later stage, the heat front of the mantle plume reached the higher parts of the crust and resulted in widespread low-pressure metamorphism and migmatization. Bimodal magmatism, with an alkaline mafic component, occurred throughout the whole of the active life of the plume. When the mantle plume cooled, the tensional tectonic regime in the crust was reversed and compressional tectonics mark the final stages of the development of the orogen.

IV.4. TECTONIC SETTING OF THE HESPERIAN MASSIF

The Hesperian Massif shows close similarities in its development to well-known continental rift structures (Burke & Dewey, 1973; Martin & Porada, 1977a; Sengor et al,, 1978). The continuous widening of the rift zone, the alkaline volcanism and the late-stage compression of the Massif are features characteristic for continental rifts (Hofman et al., 1974). In a pre-Atlantic Ocean fit of the continents (Fig. 26; Le Pichon et al., 1977), the inferred Hesperian rift zone is located at a large angle to the Appalachian collision suture and to the Late Precambrian- Paleozoic continental margin of eastern North America (Rankin, 1975; Thomas, 1977).

Fig. 26. Pre-Atlantic rift reconstruction of the continents (after Le. Pichon et al., 1977; Early Paleozoic continental margin of eastern North America after Thomas, 1977). Diagonal hatching, Hesperian rift zone.

Although no reconstruction is available of this margin north of Newfoundland, the succession of salients and recesses (Thomas, 1977) renders a location of the Hes- perian Massif at a recess north of Newfoundland likely. The large angle to a collision suture and the location at a recess are likewise characteristic features of rifts (Hofman et al., 1974). Finally, from the contemporaneity of the initial stages of the Hesperian orogeny and the opening of the Iapetus Ocean (between 700 and 600 Ma ago, Strong, 1979) it may be inferred that the Hesperian orogen devel-

- 77 - oped as an aulacogen in relation to the opening of the Iapetus Ocean and not as a collision rift or impactogen related to the closure of Iapetus (Burke, 1978; Sengor et al., 1978).

Apart from the similarities noted above, however, there are also a number of deviating features. First of all, evaporites, common deposits in rifts (Sengor et al., 1978), are unknown in the Hesperian Massif; unfavourable climatic conditions may be invoked to explain the absence of evaporites, however. Other differences are the great width (> 500 km) and the long history ( > 300 Ma) of the Hesperian orogen as compared with other aulacogens (Hofman et al., 1974; Kumarapeli, 1978). In this respect, the Hesperian Massif may be compared with the intracratonic branch of the Damaran Orogen in Southwest Africa, which is of similar width and developed over a similar time span. For this rift zone, a multiple aulacogen model has been developed (Martin & Forada, 1977a), One could imagine that the large width and long duration of ensialic orogens such as the Hesperian and Dama- ran differ from smaller aulacogens only in the size of the mantle plume, while the previous crustal structure determines the number of grabens developed over the plume (Martin & Porada, 1977b).

If the Hesperian orogen started as part of the rift system leading to the opening of the lapetus Ocean, and developed as an aulacogen or failed arm of that system within the continent east of this ocean, a correlation with areas in the Iapetus suture thought to belong to the eastern margin should be possible. Rocks regarded as forming part of this eastern margin constitute the Gander, Avalon and Megu- ma Zones of eastern Canada (Williams, 1979). The Upper Precambrian to Lower Pale- ozoic sediments of these zones (Strong, 1979; Williams, 1979) do indeed resemble those of the Hesperian Massif. The Conception Group of the Avalon Zone, consisting of siliceous shales and greywackes with local volcanic intercalations (Wil- liams, 1979), could be correlatable with the Valdelacasa group. The shales, sandstones and conglomerates of the Signal Hill Group and the subaerial felsic volcanics of the Bull Arm Formation of the Avalon Zone (Strong, 1979) resemble the Montes de Toledo group with its Olio de Sapo volcanosedimentary rocks.

A correlation even further back in geological time can be attempted by assuming that the Grenville orogeny of eastern North America reflects an event of continental rifting without the development of an ocean (Wynne-Edwards, 1976). The fact that the Grenville orogeny is not of a collision type could imply that the pre-Iapetus continuity of North America and western Europe dates back to (at least) Middle Proterozoic time. In that case lateral equivalents of the mafic rocks of the Masanteo group could possibly be found in North America. Rocks of Helikian age (1200-1400 Ma) occur on both sides and inside the Grenville orogen (Stewart, 1976; Emslie, 1978; Strong, 1979). Basaltic rocks of continental character, including plateau basalts (Seal Lake Group), were emplcaed during an event of continental rifting 120Qrl400 Ma ago in eastern Canada (Baragar, 1977; Meyers & Emslie, 1977; Emslie, ..1.978.). These rocks resemble the mafic rocks of the Masan- teo group both in tectonic setting and probably also in age. In spite of slight differences in chemical composition (Baragar, 1977; van Calsteren, 1978) it can be hypothesized that the continental basalts of eastern Canada and the oldest mafic rocks of western Galicia originated during the same event of continental extension at least some 1200 Ma ago.

IV.5. CONCLUDING REMARKS

In the foregoing it has been shown that the available data pertaining to the supra- and infracrustal history of the Hesperian Massif support an evolution as an intracontinental.rift system. In view of the age and the global tectonic setting of the Hesperian orogen, it is postulated that the Massif developed as an aulacogen related to the Iapetus Ocean. Rifts extending from the western margin of Iapetus, although of much smaller dimensions, have been described from east-

- 78 - s fa REFEREN ern Canada (Kumarapeli, 1978). Although a mantle-plume/aulacogen model could explain the geological evolution of the Hesperian Massif rather well, several problems remain to be solved. First of all, a tensional tectonic regime persisted until the Carboniferous, while the supposedly adjoining Iapetus Ocean closed in Ordovician to Silurian time (Strong, 1979; Williams, 1979). The continuous extension of the Hesperian rift zone could be related to collision rifting (Sengor et al., 1978), as has been postulated for the Rhine Graben (lilies, 1978; lilies & Greiner, 1978). On the other hand, similar problems related to the evolution of the Appalachians (i,e, the Devonian and Carboniferous orogenic events well after the closure of Xapetus; Williams, 1979) might indicate that the orogenic processes affecting the North American and West European continents during the Paleozoic were much more complex than could be described by a simple rifting/collision model. A second problem concerns the (apparent) absence of Cadomian events in the Hesperian Massif. Although it is possible that later events overprinted a Cado- mian age record, any supracrustal evidence for a Cadomian orogeny is lacking, un- ALLEGRE, conformities being of local importance only. A third, and possibly the major, 2Q7T problem concerns the relation with other Paleozoic basement areas in western Eu- rope. In most of these areas Early Paleozoic ages have been recorded, but widespread angular unconformities seem to be absent (Zwart & Dornsiepen, 1978). It may be postulated that these areas form part of more or less parallel intracontinental rift zones, comparable to the Hesperian Massif (e.g. Renouf, 1974; Krebs & Wachendorf, 1973; Krebs, 1976). In that case a major difference existed between the eastern and western margin of the lapetus Ocean, which possibly might be explained by a different structural trend in the western continent (i.e. the Gren- ville Province).

