Journal of Petrology Advance Access originally published online on July 1, 2004
Journal of Petrology 2004 45(8):1613-1629; doi:10.1093/petrology/egh026
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Journal of Petrology 45(8) © Oxford University Press 2004; all rights reserved
A Sequence of Pan-African and Hercynian Events Recorded in Zircons from an Orthogneiss from the Hercynian Belt of Western Central Iberia—an Ion Microprobe U–Pb Study
1 TECTONICS SPECIAL RESEARCH CENTRE, SCHOOL OF EARTH AND GEOGRAPHICAL SCIENCES M004, THE UNIVERSITY OF WESTERN AUSTRALIA, CRAWLEY, WA 6009, AUSTRALIA
2 CENTRE FOR MICROSCOPY AND MICROANALYSIS M010, THE UNIVERSITY OF WESTERN AUSTRALIA, CRAWLEY, WA 6009, AUSTRALIA
3 DEPARTAMENTO DE GEOLOGÍA, UNIVERSIDAD DE SALAMANCA, FACULTAD DE CIENCIAS, PLAZA DE LA MERCED, E 37008 SALAMANCA, SPAIN
RECEIVED NOVEMBER 8, 2002; ACCEPTED FEBRUARY 23, 2004
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL SETTINGPIEDRAHITA ORTHOGNEISS FORMATIONZIRCON ION MICROPROBE U-Pb...INTERPRETATIONCONCLUSIONREFERENCES |
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Sensitive high-resolution ion microprobe U–Pb dating shows that a biotite orthogneiss from the Hercynian belt of western central Iberia contains 1000–300 Ma zircon. Older, 1000–570 Ma ages within this range represent inherited, detrital material among which four age components may be recognized:
980 Ma, 830 Ma, 616 ± 10(2
) Ma and 582 ± 5(2) Ma. This inherited zircon commonly forms cores that are surrounded by rims yielding Late Pan-African ages, identical to those found in slender, prismatic, and some stubbier, bi-pyramidal, euhedral crystals. This is the predominant type of zircon with an average age of 546 ± 3(2) Ma, thought to have been formed during the main magmatic crystallization stage of the granitic protolith of the gneiss. Local deuteric replacements of magmatic zircon yield a virtually identical average age of 547 ± 5(2) Ma, suggesting rapid magmatic cooling, typical of shallow intrusive settings. Many zircon crystals have very thin, low-Th/U rims with an age of 315 Ma, suggested to represent the gneissification of the granitic rock during the Hercynian orogeny. The abundance in the gneiss body of Al-rich restitic material and inherited, detrital zircon suggests that the granitic magma was formed by anatectic melting of a meta-sedimentary source rock complex. The age of the youngest inherited, detrital zircon constrains the sedimentation age of the (youngest parts of the) anatectic source rock complex to the Late Neoproterozoic (<582 ± 5 Ma) and leaves a maximum period of 40 Myr for metamorphism and anatexis of the source rock complex, and migration and intrusion of the granitic magma. Among the inherited, detrital zircon the 616 ± 10(2) Ma and 582 ± 5(2) Ma ages are by far the most abundant, and might be derived from the West African craton, where such ages are common for main-stage Pan-African complexes. The older, 980 Ma and 830 Ma inherited zircon ages are absent or rare in the West African craton and a derivation from the Amazon craton (Grenvillian, Sunsas orogens), where such ages are common, is an alternative. KEY WORDS: SIMS U–Pb Zircon dating; inherited, magmatic, deuteric and metamorphic granitic zircon; anatectic granite; Iberian Hercynian belt
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGPIEDRAHITA ORTHOGNEISS FORMATIONZIRCON ION MICROPROBE U-Pb...INTERPRETATIONCONCLUSIONREFERENCES |
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Basement complexes generally consist of polygenetic metamorphic and magmatic complexes that are difficult to date by conventional U–Pb methods using analysis of bulk zircon samples by isotope dilution thermal ionization mass spectrometry (ID-TIMS). This approach involves homogenization of samples prior to analysis, and does not address the typically complex internal structures of the zircon crystals. As a result analytical ages may be ambiguous owing to the pooling of ages of zircon crystal domains grown during events of widely differing ages. Such results may be difficult to relate to actual geological events.
In contrast to ID-TIMS, the ion microprobe (secondary ion mass spectrometry; SIMS) method is capable of dating 10–30 µm sized domains within single zircon crystals. Analysis takes place in mounted, polished crystals and this allows for detailed cathodoluminescence (CL), backscattered electron (BSE) and secondary electron (SE) image control to position the ion beam. This allows determination of not only the age of magmatic crystallization, but also protolith ages (inherited zircon cores) and ages of post-magmatic metamorphic events (zircon overgrowths and replacements). This renders this method particularly effective in establishing a time frame for the petrogenesis of polygenetic, meta-magmatic rocks.
