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Journal of Petrology Volume 42 Number 7 Pages 1373-1385 2001
© Oxford University Press 2001
Hercynian Metamorphism in Nappe Core Complexes of the Alpine Betic–Rif Belt, Western Mediterranean—a SHRIMP Zircon Study
1INSTITUTE FOR STUDY OF THE EARTH’S INTERIOR, MISASA, TOTTORI-KEN, JAPAN
2RESEARCH SCHOOL OF EARTH SCIENCES, THE AUSTRALIAN NATIONAL UNIVERSITY, CANBERRA, A.C.T. 0200, AUSTRALIA
Received April 3, 2000; Revised typescript accepted December 15, 2000
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONREGIONAL GEOLOGIC SETTINGGRAPHITIC SCHIST 98Z75SHRIMP ZIRCON U-Th-Pb DATAINTERPRETATIONCONCLUSIONSREFERENCES |
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Ion microprobe U–Pb dating of a bimodal zircon population from a graphitic schist from the basal section of the Alpine Alpujárride nappe complex in the Betic Cordilleras, southern Spain, has yielded different ages for zircons of contrasting morphology and composition. Approximately half the grains are rounded and abraded, and have ages ranging from
2·2 Ga to 350 Ma. These are detrital grains from the schist’s sedimentary parent material, indicating its derivation from a composite basement and a Palaeozoic, possibly Carboniferous, age of deposition. The other half of the grains consist of euhedral cores of Hercynian age (305·3 ± 3·2 Ma) surrounded by Late Alpine (19·7 ± 2·2 Ma) metamorphic overgrowths (
= 0·05 errors). The Hercynian cores probably record the first, medium–high-grade amphibolite-facies metamorphism, relicts of which are seen in the schist complex, whereas the rims were formed during the final tectono-metamorphic stage in the Alpine development of the orogenic belt. The bimodality of the zircon population in age, morphology and internal structure may mainly reflect the lithological bimodality of the rock (thin quartzitic layers or schlieren alternating with graphitic micaceous folia) and the control that that exerted on fluid circulation in the metamorphic environment. KEY WORDS: SHRIMP zircon ages; Alpine belt; Western Mediterranean; Betic–Rif; Hercynian–Alpine polymetamorphism
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGIC SETTINGGRAPHITIC SCHIST 98Z75SHRIMP ZIRCON U-Th-Pb DATAINTERPRETATIONCONCLUSIONSREFERENCES |
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The Betic–Rif mountain belt forms the southwestern termination of the Alpine orogenic belt in Western Europe and North Africa. Although located in one of the very regions where the foundations of modern geology were laid (e.g. Argand, 1924
), fundamental aspects of its tectogenesis are still disputed. For example, there is no consensus on the direction of the regional collisional convergence and the location of the subduction zone. Most researchers (e.g. Egeler & Simon, 1969; Platt & Vissers, 1989; Platt & Whitehouse, 1999) have envisaged a north–south convergence within the framework of an Iberia–Africa collision. Recently, an alternative was suggested (Zeck, 1996, 1997, 1999): an east–west-directed convergence under the influence of the opening of the North Atlantic. In this model, the Alpine collisional nappe pile was derived from Tethys-realm Betic–Ligurian lithosphere in a SW–NE-trending subduction zone system obliquely underthrusting eastward-drifting Atlantic-realm Iberia.
In addition, the timing of the subduction is poorly constrained. Comparison of the scarce thermochronological data with tomographic information (Blanco & Spakman, 1993) and the sea-floor spreading record for the North Atlantic (Srivastava et al., 1990) suggest subduction activity within the period >60–30 Ma (see Zeck, 1996, 1999). Continental collision (40–30 Ma?) was followed by a final, extrusional/extensional tectonic stage that resulted in tectonic reworking of the primary nappe pile (e.g. Nijhuis, 1964; Zeck, 1968, 1996, 1999; Egeler & Simon, 1969; Michard et al., 1997) during the period 22–18 Ma (e.g. Zeck, 1996; Platt & Whitehouse, 1999; Sánchez-Rodríguez & Gebauer, 2000).
Two important questions concerning the Betic–Rif orogenesis remain to be answered:
- What is the age of the protolithic material for the metamorphic complexes that make up the basal sections of the Alpine nappes? This material is generally accepted to be pre-Permo-Triassic (e.g. Egeler & Simon, 1969; Fontboté & Vera, 1983; Sanz de Galdeano, 1997). Recent ion microprobe zircon dating (Zeck & Whitehouse, 1999) has constrained the age to be Palaeozoic. Can it be constrained further?