In any case, the following conclusions can be drawn for the Hesperian Massif: 1. The orogenic processes leading to the development of the Hesperian Massif started at least 500 Ma ago, probably before 600 Ma ago, and lasted until a- bout 280 Ma ago. This cycle is designated here as the Hesperian Orogeny. 2. The orogenic evolution of the Hesperian Massif was entirely intracontinental and may be interpreted in terms of a mantle plume induced rift zone. 3. A number of the rock units originally regarded as polymetamorphic are mono- metamorphic. The arguments from which polymetamorphism was inferred (mainly the presence of relics of a high-pressure mineral assemblage, e.g. Floor, 1966; Arps, 1970; Arps et al., 1977) are explained by steepening of geothermal gradients during one orogenic cycle due to the heat of the mantle plume. The m^t-Tmorphism in these rocks is plurifacial and the only polymetamorphic rocks appear to be the older mafic rocks in the catazonal complexes and the paragneisses of the Masanteo group. 4. The existence of a pre-Hesperian continental crust of at least 1.0 Ga old is indicated by the U-Pb systematics of mafic rocks. Zircon U-Pb data reveal the existence of still older source areas. The assumption that no continental crust older than 700 Ma existed in western Europe (Jager, 1977; Vidal, 1977) has to be abandoned in view of these data. 5. The Hesperian orogen may have evolved as an aulacogen related to the Early Paleozoic Iapetus Ocean, but several problems concerning its global tectonic setting are .=till to be solved in order to obtain a full understanding of its development.

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- 86 - #"'

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ACKNOWLEDGEMENTS

This study was completed during the tenure of a scholarship from the Netherlands Organization for the Advancement of Pure Research (ZWO). Both this Organization and the Stichting Molengraaff-Fonds provided funds for fieldwork and congress- visits.

I am indebted to Professor Dr. H.N.A. Priem, director of the ZWO Laboratory of Isotope Geology, Amsterdam, for his continuous interest and guidance during the course of the investigations. His editorial skills were of great importance during the preparation of the manuscript.

Professor Dr, E. Den Tex proposed this study, which forms part of the research program of the "Working Group Galicia" under his direction. His extensive knowledge of the geology of the Iberian Peninsula and his continuous interest in the progress of the work were a great help to me.

Dr. E.H. Hebeda and Dr. A.C. Tobi acted as co-referents. It is my pleasure to acknowledge their help through discussion and critical review of the manuscript.

The investigations were carried out in the ZWO Laboratory of Isotope Geology. I am indebted to the staff-members Dr. P.A.M. Andriessen, Dr. N.A.I.M. Boelrijk, Dr. E.H. Hebeda, Dr. E.A.Th. Verdurmen and Mr. R.H. Verschure for their continuous interest, cooperation and support. They educated me in the various fields of the analytical techniques. I benefitted very much from Dr. Hebeda's extensive knowledge of mass-spectrometry and computer programming, but I appreciate particularly the value of his critical analysis of my work. Dr. Boelrijk introduced me into the chemical procedures and the mathematical evaluation of the U-Pb data. Both he and Dr. Verdurmen spent much time on the treatment of the analytical uncertainties.

I am indebted to my colleagues Dr. C.E.S. Arps, Dr. P.W.C. van Calsteren, Mr. E.J. Keasberry and Dr. J.B.W. Wielens for fruitful discussions.

Without the cooperation of the technical staff of the ZWO Laboratory of Isotope Geology this study would have been impossible. Mr. L. IJlst introduced me into the techniques of mineral separation. He also performed a large part of the separations, assisted by Mr. E. Douwes. Mr. J.C. van Belle assisted with the chemistry, together with Messrs. P. Remkes and R.J. Smeets. Mr. B. Voorhorst performed the mass-spectrometric analyses of uranium and lead, maintaining a constant high quality. The strontium isotope analyses were made by Mr. F. Bena- vente. Mrs. J.M. Kutova performed the XRF analyses. Much support was also given by Messrs. N. Dijkstra, R. Gans and R. Scheveers.

I am very grateful to Mrs. M.J.L.H. Petit-Puts for her devotion to the typing of the manuscript.

Much assistance was also received from the technical staff of the Geological Institute of the State University at Leiden: Messrs. J.J.J. van Bergen Henegou- wen and J. Bult prepared the drawings, Messrs. W.C. Laurijssen, H. Schiet and J. Verhoeven made the photographical reproductions, Mr. J. Verhoeven provided the X-ray diffraction data, and Messrs. C.J. van Leeuwen, J. Schuilenburg and M. Deyn prepared the thin sections.

- 87 - The major element analyses of whole-rock samples were made in the chemical laboratory of the Dept, of Petrology in Leiden under the supervision of Mr. K.M. Stephan. Mineral analyses were performed by the "Werkgemeenschap voor Analytisch w> Chemisch Onderzoek van Mineralen en Gesteenten" (WACOM), Amsterdam, under the supervision of Dr. C. Kieft. During the field work fresh outcrops of some of the investigated samples were blasted by Messrs. J, Lozano Villa and E. Martin Rodriquez of Rio Xinto Fatino S.A. This study forms part of the research program of the "Stichting voor Isotopen- Geologisch Onderzoek"> supported by the Netherlands Organization for the Advance- ment of Pure Research (ZWO).

