Journal of Petrology Advance Access originally published online on April 28, 2007
Journal of Petrology 2007 48(6):1185-1218; doi:10.1093/petrology/egm014
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Duration of a Large Mafic Intrusion and Heat Transfer in the Lower Crust: a SHRIMP U–Pb Zircon Study in the Ivrea–Verbano Zone (Western Alps, Italy)
1Dipartimento Di Scienze Della Terra, Universita’ Di Trieste, Pal. N, Via Weiss 8, 34127, Trieste, Italy
2Max Planck Institut Für Chemie, Abt. Geochemie, Mainz, Germany
3US Geological Survey, Reston, VA 20192, USA
4Australian National University, Canberra, Act 0200, Australia
RECEIVED SEPTEMBER 1, 2006; ACCEPTED MARCH 7, 2007
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ABSTRACT |
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 TOP ABSTRACT INTRODUCTIONGEOLOGICAL SETTINGSAMPLESMETHODSRESULTSZIRCON MORPHOLOGICAL AND...DISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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The Ivrea–Verbano Zone in the western Italian Alps contains one of the world's classic examples of ponding of mantle-derived, mafic magma in the deep crust. Within it, a voluminous, composite mafic pluton, the Mafic Complex, intruded lower-crustal, high-grade paragneiss of the Kinzigite Formation during Permian–Carboniferous time, and is now exposed in cross-section as a result of Alpine uplift. The age of the intrusion is still debated because the results of geochronological studies in the last three decades on different rock types and with various dating techniques range from 250 to about 300 Ma. Sensitive high-resolution ion microprobe (SHRIMP) U–Pb zircon age determinations on 12 samples from several locations within the Mafic Complex were performed to better constrain the age of the igneous event. The results indicate a long history of magma emplacement and cooling, which reconciles the spread in previously published ages. The main intrusive phase took place at 288 ± 4 Ma, causing a perturbation of the deep-crustal geotherm, which relaxed to the Sm–Nd closure temperature in garnet-free mafic rocks after about 15–20 Myr of sub-solidus cooling at c. 270 Ma. These results suggest that large, deep crustal plutons, such as those identified geophysically at depths of 10–20 km within extended continental crust (e.g. Yellowstone, Rio Grande Rift, Basin and Range) may have formed rapidly but induced a prolonged thermal perturbation. In addition, the data indicate that a significant thermal event affected the country rock of the Mafic Complex at about 310 Ma. The occurrence of an upper amphibolite- to granulite-facies thermal event in the Kinzigite Formation prior to the main intrusive phase of the Mafic Complex has been postulated by several workers, and is corroborated by other geochronological investigations. However, it remains uncertain whether this event (1) was part of a prolonged perturbation of the deep-crustal geotherm, which started long before the onset of intrusion of the Mafic Complex, or (2) corresponded to the intrusion of the first sills of the Mafic Complex, or (3) was related to an earlier, independent thermal pulse.
KEY WORDS: Ivrea–Verbano Zone; thermal evolution of continental crust; underplating; U–Pb SHRIMP geochronology; zircon
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INTRODUCTION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLESMETHODSRESULTSZIRCON MORPHOLOGICAL AND...DISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Geophysical studies indicate that large plutons, up to 10 km thick and 70 km wide, occur at depths of 10–20 km within extended continental crust (e.g. Rio Grande Rift, Yellowstone, Basin and Range Province; see respectively Perry et al., 1988
; Hildreth et al., 1991; Jarchow et al., 1993). Commonly referred to as magmatic underplating (e.g. Bergantz, 1989), the ponding of large volumes of mantle-derived mafic magma in the deep crust has been cited as a potential thermal engine that drives high-grade regional metamorphism (e.g. Hansmann et al., 2001) and precedes the emplacement of large volumes of granite in the middle and upper crust (e.g. Huppert & Sparks, 1988) and the eruption of silicic magmas (e.g. Riley et al., 2001; Reagan et al., 2003; see also Schaltegger & Brack, 2007). The thermal impact of magmatic underplating on the continental crust must be a function of several factors, including the volume of injected magma, the temperature contrast between the magma and the country rock, and the duration of the magma intrusion (e.g. Annen & Sparks, 2002). Despite numerous geochemical, petrological and geophysical studies, the emplacement dynamics, rates, and duration of an underplating event remain unclear (Hawkesworth et al., 2004). Modeling geodetic data, such as provided by InSAR (Interferometric Synthetic Aperture Radar), provides insight into real-time intrusion rates and crustal response to individual intrusive pulses of the order of 100 m thick (e.g. Massonnet et al., 1995; Fialko & Simons, 2001; Lu et al., 2003; Pedersen & Sigmundsson, 2004). However, only exposed deep-crustal terranes display the time-integrated, final result of the underplating process, and thus their study is required to document the time needed to grow 5–10 km thick composite intrusions.
The Mafic Complex of the Ivrea–Verbano Zone is an excellent target for such an investigation. Now exposed by Alpine uplift, the Mafic Complex was emplaced beneath 15–25 km of continental crust (Demarchi et al., 1998), and reaches a thickness of 8 km and a length of more than 35 km in the southern part of the zone. To better constrain the timing and duration of magma emplacement in the Ivrea–Verbano Zone, we determined new sensitive high-resolution ion microprobe (SHRIMP) U–Pb zircon ages from 12 samples of the Mafic Complex collected on the basis of recently published detailed geological mapping (Quick et al., 2003) and zirconium content. During cooling under lower-crustal conditions, the U–Pb isotopic system in igneous zircon is the most resistant to resetting, and thus is the best system to unravel the intrusion history of the Mafic Complex. U–Pb zircon ages can provide insight on the duration of the intrusive process in the deep crust and the timescale of its thermal impact on the surrounding country rocks. The results of this study improve understanding of the large scatter of geochronological data published in the last 35 years for the Ivrea–Verbano Zone, which were determined using different isotopic systems and techniques.
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GEOLOGICAL SETTING |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLESMETHODSRESULTSZIRCON MORPHOLOGICAL AND...DISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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The Ivrea–Verbano Zone (Fig. 1) is a sliver of intermediate to deep continental crust, which crops out between the Insubric Line, the Cossato–Mergozzo–Brissago (CMB) Line, and the Cremosina Line. An excellent introduction to the geology of the Ivrea–Verbano Zone has been provided by Zingg et al. (1990
). Its present exposure and orientation are mainly consequences of tilting by about 90° and uplift during Alpine collision. Regional metamorphic grade and equilibration pressures increase from east to west (Zingg, 1980; Henk et al., 1997; Demarchi et al., 1998) demonstrating that progressively deeper crustal levels are exposed towards the Insubric Line.
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The Ivrea–Verbano Zone consists of two main units: the Kinzigite Formation and the Mafic Complex. Paragneiss of the Kinzigite Formation and tectonically interleaved lenses of mantle peridotite were intruded by voluminous gabbroic magmas of the Mafic Complex while resident in the intermediate to deep crust (Rivalenti et al., 1975
, 1981; Quick et al., 1995). The geology of the southern Ivrea–Verbano Zone, where the Mafic Complex is a thick, composite intrusion, has been described in detail by Quick et al. (2003, and references therein), and the geology of the northern Ivrea–Verbano Zone, where the outcrop of the Mafic Complex is thinner and consists of the Finero mafic–ultramafic body, has been described by Rivalenti & Mazzucchelli (2000, and references therein). The lower-crustal section of Val Strona di Omegna, mostly consisting of paragneiss of the Kinzigite Formation in upper amphibolite to granulite facies, has been described in detail by Bertolani (1968) and Harlov & Förster (2002a).
In the southern Ivrea–Verbano Zone (Fig. 2), the Mafic Complex is about 8 km thick and 35 km wide. Despite tilting of about 90° and uplift from more than 15 km depth, the Mafic Complex was relatively undisturbed by Alpine tectonics and its primary internal structure and relationship to the country rock are well preserved (Quick et al., 1994). Geological mapping (Quick et al., 1992, 1994, 2003) reveals a megascopic arcuate structure in the Mafic Complex, defined by mappable lithologies, high-temperature foliation, and compositional banding. The core of this structure is located in Val Sesia. As interpreted by Rivalenti et al. (1975) and subsequent workers, the contact between the Complex and the Kinzigite Formation from Val Sesia south corresponds to the roof of the intrusion. Intrusion of the Mafic Complex caused anatexis in the Kinzigite Formation near the roof (e.g. Quick et al., 1994, fig. 6; Snoke et al., 1999). Dioritic rocks form a large body in Val Sesia, at the core of the arcuate structure, and a thin seam near the roof of the complex. Small intrusions of pyroxenite and cumulus ultramafic rocks are abundant at deep levels in the complex (e.g. Rivalenti et al., 1981; Quick et al., 1992), but also occur near the roof and intrude the Kinzigite Formation.