The present study concerns SIMS (sensitive high-resolution ion microprobe; SHRIMP) U–Pb dating of zircons from a rather common basement lithology, an orthogneiss, from the Hercynian orogenic belt of the Iberian Peninsula, the geological setting of which is well known. Most of these zircons are internally complex and contain domains of different ages. The aim of the study is to demonstrate the detailed chronological information on the pre-magmatic, magmatic and metamorphic petrogenesis of the rock that can be obtained by SIMS zircon dating, provided, however, it is combined with detailed CL (BSE, SE) image analysis.
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GEOLOGICAL SETTING |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGPIEDRAHITA ORTHOGNEISS FORMATIONZIRCON ION MICROPROBE U-Pb...INTERPRETATIONCONCLUSIONREFERENCES |
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The Hercynian orogenic belt occupies a major part of the Iberian Peninsula (Fig. 1). It is surrounded and overlain by Mesozoic and Tertiary cover sequences of the Iberian Meseta, and in the south by the Alpine belt of the Betic Cordilleras. The Hercynian orogeny (
400–280 Ma) caused variable metamorphic and anatectic reworking of older basement rocks. Information on the origin and age of this pre-Hercynian Iberian basement is of particular importance for plate tectonic reconstructions involving the break-up of Rodinia, the evolution of the Pan-African orogeny (750–550 Ma; Murphy & Nance, 1991
), the subsequent assembly of Gondwanaland, and the evolution of the peri-Gondwana terranes of Avalonia, Cadomia and Iberia (Murphy et al., 2000). The present study concerns one of the pre-Hercynian basement complexes.
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Based on structural and lithological criteria, the Iberian Hercynian belt has been subdivided into a number of east–west- to SE–NW-trending zones (Fig. 1; Lotze, 1945
; Julivert et al., 1972; Díez Balda et al., 1990). The major Central Iberian Zone (CIZ; Fig. 1) has been subdivided into a northern domain characterized by large-scale recumbent folds and a southern domain characterized by vertical folds (Díez Balda et al., 1990). The boundary between these two CIZ domains is typically obscured by Late Hercynian, post-tectonic granitic intrusions (Díez Balda et al., 1990). The lithostratigraphy and tectono-metamorphic and magmatic evolution of the two domains has been discussed in a number of recent papers (e.g. Azor et al., 1992; de San José et al., 1992; Díez Balda & Vegas, 1992).
Hercynian metamorphic/anatectic reworking in the southern CIZ domain is highly variable. Large areas are occupied by Upper Neoproterozoic–Lower Cambrian sedimentary rocks, which show only very weak metamorphic overprinting (Ugidos et al., 1997a, 1997b; Valladares et al., 2000). In other areas, Hercynian metamorphism seems to have reached high grade, culminating in anatexis and development of nebulitic migmatites (Bea et al., 1990; Martinez et al., 1990; Ugidos, 1990).
The Piedrahita–Béjar area within the southern CIZ domain (Figs 1 and 2) was selected for this study because of its large domains of orthogneiss of pre-Hercynian parentage. As the largest of these bodies (50 km2) occurs west and SE of Piedrahita village, we designate these rocks as ‘Piedrahita orthogneiss formation’. The term ‘formation’ is a general, non-genetic lithostratigraphic term defining a mappable rock unit, irrespective of its origin. It is thus applicable to sedimentary, volcanic, plutonic and metamorphic rock units (American Commission on Stratigraphic Nomenclature, 1970). Younger Hercynian granites and migmatites surround the Piedrahita gneiss bodies (Fig. 2), although contacts are not exposed and most may be of tectonic nature.
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PIEDRAHITA ORTHOGNEISS FORMATION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGPIEDRAHITA ORTHOGNEISS FORMATIONZIRCON ION MICROPROBE U-Pb...INTERPRETATIONCONCLUSIONREFERENCES |
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The Piedrahita formation consists of calc-alkaline (Irvine & Baragar, 1971
; see Bea et al., 1990, for chemical data), granitic to granodioritic gneisses of variable appearance. Grey biotite augen gneisses are widespread, and contain up to several centimetre-sized K-feldspar augen set in a foliated matrix that has 15–20 vol. % biotite; sillimanite- and cordierite-bearing types are common (present in about a third of the thin sections). More even- and finer-grained varieties are common, as are more leucocratic types with only 5–10 vol. % biotite. The foliation is defined by an alternation of 1–3 mm wide, biotite-rich schlieren with 1–2 cm wide, quartzo-feldspathic lenses; individual biotite crystals are parallel to, and further sustain, the foliation. Finer-grained varieties may develop foliation-parallel layering. The poor degree of exposure does not allow information on contact or volumetric relations between the various rock types within the gneiss formation. Thin sections reveal a general grain size of a few millimetres and the following mineralogy (estimated vol. % in parentheses): apatite (<1), biotite (5–25), chlorite (0–<1), cordierite (0–8), K-feldspar (5–35), opaque material (<1), plagioclase (20–50), quartz (20–35), sillimanite (0–3), white mica (1–5), zircon (<1). Metamorphic recrystallization related to the gneissification resulted in an overall xenoblastic texture with only the mica crystals commonly showing subhedral habit. Locally in the quartzo-feldspathic schlieren, magmatic textures are preserved, consisting of larger quartz crystals accommodating euhedral plagioclase habits, and blasto-megacrystic ‘augen’ that comprise centimetre-sized, subhedral, perthitic K-feldspar crystals surrounded by a mosaic consisting mainly of smaller K-feldspar crystals. The latter represent the recrystallized rims of magmatic megacrysts. Plagioclase crystals are mostly metamorphic, anhedral and texturally concordant to the foliation. Some biotite-rich schlieren may contain dense fibroblastic aggregates of sillimanite. Elongated, anhedral cordierite crystals grew at the expense of these biotite–fibrolite aggregates; many cordierite crystals enclose swarms of fibrolite needles, reflecting local excess of sillimanite. In some rocks, a different type of cordierite occurs, forming independent, equant crystals, commonly also with inclusions of sillimanite needles. These two types of cordierite appear to represent two different genetic stages; the first might be related to the restitic biotite–sillimanite schlieren, the second to the main metamorphic stage (gneissification). White mica occurs in two modes. In some rocks it forms millimetre-sized crystals whose textural setting suggests an origin during the magmatic or main metamorphic stage. More commonly, white mica (plus ilmenite) is in much smaller crystals replacing biotite; these replacive aggregates may have formed during late-stage alteration. Obvious low-temperature, late-stage alteration products are sericite replacing plagioclase (and K-feldspar), pinite replacing cordierite, and chlorite replacing biotite.