- 88 - j APPENDIX

EXPERIMENTAL PROCEDURES AND UNCERTAINTY ESTIMATES

I. SAMPLE PREPARATION

1.1. Mineral separation Samples of about 20 kg were crushed and pulverized to < 250 Wm using a jaw crusher and a Bico-pulverizer, after which a representative whole-rock sample was taken. Zircons and monazites were separated using stabilized bromoform and diiodomethane (IJlst, 1973a) and Clerici solution in a large laboratory overflow centrifuge (a modified and enlarged version of the type described by IJlst, 1973b), and a modified Frantz isodynamic separator (Verschure & IJlst, 1969), Before splitting the non- to slightly magnetic zircon fractions into size fractions, a small representative aliquot was taken for optical and X-ray study. For the sizing sieves were used with oblong apertures in order to facilitate the passing of the elongated prismatic crystals. The 2 to 20 mg zircon and monazite aliquots of different size and magnetic fractions were handpicked to 99,9% purity. A few fractions containing a large proportion of sulphide minerals and/or rutile were also purified by dielectric separation, using two bent needles mounted in an insulating handle (Ralston, 1961).

1.2. Whole-rock samples used for U-Pb analysis Samples of about 8 kg were crushed, after which pieces showing alteration and/or retrogradation along joints and veinlets were discarded. The remaining fresh chips were pulverized using a disk pulverizer and a ball-mill, and whole-rock aliquots were taken by means of a rotary sample-splitter. Samples of Cabo Ortegal eclogites were taken from the aliquots prepared for the Rb-Sr whole-rock investigations by van Calsteren et al. (1979).

II. MAJOR AND TRACE ELEMENT ANALYSIS

Major element analyses were made in the chemical laboratory of the Department of Petrology, Mineralogy and Crystallography of the State University at Leiden, The Netherlands, following the rapid-rock-analysis-method of Shapiro (1967). Results of replicate analysis of U.S.G.S. and "inhouse" standard rocks havt been published by van Calsteren (1978) and Dekker (1978).

Trace elements (Rb, Sr, Th, U, Pb, Y, Zr) were analyzed by X-ray fluorescence spectrometry on pressed-powder pellets in the ZWO Laboratory of Isotope Geolo- gy, Amsterdam.

Mineral analyses have been carried out by electron probe microanalysis by the Working Community for Analytical Research of Minerals and Rocks (WACOM).

III. Rb-Sr ANALYSIS

Rb/Sr ratios and Rb and Sr contents of whole-rock samples were measured by X-ray fluorescence spectrometry (Verdurmen, 1977). The 87Sr/86Sr ratios were measured directly on unspiked Sr, using a computer-controlled VARIAN CH5 mass- spectrometer with Faraday cage collector and digital output. Estimated percentage errors are 1% for Rb/Sr and Rb and Sr contents, and 0.05% for aVSr/ 8SSr. Rb-Sr isochron ages and errors were computed according to York (1966, 1967).

- 89 - IV. U-Pb CHEMISTRY

IV,1, Minerals Chemical decomposition and separation of U and Pb were essentially according to Krogh (1923). The zircon and monazite samples were washed in hot nitric acid (1:1), rinsed thrice with water, pulverized in an agate mortar, and split into an aliquot for analysis of the U and Pb contents and another for analysis of the Pb isotopic composition. Both splits were transferred to teflon capsules, after addition of a mixed 208Pb + 23 U spike to the capsule for concentration analysis. They were digested under hydrothermal conditions in 1.5-2 ml hydrofluoric acid (48%) and three drops of concentrated nitric acid. Uranium and lead were extracted by means of ion exchange columns, using 1.0 ml AG 1X8 200-400 mesh as a resin. The lead concentrates were further purified by anodic deposition (Arden & Gale, 1974; Barnes et al., 1973a), Ultrapure reagents were obtained by sub-boiling destination in teflon or quartz stills. Processing blancs were 1.6-5.4 ng for Pb and 0.05-0.3 ng for U (Table Al). A detailed description of the procedures is given by Wielens et al. (1979).

Table A.I, Contamination levels ng/g Pb Reagents (mean) individual values Hydrofluoric acid 0.131 0.167, 0.061, 0.201, 0.196, 0.031 Hydrochloric acid 0.057 0.047, 0.052, 0.100, 0.064, 0.044 0.095, 0.32, 0.34 Water 0.113 0.407, 0.232, 0.014, 0.055, 0.037 0.078, 0.035, 0.043 Nitric acid 0.072 0.091, 0.053 Theoretical processing blanc: < 1 ng

Processing blancs range (ng) mean (ng) Pb 1.60 - 5.36 2.47 U 0.05 - 0.28 0.15 Maximum estimated loading blanc range (ng) mean (ng) Pb 0.12 - 0.21 0.15

IV.2. Whole-rock samples About 100 mg of samples was digested under hydrothermal conditions in teflon capsules, using 3 ml hydrofluoric acid (48%) to which 6 drops of concentrated nitric acid were added. Spikes were prepared from the lead-206 spike assay NBS SRM 991 and the uranium-235 spike assay NBS SRM 993. After digestion the solution was taken to dryness. The residue was then taken up in 3 ml hydrochloric acid 7 M, after which the capsule was closed again and heated overnight at 200°C. This was repeated once, after which the cooled hydrochloric acid solution was loaded directly on the anion exchange resin.

A three-step procedure was followed for the separation of U and Pb. In the first step U + Fe were separated from Pb and other elements, a second step concerned the separation of Pb, while in the third step U was separated. V + Fe/Pb + bulk separation A polypropylene column (Biorad Econo Column, I.D. 7 mm, volume 1.5 ml) was filled with 1.5 ml (40 mm bed) anion exchange resin (AG 1X8 200-400 mesh). The resin was purified by elution with two reservoir volumes (about 8 ml) of hydrochloric acid 3 M and 2 volumes of water. The resin was conditioned by elution of one reservoir

- 90 - volume of hydrochloric acid 7 M. The sample solution of 3 ml hydrochloric acid 7 M was loaded directly on the column and Pb + other elements were eluted with 7 ml hydrochloric acid 7 M (subsequently 1 ml and three times 2 ml, using 4 micro-pipet). U + Fe were eluted with two times 3 ml water. Both solutions were taken to dryness.