- Can a pre-Alpine metamorphic imprint be recognized in these metamorphic complexes? Most workers, for example, Sanz de Galdeano (1997), Azañón et al. (1998), Platt et al. (1998), Soto & Platt (1999), Argles et al. (1999) and García-Casco & Torres-Roldán (1999), have proposed an exclusively Alpine scenario. Others, for example, Nijhuis (1964), Zeck (1968), Egeler & Simon (1969), Gómez-Pugnaire & Franz (1988), Michard et al. (1997), Bouybaouène et al. (1998) and Zeck & Whitehouse (1999), have suggested models that involve medium- to high-grade pre-Alpine tectono-metamorphism in the basal section of the nappe stratigraphy.
The present study addresses these problems with the use of ion microprobe U–Th–Pb zircon dating. The very high closure temperature (>900°C, Lee et al., 1997) of this thermochronometer makes it particularly suited to record events that pre-date the Alpine orogeny.
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REGIONAL GEOLOGIC SETTING |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGIC SETTINGGRAPHITIC SCHIST 98Z75SHRIMP ZIRCON U-Th-Pb DATAINTERPRETATIONCONCLUSIONSREFERENCES |
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The Alpine collisional nappe pile in the Betic–Rif belt consists of two major nappe complexes, the tectonically higher Alpujárride complex and the tectonically lower Nevado–Filábride complex. Both complexes consist of a basal section, comprising mainly medium–high-grade schists and quartzites, overlain by a (Permo-)Triassic cover sequence consisting mainly of schists, phyllites and quartzites, topped by carbonate rocks. The primary, collisional nappe pile has undergone tectono-metamorphic reworking during a later Alpine extrusional/extensional phase. Most present-day lithological contacts are in fact such late-stage extensional features (e.g. Egeler & Simon, 1969
; Azañón & Goffé, 1997; Sanz de Galdeano, 1997).
The graphitic schist sampled for the present study is from the Sierra Alhamilla, an east–west-trending mountain range located in the eastern part of the Betic Cordilleras, NE of Almería (Fig. 1). The central part of the Sierra consists mainly of medium–high-grade metamorphic schists and quartzites of the Palaeozoic basal section. Towards the margins of the Sierra these are overlain by the low-grade metamorphic, (Permo-)Triassic cover section phyllites, quartzites and marbles. The present study had the benefit of detailed earlier studies of the geological relations within the area (Zeck, 1968). As we set out to solve two straightforward and well-defined geochronological problems within a well-established geological framework (Fig. 1), we could restrict the investigation to one graphitic schist sample from the basement complex. This sample was selected because of its intimately interlayered bimodal mineralogy and lithology, which could be expected to record both protolith ages and metamorphic history.
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GRAPHITIC SCHIST 98Z75 |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGIC SETTINGGRAPHITIC SCHIST 98Z75SHRIMP ZIRCON U-Th-Pb DATAINTERPRETATIONCONCLUSIONSREFERENCES |
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The rock selected for zircon analysis (98Z75), a graphite-bearing quartzitic chlorite–garnet–plagioclase–mica schist from the basal nappe section, is from the southern slopes of the Sierra Alhamilla. The sample site (Fig. 1) is at the western side of the Níjar–Lucainena de los Torres road
5·5 km from Níjar (X/Y UTM coordinates: 574.020/4093.440). (Permo-)Triassic phyllites and quartzites of the cover section are exposed 200–300 m SSE of the locality [Fig. 1; for more details see Zeck (1968)]. Sample 98Z75 is representative of a laminated/foliated rock with greyish, quartz-rich/lamellae or schlieren of 1–10 mm width alternating with black, graphitic micaceous folia. Very fine red–brown to ochre iron compounds are conspicuous, although quantitatively minor, forming impregnations along schistosity planes and cross-cutting veinlets. The rock is characterized by a well-developed schistosity defined by the parallel orientation of phyllo-silicate crystals.