The mineralogical and textural relations suggest the following petrogenetic, paragenetic scheme: (1) anatectic stage (restitic material: biotite, sillimanite and ?cordierite); (2) magmatic stage (crystallization products of the melt of the intruded anatectic magma: biotite, K-feldspar, plagioclase, quartz, ?white mica); (3) main metamorphic stage (gneissification: biotite, cordierite, K-feldspar, plagioclase, quartz, white mica); (4) low-temperature alteration (chlorite and white mica, mainly pinite and sericite).
The sample that was selected for the present study (sample 01Z08; 40°27'4''N, 5°20'19''W; Fig. 2) was collected along the road from Piedrahita to Barco de Avila. It is a grey biotite gneiss containing <2 cm sized feldspar augen set in a foliated matrix that consists mainly of 1–3 mm crystals of quartz, feldspar and biotite. The sample has neither sillimanite nor cordierite, suggesting that it represents a less pelitic part of the pre-anatectic source rock complex.
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ZIRCON ION MICROPROBE U–Pb AGES |
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Sample preparation and analytical procedure
Sample 01Z08 represents an
75 kg block of gneiss that was broken up and cleaned by hammer chipping in the field. The resulting 7 kg sample consisted of 2–10 cm fragments, which were cleaned in an ultrasonic bath with de-ionized water, then dried and crushed in a jaw crusher to a size of 3–10 mm. Approximately 1 kg of this material was crushed in a swing mill in several steps, interspersed with screening to <0·5 mm. These <0·5 mm crushings were panned, and the resulting heavy fraction was purified by heavy liquid (methylene iodide) and magnetic separation. Care was taken not to bias the zircon separates by removing the more magnetic grains. The yield was 1500 zircon grains, all of which were cast in epoxy resin (Fig. 3); no hand sorting was applied. The mount was polished to expose the crystals' interiors. Prior to analysis, the mount was cleaned and vacuum-coated with gold.
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Measurements of U, Th, and Pb were made on the SHRIMP II ion microprobe at Curtin University, Perth. Decay constants employed are those recommended by Steiger & Jäger (1977)
. U–Th–Pb ratios and absolute abundances were determined relative to the CZ3 standard zircon (564 Ma, 206Pb/238U = 0·09143, 550 ppm 238U; Pidgeon et al., 1984; Nelson, 1997), using operating and data-processing procedures similar to those described by Compston et al. (1984), Claoué-Long et al. (1995) and Nelson (1997). The data were collected in three analytical sessions with zircon standard calibration uncertainties (1) of 0·29, 0·33 and 0·29%, respectively, which are included in the uncertainties on the ages listed in Table 1 and elsewhere in the paper.
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Measured compositions were corrected for common Pb using non-radiogenic 204Pb. Prior to analysis, each site was cleaned by rastering the ion beam over the area for 3 min. Subsequent 204Pb counts for most analyses remained low and showed no tendency to decrease over the course of a 15 min analysis, suggesting that common Pb in these crystals is inherent to the mineral rather than surface-related. In most cases, corrections are sufficiently small to be insensitive to the choice of common Pb composition, and an average crustal composition (Cumming & Richards, 1975
) appropriate to the age of the mineral was assumed.
Zircon morphology, cathodoluminescence and backscattered electron zoning
Cathodoluminescence (CL), backscattered electron (BSE) and secondary electron (SE) imaging techniques were employed to locate the ion beam within the zoning pattern of the zircon crystals. A total of 700–800 crystals, about half of the total mount, were imaged using the Jeol 6400 SEM at the Centre for Microscopy and Microanalysis at the University of Western Australia. After each analytical session, the actual positions of the analytical craters were verified using additional BSE and SE imaging.