Pb-separation The Pb residue was taken up in 2,5 ml hydrochloric acid 3 M and loaded on a column which was prepared in the same way as for the zircon U-Pb separation. The column was eluted with 3 ml hydrochloric acid 3 M after which the fraction containing the Pb was eluted with 3.5 ml hydrochloric acid 6 M. This Pb fraction was taken to dryness, converted to nitrate and purified by anodic deposition,

U-sepcwation ..-.'..,_:_• For the U-separation the procedure described by Tera & Wasserburg (1975), slightly adapted for larger samples, was followed. The U residue was taken up in 1 ml > hydrochloric acid 4 M and taken to dryness. This residue was dissolved in 1 ml ~-"~ hydrochloric acid 4 M containing 1 gram of ascorbic acid per 20 ml hydrochloric acid 4 M. The solution was then allowed to stand for about 15 minutes until the yellow colour of iron disappeared. The column, containing a 20 mm bed of AG 1X8 200-400 mesh resin, was washed with 1.5 reservoir volumes of hydrochloric acid 4 M and 1.5 reservoir volumes of water. Preconditioning of the resin was done by elution of 3 ml hydrochloric acid 1 M followed by 2 ml hydrochloric acid 4M/a- scorbic acid. The sample solution was loaded on the resin bed, after which Fe was removed by elution with 3 ml hydrochloric acid 4 M/ascorbic acid. The ascorbic acid was removed by elution with 2x4 ml hydrochloric acid 6 M. Finally, the U was eluted with 3 ml hydrochloric acid \ M. This solution was taken to dryness and the residue converted to nitrate.

V, MASS-SPECTROMETRY

Uranium and lead were analyzed on a computer-controlled Teledyne SS 1290, 12 inch, 90° sector mass-spectrometer with Faraday cage collector and digital output. Lead was loaded as nitrate on a single zone-refined rhenium filament, with silica-gel and phosphoric acid (Barnes et al.,_ 1973a) _._ During,analysis the_procedures of the same authors were followed. The quantities of lead loaded were typically between 200 and 500 ng for the mineral samples and 50-200 ng for the whole-rock samples. Due to difficulties in measuring the extreme isotopic ratios of the spike used for the determination of loading blancs for lead, these can only be estimated to be less than 0.15 ng (Table A.1.). Uranium was mounted as nitrate on triple rhenium filaments and analyzed according to the low-temperature procedure described by Shields (1966). The amount of U loaded was of the order of 500 ng for the mineral samples and 250-350 ng for the whole-rock samples.

VI. DATA REDUCTION

VI.1. Bias factors All measured isotopic ratios were adjusted using bias factors (defined as Ratiomeasure^/Ratiostan(jar^) obtained from replicate runs of NBS lead standards SRM 981 and SRM 982 and NBS uranium standards U-150, U-500 and U-800 (Table A.2). Analysis of small smounts of the NBS uranium standard U-500 indicated that the bias factor for small amounts of U, as is the case with the whole-rock samples, is slightly higher than for the zircon samples: 1.005 for the measured 23su/238U ratio.

- 91 - Table A,2, Isotopia ratios of lead and uranium standards measured during this All study 0.9 of LEAD ISOTOPIC STANDARDS VII. NBS SRM 981 208pb/206pb 207pb/206pb NBS* 2.1681± 0.0004 0.91464+0.00016 0.059042+0.000018 Measured (18 analyses) 2.1632+0.0013 0.9138+0.0003 0.05911 ±0.00006 (Rra/Rs)*** 0.99775 + 0.00067 0.99904+0.00033 1.0011

NBS SRM 982 NBS* 1.00016 + 0.00018 0.46707 + 0.00010 0.027219 + 0,000014 Measured (17 analyses) 0.9985 + 0.0005 0.4666 + 0.0002 0.02726+0.00006 (Rm/Rs) 0.99837 ±0.000A9 0.99889 ±0.00041 1,0014

URANIUM ISOTOPIC STANDARDS NBS U-500 N3S** 0.9997 + 0.0010 Measured 1.0040 ±0.0006 (22 analyses) 1.00434 + 0,00059 (Rm/Rs)*11* NBS U-800 4.266 ±0.004 NBS** Measured 4,285 + 0.004 (9 analyses) 1,00453 ±0.00083 (Rm/Rs)*** NBS U-150 0.18109 + 0.00018 * NBS (Catanzaro et al., 1968) NBS** ** NBS certificate Measured 0.1817*0.0003 *** Bias factor. Rm = measured ratio, (5 analyses) Rs = standard ratio. (Rm/Rs)*** 1.00319 ±0.00161

VI.2. Blanc and common lead correction '• ;.! VI.2.1. Minerals The isotopic ratios of the radiogenic lead were obtained after correction of the measured ratios for, successively, 1 - The processing blancs (chemistry and loading) determined for each set of samples, assuming that the composition of the contaminant lead is the same as that of average modern lead (AML, chapter I.4.); 2. The initial common lead, obtained by assuming that the remaining 201|Pb is initial. The isotopic composition of this lead has been calculated using a 207Pb/ zosPb model age and a supposed development with time of common lead with ]i - 238u/20>.pb = 9-74 and w = 232Th/2D'tpb = 3684 (Stacey& Kramers, 1975), extra- polated backwards from AML. The2tt7 Pb/206Pb model age was calculated assuming that all20 "Pb is AML. In view of the very low20l *Pb/206Pb ratios of most samples, the uncertainties arising from this procedure are very small.

- 92 - VI.2.2. Whole-rock samples All concentrations and isotopic ratios were corrected for the processing blanc of 0.9 ng U and 3.7 ng Pb. For the laboratory contaminant lead again the composition of Average Modern Lead was taken (chapter 1.4.).

VII. ESTIMATION OF UNCERTAINTIES OF U-Pb MINERAL DATA

Analytical uncertainties arise from two main sources of errors: systematic and random. The random errors limit the precision (reproducibility), whereas both systematic and random errors affect the accuracy. An attempt will be made to estimate the uncertainties of the U-Pb mineral data of this study, or at least to outline the policy adopted for the estimation of the uncertainties.

"Uncertainty" is used here in a general sense and relates both to precision and accuracy. The term error relates to the accuracy, whereas the standard deviation indicates the precision. The error is expressed as an absolute value or as a percentage. The standard deviation is expressed as an absolute value or as a decimal fraction of the measured value (relative standard deviation). When a sufficient number of measurements is available, estimates of the standard deviations can be made according to the theory of statistics. In the present study this is only possible for the NBS isotopic standards, from which the standard deviations of the mass-spectrometric bias factors are statistically estimated.