Petrography
The minerals visible in thin section (and estimated volume percentages) are: apatite (<1), biotite (3), chlorite (5), garnet (6), graphite (3), leucoxene (1), plagioclase (10), quartz (50), red–brown iron compounds (1) and white mica (20); zircon was sought but not found in the three thin sections of 2 cm x 4 cm that were studied. The quartz-rich lamellae or schlieren consist of 50–70 vol. % of quartz, plus mostly garnet, white mica, chlorite and biotite. The texture is schistose, defined by up to 0·5 mm, platy crystals of white mica, biotite and chlorite, with quartz passively filling the space between them. Quartz is thus in platy, mosaic textured aggregates of up to 350 µm crystals, which are well crystallized, without signs of sub-graining or wavy extinction. Chlorite replacement of biotite is widespread, and locally complete. Apart from the thin, schistosity-supporting crystals, chlorite, usually with parallel intergrown remnants of biotite, forms much thicker crystals that cross-cut the schistosity. Some of these crystals appear micro-kinked along intersecting schistosity planes. Locally, aggregates of biotite–chlorite intergrowths form micro-augen trapped between schistosity planes. Garnet crystals are generally euhedral and vary from 30 to 400 µm in size. Cores of garnet crystals, especially larger ones, may appear turbid because of many small inclusions. In some crystals these centres are preferentially altered to aggregates containing chlorite and opaques.
The graphitic micaceous lamellae contain the same minerals as the quartz-rich domains, in similar genetic relationships but different proportions. Most of the graphitic micaceous domains consist of >50 vol. % white mica, with biotite and chlorite being subordinate. The texture is that of a stringent schistosity, with quartz forming slender lenses in the white mica mats. Locally some millimetre-sized elongated quartz lenses are present with enclosed polygonal arcs outlining tight isoclinal fold hinges with axial planes parallel to the external schistosity. Graphite forms very fine (10–50 µm) crystals. Garnet crystals in the graphitic micaceous lamellae are usually less abundant (or even absent) and smaller (10–150 µm) than in the quartzitic lamellae. Biotite crystals are usually replaced by chlorite, in some places also by leucoxene (locally grading into somewhat coarser aggregates of Ti minerals). Cross-cutting veinlets of quartz and red–brown iron compounds occur throughout, often showing series of small offsets indicating small late-stage movements along the schistosity planes. Limonite replacements and impregnations, often along small cracks, in garnet, biotite and chlorite are common.
Paragenetic relations and metamorphic conditions
The main paragenesis of the rock comprises quartz, white mica, biotite, plagioclase and garnet. Preserved tight isoclinal fold hinges within the strongly schistose micaceous domains indicate that the stringent schistosity is not a first-stage feature, but that the rock has had a more complex history than its deceptively simple field appearance might suggest. The biotite–chlorite aggregates trapped as small augen between schistosity planes may pre-date the main schistosity. Relict crystals of staurolite and kyanite up to 1 cm in diameter have been described within nearby rocks of the same graphitic schist sequence (Zeck, 1968). It is possible that the biotite–chlorite aggregates represent deformed pseudomorphs after staurolite.
The main paragenesis is medium-grade amphibolite facies. Probably the paragenesis was produced by reworking of a higher-grade amphibolite-facies rock. Cross-cutting chlorite crystals containing remnants of biotite may suggest a phase of mineral growth post-dating the synkinematic main paragenesis. Final mineral growth comprises low-temperature retrogression with formation of chlorite and probably quartz veinlets with red–brown iron compounds.
Goffé et al. (1989) reported parageneses with relict Mg–Fe carpholite from the Permo-Triassic cover section of this part of the Sierra Alhamilla, implying Alpine collisional stage metamorphic conditions of 6–7 kbar and 280–330°C. The cover section is in contact with the graphitic schist unit very close to the location of sample 98Z75 (Fig. 1) and therefore, although the contact is tectonic, a high-pressure metamorphic imprint in the graphitic schist was expected. However, no indications of carpholite were found despite careful inspection, particularly of the many quartz veins.
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SHRIMP ZIRCON U–Th–Pb DATA |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGIC SETTINGGRAPHITIC SCHIST 98Z75SHRIMP ZIRCON U-Th-Pb DATAINTERPRETATIONCONCLUSIONSREFERENCES |
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Sample preparation; analytical and data reduction procedures
Sample 98Z75 represents an
80 kg block of schist, which was broken up and cleaned by hammer chipping in the field. The resulting sample consisted of 8 kg of fragments of 2–10 cm size. These were reduced to chips of 3–10 mm in a jaw crusher. About 1·5 kg of the chips were crushed in several steps in a swing mill and screened to <0·5 mm. The 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 sample by removing the more magnetic grains. Final purification was by hand sorting. The total zircon yield consisted of 41 small crystals, which were cast in epoxy, along with crystals of reference zircons AS3 and SL13, then polished to their centres. After optical photomicrographs were taken, the zircons were imaged by SEM cathodoluminescence (CL) to document their internal structures. Before ion microprobe analysis the mount was washed in petroleum spirit, hot detergent solution and high-purity water to reduce surface contaminants, particularly common Pb, then coated with 5 nm of high-purity gold to prevent surface charging under ion bombardment.