The mounted zircons cover a rectangular area of 11 mm x 2 mm, about half of which is portrayed in Fig. 3. To estimate the proportions of different zircon grain types, the CL–BSE–SE survey included a continuous band along one side of the mount, containing about 600 zircon grains (40% of the total). Zircon grains in sample 01Z08 vary in size from 30 to 500 µm, and may be equant, bi-pyramidal to short prismatic, or long prismatic (aspect ratios up to seven). In transmitted light, most grains are colourless to pinkish translucent. A few are slightly turbid in their central parts. Discrete cores were not observed in transmitted light. In contrast, CL and BSE imaging, typically at 200x to 500x magnification, shows a wide variety of morphologies and internal structures (Table 1, column 2, and Fig. 4; grain numbers given in Fig. 4 are those listed in column 1 in Table 1). Many crystals are broken, obviously during the zircon sample preparation (Fig. 3; Fig. 4, grains 20 and 46). Unbroken crystals are typically euhedral to subhedral, showing an overall euhedral shape, with both euhedral and some rounded crystal outlines (Fig. 4). Perfectly euhedral and entirely anhedral crystals are rare. The latter are typically rounded and only a few more irregularly shaped grains have been observed.
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Crystals less than 50 µm in size make up only a few per cent of the mount. All appear equant on the polished surface (Fig. 3). Zircons between 50 and 100 µm (
30% of the number of crystals in the mount) are equant to short prismatic with aspect ratios of up to two; a typical dimension would be 80 µm x 50 µm (Fig. 3; Table 1; Fig. 4, grains 4, 18 and 35). In the grain size range of 100–200 µm (60%), prismatic crystals with aspect ratios of 3–4 dominate; a typical dimension is 150 µm x 40 µm; other crystals are shorter prismatic or bi-pyramidal, with typical dimensions of 180 µm x 70 µm (Fig. 3; Table 1, column 2; Fig. 4, grains 6, 7, 8, 9, 12, 14, 15, 28, 29, 31, 49, 50). Crystals in the range 200–500 µm make up 5–10%. Most of these larger grains are slender prismatic with aspect ratios of 4–7; a typical dimension would be 300 µm x 60 µm (Fig. 3; Table 1, column 2; Fig. 4, grains 20, 24, 30, 38, 40, 51, 52, 53). Fragments of broken prismatic crystals are present also in this size fraction (Fig. 3; Fig. 4, grain 20), indicating that still larger crystals are present in the gneiss, but were excluded from the mount by the final sieve screening at 500 µm during sample preparation. CL images show a large variety of structures within the zircon crystals (Fig. 4; Table 1, column 3). Many crystals, both small and large, contain well-defined cores, surrounded by rather wide rims that typically show euhedral, oscillatory zoning (Fig. 4, grains 4, 6, 7, 8, 9, 14, 15, 18). The zoning pattern in the rim is discordant to the outline and zoning of the cores. On irregular, angular cores the rim growth starts by filling out the irregularities before building the euhedral oscillatory zones (Fig. 4, grains 6, 7, 15). The shapes of the cores vary from irregular, angular (Fig. 4, grains 4, 6, 7, 8, 15, 18) to rounded (Fig. 4, grains 9, 14; the latter crystal appears to have been broken after rounding, but before formation of the rim), and in some cases they are roughly euhedral (Table 1, column 3, grains 2, 10, 17). These discordant cores were seen in equant and stubby prismatic crystals, whereas they were not found in more elongate prismatic crystals (Fig. 4, grains 20, 24, 30, 38, 51). Examples of more equant crystals without a discordant core are shown in Fig. 4 (grains 28, 29, 31, 35, 49, 50); some of these crystals (grains 28, 31, 49) have a concordant core.
Many crystals exhibit very thin, <5–30 µm wide, discontinuous, dark CL outer rims (Fig. 4, grains 6, 7, 14, 15, 24, 28, 29, 38, 46). A few crystals have domains of very dark and homogeneous CL zircon replacing oscillatory zoned zircon (Fig. 4, grains 40, 52, 53). This type of dark CL zircon does not form rims, but occupies crystal tips and other irregular domains, typically leaving the earlier zircon with an embayed morphology. On the basis of CL imagery alone it may be difficult in some crystals to discriminate between these two types of dark-CL rim or overgrowth, e.g. in (Fig. 4) crystal 6, upper and lower tip, and crystal 7, lower tip; crystal 46 may have both types. However, there is a distinctive difference in age (see below).
Analytical results
Fifty-four crystals were analysed, nine of them in several places, resulting in 71 analyses. The data, including a summary of morphological and zoning characteristics of each analysed crystal, are listed in Table 1 and plotted in Tera–Wasserburg concordia diagrams (Figs 5 and 6) and a probability density diagram (Fig. 7). Grain numbers listed in Table 1 (column 1) are followed by a decimal point and the analytical spot number in that specific grain. Figure 4 shows CL images of 24 representative crystals featuring location of the analysis craters and resulting ages; grain numbers given are those listed in Table 1. Spots for analysis in the zircons were selected to avoid the abundant cracks, holes and inclusions that may act as loci for Pb loss. The SE and BSE images employed to guide the analyses highlight these crystal imperfections and show that only a small number of crystals have sufficient ‘clean’ areas available to analyse both cores and rims. CL images (Fig. 4) do not reveal these crystal defects.