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Thin paragneiss layers, termed septa, were derived from the roof and incorporated into the complex (Quick et al., 1992
). Septa are concentrated in a zone identified as the ‘paragneiss-bearing belt’ by Sinigoi et al. (1995; see Fig. 2), although individual septa flare upward from the belt into higher levels of the intrusion. According to Sinigoi et al., (1995) the paragneiss-bearing belt marks the rough boundary between the lower Mafic Complex, dominated by granoblastic amphibole-gabbro, and the upper Mafic Complex, mainly composed of norites, gabbros and diorites.
The Kinzigite Formation contains a diverse assemblage of schist, paragneiss, amphibolite, and calc-silicate rocks. The metamorphic grade of these rocks is amphibolite facies in the eastern Ivrea–Verbano Zone and granulite facies in the western Ivrea–Verbano Zone (Zingg, 1980). The bulk composition of the granulite-facies rocks can be explained by extracting 20–40% anatectic melt from a protolith similar to the amphibolite-facies rocks (Sighinolfi & Gorgoni, 1978; Cavallini et al., 1994; Schnetger, 1994; Harlov & Förster, 2002b). Although the Mafic Complex is an attractive potential source for the heat required to drive the granulite-facies metamorphism, Barboza et al. (1999) have shown that, in the southern Ivrea–Verbano Zone, granulitization predates the emplacement of the upper Mafic Complex, and that the thermal impact of the intrusion was relatively small and localized to within about 1 km of the roof. Also, stable isotope studies indicate that a link between granulitization of the crust and intrusion of the Mafic Complex is improbable (Ahrendt et al., 1989; Baker, 1990).
The Mafic Complex has been the subject of geochronological investigations for more than 30 years. Figure 3 shows a schematic compilation of the age data that are relevant to this study, and a more detailed compilation is contained in Fig. 2E of the Electronic Appendix (available for downloading from http://petrology.oxfordjournals.org). Age determinations and interpretations have been reviewed by Gebauer (1993), Schaltegger & Gebauer (1999), Vavra et al. (1999) and Mayer et al. (2000). Attempts to date the complex have utilized Rb–Sr, Sm–Nd, and Re–Os whole-rock methods and mineral isochrons, Ar–Ar methods, conventional U–Pb zircon dating, the Pb–Pb zircon evaporation method, and SHRIMP U–Pb zircon dating. Attempts to constrain the age of intrusion have also been made by conventional U–Pb analysis of zircon and monazite from the country rocks of the complex, and by SHRIMP U–Pb analysis of zircons from a paragneiss septum (Vavra et al., 1996, 1999; Vavra & Schaltegger, 1999). Collectively, these ages range from over 600 Ma (Hunziker & Zingg, 1980; Pin & Sills, 1986; Voshage et al., 1987; Gebauer et al., 1992) to 30 Ma (Gebauer et al., 1992). Ages older than 400 Ma generated by Rb–Sr and Sm–Nd whole-rock methods are best interpreted as the slopes of mixing lines with no age significance (Voshage et al., 1990). U–Pb zircon ages older than Permo-Carboniferous reflect inherited-core ages, and as such have to be interpreted as useful information recorded within this continental basement prior to the emplacement of the Mafic Complex. Finally, ages in the range of 30 Ma date Alpine events that clearly postdate the igneous crystallization age of the Mafic Complex (Mayer et al., 2000). Although the remaining U–Pb, Sm–Nd and Re–Os geochronological determinations on mineral phases point to emplacement of the Mafic Complex in the southern Ivrea–Verbano Zone between 300 and 210 Ma, the age relationships between the principal units of the complex are unknown and the age data published to date are insufficiently comprehensive or insufficiently precise to constrain modeling of the duration of the intrusion and its thermal impact on the crust.
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SAMPLES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLESMETHODSRESULTSZIRCON MORPHOLOGICAL AND...DISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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In this study, 10 samples of gabbroic rock were chosen to document the range in ages at different depths in the Mafic Complex and along a traverse parallel to its roof that cuts its internal structural grain (Fig. 2; samples 1d–8gn and 11ag–12ag). Sample locations and paragenesis are reported in Fig. 2 and Table 1. Although samples were chosen to be representative of major units in the Mafic Complex, collection was also based on Zr content, with the assumption that this would correlate with zircon abundance. In the lower Mafic Complex, the choice of samples was primarily influenced by this geochemical information, because our geochemical database, which contains about 600 whole-rock X-ray fluorescence analyses performed at the Università di Trieste, records Zr abundance <40 ppm for most of the mafic intrusive rocks in this unit. Within the paragneiss-bearing belt, two different rock types were also sampled in Val Sessera (Fig. 4). Samples 9mp and 10ck represent respectively a restitic paragneiss septum and the associated leucocratic (charnockitic) counterpart. Within the lower Mafic Complex, samples 11ag and 12ag were collected from gabbronoritic bands within the amphibole gabbro in the upper Val Sessera.
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To unravel the magmatic ages, we chose to rely on the U–Pb system in zircon because it is resistant to resetting, and on the SHRIMP dating method because of possible complications caused by multiple events and inheritance. The precision provided by conventional single-grain U–Pb dating would be about one order of magnitude higher than that obtained by SHRIMP, and modern high-precision thermal ionization mass spectrometry (TIMS) dating would guarantee the resolution of multiple events. Nonetheless, we preferred an in situ dating technique, so that the age determinations can be interpreted in direct relation to the characteristics of the zircon as revealed by cathodoluminescence (CL) imaging.
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METHODS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLESMETHODSRESULTSZIRCON MORPHOLOGICAL AND...DISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Sample preparation
Standard mineral separation techniques based on density and magnetic properties were used to concentrate zircons from samples that on average weighed 12 kg. Crystals used for U–Pb analyses were not air-abraded or acid-washed at any stage of the preparation. For each sample, c. 100 zircons were hand-picked from the diamagnetic fraction at 1·2 mA, following heavy liquid separation at
3·1 g/cm3. Cathodoluminescence investigation of each mount was carried out on a JEOL 5600LV scanning electron microscope at Stanford University, USA, and on a Hitachi S2250-N at the Electron Microscope Unit at the Australian National University (ANU) prior to SHRIMP analyses. The exact position of the analytical spots within the grains was determined following SHRIMP analysis with optical images at ANU and with back-scattered electron images obtained using a Hitachi S450 at the Max-Planck Institut für Chemie in Mainz (Germany). In many cases, grey-scale image enhancement was necessary to highlight internal features. Thus, the images in subsequent figures emphasize textures but are not suitable for direct comparison of cathodoluminescence brightness as an indicator of trace element content.
Data acquisition
Zircons were analyzed for U, Th and Pb using the US Geological Survey–Stanford SHRIMP II RG located at Stanford University, and the SHRIMP II at the Research School of Earth Sciences at ANU. The protocol for data acquisition consisted of nine scans at Stanford University and six at ANU, with count times per scan ranging from 2 to 30 s for each peak (Fig. 1E of the Electronic Appendix, available for downloading from http://petrology.oxfordjournals.org).
SHRIMP II RG (Stanford). Prior to analysis, mounts were acid-rinsed and coated with 100 nm of Au. All data were collected during a single week in six analytical sessions, using the routine procedures described by Dusel-Bacon et al. (2001). Using a 5·7–12 nA primary O2 ion beam, a 50 µm x 70 µm region was rastered for 90 s prior to each analysis. Although intensities recorded for the primary beam were apparently highly variable, the secondary beam was very stable. During data collection, the primary beam produced clean, sharp, 15–30 µm-diameter, 1 µm deep, flat-bottomed analytical pits (Fig. 1E of the Electronic Appendix).