Repeats of the mineral analysis cannot be considered true duplicates, as they are performed on small aliquots from inhomogeneous mineral fractions. A statistical estimation of the standard deviation is therefore not possible, and maximum errors will be assigned to the various data from which the Pb/U ratios for the concordia diagrams are calculated. The maximum errors cover both the systematic and the random errors and will be expressed as a percentage (maximum percentage error or m.p.e.)

The combined effect of the maximum errors of the single measurements has been estimated using the statistical error propagation formula, which for a function x = - i f(u,v) is:

Gx == 0u0 (2 (Anderson, 1978). Whera data for the statistical estimation of the standard deviation are absent, the standard deviation will be replaced by an assumed standard deviation, rather arbitrarily set at 0,4 times the maximum errors as assigned above.

The formulas used in calculating the Pb/U ratios are (Boelrijk et al., 1979):

2O7 235 pb/ u= -. (1 + H.) (N7 - Pb ss "

D5 " - D, 20SPb/238U N ) 5 *Pb F' S U s -

in which F!Pb related to the concentration of EE. M. ir l

the spike by F"

- 93 - v . c Cc - EQ EN. M. ite - , related to the concentration of the spike

ESirMi by Fu "c c ' -TT • where 235 238 D5, S5, N5, C5 and E5 = U/ U ratios of the mixture for the determination of the U content, the spike, the natural element, the mixture for calibration of the spike and the calibration solution; 208 206 DB, SB, N8, C8 and E8 = corresponding Pb/ Pb ratios; 207 206 N7 = Pb/ Pb ratio in the natural element; N,, = 201tPb/ZO6Pb ratio in the natural element; ve vc - weights of the calibration solution and the spike; ce cc = concentrations of the calibration solution and the spike; M^ = atomic mass of uranium and isotopic masses of the U and Pb isotopes; 206 201 207 zol c» <*&

Fpb was calculated from a one-step calibration with the lead-206 spike assay NBS SRM 991, hence ce in the formula is directly derived from the NBS value. Because Fy was obtained by a two-step calibration with the uranium-235 spike assay NBS SRM 993 and a laboratory standard, ce in the formula is the concentration of the laboratory standard; this is calculated from a formula similar to that of F'.

In the formulas given above, each of the measured quantities (isotopic ratios, weights and concentrations) is subject to error. These are assumed to be statistically independent, hence in the error propagation formula the term containing the 206 correlation coefficient puv is cancelled. The correlation between the Pb/ U and 207Pb/235U ratios, which is required for the discordia calculation, is dealt with by Boelrijk et al. (1979).

VII.I. Measured isotopic ratios Assumed standard deviations of measured isotopic ratios are derived from the spread in the results of all analyses of NBS isotopic standards (Table A.2.). Although a standardized procedure is followed during mass-spectrometric data collection, the sample-runs will deviate from the standard-runs both in the amount of element loaded as well as in the range of isotopic ratios. To allow for these differences, a value of at least twice the standard deviations of the isotopic ratios of the standard analyses (Table A.2.) is assigned to the assumed standard deviations of the measured isotopic ratios. The 208Pb/206Pb and 235U/238U ratios of spiked runs, however, are mostly comparable to those of the NBS standards. The maximum percentage errors become thc.,\ 0.5% for 235U/238U, 0.3% for 208Pb/206Pb in the spiked runs, '1 and in the unspiked runs 0.5% for 208Pb/206Pb and 207Pb/206Pb and 2.5% for 2OltPb/ 206Pb.

As can be seen from the formulas, the effect of the m.p.e.'s of the isotopic com- _ position of the spike, i.e. the S5 and SB ratios, will be strongly diminished in the Pb/U ratio calculation, as these ratios occur both in the calculation of Fy and Fp, , respectively, and in the Pb/U ratio calculation.

VII.2. Spike calibration factors, F' and F" Using the m.p.e.'s of the isotopic ratios and weights (m.p.e. estimated at 0.15%), and the m.p.e.'s given by NBS for various ratios and concentrations of their as- says used in the spike calibration, the maximum percentage error is assumed to be

- 94 - 1.7% for Fy and 1% for Fpb. The difference between the m.p.e.'s of both spike factors is due to the difference in calibration of the respective F-values (see above).

The results of replicate spike calibration were as follows:

Table A.3, Results of spike calibration F1 F lab-standard U 9.36452 E-5 84.8837 0.500209 9.35992 E-5 84.7956 0.499477 9.35859 E-5 84.7643 0.500259 9.36160 E-5 84.7881

Four values determined for Fp'b gave an average of (9.3612 + 0.0026) x E-5. For the laboratory standard used in the calibration of the U-spike, four values gave an average of 84.81+0.05. Three values determined for Fy gave an average of 0.5000 + 0,0004.

VII.3, Pb/U ratios The uncertainty of the Pb/U ratios can be estimated by combining the m.p.e.'s of the isotopic ratios and the F-values. The result is an assumed maximum percentage error of 2.2% for 2Q7Pb/235U and 2.1% for 206Pb/2:i8U. The larger part of the uncertainty of the Pb/U ratios is due to the uncertainty of the F-values. Uncertain- ties in the composition of the common lead have not been included in the estimation of the m.p.e.'s (see below).

VII.4. Pb and U concents Combining the estimated m.p.e.'s of the isotopic ratios, weights and F-values results ir» an assumed maximum percentage error of 1.8% for the U contents and 1.1% for the Pb contents. Duplicate analysis of the NBS TEG-500 ppm glass standard (Barnes et al., 1973b) gave values of 430.1 and 429.2 ppm Pb (NBS: 426.15±0.41) and 463.7 and 463.8 ppm V (NBS: 461.5 + 0.4).

VII.5. .Uncertainty ellipses The m.p.e.'s given above are for "average samples", i.e. the isotopic ratios in the common range of the investigated samples. To account for the variation in isotopic ratios, for different amounts of common lead, and for different sample/ spike ratios, the assumed standard deviations of the Pb/U ratios used in calculating the uncertainty ellipses were computed for each data-point following the formulas given by Boelrijk et al. (1979). Uncertainties in the composition of the common lead are not included in the ellipse calculations. The uncertainties of the F-values, which introduce a systematic error, are taken up in the ellipse calculations, however. For these calculations the assumed standard deviations (0.4 times the m.p.e.'s) were used and a 95% confidence ellipse is calculated, which thus approximates the maximum percentage error as defined above. The results are identical to the above estimated m.p.e.'s for "average samples" (ellipse- m.p.e.'s around 2.2% for the 207Pb/235U ratios), but can be much larger (ellipse- m.p.e.'s for the 207Pb/ 35U ratios up to 11%) for samples with less favourable common lead contents and sample/spike ratios. It must be emphasized that the uncertainty ellipses are calculated on the basis "of the assumed standard deviations, hence the 95% confidence region only applies to these tentatively assigned values.