The zircons were analysed for U–Th–Pb isotope compositions using the SHRIMP II ion microprobe at the Australian National University. The procedures for analysis were based on those described by Williams & Claesson (1987) and Williams et al. (1996). A 10 kV primary beam of negative O2 ions was focused to an 10 µm diameter spot on the sample surface. Positive secondary ions were extracted at 10 kV and transferred to a high mass resolution (5000R) double focusing secondary mass analyser, by which the relative abundances of the Zr, Pb, U and Th isotopes of interest were determined by pulse counting using a single electron multiplier and cyclic magnet field switching. Each five-scan analysis took about 13 min.
Pb isotopic ratios were measured directly, without correction for mass-dependent mass fractionation (<2·5
/a.m.u.). Inter-element fractionation was calibrated against zircon standard AS3 (radiogenic 206Pb/238U = 0·1859: Paces & Miller, 1993) assuming the relationship Pb/U = a(UO/U)2, where a is a constant for a given analytical session (Claoué-Long et al., 1995). Common Pb contents were mostly very low, so it was generally assumed that the common Pb was laboratory-derived contamination of the sample surface. The few analyses that showed significantly higher common Pb contents were corrected assuming a common Pb composition commensurate with the target grain’s age (Cumming & Richards, 1975). Most common Pb corrections were made using the measured 204Pb, but for several analyses, particularly those with low Th/U, significantly better precision could be achieved using 208Pb and Th/U. In those cases, radiogenic 208Pb/206Pb and 208Pb/232Th were not determined independently.
Zircon morphology and zoning—optical microscopy and cathodoluminescence
The zircon grains recovered from the graphitic schist are small (40–110 µm), equant to short prismatic (ratiolength:width <3) with blunt or rounded pyramidal terminations. In transmitted light the zircon sample is bimodal in appearance. About half of the grains are rounded and pitted to variable degrees, uniformly clear with relatively few inclusions. The remainder appear more euhedral and consist of a clear, inclusion-free core surrounded by an inclusion-rich overgrowth. In many cases the innermost layer of the overgrowth is particularly inclusion rich.
CL imaging confirms the bimodality of the population, but reveals two grains (22 and 32) with transitional properties. Most of the clear, rounded and pitted grains luminesce strongly to reveal a range of internal structures (8, 10, 15, 17 and 35 in Fig. 2). Some are simple grains, euhedrally growth zoned throughout, but many consist of a core, which may be structureless or euhedrally zoned, rounded or irregular in outline and surrounded by a thin (25 to <5 µm), relatively weakly luminescent rim. Internal zoning is generally discordant to grain outline. All other grains (except 32) are very weakly luminescent throughout and have a consistent structure: a sharply euhedral core (50 µm) with simple sector zoning surrounded by an unzoned overgrowth (15–30 µm wide) containing 50% by volume of non-luminescent inclusions (7, 20 and 26 in Fig. 2). Grain 32 (Fig. 2) lacks such clear spongy overgrowth, but is sector zoned like the sharply euhedral cores. Grain 22 has a spongy overgrowth, but its core consists of a rounded crystal fragment with internal zoning discordant to its external outline, similar to that seen in the rounded and pitted grains.
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Thirty-four zircon grains were analysed, five of them on both core and rim. Despite the use of a small-diameter primary beam (10 µm), the narrowness of the overgrowths and abundance of inclusions left very few overgrowth areas accessible to analysis.
Analytical results
The analytical results are listed in Table 1 and plotted in Tera–Wasserburg concordia diagrams in Figs 3–5 . The uncertainties listed and plotted are precision estimates of one standard error, calculated from counting errors and the propagation of uncertainties introduced by the common Pb correction. The uncertainty in Pb/U standardization is an additional 0·26%. The list of interpreted ages gives the best estimate of the age inferred from each analysis, taking into account the age, degree of discordance, and size of the common Pb correction. For zircons with concordant ages younger than 1500 Ma, the interpreted age is based on radiogenic 206Pb/238U and assumed concordance. For discordant zircons, and those older than 1500 Ma, the interpreted age is based on radiogenic 207Pb/206Pb. By these means the inaccuracies introduced by discordance and imprecise common Pb corrections have been minimized. The ages were calculated using the constants recommended by the IUGS Subcommission on Geochronology (Steiger & Jäger, 1977) and, except in the table, the uncertainties cited are 95% confidence limits, namely t
, where t is Student’s t.