Zircon standard R33 from the Braintree Complex in Vermont (419 Ma, Black et al., 2004, and references therein), was analyzed every fourth analysis. Zircon SL13 from a Sri Lankan pegmatite was used as the standard for U (238 ppm, Williams, 1998). Although some individual spot analyses are characterized by large imprecision, in particular for 207Pb, we chose to work with a small spot size, accepting large uncertainties on isotopic ratios rather than the risk of sampling multiple microdomains within the grain.
SHRIMP II (Canberra). FC1 and SL13 were used as reference zircons (Paces & Miller, 1993; Williams, 1998). Pb–U ratios were normalized relative to a value of 0·1859 for the 206Pb/238U of the FC1 reference zircons, equivalent to an age of 1099·0 ± 0·6 Ma (Paces & Miller, 1993).
The 5 nA primary, O2 beam produced clean, sharp, ellipse-shaped, flat-bottomed pits. Pit sizes range from 15 µm in diameter for standard analyses, to 35 µm for low-U grains. Acquisition in the latter case was performed with an enlarged beam and longer acquisition times (Fig. 1E of the Electronic Appendix). Despite the modification of the run table and the enlarged spot size, the precision on some samples was low and the relative, individual spot analyses display a large uncertainty. The full dataset is available upon request from the first author.
Data reduction
Preliminary data reduction was performed using the SHRIMP automation software based on PRAWN (Playing Retrospectively Around With Numbers, unpublished program routine of T. R. Ireland). Final data reduction and age calculations were performed using the Excel add-in SQUID (Ludwig, 2001). Results are represented graphically using ISOPLOT (Ludwig, 2003). Age calculations are based on IUGS recommended values for U decay constants (Steiger & Jäger, 1977).
Correction for Pbcom can have dramatic effects on the calculated ages, in particular when the counts on 204Pb are too low for meaningful extrapolation of the Pbcom composition. In zircons from the upper Mafic Complex, measured 206/204 ratios are below 1000 with few exceptions. Corrections for Pbcom were performed using different approaches, based both on the Stacey and Kramers model (Stacey & Kramers, 1975), and on lower-crustal Pbcom values from the literature (Cumming et al., 1987; A. Mayer, unpublished data). The latter gave mostly random results, which were thus discarded from further consideration.
For samples in which counts on 207Pb were insufficient for concordia calculations, or the spread of spot ages too large, we relied either on the calculation of an intercept, or on the weighted average of the 206Pb/238U ages. In these cases, Pbcom corrections were based on the isotope that led to higher precision in the calculated age, as low counts on the Pb masses imply an imprecise Pbcom correction.
Ages reported in Table 4 are calculated on the population of spot analyses for each sample with the 2
error of the mean age. When possible, concordia ages or discordia–intercept ages calculated through a regression to a 300 Ma Pbcom (Stacey & Kramers, 1975) are preferred to the average on the 206Pb/238U ages. When the datum is only slightly discordant, the choice not to force 206Pb/238U–207Pb/235U concordance does not significantly affect the calculated ages, but only the precision of the result. We choose to discard the strongly discordant data, as a record either of Pb loss or inheritance, or of the presence of cracks or inclusions that were not detected optically.
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RESULTS |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLESMETHODSRESULTSZIRCON MORPHOLOGICAL AND...DISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Data for the main isotopic ratios and calculated ages are reported in Tables 2 and 3 and plotted in Figures 5–9
together with representative cathodoluminescence images for each sample. Data for inherited cores are reported in the Electronic Appendix only (available for downloading from http://petrology.oxfordjournals.org).
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Zircons separates that were complex were hand-picked for splits separated on the basis of optical features such as grain size, color and shape. These splits have been reanalyzed using the SHRIMP II at ANU. Therefore, samples 7n, 8gn, 10ck and 11ag have a composite dataset of SHRIMP II RG (Stanford) and SHRIMP II (Canberra) analyses. Concordia plots show that error sizes are similar (Figs 6b–e, and 7b, c), and ages are calculated after pooling the two datasets.
In this study, the correlation observed by Vavra et al. (1999) between Th/U and 206Pb/238U age was not found. Th/U << 0·1, generally regarded as indicative of metamorphic zircons (Hoskin & Black, 2000; Hermann & Rubatto, 2003; Hartmann & Santos, 2004), was found only in some zircons from the paragneiss-bearing belt. Samples 11ag and 12ag, from the Amphibole Gabbro unit in the Val Sessera section, display recurrent values of Th/U >1, compatible with zircons crystallized from a mafic magma (see discussion by Amelin, 1998; see also Pidgeon, 1992; Vavra et al., 1999), and possibly indicative of late-stage crystallization from a mafic magma (Heaman et al., 1990).
With the exception of samples 7n and 8gn, which yielded heterogeneous populations of zircon grains in terms of shape, size and color, all samples yielded homogeneous populations that were distinct under optical microscope observation, and correspond to distinct cathodoluminescence patterns. Zircons from all samples are at least slightly resorbed (e.g. Fig. 8a), a feature previously described for zircons from the Ivrea–Verbano Zone by Vavra et al. (1996).
The SHRIMP II RG (Stanford) was unable to produce high-quality data on more than one spot per grain without repolishing and recoating the sample. Thus, analysis of multiple spots per grain to investigate core–rim age relationships was limited to only a few crystals. As a result, few core–rim comparisons exist on the same grain in the Stanford dataset, and the interpretation of multiple ages in complex samples is based on statistical treatment of the data and on the cathodoluminescence imaging.
Multiple-spot investigations performed at Canberra are indicated by the arrows in Fig. 6b and show that there is no consistent systematic core-to-rim age relation for samples 7n and 8gn. Core–rim age relationships are instead preserved in samples 9mp and 12ag (Fig. 7a and d, respectively).
Upper Mafic Complex (Val Sesia)
Val Sesia cuts across the core of the arcuate structure of the Mafic Complex. In this area, the complex consists of relatively homogeneous gabbroic and dioritic rocks. Igneous textures are locally preserved and paragneiss septa are scarce. Samples from this area include numbers 1–5. Sampling localities are shown in Fig. 2, and concordia plots, cathodoluminescence images and spot locations are reported in Fig. 5.
Zircons from these samples are characterized by uniformly low U contents (on average 107 ppm), low measured 206/204 (on average 660), and uniquely, albeit imprecisely, defined ages. Inherited cores are recognized in only one case (sample 5g, Fig. 5e). These excluded, the spread in spot ages for each sample is typically within analytical error, and therefore interpreted to represent statistical dispersion of a single age for the sample. Spot ages characterized by strong discordance or by errors >10% on the calculated ratios were discarded from age calculations, being considered results of analytical problems such as low counts on Pb and U masses as described in the Electronic Appendix (available for downloading from http://petrology.oxfordjournals.org).
Zircon morphology is mostly simple and typical of high-temperature magmas, with short, stubby prisms as the prevalent type (e.g. Fig. 5a, Corfu et al., 2003). Th/U ratios measured with the SHRIMP are always >0·2, and most fall between 0·4 and 0·7, a range considered typical for zircons grown from a melt (Hoskin & Schaltegger, 2003). Zircons from this part of the complex have either a uniform cathodoluminescence pattern or a broad compositional zoning, from a dark interior with an average [U] of 100–400 ppm, to a bright outer zone, with an average [U] of 10–50 ppm (e.g. Fig. 5a and d). The meaning of this zoning remains unclear, but could reflect either the appearance of amphibole on the solidus during final crystallization of the rock, or a change in composition of the magma from which zircon crystallized. Whatever the case, the decrease in U content in the crystal rims is not correlated with a decrease in Th/U, nor with different age results, thus excluding a later metamorphic event.
In general, Th/U, morphological and cathodoluminescence features of the zircons from the upper Mafic Complex are compatible with, and indicative of, zircons grown from a (mafic) magma.
Sample 1d—diorite from Varallo (Fig. 5a)
This sample has a more or less uniform population of subhedral to euhedral zircons that are clear, generally inclusion-free, and pink to pale pink. The zircons selected for analysis range in size from 720 µm x 140 µm to 140 µm x 70 µm. In cathodoluminescence, zircon crystals are uniform, with medium luminescence. They display no internal pattern (e.g. Fig. 5a, grains 5 and 6) or have only a weak magmatic zoning consisting of broad bands with brightness increasing toward the crystal edges (e.g. Fig. 5a, grains 2 and 10), corresponding to a decrease in U (on average, 71 ppm for the internal darker domains; 34 ppm for the external brighter domains). Th/U values span 0·37–0·80 (except for spot 3; see Electronic Appendix, available for downloading from http://petrology.oxfordjournals.org). The weighted average on the 206Pb/238U ages is 287·9 ± 3·2 Ma (2, MSWD = 1·3, P = 0·22, n = 16). Regression to common lead according to Stacey & Kramers (1975) gives an intercept of 287·8 ± 3·1 Ma (MSWD = 1·12, P = 0·33, n = 16).