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1000–570 Ma, discordant cores and rounded crystals
Analyses were conducted of 16 discordant cores (Table 1, grains 1, 2, 4–11, 13–18, of which grains 4, 6, 7, 8, 9, 14, 15, 18 are shown in Fig. 4), the zoning pattern of which is truncated by rather wide oscillatory rims, and of two subhedral-rounded whole crystals (Table 1, grains 3, 12; the latter is shown in Fig. 4). Some cores or crystals were analysed in several places, resulting in 21 ages, varying between
1000 and 500 Ma (Table 1). Among 18 of these ages, four distinct groups can be distinguished (Figs 5, and 7a and b) within which all 238U/206Pb values agree to within analytical precision (Fig. 7a and b). One age component is around 980 Ma [978 ± 11(1) Ma; Table 1, spots 17.1, 18.1; grain 18 shown in Fig. 4], another around 830 Ma [829 ± 15(1) Ma; Table 1, spots 1.2, 16.1], a third has a weighed average of 616 ± 10(2) Ma, MSWD = 0·96 (n = 5; Table 1, spots 11.1, 12.1, 13.1, 14.1, 15.1; grains 14 and 15 shown in Fig. 4) and a fourth has a weighed average of 582 ± 5(2) Ma, MSWD = 1·04 (n = 9; Table 1, spots 3.2, 3.3, 4.1, 5.1, 6.1, 7.1, 8.1, 9.1, 10.1; grains 4, 6, 7, 8 and 9 shown in Fig. 4). Measured U and Th contents confirm the heterogeneous character of this type of zircon. U ranges from 35 to 1000 ppm, Th from 12 to 1200 ppm and Th/U varies between 0·1 and 1·9, without apparent systematic differences between the four age groups. Common Pb is very low; the proportion of common 206Pb in total measured 206Pb (f206 in Table 1) ranges from 0·02 to 0·5%.
Three analyses (spots 1.3, 2.1, 3.1; marked with an asterisk in Table 1, column 1) yielded 238U/206Pb ages that are significantly younger than the youngest, 582 ± 5(2) Ma age cluster. These results are believed to reflect loss of radiogenic Pb from the analysed sites; post-analytical imaging revealed cracks transecting two of the craters. These three analyses will not be considered further.
546 ± 3(2) Ma, mainly euhedral, prismatic grains
Forty analyses were conducted of euhedral to subhedral, oscillatory zoned zircons of highly variable shape and size (Table 1; grains 19–51; Fig. 4 shows 11 of these crystals, grains 20, 24, 28, 29, 30, 31, 35, 38, 49, 50, 51). Most are prismatic and 200–500 µm long with aspect ratios of 3–7. Smaller prismatic crystals are 100–150 µm long and stubbier with aspect ratios of 2. A few equant crystals are 50–120 µm across. A further two analyses are of oscillatory rims surrounding older cores (Table 1, spots 1.1, 7.2; grain 7 shown in Fig. 4).
Of these 42 analyses, 38 have 238U/206Pb ratios that agree to within analytical precision and yield a weighted average age of 545·6 ± 2·6(2) Ma (MSWD = 1·15) (see Figs 6, and 7a and c); the remaining four analyses (see below) are marked with an asterisk in Table 1, column 1. U and Th contents are variable (95–510 ppm and 15–400 ppm, respectively) resulting in very variable Th/U values (0·06–1·13) suggesting growth in a chemically heterogeneous environment; U and Th contents cannot be related to structures as seen on grains in Fig. 4. Common Pb is low; values for f206 range up to 0·8%, and have a median of 0·12%. No correlation exists between 238U/206Pb age and Th concentration or Th/U, although there is a weak negative correlation between 238U/206Pb age and 238U concentration (R = 0·51, Fig. 8), which suggests that some of the higher-U zircons have lost minor amounts of radiogenic Pb. However, if 11 analyses above an arbitrary limit of 300 ppm 238U are excluded, the mean age becomes 547·1 ± 3·2 Ma (2), which is not a significant increase. Therefore, although we recognize the possibility that minor Pb loss may have occurred in some grains of this age sample, we prefer the average of all accepted data, at 546 ± 3(2) Ma, as the best estimate of the age of this type of zircon.
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Four analyses (Table 1, spots 20.1, 21.1, 34.1, 44.1, marked with an asterisk in column 1; grain 20 shown in Fig. 4) yielded 238U/206Pb ages that are significantly different from the 546 ± 3(2
) Ma average for the group of 38 ages. One age (spot 20.1) is significantly younger and may reflect radiogenic Pb loss; however, SEM images do not explain this and the CL image in Fig. 4 shows a crystal in no way different from other crystals of this age group. Three ages are significantly older (spots 21.1, 34.1, 44.1); it is possible that parts of an older zircon core were included in the analysed material, although this could not be discerned in CL images, or these analyses might record some local chemical or isotopic redistribution. These four analyses will not be considered further.