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Rounded and pitted grains
The centres/cores of the rounded and pitted grains (including grain 22) have a wide range of U (120–1690 ppm) and Th (20–570 ppm) contents, Th/U ratios (0·037–0·9) and isotopic compositions (Table 1, Fig. 3). The apparent ages range from 350 to 2200 Ma with age clusters at 500, 600–700 and 1000 Ma. Only two of the rims of rounded grains were wide enough (20 µm) to analyse. Both are relatively weakly luminescent, with a hint of broad euhedral zoning. They have very different U (1870 and 231 ppm) and Th (101 and 4 ppm) contents, but both show the extremely low Th/U (<<0·1) that is a common feature of metamorphic zircon. Both isotopic compositions are concordant within analytical uncertainty, but yield very different ages, 1045 and 630 Ma, respectively. These ages are both less than the ages of the cores of the same grains, 1380 and 1900 Ma, respectively.
Sector zoned and spongy grains
In contrast to the rounded and pitted grains, the sector zoned cores of the other type of grains (including grain 32) have a much more restricted range of compositions and ages. Their U contents range from 265 to 640 ppm, Th contents from 100 to 410 ppm and Th/U from 0·38 to 0·68. In general, the higher the U content, the higher the Th/U. Plotted in a Tera–Wasserburg concordia diagram (Fig. 4), 13 of the 15 uncorrected analyses lie within analytical uncertainty [mean square weighted deviation (MSWD) = 0·3] on a mixing line between Hercynian common and radiogenic endmembers. The exceptions are two cores (6 and 36) that have high common Pb and appear to have lost 5–8% of their radiogenic Pb. When the uncertainty of 0·26% on the Pb–U calibration is taken into account, the weighted mean 206Pb/238U age for the group of 13, calculated by subtracting common Pb using the radiogenic 207Pb/206Pb expected for concordance, is 305·3 ± 3·2 (t) Ma.
Most of the weakly luminescent, inclusion-rich overgrowths on the euhedral zircons were either too thin or too inclusion rich to have areas large enough to be analysed with the 10 µm diameter primary beam. Three analyses were possible, however. Each has a very high U content (4600–7600 ppm) and low Th (50–60 ppm), giving consistently low Th/U ratios (0·01). Common Pb contents in all areas are much higher than average, between five and 20 times as high as those measured on the sector zoned cores or rounded zircons. Plotted in a concordia diagram (Fig. 5) before common Pb correction, the analyses disperse along a common Pb mixing line, but with considerable scatter. Recent work on high-U zircon from Tasmanian dolerites (Williams & Hergt, 2000) suggests that matrix effects inflate SHRIMP measurements of zircon 206Pb/238U in proportion to U content when U exceeds 2500 ppm. Applying a correction of 1·8%/1000 ppm U to the overgrowth data improves the line somewhat, but the scatter remains significant. With so few analyses and large scatter, it is difficult to compound these data. If the two analyses with highest common-Pb corrections are assumed to be isotopically disturbed, then the best estimate of the age is given by the remaining analysis: 22 Ma. This seems unlikely, however, because the crystal domain with highest common Pb (analysis 20.2) has the lowest U, and therefore should be least susceptible to isotopic disturbance. The alternative interpretation is that the analysis with lowest common Pb (analysis 13·2) contains a small amount of radiogenic Pb from the adjacent core, and that the age of the overgrowth itself is closer to 20 Ma. As there is no objective way to choose between these options, the best solution is to take the weighted mean radiogenic 206Pb/238U for the three analyses (0·00307) and assign an uncertainty that takes into account the observed scatter (0·00008), yielding an age estimate of 19·7 ± 2·2 Ma (t).
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INTERPRETATION |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGIC SETTINGGRAPHITIC SCHIST 98Z75SHRIMP ZIRCON U-Th-Pb DATAINTERPRETATIONCONCLUSIONSREFERENCES |
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The wide range of ages revealed by the present ion microprobe survey of zircon from a graphitic schist from the Betic–Rif belt reflects two main influences: the age spectrum of the source rock complexes for the sedimentary protolith and the Hercynian–Alpine polymetamorphic history. An additional factor also has to be considered: differences in the development and age of the zircon overgrowths suggest that conditions for metamorphic zircon growth differed from grain to grain within the rock. Such differences may be characteristic of the metamorphic environment in which penetrative fluid circulation mainly follows intergranular interstices and intragranular microcracks, such that different zircon grains may come to experience different fluid access histories.