For this sample, although 207Pb values are not precise, spot ages are mostly concordant and the concordia age calculation corresponds to 287·8 ± 3·2 Ma (2, MSWD = 3·4, P conc. = 0·06, n = 16). Data rejection, described in detail in the Electronic Appendix, leads to less precise ages around the same value; thus we consider the best age determination for sample 1d to be the concordia age calculated on all analyzed spots at 287·8 ± 3·2 Ma (2).
Sample 2g—gabbro from Valmaggia (Fig. 5b)
This sample is more altered than the others and had a low zircon yield. Zircons are prismatic with rounded edges and are relatively large, ranging in size from 540 µm x 180 µm to 120 µm x 80 µm. All are milky, ranging in color from white to yellow. A network of cracks cutting across the grains (Fig. 5b) has probably led to fragmentation during mineral separation, and might be responsible for the milky aspect observed under the optical microscope. In cathodoluminescence, zircons display no zoning and appear uniformly bright (Fig. 5b, after tone enhancement; dark areas are laser-ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) craters unrelated to SHRIMP analyses. The U content is low (<41 ppm), resulting in poor counting statistics, with errors on the measured isotopic ratios generally >10%. Consequently, this sample did not yield well-constrained spot ages. Using only spot ages that show a Pbcom trend (spots 4, 5, 7, 8), the intercept is 285 ± 11 Ma (MSWD = 0·082) with a concordia age at 287 ± 11 Ma (2, MSWD = 2·6). The sample was collected close to an Alpine brittle fault, and a Tertiary overprint is suggested by the brittle fractures visible in cathodoluminescence, although it remains obscure in the U–Pb zircon system. We regard the best estimation of the age of the sample as the concordia age of 287 ± 11 Ma (2).
Sample 3g—gabbro from Val Mastallone (Fig. 5c)
Zircons are clear or only slightly milky, pale pink, with sharp or slightly rounded edges, and a variety of sizes and shapes. In cathodoluminescence they display a zoning consisting of broad bands with brightness increasing toward the crystal edges. The oscillatory zoning is truncated by an external brighter rim, corresponding to a decrease in U (on average, 405 ppm for the internal dark domains; 56 ppm for the external bright domains), but with no difference in the age-record. The weighted average of the 204Pbcom-based 206Pb/238U ages is 286·5 ± 6·4 Ma (95% confidence, MSWD = 5·5, n = 11), and the discordia intercept returns a value of 287·2 ± 6·6 Ma (MSWD = 5·8). After data rejection based on low U and Pb counts (spot 11, 206Pb/204Pb = 226; spot 3, 7 ppm U), the weighted average 206Pb/238U age is 286·8 ± 6·8 Ma (95% confidence, MSWD = 5·6, n = 9) and the regression to common lead is 287·7 ± 7·1 Ma (MSWD = 6·0). From the cathodoluminescence and the overall age spread from about 275 to 300 Ma, we suspect that older cores are present within the zircons from this sample, although they were not detected as ages older than Carboniferous. Thus, we prefer a conservative estimate for the age of sample 3g, given by the 286·8 ± 6·8 Ma (95% confidence) weighted average age on nine spot analyses.
Sample 4g—gabbro-diorite from Val Duggia (Fig. 5d)
The sample has a homogeneous population of elongated zircons with slightly rounded edges, and a range in size from 60 µm x 130 µm to 100 µm x 340 µm. The most elongated grains are clear and pink, but in general contain inclusions. In cathodoluminescence, they display either no internal pattern or a faint zoning with low contrast; in the latter case, a bright external rim truncates the internal oscillatory zoning. Cores and rims were measured on the same grain. In these cases, the ages of cores and rims agreed within error (e.g. spots 17 + 24, 15 + 23), supporting the interpretation that U zoning in these zircons is magmatic (Fig. 5d). This is further discussed in the Electronic Appendix (available for downloading from http://petrology.oxfordjournals.org).
The age results reflect the low U content of most analyzed spots (Fig. 5d). The weighted average of the 204Pbcom-based 206Pb/238U ages is 287·9 ± 3·7 Ma (95% confidence, MSWD = 3·5, n = 26), and the intercept results at 289·4 ± 3·3 Ma (95% confidence, MSWD = 2·5, n = 22). Rejection of some analyses from the rim portions of six grains on the basis of large errors related to low counts on 204Pb (206/204: spot 5: 531, spot 8: 545, spot 10: 375, spot 17: 364, spot 21: 363, spot 22: –537), yields a weighted average 206Pb/238U age of 289·5 ± 3·9 Ma (95% confidence, MSWD = 3·37, n = 16). Although a concordia age based on the only five concordant analyses with small 207Pb errors yields 286 ± 3 Ma (2, MSWD = 2·3, P = 0·13), our preferred age for the sample is 289·4 ± 3·3 Ma, resulting from a regression to Pbcom.
Sample 5g—gabbro from Morca (Fig. 5e)
Zircons in this sample constitute a heterogeneous population in terms of shape, color and dimensions. Large grains are milky regardless of color, which can be white, pale or dark yellow, or orange. All crystal edges are rounded, and most of the grains are intensely broken, which might cause the milky aspect of the crystals. Of the 19 spot analyses, two are non-reducible because of extremely high 204Pb content (spots 14 and 15: 206/204 = 147 and 70, respectively), and three analyzed inherited cores with 206Pb/238U ages of 623 ± 12, 945 ± 17, and 1710 ± 26 Ma (Fig. 5e: spots 2, 17 and 3; data in Table 1E of the Electronic Appendix).
The weighted average of the 204Pbcom-based 206Pb/238U ages is 281·2 ± 6·6 Ma (95% confidence, MSWD = 1·9, P = 0·02, n = 14). After data rejection based on error size (spots 5, 6, 8 and 11, errors on ratios >15%), the weighted average 206Pb/238U age is 283·4 ± 4·7 Ma (2, MSWD = 1·15, P = 0·32, n = 10). Rejection of errors >10% yields a concordia age of 284·6 ± 5·2 (2, MSWD = 0·52, P = 0·47, n = 5).
A selection based on the best U content yields 286·4 ± 5·8 Ma (2, MSWD = 1·2, P = 0·27, n = 3, spots 16, 18 and 19: 444, 210 and 323 ppm, respectively; average on remaining 11 spots: 7 ppm). Considering that this datum is less affected by analytical errors, this is our preferred age for this sample. Although zircons from sample 5g display cathodoluminescence patterns typical of recrystallized grains, the best analytical results on zircons 16, 18 and 19 indicate an age similar to that of the other samples from the upper Mafic Complex.
Paragneiss-bearing belt and lower Mafic Complex
Samples 6–12 were collected in Val Sessera from the paragneiss-bearing belt and the lower Mafic Complex, which include the deepest exposed rocks of the Mafic Complex (Fig. 2). Here, the relatively homogeneous, granoblastic amphibole gabbro of the lower Mafic Complex is overlain by the paragneiss-bearing belt, within which thin, restitic paragneiss septa are interlayered with norite and leuconorite. Charnockitic rocks are common within and in proximity to the septa. Two samples (11ag and 12ag) were collected from the lower Mafic Complex; the other samples are norites, and one restitic metapelite and associated charnockite within the paragneiss-bearing belt (Fig. 4).
Zircons in all but one sample from this area display a complex patchwork of cathodoluminescence patterns and have age distributions that are too broad to be interpreted as statistical dispersion of a single age (Figs 6 and 7). Inherited cores, identified in cathodoluminescence by either discordant or well-distinguished zoning patterns, are common in most samples from the paragneiss-bearing belt (samples 7–10). Up to 100 µm thick, inclusion-free, bright rims without zoning in cathodoluminescence are displayed by zircons in some samples (7n, 8gn, 11ag and 12ag, the last displays the maximum development). In general, zircons from the paragneiss-bearing belt have U and Pb contents higher than those of the upper Mafic Complex, resulting in smaller 207Pb errors so that concordia diagrams and ages may be used with less uncertainty.