547 ± 5(2) Ma, low Th/U replacements
Four analyses (Table 1, spots 40.1, 52.1, 53.1, 54.1) were made in very dark CL zircon, which appears to have replaced earlier zircon [p. 1622; Fig. 4, grains 40, 52, 53; grain 46 shows a very dark CL replacement (no age available) in a 553 ± 10(1) Ma core]. The weighed average for these four ages is 546·5 ± 5·2(2) Ma (MSWD = 1·86). This average age is statistically indistinguishable from the 546 ± 3(2) Ma age derived above for the mainly euhedral, prismatic grains. The U content in this type of zircon is high (800–1000 ppm) and Th content low (7–21 ppm) resulting in very low Th/U values (0·007–0·02). In one crystal (Fig. 4, grain 40) the replaced, oscillatory zoned, relict zircon could be analysed (Table 1, 40.2) yielding a 550 ± 7(1) Ma age, indicating it belongs to the 546 ± 3(2) Ma zircon group.
315 Ma, thin, discontinuous, discordant, low Th/U overgrowths
Although many zircon crystals show thin, dark CL outer rims (p. 1622; Fig. 4, grains 6, 7, 14, 15, 24, 28, 29, 38 and 46), only in one crystal (Fig. 4, grain 46) was such a rim sufficiently wide and free of inclusions, cracks and holes to allow analysis. The CL image in Fig. 4 shows that the crystal consists of a subhedral, oscillatory-zoned core [553 ± 10(1) Ma, spot 46.2, reported above], which shows a homogeneous, dark CL replacement (no age available) of the type described in the preceding section. The discontinuous, very dark CL rim of the crystal is <30 µm wide and shows irregular, schlieric zoning (Fig. 4). Two analyses were obtained (spots 46.3, 46.4, Table 1, Fig. 4) yielding a weighed average age of 315 Ma [316 ± 2(1) Ma, MSWD = 0·71]; the remaining rim analysis (spot 46.1, Table 1, Fig. 4) produced a mixed age, as described below. This overgrowth has high U (700–1000 ppm) and very low Th (5 ppm) resulting in very low Th/U values (0·005). Common Pb is 0·02 for one analysis, and 0·19 for the other.
Mixed ages
BSE and SE imaging after analysis showed that inadvertently two shots (Table 1, spots 22.2, 46.1) had straddled a core–rim interface. Spot 46.1 (328 ± 3 Ma) included part of a 553 ± 10 Ma core (Table 1, Fig. 4, grain 46.2) and an 315 Ma metamorphic rim (Table 1, Fig. 4, spots 46.3, 46.4). Spot 22.2 (415 ± 8 Ma) included part of a magmatic core (Table 1, spots 22.1, 22.3) and a thin younger rim. These ages are not considered further.
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INTERPRETATION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGPIEDRAHITA ORTHOGNEISS FORMATIONZIRCON ION MICROPROBE U-Pb...INTERPRETATIONCONCLUSIONREFERENCES |
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Interpretation of the SHRIMP U–Pb ages hinges on the internal structures in the analysed zircons as revealed by CL (and BSE) imaging. Information on these growth structures combined with SHRIMP age data has allowed recognition and dating of four stages of zircon generation (Fig. 7; the older, inherited zircon, was subdivided further into four age clusters, Fig. 7a). Because of the uncertainties on individual zircon ages (
5–35 Ma, 1; Table 1), the electron optical recognition of genetic zircon types is crucial, as it provides a guide to pooling individual ages statistically, and to excluding statistical outliers (seven among 71 ages obtained; see Table 1, column 1).
The relatively large uncertainties on the individual zircon ages (1–7%) suggest heterogeneity in the analysed zircon. As probability density curves and histograms for the various age clusters are symmetrical (Fig. 7), and both younger and older outliers were recognized, this is not merely a matter of Pb loss. Rather it seems that zircon has undergone local, small-scale Pb redistribution (see Zeck & Whitehouse, 2002a).
Late Pan-African intrusion age of the Piedrahita granitic magma
The 546 ± 3(2) Ma weighted average age of 38 analyses of the oscillatory zoned zircon (Figs 6 and 7c) is interpreted as the best estimate for the age of the main magmatic zircon crystallization. The well-defined, symmetrical peak in the relative probability diagram (Fig. 7c) suggests a simple, non-composite age event. The 547 ± 5(2) Ma weighted average of four ages for the dark CL zircon replacing this magmatic zircon is indistinguishable from the main magmatic age. The chemistry of these two zircon phases, however, is distinctly different. The magmatic zircon has U, Th and Th/U values of 95–410 ppm, 15–400 ppm, and 0·06–1·13, respectively, whereas the corresponding values for the replacive phase are 800–1000 ppm, 7–21 ppm, and 0·01–0·02. The pronounced replacement structures, and the marked increase in U and decrease in Th, indicate a post-magmatic, deuteric origin. The very low Th/U values agree with such metamorphic origin (Williams & Claesson, 1987; Zeck & Williams, 2001, 2002; Zeck & Whitehouse, 2002a). The indistinguishable ages for main magmatic and deuteric zircon growth suggest a rapid cooling typical for shallow intrusion levels. This would imply that the 546 ± 3(2) Ma age is identical to or only slightly older than the age of intrusion.