VH'6. influence of uncertainties in the common lead composition In this section, the influence of the uncertainties in the common lead compositions will be estimated. For a numerical example deviations are estimated which are comparable to assumed standard deviations.

- 95 - The uncertainty in the composition of the non-radiogenic lead (ac,,8c) used in the calculation of the amounts of radiogenic lead gives rise to a systematic error in the Pb/U ratios. This uncertainty is usually composed of two parts: a contaminant lead introduced during the analyses, and an initial lead incorporated by the zircon during its crystallization.

The contaminant lead is introduced in the processing of the samples. Although no data about its composition are available, it is assumed that the composition of average modern lead (AML, chapter 1.4.) is a fair approximation. Eight recently published compositions of contaminant lead in a number of other laboratories, which in some cases are measured values, give averages of 18.549±0.525 for a, 15.700 4 0.157 for fl and 37.62+ 1,45 for Y- From a comparison with these data a deviation of 37. is assigned to the 206Pb/2l"*Pb and the 207Pb/2t"*Pb ratios of the laboratory contaminant lead.

When estimating the influence of the uncertainty in the contaminant lead, it is assumed that the contaminations of blanc and samples are roughly the same. This is only the case, however, when the contaminant derives entirely from the chemicals.

All 201*Pb remaining after correction for the processing blann' is assumed to be part of a lead component incorporated by the zircon at the time of crystallization. Such initial lead will usually occur in minor inclusions within the zircon. The composition of the initial lead is in most cases supposed to be close to that of average common lead at the time of crystallization, and is calculated using Stacey & Kramers' (1975) two-stage models (Appendix VI.2,1,). Because the composition of the initial lead depends on the lead evolution in the rock system prior to the zircon crystallization, it is very difficult to estimate the corresponding systematic er- or for the zircons in a co-genetic suite of samples. For the zircons analyzed in this study, which are all Paleozoic in age and from rocks thought to be derived from crustal protoliths, the uncertainties in the isotopic composition of the initial lead can be estimated by comparing the published values of the isotopic compositions of common lead in Precambrian and Paleozoic crustal rocks (Doe, 1970) with those calculated here. This leads to a deviation of 5% in the Z06Pb/ 20l

As the initial lead composition, is calculated for each size fraction on the basis of the age obtained when assuming that all common lead is a processing blanc, the calculated composition depends on the measured Pb/U and 2OI*Pb/2O6Pb ratios. The random error introduced in this way is neglected.

To demonstrate the influence of the uncertainties in the composition of the correction lead, the displacement of the center of the uncertainty ellipse is calculated for two samples, one with a very low amount of non-radiogenic lead and another with a proportion of common lead constituting about one third of the total lead content. For the non-magnetic fraction > 160 ym from sample 76-Gal-l, a calculation of the Pb/U ratios taking into account the deviations in the compositions of the contaminant and initial lead results in shift of less than 1.0% for the z0s 238 207pb/235u rati0> whereas the Pb/ U ratio remains the same (Table A.4.). The displacement of the uncertainty ellipse (Fig. A.I.) is thus negligible in this case and the error introduced with the common lead correction is very small in comparison with the errors from other sources. For the non-magnetic fraction <30 pm from sample 76-Gal-4, however, variation in the common lead composition by plus or minus the deviation results in a shift of 11% in 207Pb/235U and 1% in 206Pb/238U (Table A.4.). The systematic error introduced by the common lead correction is thus substantial in comparison with the errors from other sources. As can be seen from Fig. A.1., the line connecting the two extreme points makes a small angle with the longest axis of the uncertainty ellipse, but a large angle with the discordia; this corresponds to a considerable displacement of the data-point. The error calculations for slope and intercept are not valid in such a case, and should be re-

- 96 - Table A. 4. Influence of uncertainties in the common lead composition

76 Gal-1 NM > 160 vim 1.500 0.1117 Adding deviation of blanc 1.499 0.1117 Subtracting deviation of blanc 1.501 0.1117 76-Gal-4 NM < 30 pro 0.5065 0.06464 Adding deviation of initial lead 0.4525 0.06421 Subtracting deviation of initial lead 0.5607 0.06509

UNCERTAINTY ELLIPSE

UNCERTAINTY ELLIPSE OF DEVIATIONS 95% CONFIDENCE REGION

Fig. A.I. Uncertainty ellipses of fraction M > 160 ym of sample 76-Gal-l (above) and of fraction NM < 30 \m of sample 76-Gal-4 (below). Dashed line in lower ellipse: shift of ellipse center when another initial lead composition is chosen.

- 97 - f~~{

placed by a calculation which takes into account the uncertainties in ac and these uncertainties are correlated for all data-points.

VIII, DISCORDIAS AND INTERCEPT AGES

Using the correlation coefficients and variances calculated for each data-point (Boelrijk et al., 1979), lines through the discordant data-points were fitted according to York (1969). The deviation of the slope was derived from the scatter of the points about the discordia. The deviations of the discordia/concordia intersections were calculated by allowing the slope of the discordia to vary about the center of gravity of the data-points by plus and minus its one standard deviation. As the scatter of the data-points about the discordia is used to estimate the standard deviation of the slope, the calculated uncertainties of the intercept ages are in some cases lower than when calculated from the (assumed) analytical uncertainties. In such cases the quoted "standard deviations" of the intercept ages are based on the analytical uncertainties. A more rigorous treatment to obtain the limits of error for the intercept ages in is preparation. This treatment will take into account that an error in the spike calibration factors (F-values) causes a parallel displacement of the regression line.

The parallel displacement of the discordia due to uncertainties in the common lead composition depends on the proportion of non-radiogenic lead. For samples low in common lead, the error in the discordia and the intercept ages will be within the analytical uncertainties. For samples containing high proportions of common lead, the systematic error will be larger than the analytical uncertainties and the accuracy has to be estimated in order to produce reliable age estimates.