Sample 6n—norite from the paragneiss-bearing belt (Fig. 6a)
This sample yielded a homogeneous population of pink zircons, with prisms prevailing over pyramids and rounded edges, and with no inherited cores. Cathodoluminescence brightness is uniformly medium, and some crystals display oscillatory zoning patterns (Fig. 6a). The CL pattern is either flat (grains 6–8) or that of a grain crystallized from a less evolved melt (Hoskin, 2000; e.g. grains 5, 9 and 13). Most grains display a thin (<5 µm) bright rim. The concordia age calculated on all 13 spots yields 287·2 ± 2·6 Ma (2, MSWD = 0·0023, P = 0·96), coincident with the weighted average on 207Pbcom-corrected 206Pb/238U spot ages of 286·9 ± 2·5 Ma (2, MSWD = 1·6, P = 0·08, n = 13). The data define a unique age peak at 287–288 Ma. Rejection of spots 1, 2, 3, 10, and 12 based on large errors (>10%) on the measured 206Pb/238U yields a concordia age of 288·2 ± 3·1 Ma (2, MSWD = 0·0077, P = 0·93, n = 8). Our preferred age for the sample is the more precise concordia age of 287·2 ± 2·6 Ma (2), which is statistically equivalent to the 288·2 ± 3·1 Ma age. Within the paragneiss-bearing belt, sample 6n is unique in having zircons with Th/U typical for magmatic zircons (on average, 0·8) and a Gaussian distribution of spot ages similar to that of the upper Mafic Complex.
Sample 7n—norite from the paragneiss-bearing belt (Fig. 6b and c)
Zircons from this sample are characterized by elongated prisms with an aspect ratio up to 10:1 and short pyramids. Sizes range from 750 µm x 220 µm to 70 µm x 40 µm. Most grains are pink, some are yellow or orange, and a few are colorless. Cores are visible under the optical microscope, but no 206Pb/238U ages older than 323·5 ± 4·7 Ma (1, spot 3) were detected. In cathodoluminescence, the zircons display bright rims that strongly contrast with dark interiors (Fig. 6b and c).
Results for this sample are complex. Spots 6 and 16 (Fig. 6c, SHRIMP II RG data) were discarded from the calculation of the magmatic age because their younger ages are most probably a result of Pb loss. After rejecting five additional analyses on the basis of large errors (spots 8, 11, 9, 7 and 12, not reported on the plot) and two analyses on the basis of discordance (5 and 14) the remaining 15 analyses range from 277 ± 4 Ma to 308 ± 4 Ma. The results of double-spot investigations are not definitive, with calculated ages that have no core-to-rim progression.
Twenty-two additional spot analyses were collected on the SHRIMP II at Canberra with multiple spots per grain (Fig. 6b, top). These data reproduced the results of the SHRIMP II RG at Stanford, confirming that no systematic core–rim age relationships are preserved, and confirming the existence of the spread of ages observed with the SHRIMP II RG, with results ranging from 283 ± 4 Ma to 312 ± 4 Ma for individual zircons.
Sample 8gn—gabbronorite from the paragneiss-bearing belt (Fig. 6d and e)
This is a large, composite sample which consists of a quartz-norite (8n, 53% SiO2, 2·38% MgO) interfingered with a more mafic, fine-grained gabbro (8g, 47% SiO2, 4·7 MgO). Zircons were extracted from both portions 8n and 8g. They are characterized by a variety of colors, rounded edges, and elongated prisms as a prevalent shape. When a core is visible optically, the zircon displays a weak yellow–orange pleochroism. In cathodoluminescence, all crystals display a bright rim of variable thickness, which can constitute a large part of the grain. Some grains preserve a high-frequency, magmatic zoning; these are typically subhedral, elongated, and have only a thin bright rim (Fig. 6d and e). Zircons from the two fractions 8n (norite) and 8g (gabbro) were analyzed separately at the SHRIMP II (ANU), and the relative cathodoluminescence images are reported respectively at the top and bottom of the Tera–Wasserburg plot in Fig. 6d. In 8n, although cathodoluminescence images indicate the existence of core–rim relationships and resorption surfaces within a grain (e.g. grains 6–8), age relationships are not preserved. In 8n, the two populations of bright, recrystallized zircons (grains 2 and 4) and of large, oscillatory zoned, inclusion-rich zircons (grain 1) do not yield distinguishable ages. Thus, overall, in sample 8gn (Fig. 6e), a variety of cathodoluminescence patterns apparently indicates a variety of processes leading to the development of core–rim relationships; for example, recrystallized rim on a preserved core (grain 3), resorbed core overgrown by oscillatory zoned rim (grain 15/28). There is no systematic relationship between cathodoluminescence aspect, U content, Th/U, and age (see also Fig. 8a, discussed below).
Single spot analyses indicate the presence of inherited cores as old as 586 ± 14 Ma (grain 21) and 400 ± 6 Ma (grain 14). Ignoring the inherited cores (Fig. 9d, Table 1E of the Electronic Appendix), the cumulative probability plot (Fig. 9c) shows that the two portions, 8n and 8g, are distinct. Zircons extracted from 8n define a 206Pb/238U age peak at 285 Ma, whereas those from the more mafic, fine-grained portion, 8g, define a 206Pb/238U peak around 295 Ma. Both fractions contain a few zircons with older ages around 305–310 Ma. Although these differences are at the limit of significance, they might shed light on the meaning of the multiple cumulative-probability age peaks detected in the paragneiss-bearing-belt samples, and will be discussed below.
Sample 9mp—metapelite from paragneiss-bearing belt (Fig. 7a)
This is a restitic metapelite (stronalite) collected from the thickest septum of the paragneiss-bearing belt (Fig. 4). Three grains were analyzed (Fig. 7a): zircon cores are inherited with ages of 442 ± 12; 673 ± 8 and 588 ± 7 Ma (spots 1.1, 2.1 and 3.2, Table 1E of the Electronic Appendix), whereas rims (spots 1.2 and 3.1, Table 3) define a weighted average 204Pbcom-corrected 206Pb/238U age of 310·9 ± 6·6 (MSWD = 1·3, P = 0·25), and a concordia age of 310·4 ± 6·5 Ma (MSWD = 0·92, P = 0·34). Although a larger scatter may possibly remain undetected because of the restricted number of points, we tentatively interpret these age data as indicating a 
310 Ma melting event.
Sample 10ck—charnockite from paragneiss-bearing belt (Fig. 7b)
This is the most silicic of the analyzed samples and was collected close to sample 9mp. Zircons are homogeneous in size and shape, and tend to be rounded and equant, with diameters less than 200 µm. Within the zircon population, grains are milky, and exhibit many colors, including white clear, white milky, pale yellow, dark yellow, orange, pale pink and dark pink. Most grains display a thick bright rim in cathodoluminescence (Fig. 7b). The age of magmatic growth of these zircons is difficult to determine, because a high fraction of the data yields information relative to older events (Fig. 9d and Table 1E of Electronic Appendix, available for downloading from http://petrology.oxfordjournals.org).
Twenty-seven spot analyses define a large spread with two clusters of different age. The older age is defined by six concordant spot ages with a concordia age of 304·8 ± 2·4 Ma (2, MSWD = 0·24, P = 0·62), and by the three analyses on grain 10ck-8. The latter shows an age pattern similar to that of sample 9mp, with the rim analyses partially reset towards a Carboniferous event (indicated by the arrows in Fig. 7b). We interpret this age cluster as given by zircons inherited from the country rock that were largely reset during its anatexis, possibly enhanced by incorporation into the charnockitic melt. This is consistent with the rim ages of the companion metapelite sample 9mp. Two grains defined a younger concordia age of 236·4 ± 3·6 Ma (2, MSWD = 0·021, P = 0·88, n = 3). Of these two zircons, grain 3 has a magmatic, oscillatory zoned CL pattern, but Th/U 0·1, indicative of metamorphic growth. The result for this sample is a rough spread of ages from 275 Ma to 320 Ma and a minor event at 236·4 ± 3·6 Ma.