The 0·5–1 cm biotite–sillimanite ± cordierite restitic schlieren occurring in the Piedrahita gneisses formation indicate a sedimentary, pelitic anatectic parent rock (Zeck, 1970, 1992; Munksgaard, 1984). The absence of this Al-rich material at some localities of the gneiss formation suggests that either the restitic material was removed during magma migration or the source rock complex locally was of more quartzo-feldspathic composition. The overall lithological variation of the gneiss formation likewise supports both hypotheses. The large variation in U and Th content (95–510 ppm and 15–400 ppm, respectively) and Th/U (0·06–1·13) in the magmatic zircon crystals indicates a chemically variable environment of zircon crystallization, suggesting a magmatic melt of variable composition. This supports the second option and would imply mingling or mixing of magma batches (just) before intrusion, such as is known from other anatectic complexes (Zeck et al., 1998, 1999; Zeck & Whitehouse, 2002b; Zeck & Williams, 2002).
The 546 ± 3(2) Ma, Latest Neoproterozoic, intrusion age proposed here for the protolith of the Piedrahita orthogneiss formation differs from previously published, conventional U–Pb zircon ages for similar pre-Hercynian granitic bodies from the Central Iberian Zone (CIZ, Fig. 1). Both Middle Ordovician (460–480 Ma; Lancelot et al., 1985; Valverde-Vaquero & Dunning, 2000) and Pan-African ages (600–650 Ma; Lancelot et al., 1985; Wildberg et al., 1989) have been claimed, but no Late Neoproterozoic–Early Cambrian ages (the Precambrian–Cambrian boundary is estimated at 543·3 ± 1 Ma; Bowring & Erwin, 1998). This difference reflects either larger variability in the basement than earlier assumed, or some bias in the previous results. More data are needed to resolve this issue.
Hercynian tectono-metamorphicreworking
Very thin, discontinuous rims that yield ages of 315 Ma (Figs 5 and 7a) represent metamorphic zircon precipitated during the Hercynian tectono-metamorphic reworking. Their very low Th/U (<0·01) agrees with the suggested metamorphic origin (Williams & Claesson, 1987; Zeck & Williams, 2001, 2002; Zeck & Whitehouse, 2002a). Apparently some Zr was mobilized and new zircon was formed during the Hercynian gneissification of the Late Pan-African granitic rock complex. Because only two analyses were obtained, we do not quote a firm uncertainty. The age is in agreement with the 306 Ma age obtained for post-tectono-metamorphic Hercynian, granitic–gabbroic rocks in the area (Zeck et al., in preparation).
Inherited, detrital zircon—time constraints for magma generation and provenance of anatectic source rocks
Texturally discordant zircon cores have outlines and internal zoning patterns that are truncated by the zoning in the surrounding 546 ± 3(2) Ma magmatic zircon, suggesting that these cores are exotic to the magmatic stage. This is confirmed by their clearly older, 1000–570 Ma ages. It is therefore concluded that these cores are inherited from the pre-anatectic parent rock. The few rounded, unrimmed grains in the same age bracket (Table 1, grains 3.1, 12.1; Fig. 4, grain 12) are interpreted as crystals that were shielded from later zircon precipitation by being enclosed within earlier formed mineral grains. This type of zircon has been subdivided into four age groups (± 2): 582 ± 5 Ma, 616 ± 10 Ma, 830 Ma and 980 Ma (Fig. 7a and b). The last two averages are based on only two analyses each, and we have not calculated uncertainties. The two younger average ages are rather well defined (Fig. 7b), but 2 uncertainties still are in the 1–2% range. Because consistent CL characteristics to guide statistical evaluations are lacking for these inherited cores or grains, a relatively large number of analyses is needed to place effective constraints on inherited zircon age clusters. This is a general problem in zircon provenance studies (see Zeck & Whitehouse, 2002a).
The abundance of inherited, detrital cores or grains in the Piedrahita gneiss argues against derivation by assimilation of country rock. The amount of assimilation required would demand very high degrees of superheating of the magma. The wide age spectrum of the inherited zircon, and the rounded and angular shapes of the zircon cores or grains (Fig. 4, grains 4, 6, 7, 8, 9, 12, 14, 15, 18) strongly favour a sedimentary parentage for the pre-anatectic protolith. Deposition of the sedimentary material must postdate the youngest inherited, or detrital zircon. That would imply an age of deposition younger than 582 ± 5 (2) Ma (the youngest inherited zircon component) leaving a maximum time window of 40 Myr for the subsequent metamorphism and anatexis of the source rock complex, and migration and intrusion of the granitic magma.