Detrital zircon
The variety of compositions, internal growth structures and ages of the centres or cores of the rounded grains, their rounded shapes, and evidence of surface abrasion, combine to indicate that these are detrital zircons from the sedimentary parent material of the graphitic schist. Most have relatively strong luminescence, and their range of growth structures indicates derivation from a variety of lithologies. The obvious explanation is that the detrital zircon, and hence the sedimentary protolith, was ultimately derived from a composite basement containing a range of source rocks of different ages. The two detrital/inherited crystals with rims large enough to analyse show Grenvillian (1045 Ma) metamorphic overgrowth on a 1380 Ma zircon (grain 27), and Pan-African (630 Ma) overgrowth on a 1900 Ma zircon (grain 10). These, and the core ages at 500–700 and 1000 Ma, suggest that the basement that supplied the sedimentary clastic material was affected by both Grenvillian and Pan-African magmatic–metamorphic events.
It could be argued that some of these U–Pb ages have been partly or completely reset. Such an effect has been described by Pidgeon (1992), Zeck & Whitehouse (1999) and Vavra et al. (1999). Vavra et al. reported resetting of U–Pb ages to be related to alteration domains characterized by strong luminescence and/or isometric zoning. Such domains are not present where ‘young’ ages were measured on zircon crystals studied here, but that does not rule out partial resetting as an explanation of some of the dispersion in the dataset. The detrital zircon age pattern for the graphitic schist is shown in a Tera–Wasserburg concordia diagram (Fig. 3) and relative probability histogram in Fig. 6. This latter figure offers a comparison with the age pattern recently reported by Zeck & Whitehouse (1999) for an anatectic granitic rock within the Alpujárride nappe complex, 150 km west of the graphitic schist locality in the Sierra Alhamilla. Although there are too few grains from either sample to define specific age groups or their relative abundances quantitatively, the similarity in the two age patterns suggests that the sedimentary parent material for the graphitic schist and the metamorphic source rock of the anatectic granitic magma were derived from similar, or the same, source regions.
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A maximum age for the deposition of the parent sedimentary material for the graphitic schist is given by the youngest well-defined detrital zircon ages, 500 Ma (Figs 3 and 6). There are two grains with younger apparent ages, 360 Ma and 335 Ma (Table 1, Fig. 3). These ages either constrain the deposition age further—to Carboniferous time—or record a later isotopic resetting event. Resolution of this question will have to await further work.
Metamorphic zircon
With one exception, the Proterozoic detrital or inherited zircon crystals do not show indications for overgrowths of Hercynian and/or Alpine age. Grain 22 shows a spongy overgrowth identical in appearance to those dated as Alpine in other crystals; however, its small width (10 mm) did not allow analysis. Apparently, zirconium-bearing fluids had only very restricted access to the detrital/inherited crystals during the Hercynian and Alpine metamorphic episodes. The crystals have either been enclosed in older crystals during these metamorphic episodes, or are from parts of the rock where zircon growth did not take place, possibly because zirconium was not available at those sites. It has been suggested (Vavra et al., 1999) that the source of zirconium for metamorphic zircon growth in metasedimentary rocks is very fine clastic zircon (‘zircon dust’), which by Ostwald ripening during the metamorphic evolution would give rise to growth of larger zircon crystals. As the graphitic schist studied here consists of two lithological domains, quartzitic lamellae/schlieren and graphitic micaceous folia, we tentatively suggest that the quartzitic lamellae or schlieren, which represent the sandy layers in the original sedimentary pile, were poor in ‘zircon dust’ but richer in coarser, clastic zircon grains, compared with the graphitic micaceous lamellae that represent the pelitic layers. Subsequent metamorphic recrystallization could then produce the observed bimodality of the zircon population, by growing euhedral Hercynian grains with later Alpine overgrowths in the graphitic lamellae while at the same time preserving the clastic grains in the quartzitic domains. Grain 22 may represent the exceptional clastic grain featuring an Alpine overgrowth. The high Th/U, sector zonation and simple, sharply euhedral crystal shape of the Hercynian zircon cores are features often seen in magmatic rocks, but are also fairly common in metamorphic zircon (R. T. Pidgeon, personal communication, 2000).