Lower Mafic Complex: amphibole-gabbro
Two samples were selected for zircon analyses from the deepest levels of the lower Mafic Complex (samples 11ag and 12ag; see Fig. 2 for sample location). In this area, the dominant lithology is a virtually zircon-free amphibole-gabbro, easily distinguishable from the gabbronorite of the upper Mafic Complex by its granoblastic texture and higher color index. Zircons were separated from noritic bands interlayered with the amphibole-gabbro. Th/U >1 was measured on most of these zircons. The cathodoluminescence patterns are complex (Fig. 7c and d) and some crystals display the ‘soccerball’ habit as defined by Schaltegger et al. (1999) for some granulite-facies zircons.
Sample 11ag—amphibole-gabbro from Val Sessera (Fig. 7c)
This sample has a uniform population of rounded zircons that are clear and bright. Th/U >1 was measured in more than 50% of the analyzed spots. Spot locations and relative concordia plots are reported in Fig. 7c. Some grains display domains with patchy zoning that we term ‘mantle’, shown in Fig. 8b and c. The ‘mantle’ is either situated between two oscillatory zoned domains or directly overgrown by a thick bright rim. Clearly, grains such as 11ag-9 record a series of events, superimposed on the primary igneous growth stages, and are, thus, formally metamorphic (Fig. 8b). The cathodoluminescence features of sample 11ag are described in more detail below.
Twenty-two spots were measured, of which two (spots 1 and 18) were rejected from further consideration because of large errors (>15%). Spots 3, 5, 12 and C7-1.2 (plus 1 and 18) were measured on bright-rim domains. These have the same Th/U ratio as the cores, but systematically show lower Th and U abundances. The ages that result from analyses on this domain are not distinguishable from those obtained on other domains, but are inevitably affected by larger analytical errors, and are not plotted in Fig. 7c. Three analyses returned discordant ages (spots 8, 10 and 14). The remaining 13 data points (Fig. 7c) show a spread of nearly concordant ages spanning 287–265 Ma, with a single older age recorded at the center of grain 16 (296·6 ± 4·4 Ma, 1). Emphasizing the equivocal nature of the determination, the age of this sample seems to be best estimated by the cluster of six spot analyses (spots 2, 6, 11, 13, C1-2.1 and C7-1.1) with relatively high U contents and low percentage of discordance, which return a weighted average 206Pb/238U age of 281 ± 3 Ma (2, MSWD = 0·73, P fit = 0·60) and a concordia age of 282 ± 3 Ma (2, MWSD = 0·016, P conc. = 0·90). However, the most important information is the spread from 287 Ma to 265 Ma, the meaning of which will be addressed below.
Sample 12ag—amphibole-gabbro from Val Sessera (Fig. 7d)
The sample yielded a composite zircon population. Most zircons are rounded and bright, with sizes less than 300 µm. Some zircons are yellow or pink, and a fraction of elongated thin prisms is also present. Eight spots were analyzed on the pink zircon split. Measured Th/U values are close to or greater than one (except for spot 4.2, and spot 12ag-1.1, which contained a thorite inclusion). Choice of spot locations was limited by the low U content of the high-luminescence domains, with consequent imprecise analyses, and by the presence of disseminated inclusions in the internal darker domains (see also Fig. 8b and c).
The weighted average age on the 208Pbcom-corrected 206Pb/238U spot ages is 280·6 ± 5·4 Ma (2, MSWD = 0·60, P = 0·76, n = 8), and the regression to common lead gives an age of 281·1 ± 3·6 Ma (2, MSWD = 0·61, P = 0·66, n = 5). Cores yield ages older or closer to concordance than the rims (Fig. 7d). Only two of the analyzed spots with U >30 ppm (spots 2.2 and 3.1) are concordant, defining a concordia age of 278·6 ± 4·9 Ma (2, MWSD = 0·53, P = 0·47), which is interpreted to be the best determination of the age of the sample.
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ZIRCON MORPHOLOGICAL AND ISOTOPIC FEATURES |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLESMETHODSRESULTSZIRCON MORPHOLOGICAL AND...DISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Whereas zircons from the upper Mafic Complex show relatively simple igneous zoning (Fig. 5), those from the paragneiss-bearing belt and the lower Mafic Complex are characterized by more complex CL patterns with features that include resorption surfaces, thick rims, and heterogeneous domains. Figure 8 shows examples of some peculiar features of these crystals, as described below.
Resorption surfaces (Fig. 8a). Within the paragneiss-bearing belt, some grains from sample 8gn display oscillatory zoning, reflecting continuous growth with depletion–enrichment events. Oscillatory zoned domains are separated by rounded surfaces (Figs 6d, e and 8a), which are a common feature for natural zircon, and have been interpreted by Schaltegger et al. (1999) as representing phases of resorption as a result of Zr-undersaturation of the melt. In general, domains that are more internally located do not preserve older ages. For grains on which both core and rim ages were measured, most of the core–rim pairs agree within error. We conclude that inherited zircon cores are not abundant or have been almost entirely reset.
‘White pest’ (Fig. 8b). Although common in most samples in this study, a thick bright rim is mainly developed in the amphibole-gabbro zircons of the lower Mafic Complex, where the rim reaches a thickness of over 100 µm (Fig. 8b). This kind of cathodoluminescence pattern was named ‘white pest’ by Zeck & Whitehouse (2002), who interpreted the bright rims to be annealed or recrystallized regions produced by a metamorphic process. According to those workers, lead loss would be prevented in a fluid-free metamorphic environment, whereas in the presence of interstitial fluid(s), the ‘white pest’ domains would be reset and record the age of such an event. The zircons analyzed by Zeck & Whitehouse (2002) do not show any isotopic or compositional record of the ‘white pest’ event. In our case, the ‘white pest’ domains are characterized by lower Th and U contents [compare Pidgeon (1992)], but maintain the same Th/U as the other domains [in contrast to Pidgeon (1992)]. Assuming that cathodoluminescence brightness is a direct indicator of crystallinity of the zircon (Kempe et al., 2000), the ‘white pest’ bright rims can be interpreted as the result of slow recrystallization, occurring from the outside inward to the interior of the crystal. This process is similar to the ‘surface-controlled alteration’ mechanism proposed by Vavra et al. (1999), except that the Th/U and isotopic ratios are not disturbed. Carson et al. (2002) have shown that significant grain-scale isotopic modifications in zircon do not take place under static high-grade metamorphic conditions. Such modifications would rather be produced in presence of pervasive deformation, as is the case for the granoblastic amphibole-gabbro and for the whole lower Mafic Complex (e.g. Quick et al., 1992). Thus we interpret the ‘white pest’ as resulting from recrystallization under long-lasting high-T deformation that occurred in the presence of late-stage melts, probably as a final phase of a prolonged magmatic growth.
Mantle domain (Fig. 8b and c). Figure 8b shows that zircons from the amphibole-gabbro consist of an inner part that is a mixture of relics of magmatically zoned zircon and domains that appear patchy and strongly heterogeneous, which we term ‘mantle domains’. This composite inner part, when present, is surrounded either by oscillatory zoned domains or by a white-pest rim (Fig. 8b).
Preliminary LA-ICP-MS rare earth and trace element data obtained at the Istituto di Geoscienze e Georisorse del CNR, Università di Pavia, indicate that the mantle domains are characterized by positive anomalies of Ca and P (both values up to 700 ppm). The strong Ca/P correlation indicates a clear ‘apatite-signature’, rather than a secondary Ca gain (Geisler & Schleicher, 2000). We suggest that the mantle domain represents a stage of rapid growth of zircon that incorporated micron-sized crystals of apatite (Fig. 8c) and possibly other phases (e.g. uraninite and monazite, indicated by high U, Th, Hf, Ce and P).