The two major inherited zircon age components, 616 ± 10(2) Ma and 582 ± 5(2) Ma, correspond to ages found for common igneous rocks related to the late collisional stage of the Pan-African orogeny of the West African craton (e.g. Villeneuve & Cornée, 1994). This part of the North Gondwana stable continental margin has been considered the most plausible source for the voluminous Upper Neoproterozoic–Lower Cambrian sedimentary successions of the Central Iberian Zone (Ugidos et al., 1997a, 1997b; Valladares et al., 2002). The Piedrahita granitic magma might thus have had a pre-anatectic source rock comprising sedimentary material of similar derivation, and its age of 546 ± 3(2) Ma, belonging to the closing stage of the Pan-African orogeny, suggests that it may have intruded into a Pan-African basement consisting of Upper Neoproterozoic–Lower Cambrian sedimentary rocks.
Ages of 980 Ma and 830 Ma found for older inherited zircon in the Piedrahita gneiss body are absent or rare in the West African craton [for support for such older ages, see Goodwin (1991), Murphy & Nance (1991), Saqueque et al. (1992) and Villeneuve & Cornée (1994)]. An alternative, and perhaps more likely, derivation of these older detrital zircon grains was suggested by a study of Fernández-Suárez et al. (2000) of the Neoproterozoic–Lower Cambrian sedimentary sequence in the West Asturian Leonese Zone in NW Iberia (Fig. 1). Those workers reported zircon ages between 0·9 and 1·2 Ga, and suggested provenance from Grenvillian or Sunsas orogenic complexes from the Amazon craton where similar ages seem common (e.g. Keppie et al., 1998).
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CONCLUSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGPIEDRAHITA ORTHOGNEISS FORMATIONZIRCON ION MICROPROBE U-Pb...INTERPRETATIONCONCLUSIONREFERENCES |
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The large variation in size, shape and growth structure of the zircons from the Piedrahita orthogneiss reflects its extended history of zircon generation. Detailed CL, BSE and SE imaging of zircon structures followed by SHRIMP U–Pb analysis allowed discrimination and dating of four genetic stages of zircon growth (±2
): pre-magmatic (980 Ma, 830 Ma, 616 ± 10 Ma and 582 ± 5 Ma), magmatic (546 ± 3 Ma), deuteric (547 ± 5 Ma) and metamorphic (315 Ma). This sequence of ages covers the complete evolution of the Hercynian basement in the western Central Iberian Zone. The study of this rather common basement lithology demonstrates the efficiency of the SIMS method to establish the detailed geochronology of polygenetic metamorphic, magmatic rock complexes. The precision of the weighted average ages obtained is mainly dependent on the number of individual analyses, which is constrained by the analytical time available and the number of zircon domains suitable for analysis. In the present study the relatively large uncertainties on the pre-magmatic, inherited zircon ages are explained by the lack of analytical time, and the rather large uncertainty on the Hercynian age of metamorphic overprinting is due to the scarcity of suitable analytical sites.
The oldest zircon (1000–570 Ma) represents inherited grains from the pre-anatectic rock complex, which melted to produce the granitic magma that intruded to form the granitic protolith for the gneiss formation. Of the four age components recognized (Fig. 7a), the older ones, 980 Ma and 830 Ma, are rare and might be exotic to the West African craton, but the more common, 616 ± 11(2) Ma and 582 ± 5(2) Ma zircon components are compatible with derivation from Pan-African rock complexes of the West African craton.
The 546 ± 3(2) Ma magmatic age represents the main crystallization stage of the anatectic granitic melt. The nature of the restitic material in the granitic rock—mainly inherited zircon cores and biotite–sillimanite (±cordierite) schlieren—indicates a pre-anatectic protolith of sedimentary parentage. As anatectic melting must postdate the youngest detrital zircon component (582 ± 5 Ma), a maximum of 40 Myr remains available for metamorphism and anatexis of the source rock complex and migration and intrusion of the granitic magma.
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ACKNOWLEDGEMENTS |
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The study was supported by the Carlsberg Foundation (grant 45/10-800), the Danish Research Council (SNF grant 9800918), the Spanish Ministry of Science and Technology (MCYT project BTE2002-04241-CO2-O2) and the Gledden Foundation (University of Western Australia). Analyses were conducted using the SHRIMP-II ion microprobe in Perth, operated by a university–government consortium supported by the Australian Research Council. TSRC director John Dodson, UWA Pro Vice-Chancellor Michael Barber, Syd Hall (UWA), John de Laeter, Kevin Rosman, Bob Pidgeon and Allen Kennedy (Curtin University) are thanked for hospitality and gracious support in a difficult SHRIMP situation. We thank John Murphy and Bob Pidgeon and Richard J. Arculus (Canberra) for thorough and constructive reviews. This is Tectonics Special Research Centre publication number 236 and a contribution to IGCP Project 440.
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FOOTNOTES |
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* Corresponding author. Present address: Geological Museum and Institute, Copenhagen University, ØsterVoldgade 10, 1350K Copenhagen, Denmark. E-mail: zeck.frappier{at}wanadoo.fr
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