IX. ESTIMATION OF UNCERTAINTIES OF WHOLE-ROCK U-Pb DATA

In view of the small amounts of U and Pb analyzed,, the maximum percentage errors of the isotopic ratios were assumed to be twice those used in the study of the zircons: 1% for the 235U/23aU, zoaPb/2oePb (spiked and unspiked) and 207Pb/206Pb, and 2.5% for the 20ttPb/26i?b ratios.

On the basis of the above estimates the maximum percentage errors of 206Pb/20'*Pb and 207Pb/20l*Pb in the lead-lead diagram are taken at 3%. In the modified concordia diagram, the maximum percentage error of (206Pb/238U)* is 7% and of (207Pb/ 235U)* 12%. In the isochron plots, the maximum percentage errors were assumed to be 3% for both U/Pb ratios. The maximum percentage error of Pb and U contents are taken at 1.5%. For each sample maximum errors were calculated to allow for variations in isotopic ratios and error enlargement factors for single isotope dilution. For the data-points in the modified concordia diagram, covariances and uncertainty ellipses have been calculated analogous to Boelrijk et al. (1979). - ii

From four of the samples two aliquots of the whole-rock powder were analyzed separately for their lead isotopic composition. Differences between the isotopic ratios of the duplicate runs varied between 0.06% and 1.6%, the highest values for the 2oaPb/2OSPb ratios of the samples with low lead contents. Most of the difference can be accounted for by a mass discrimination effect. In view of these differences, the maximum errors of the isotopic ratios may be slightly optimistic for those samples containing less than 1 yg/g of lead.

Lines through the data-points were calculated in the same way as for the zircons.

- 98 - £.„.. ^1

REFERENCES ANDERSON, G.M.. 1978: Uncertainties in calculations involving thermodynamic data. In: Short course in application of thermodynamics to petrology and ore deposits. (Ed,: H.J. Greenwood). Miner. Ass. Canada, 199-215. ARDEN, J.W. & GALE, N.H., 1974: New electrochemical technique, for the separation of Pb at trace levels from natural silicates. Anal. Chem. 46, 2-9. BARNES, I.L., MURPHY, T.J., GRAMLICH, J.W. & SHIELDS, W.R., 1973a: Lead separation by anodic deposition and isotope ratio mass spectrometry of microgram and smaller samples. Anal. Chem. 45, 1881-1884. BARNES, I.L., GARNER, E,L,, GRAMLICH, J.W., MOORE, L.J., MURPHY, T.J., MACKLAN, L.A., SHIELDS, W.R,, TATSUMOTO, M. & KNIGHT, R.J., 1973b: Determination of lead, uvanium, thorium and thallium in silicate glass standard materials by isotope dilution mass spectrometry. Anal. Chem. 45, 880-885. BOELRIJK, N.A.I.M., KUIJPER, R.P. & WIELENS, J.B.W., 1979: Expressions for the calculation of error ellipses. Appendix 3, Verhandeling Nr. 4 ZWO Lab, voor Isotopen-Geologie, Amsterdam, 82-87. CALSTEREN, P,W.C. VAN, 1978: Geochemistry of the polyraetamorphic mafic-ultramafic complex at Cabo Ortegal (NW. Spain), Lithos 11, 61-72. CALSTEREN, P.W.C, VAN, BOELRIJK, N.A.I.M., HEBEDA, E.H., PRIEM, H.N.A., DEN TEX, E., VERDURMEN, E.A.TH. & VERSCHURE, R.H., 1979: Isotopic dating of older elements (including the Cabo Ortegal mafic-ultramafic complex) in the hercynian orogen of NW, Spain: manifestations of a presumed early Paleozoic mantle-plume. Chem. Geol. 24, 35-56. CANTAZARO, E.J., MURPHY, T.-J., SHIELDS, W.R. & GARNER, E.L., 1968: Absolute isotopic abundance ratios of common, equal-atom and radiogenic lead isotopic standards. J. Res. NBS-A, 72-A, 261-267. DEKKER, A.G.C., 1978: Amphiboles and their host rocks in the high-grade metamorphic Precambrian of Rogaland/Vest-Agder, SW. Norway. Ph.D. thesis, State University at Utrecht, The Netherlands. IJLST, L., 1973a: New diluents in heavy liquid mineral separation and an improved method for the recovery of the liquids from the washings. Amer. Miner. 58, 1084-1087. IJLST, L., 1973b: A laboratory overflow-centrifuge for heavy liquid mineral separation. Amer. Miner. 58, 1088-1093. KR0GH, T.E., 1973: A low-contamination method for hydrothermal decomposition of zircon and extraction of U and Pb for isotopic age determinations. Geochim. Cosmochim. Acta 37, 485-494. RALSTON, O.C, 1961: Electrostatic separation of mixed granular solids. Elsevier, Amsterdam. SHAPIRO, L., 1967: Rapid analysis of rocks and minerals by a single-solution method. U.S. Geol. Survey Prof. Pap. 575B, 187-191. SHIELDS, W.R., 1966: Nat. Bur. Stand. (U.S.) Technical Note No. 227. TERA, F. & WASSERBURG, G.J., 1975: Precise isotopic analysis of lead in picomole and subpicomole quantities. Anal. Chem. 47, 2214-2220. VERDURMEN, E.A.TH., 1977: Accuracy of X-ray fluorescence spectrometric determination of Rb and Sr concentrations in rock samples. X-Ray Spectr. 6, 117-122. VERSCHURE, R.H. & IJLST, L., 1969: An asymmetrically vibrating unit for the Frantz magnetic separator. Reports on investigations 1968/69, ZWO Laborato- rium voor Isotopen-Geologie, Amsterdam, 90. WIELENS, J.B.W., BELLE, J.C. VAN, BOELRIJK, N.A.I.M. & KUIJPER, R.P., 1979: Chem- ical procedures for U-Pb dating. Appendix 2, Verhandeling No. 4, ZWO Labora- torium voor Isotopen-Geologie, Amsterdam, 80-81. YORK, D., 1966: Least-squares fitting of a straight line. Can. J. Phys. 44, 1079-1086. YORK, D., 1967: The best isochron. Earth Planet. Sci. Lett. 2, 479-482. YORK, D., 1969: Least-squares fitting of a straight line with correlated errors. Earth Planet. Sci. Lett. 5, 320-324.