We suggest that resorption surfaces, thick white-pest rims and mantle domains, which are observed only in the lower Mafic Complex, are the consequence of repeated resorption–crystallization–recrystallization episodes, possibly fluid-mediated, during prolonged deformation in a high-T environment. This would be consistent with the growth model for the Mafic Complex proposed by Quick et al. (1992, 1994) according to which the Mafic Complex grew within a dilatant region in an extending continental crust as crystal mush flowed downward and outward away from a small, episodically replenished magma chamber located near the core of the complex in Val Sesia. This growth mechanism produced an arcuate structure, distribution of strain, and distribution of magmatic and metamorphic textures consistent with the ‘gabbro glacier’ model (Quick & Denlinger, 1992, 1993) for growth of layer 3 of the oceanic crust from small magma chambers. This model explains why igneous textures are well preserved in the core of the upper Mafic Complex whereas rocks are increasingly foliated with depth. During growth of the complex, high-temperature, synmagmatic deformation persisted for a prolonged period, and late-stage melts eventually crystallized as patches of undeformed gabbro (Quick et al., 1994). Consistent with this process, zircons from the upper Mafic Complex have habits and internal characteristics consistent with simple igneous growth, whereas those from the lower Mafic Complex have habits and complex CL patterns, which are consistent with continuous recrystallization as a result of pressure-dissolution and precipitation of the outer parts of zircon grains in such a dynamic environment. In this case, the U–Pb systematics of zircons at the deepest levels would have been continuously reworked during the growth of the complex and their ages would be skewed toward the age of the final cooling.
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DISCUSSION |
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLESMETHODSRESULTSZIRCON MORPHOLOGICAL AND...DISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Upper Mafic Complex
Samples from the upper Mafic Complex in the Val Sesia area yield age data consistent with simple and relatively uniform zircon populations with little evidence of inheritance. The ages for gabbro samples (2g, 3g, 4g, and 5g) range from 286 ± 6 Ma to 289 ± 3 (Fig. 9a). These are indistinguishable from the age of the diorite (1d, 288 ± 3 Ma), and agree well with zircon ages published by Pin (1986
), who measured a multigrain U–Pb upper intercept of 287 + 7/–5 Ma for a diorite sample collected approximately 3 km north of our sample 1d, and by Garuti et al. (2001), who obtained Pb–Pb zircon evaporation ages of 292 ± 4, 288 ± 3, and 287 ± 3 Ma for a diorite and two mafic dikes collected approximately 2 km south of our sample 1d. Collectively, the diorites and gabbros from the upper Mafic Complex in this study have identical ages within the uncertainty for each sample, and they may be regarded as a homogeneous population for which the individual age measurements can be pooled. Figure 9a shows the result of pooling the ages of samples 1–5 (Table 4); this yields a weighted average age of 288 ± 4 (2, MSWD = 0·074, P fit = 0·990). We conclude that gabbros and diorites of the upper Mafic Complex crystallized within the resolution of the SHRIMP method at 288 ± 4 Ma (2).
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Paragneiss-bearing belt and lower Mafic Complex
The simple and well-defined igneous age of the upper Mafic Complex contrasts with the results from the paragneiss-bearing belt and the lower Mafic Complex, which spread over 40 Myr (Table 4). This difference is not an artifact of analytical errors, because zircons from the lower Mafic Complex have generally higher U and Pb concentrations, and the spread of data is larger than the 2
uncertainty of individual spot analyses. Within the paragneiss-bearing belt, zircons from norite sample 6n show an igneous cathodoluminescence pattern without inherited cores and provide the same age, within error, as the upper Mafic Complex (287 ± 3 Ma). This age is very close to the youngest peak shown in the cumulative probability plots of norite samples 7n and 8n, whereas the ‘youngest peak’ of the more mafic portion 8g is about 295 Ma. These samples were collected close to paragneiss septa and, excluding 6n, contain inherited zircons. The distribution of spot ages suggests that most inherited zircons were partially reset at about 310 Ma.
Carboniferous ages
In charnockite 10ck, excluding inherited ages older than 350 Ma, spot analyses span a range of about 315–235 Ma, with a cluster of data defining an age of 305 ± 3 Ma, which is compatible with the ‘310 Ma event’ recorded in the norites. In the associated restitic paragneiss (9mp), zircon cores are obviously detrital, whereas the rims give an age of 310 ± 7 Ma, which also corresponds to the age of the reset zircons in the norites and in the charnockite. Considered collectively, 18 analyses performed with the SHRIMP II and the SHRIMP II RG on samples from the paragneiss-bearing belt (7n, 8gn, 9mp, and 10ck) suggest the presence of a meaningful age at 310 Ma. The distribution of 206Pb/238U ages is shown in Fig. 9b. The weighted average on the 206Pb/238U ages is 309·2 ± 2·7 Ma (95% conf., MSWD = 1·7, P = 0·040, n = 18), indicating a Carboniferous event that affected the Kinzigite Formation more than 20 Myr before the final crystallization of the upper Mafic Complex. A regional granulite-facies metamorphic event between 320 and 300 Ma was previously inferred by Handy et al. (1999), based on the presence of 310–290 Ma monazite ages in the Kinzigite Formation (Henk et al., 1997; see also Graeser & Hunziker, 1968; Köppel, 1974; Köppel & Grünenfelder, 1979; Ragettli et al., 1992) and on 206Pb/238U-weighted average SHRIMP ages on 10 overgrowth rims at 296 ± 12 Ma on zircons from a metapelite in Rimella (Vavra et al., 1996). Vavra et al. (1999) found an age difference between a stronalite sample (STR1r, mean age 310·6 Ma) and all the other five restite and leucosome samples (299 ± 5 Ma) collected on the same traverse; those workers concluded that this difference ‘is on the edge of significance and must be verified by further investigations’. The re-occurrence of the 310 Ma age in our dataset provides additional support for a late Carboniferous granulite-facies metamorphic event in the Kinzigite Formation of the Ivrea–Verbano Zone.
It is possible that the 310 Ma age corresponds to the onset of the magmatic intra-plating process that generated the Mafic Complex, and that the growth of the entire gabbro body occurred by episodic injections of mafic magma from about 310 to 288 Ma. With the present dataset, however, we cannot exclude the possibility that the 310 Ma event is completely unrelated to the intrusion of the Mafic Complex, although in this case the heat supply for the metamorphic event is unidentified.
Permian records in the lower Mafic Complex
In the amphibole-gabbro, most zircons are rounded, largely affected by the ‘white pest’, and have a peculiar Th/U >1. The features of the zircons are consistent with the granoblastic texture of their host rocks (Müller, 2003). Spot analyses range from about 300 Ma to 265 Ma. A few euhedral pink crystals define an age of 279 ± 5 Ma. In any case, the range of spot analyses in the lower Mafic Complex overlaps that of the upper Mafic Complex, suggesting that the upper and lower Mafic Complex are part of a single prolonged episode of sequential emplacement rather than genetically independent intrusions.
The lower Mafic Complex constitutes the deepest part of the Mafic Complex, and features within it are best interpreted as the consequence of extensional shear of partially molten cumulates during slow cooling, as might be expected during the growth of such a deep-seated intrusion (Quick et al., 1992, 1994). Under these conditions, zircons that are either interstitial or included in high-T phases would be affected by fluids or melts for a prolonged period, resulting in a spread of ages, possibly analogous to the behaviour of zircon in intermediate to acidic melts, in which zircon is one of the first phases on the solidus (e.g. Sergeev et al., 1995; Klötzli & Parrish, 1996). A similar process might be responsible for the spread in ages recorded in sample 3g from the core of the upper Mafic Complex. The habit and internal textures of zircons in the lower Mafic Complex and their broad age distribution can be interpreted as the result of progressive growth and recrystallization during the growth of the complex. If so, the age of onset of the intrusion of the lower Mafic Complex might have been largely obliterated by the recrystallization processes, and may be approximated only by the oldest non-inherited age determination.
The 295 Ma age-peak of the mafic portion of sample 8g might record intrusion of mafic magma into the paragneiss-bearing belt and consequent anatectic melting, followed at 285 Ma by the crystallization of the most acidic quartz-norite 8n. In this view, an age of 295 Ma is close to the onset of the intrusion of mafic magma into the paragneiss-bearing belt, and the paragneiss remained partially molten for about 10 Myr. Consistent with this limit, the lower Mafic Complex may have been intruded at about 310 Ma, thus representing a possible heat source for the granulite-facies metamorphic event in the Kinzigite Formation, and resided in the lower crust at high temperature as partially molten, deforming cumulates until at least 285 Ma.
Heat flow and cooling Sm/Nd ages
Sm–Nd mineral isochron ages published for the Mafic Complex range from 274 ± 11 to 244 ± 4 Ma (Voshage et al. 1987; Mayer et al., 2000). As shown in the compilation of Fig. 3, the youngest Triassic ages are obtained on garnet-bearing assemblages, whereas the Permian ages are obtained on garnet-free assemblages. We interpret these ages to date the cooling of mineral assemblages below the blocking temperature for the Sm–Nd system, which is probably lower in garnet-bearing assemblages (i.e. <750°C; see Mezger et al., 1992; Mayer et al., 2000). The span of at least 10 Myr between Sm–Nd internal isochrons and U–Pb zircon ages indicates slow cooling of the complex, which would be consistent with thermal conduction as the principal mechanism for heat dissipation.