The two outliers (grains 6.1 and 36.1; Fig. 3, Table 1) at the lower end of the age spectrum of the 15-point Hercynian dataset (Fig. 4) are interpreted as slightly reset during the Alpine event. The weighted mean age of 305·3 ± 3·2 (t) Ma for the remaining 13 ages for the euhedral, sector zoned cores, including grain 32, is regarded as the best estimate of the Hercynian metamorphic age.
The 19·7 ± 2·2 Ma (t) age estimate derived for the spongy overgrowths on the euhedral cores is interpreted as the age of the final Alpine metamorphic imprint in the Alpujárride nappe complex in the Sierra Alhamilla in the eastern part of the Betic Cordilleras. This age is similar to a large number of age estimates for the final stage of Alpine metamorphism in the Alpujárride nappe complex that have become available in recent years. These define an age interval of 22–18 Ma and include palaeontological ages of nappe sealing sedimentary rocks [both nannoplankton (Aguado et al., 1990) and planktonic Foraminifera (González-Donoso et al., 1982; Serrano, 1990)] and a range of thermochronometric ages (Priem et al., 1979; Zeck et al., 1989a, 1989b, 1992; Monié et al., 1991, 1994; Andriessen & Zeck, 1996; Platt & Whitehouse, 1999; Sánchez-Rodríguez & Gebauer, 2000) of highly variable nature (Rb–Sr muscovite and biotite whole rock; K–Ar biotite; 40Ar/39Ar muscovite, biotite and hornblende; fission tracks in apatite and zircon; U–Th–Pb zircon ion-microprobe). At an early stage it was pointed out that the tight clustering of these ages, in spite of the very different closure temperatures of the different thermochronometers employed, indicates a very high regional cooling rate during this late-stage Alpine development (Zeck et al., 1989b, 1992; Zeck, 1996). A detailed thermochronological study in a small area in the central part of the Alpujárride nappe complex established a cooling rate of not less than 500°C/my during the age interval of 19·5–18·5 Ma (Zeck, 1996). Subsequent studies (e.g. Platt & Whitehouse, 1999; Sánchez-Rodríguez & Gebauer, 2000) have confirmed similar, exceptionally high cooling rates for the Betic Cordilleras during the period 22–18 Ma.
Alpine plate tectonic evolution
The plate tectonic evolution of the Alpine Betic–Rif belt involves a collisional stage (40–30 Ma?) and a second, final stage, which, as argued above, occurred during the age interval of 22–18 Ma. Ages of 12–1 Ma have been reported for post-collisional volcanism (Zeck et al., 1998, 2000).
To explain the extremely rapid cooling that characterizes the final tectono-metamorphic stage, a sinking slab model has been proposed (Zeck, 1996, 1997, 1999). This model builds on seismic studies of the very deep earthquakes of the region (Chung & Kanamori, 1976; Grimison & Chen, 1986), seismic tomography studies (Blanco & Spackman, 1993) and a compilation of the 1991–1994 seismicity in the Alborán area (Seber et al., 1996). The slab break-off implied by this model caused in-flow of high-temperature, low-density asthenosphere into the widening gap above the sinking slab. The resulting isostatic effect may have triggered the very rapid uplift of the overlying lithosphere (Zeck, 1996). The main uplift effect, however, was probably due to another factor: thermal softening in the stationary slab and the collisional nappe pile, caused by the high-temperature regime induced by the asthenospheric in-flow. The rise in temperature would decrease the viscosity of the rock mass, and in a compressional setting this may lead to very rapid upward extrusion, ‘like toothpaste in a tube’ (Thompson et al., 1997a, 1997b).
The graphitic schists studied here have a complex structure and mineralogy. The rocks show an older, relict metamorphic stage with staurolite and kyanite whereas the main paragenesis is of medium-grade amphibolite facies. Considering the 305 Ma metamorphic zircon age suggested by the present study it seems feasible that the relict paragenesis is of Hercynian age. The main mineral paragenesis could represent metamorphism during the extrusional tectonic stage with a zircon age of 19·7 ± 2·2 Ma. Although this interpretation fits both our present observations and the available information in the literature, great caution is inspired by the structural and metamorphic complications that might be hidden in these highly schistose rocks. Relict structures suggest various generations of parallel schistosity planes, and unravelling their age relations will demand very careful field and petrographic work, supported by electron microprobe thin-section mapping.