- 99 - SAMENVATTING

Westelijk Galici'è maakt deel uit van het noordelijk deel van de centrale zone van het Hesperisch Massief, dat voornamelijk bestaat uit supracrustale successies van Laat Precambrische tot Laat Paleozoische ouderdom, en uit Vroeg en Laat Paleozo- ische granieten en migmatieten. Vroeg Paleozoische katazonale gesteentecomplexen, door breuken begrensd, worden uitsluitend aangetroffen in westelijk Galicië en noordelijk Portugal, In dit proefschrift worden de resultaten van een onderzoek aan de U-Pb systemen van zirkonen, monazieten en "whole-rock" monsters uit het Pa- leozoische grondgebergte van westelijk Galicië behandeld.

De U-Pb daca, verkregen van zirkonen en monazieten uit zes orthogneizen, een paragneis en twee granieten, worden behandeld in hoofdstuk II. In het concordia-diagram vertonen de U-Pb systemen van de zirkoon-series een grote spreiding in bovenste en onderste snijpunten. Aan de meeste snijpunten wordt géén geochronologische beteke- nis toegekend. Zij worden geïnterpreteerd met behulp van een tnodel waarin de U-Pb evolutie verschillende stadia heeft doorlopen: variabel lood-verlies uit zirkonen met discordante U-Pb systemen. De primaire discordantie wordt toegeschreven aan menging van een nieuwe zirkoongeneratie met een geringe zirkooncoraponent van hoge- re ouderdom. De ouderdom van beide zirkooncomponenten is alleen in de zirkonen van de twee orthogneis-monsters van het Mellid complex bewaard gebleven, welke klaar- blijkelijk in zeer geringe mate verstoord zijn door later lood-verlies. De zirkoon- series van deze monsters geven bovenste snijpunten bij 2548 t\l\ Ma en 2267 ±32 Ma, welke worden geïnterpreteerd als een benadering van de ouderdom van de zirkooncom- ponent die is geërfd van de metasedimenten waaruit de granitische oorsprongsgesteenten ontstonden. De onderste snijpunt-ouderdommen van 482± 12 Ma en 459± 11 Ma worden geïnterpreteerd als een benadering van het tijdstip van zirkoon-groei tijdens de plaatsname en stolling van het granitisch magma. De zirkoon-series van de overige vier orthogneizen en de twee granieten vertonen een variabele mate van later lood- verlies; de intensiteit hiervan werd mogelijk beïnvloed door de chemische samen- stelling van het gesteente. De ouderdom van dit lood-verlies kon niet worden vast- gesteld, maar het moet jonger zijn dan 300 Ma; mogelijk is lood-verlies opgetreden als gevolg van (sub)recente verwering.

De zirkoon-serie van het paragneis-monster geeft een bovenste snijpunt bij 2272 ty9 Ma. Dit wordt geïnterpreteerd als de ouderdom van het achterland van de Masanteo groep, de oudste sedimentaire successie. Een maximale sedimentatie-ouderdom van ongeveer 1.5 Ga kan worden afgeleid uit de Rb-Sr data. De ouderdom van het onderste snijpunt, 476 ±12 Ma, wordt beschouwd als een benadering van het begin van de eerste fase van granuliet-facies metamorfose, die de gesteenten van de katazonale complexen heeft beïnvloed en waarmee het ontstaan van de granitische oorsprongsgesteenten van de orthogneizen wordt gerelateerd.

Een concordante monaziet-ouderdom van 471 Ma van hetzelfde paragneis-monster wordt eveneens geïnterpreteerd als overeenkomend met het begin van de granuliet-facies metamorfose. Monazieten van de twee monsters van Laat Paleozoische granieten vertonen zowel concordante als discordante U-Pb systemen; de concordante monazieten geven ouderdommen van ongeveer 300 Ma.

Rb-Sr gegevens van "whole-rocks" van vier van de orthogneizen waarvan zirkoon U-Pb data zijn verkregen, worden behandeld in hoofdstuk II. De gegevens komen overeen met de eerder gepubliceerde Rb-Sr ouderdommen van deze gesteenten: twee van 470-460 Ma en één van ongeveer 400 Ma. Daarenboven is een nieuwe Rb-Sr "whole-rock" isochron ouderdom van 450±25 Ma bepaald voor de orthogneizen van het Sobrado/Teijeiro complex. De ouderdommen van ongeveer 470 Ma (Rb-Sr en onderste snijpunten van U-Pb zirkoon-systemen) worden geïnterpreteerd als een benadering van het tijdstip van vorming en plaatsname van de granieten. Voor de Mellid gneizen komen de onderste snijpunten van de U-Pb zirkoon-systemen overeen met een ouderdom van 480-460 Ma, in tegenstelling tot de corresponderende Rb-Sr ouderdom van 400 Ma; dit verschil

- 100 - wordt geïnterpreteerd als een indicatie voor langdurige migratie van Sr isotopen na de intrusie van de gesteenten, die op een dieper niveau van de korst heeft plaatsgevonden dan die van de granitische voorlopers van de overige orthogneizen.

De petrografie van de hooggradig metamorfe gesteenten van het Sobrado/Teijeiro complex wordt beschreven in hoofdstuk III, tesamen met de U-Pb data van "whole- rocks" van eclogieten en granulieten van de Sobrado/Teijeiro en Cabo Ortegal complexen. De mineraal-relaties» met name de relatieve datering van clinopyroxeen- ontmenging, geven argumenten voor een tijdsinterval tussen de condities van de eclogiet en die van de granuliet-facies. Tegen de achtergrond van de metamorfe geschiedenis worden de U-Pb data geïnterpreteerd als het resultaat van een uit vier stadia bestaande ontwikkeling. Voor het tijdstip van plaatsname van deze gesteenten is een minimale ouderdom van ongeveer 1.0 Ga berekend, hetgeen een minimum-ouderdom voor de sedimentatie van de Massnteo groep geeft.

De supra- en infracrustale geschiedenis van het Hesperisch Massief is samengevat in hoofdstuk IV. Op grond van deze gegevens wordt geconcludeerd dat de geologische evolutie zich geheel intracratonisch heeft voltrokken tijdens één dynamo- thermale cyclus, die ten minste 600 Ma geleden is begonnen en tot ongeveer 280 Ma geleden heeft voortgeduurd. Eenraantel-pluim/aulacogeen mode l wordt uitge- werkt als verklaring voor de geologische evolutie van het Hesperisch Massief.

- 101 -