The present study does not provide constraints on the age of the collisional stage of the orogenic evolution. High-pressure mineral parageneses have been found in the phyllitic rocks that tectonically overlie the graphitic schists studied here (Goffé et al., 1989), but not in the graphitic schist complex itself. Our preferred explanation for this discrepancy is that the graphitic schists had been metamorphically processed during the Hercynian orogeny, which left the rocks partially dehydrated and less reactive than the pristine Permo-Triassic pelitic sedimentary rocks during the Alpine event. A recent paper (Sánchez-Rodríguez & Gebauer, 2000) inferred a 19·7 ± 3·4 (2) Ma age for the high-pressure collisional metamorphism in the Alpujárride nappe complex based on a U–Pb ion microprobe concordia intersection age for zircon grown in ‘eclogites’ from the Alpujárride nappe complex in the western part of the belt (Sierra Bermeja–Sierra Alpujata). However, this may not be so, as the age of the sedimentary rocks overlying the extensional allochthons in which these rocks are found is 22–20·5 ± 1·5 Ma (nannoplankton biozone NN2, Aguado et al., 1990; compare Zeck, 1996). Further, there is good petrographic and field evidence (Tubía & Ibarguchi, 1991) that the ‘eclogites’ have a complex petrogenesis, forming rare, small pods in amphibolitic rock units, and containing abundant amphibole, with their high-pressure parentage revealed by relict pyroxene, garnet and rutile. The 20 Ma zircon age for these rocks therefore seems better interpreted as the age of the amphibolite-facies reworking during the extrusional tectonic stage (see Zeck, 1996).
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CONCLUSIONS |
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TOPABSTRACTINTRODUCTIONREGIONAL GEOLOGIC SETTINGGRAPHITIC SCHIST 98Z75SHRIMP ZIRCON U-Th-Pb DATAINTERPRETATIONCONCLUSIONSREFERENCES |
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Zircon U–Th–Pb ion microprobe dating reveals ages of
2200–350 Ma for detrital zircon from the meta-sedimentary basement section of the Alpujárride nappe complex in the eastern part of the Betic–Rif belt. This suggests a clastic provenance pattern for the sedimentary parent material similar to that suggested by Zeck & Whitehouse (1999) for Alpujárride nappe rocks 150 km further west. This detrital age pattern indicates that the source for the clastic sedimentary parent material of these rocks was Gondwana crust (see Gebauer, 1993). The maximum age for the sedimentation of this parent material is given by the age of the youngest concentration of ages within the detrital age spectrum: 500 Ma. As the nappe stratigraphy provides a minimum Permo-Triassic age, a Palaeozoic sedimentation age is indicated. This conclusion confirms earlier suggestions by Zeck & Whitehouse (1999). Two grains with ages of 350 Ma might constrain the deposition age to Carboniferous time, but more work is required to confirm this.
This zircon study shows that the graphitic schist complex in the Sierra Alhamilla has metamorphic zircon grains showing Hercynian (305·3 ± 3·2 Ma) cores and Alpine (19·7 ± 2·2 Ma) overgrowths ( = 0·05 errors). This age information, combined with the petrographic and regional geological constraints outlined above, leads to the conclusion that the nappe basement complexes are of a Hercynian–Alpine polymetamorphic character. Hercynian amphibolite-facies metamorphism was followed by substantial Alpine (20 Ma) tectono-metamorphic reworking during the extrusional tectonic stage. This result confirms earlier conclusions (e.g. Nijhuis, 1964; Zeck, 1968; Egeler & Simon, 1969; Gómez-Pugnaire & Franz, 1988; Michard et al., 1997; Bouybaouène et al., 1998; Zeck & Whitehouse, 1999), which have been challenged in a series of recent papers (e.g. García-Casco et al., 1993; García-Casco & Torres-Roldán, 1996, 1999; Platt et al., 1998; Soto & Platt, 1999; Argles et al., 1999) that claim an exclusively Alpine metamorphic development.
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ACKNOWLEDGEMENTS |
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This work was supported by the Danish Research Council (SNF grant 9800918/28808/1001) and the Carlsberg Foundation (grant 950407/70). The senior author thanks the Japanese Ministry of Education for funding (Monbusho fellowship), ISEI directors and Ikuo Kushiro and Masaru Kono for hospitality and gracious support at the Misasa Institute. We thank Mr J. Mya for assistance with mineral separation, and Mr N. Gabbitas and the staff of the ANU Electron Microscopy Unit for assistance with CL imaging. We thank Bob Pidgeon and Claudio Faccenna for very thorough and constructive reviews.
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FOOTNOTES |
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*Corresponding author. Present address: Geological Institute, Copenhagen University, Öster Voldgade 10, 1350K Copenhagen, Denmark. E-mail: zeck{at}geo.geol.ku.dk
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