Figure 10 presents the results of a one-dimensional, finite-difference, conductive heat-flow model for the Mafic Complex using temperature-dependent thermal conductivities from Clauser & Huenges (1995) and heat-capacity data from Robie & Hemingway (1995). This calculation suggests that tens of millions of years are required to cool conductively the center of the Mafic Complex from near-liquidus temperatures to the blocking temperature for the Sm–Nd system in gabbroic assemblages. Because S-type granitic melts, similar to those that crystallized within the Kinzigite Formation, have solidus temperatures of the order of 750°C (e.g. Miller et al., 2003), it follows that the deep crust could remain partially molten for tens of millions of years following a large under- or intra-plating event.
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The prolonged presence of late-stage melts would be consistent with the sporadic presence of zircons that yield ages in the range 265–235 Ma (Fig. 3). Examples include the U–Pb zircon age of 251 ± 2 Ma reported by Wright & Shervais (1980
) for a single zircon from a pegmatitic albite granite from a deep level of the complex in Val Sesia, the 249 ± 7 Ma age reported by Vavra et al. (1999) for zircon rims (ZCA, Fig. 3), the 261 ± 4 Ma age reported by Vavra et al. (1996) for zircon overgrowths in the psammitic layer of a metasedimentary granulite from Rimella (Val Mastallone), and the 265 + 4/–5 Ma age reported by Gebauer et al. (1992) for metamorphic domains in zircon from a (meta-) pyroxenite in the vicinity of the Balmuccia peridotite, and possibly also the event that affected sample 10ck at about 235 Ma.
These results are consistent with those of Annen & Sparks (2002), who showed how heat is advected into the zone of intrusion at a much faster rate than it can be conducted away, and that reverse geothermal gradients are a characteristic feature of under- or intra-plated zones. Clearly, the thermal anomaly caused by the intrusion of the Mafic Complex in the Kinzigite Formation required tens of millions of years to decay back to the background geothermal gradient after the end of gabbro-glacier growth.
The cooling front was most probably irregular, and it is likely that it was locally perturbed by minor intrusive episodes such as those described above (Wright & Shervais, 1980; Gebauer et al., 1992; Vavra et al., 1996). Thus, it is possible that most of the geochronological determinations performed in the last 35 years on samples from the Ivrea–Verbano Zone do represent meaningful age determinations (closure temperatures included), that record discrete events during the slow cooling of the Ivrea–Verbano Zone, and should not be discarded as unreliable analytical results.
No magmatic events younger than 235 Ma were detected in the zircons from the southern Ivrea–Verbano Zone in this study. This contrasts to the northern Ivrea–Verbano Zone, where, in the Finero area, numerous zircon U–Pb ages as young as 208–225 Ma have been reported (e.g. Stähle et al., 1990, 2001; von Quadt et al., 1993; Lu et al., 1997; Oppizzi & Schaltegger, 1999; Grieco et al., 2001; Peressini et al., 2004). In the southern Ivrea–Verbano Zone, the only Triassic U–Pb zircon ages are the age of 236·4 ± 3·6 Ma determined on two grains from charnockite 10ck, and similar conventional U–Pb ages on zircons from a metapelitic septum near the village of Isola (Saller et al., 1991; T. Reischmann, personal communication, 2006). Triassic ages determined by Bürgi & Klötzli (1990), Biino & Meisel (1996) and Brodie et al. (1989) are based on Re–Os, Ar–Ar and Rb–Sr dating, which are isotopic systems that might have been partially reset during Triassic events. Vavra & Schaltegger (1999) documented monazite and zircon rejuvenation down to c. 210 Ma in granulite-facies rocks from the Ivrea–Verbano Zone, and interpreted these age-records as the product of local, Late Triassic or Early Jurassic fluid activity. Therefore, we cannot exclude the possibility that the Mesozoic thermal pulses that affected the northern Ivrea–Verbano Zone in the Finero area may have also affected the southern Ivrea–Verbano Zone, although the signal in the south appears to be much weaker.
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLESMETHODSRESULTSZIRCON MORPHOLOGICAL AND...DISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Whereas the zircons from the upper Mafic Complex collectively provide a simple and well-defined magmatic age of 288 ± 4 Ma, those from the lower Mafic Complex show more complex cathodoluminescence patterns and a large spread of ages. This spread is not an artifact of larger errors, but rather reflects inheritance and/or long-lasting episodes of zircon growth. Rocks deep in the complex have high-T foliation and recrystallization textures, and it is likely that diffusion and recrystallization during prolonged high-T deformation attending the growth of the complex affected the zircons. Disregarding inherited ages older than 350 Ma, the 206Pb/238U ages in the lower Mafic Complex cluster at an age of about 310 Ma and form a somewhat less well-defined group of ages between about 295 and 285 Ma. The cluster of data at about 310 Ma suggests a high-grade metamorphic event that affected the Kinzigite Formation about 20 Myr before the crystallization of the upper Mafic Complex, and would be in agreement with the observations of Barboza et al. (1999
), who concluded that the granulite-facies metamorphism of the Kinzigite Formation pre-dates the emplacement of the upper Mafic Complex. We interpret the group of ages between 295 and 285 Ma to be an indication of the ages of intrusion and(or) recrystallization related to the emplacement of the lower Mafic Complex, although more work is required to better date these events.
Collectively, our data constrain with good precision the final crystallization of the Mafic Complex, whereas the timing of the onset of intrusion might have been partially obliterated by recrystallization processes that affected the lower Mafic Complex during the growth of the intrusive body. U–Pb data on zircons preserve age-records that are probably closer to the emplacement age of the pluton than any other applicable isotopic system. In the lower Mafic Complex, zircons may have lost their primary magmatic memory and provide a large scatter of cooling ages. In contrast to the upper Mafic Complex, we interpret the younger age of samples 11ag and 12ag (280 Ma) to be indicative of a cooling front that descended slowly through the crust after the growth of the Mafic Complex.
The range in ages from 288 Ma defined by zircons to 274 ± 11 Ma defined by Sm/Nd mineral isochrons indicates that the complex took about 10–15 Myr to reach the closure temperature of the Sm–Nd system in garnet-free mafic rocks. The time range is about 20 Myr longer to reach the closure temperature of garnet-bearing mafic assemblages, which give ages close to 245 Ma. Although the latter age may be an artifact of disequilibrium between garnet and plagioclase, it is very close to the 251 ± 2 Ma age of Wright & Shervais (1980), and to the Pb-loss age of sample 10ck. We conclude that the timescale for cooling of the complex was of the order of tens of millions of years, in agreement with the results of one-dimensional thermal modeling based on conductive heat transfer.
In conclusion, our data describe the intra-plating event in the Ivrea–Verbano Zone as follows. High-grade metamorphism of the Ivrea–Verbano-Zone lower crust began at about 310 Ma. The relationship of this metamorphic event to the intrusion of the Mafic Complex remains uncertain. We cannot exclude the possibility that high-grade metamorphism was driven by heat released by injections of mantle-derived magmas as dikes and sills (e.g. Dufek & Bergantz, 2005) during the onset of magmatism that ultimately formed the Mafic Complex. However, it remains possible that high-grade metamorphism and intrusion of the Mafic Complex were distinct events. In any case, it is clear that intrusion of mantle-derived magma culminated by about 288 Ma with the formation of the voluminous upper Mafic Complex, and an additional 10 Myr was required for the Mafic Complex to cool below the closure of the Sm–Nd system in mafic rocks. Clearly, the emplacement of substantial volumes of mafic magma in the lower crust, as exemplified by the Mafic Complex of the southern Ivrea–Verbano Zone, could produce a thermal anomaly large enough to maintain partially molten rocks in the lower crust for tens of millions of years. The presence of large, long-lived, partially molten bodies in the deep crust should be considered when interpreting the petrogenesis of volcanic rocks and the rheology of the crust in large, long-lived volcanic provinces, such as the Rio Grande Rift, Yellowstone caldera, and the Basin and Range Province of the western United States (Gans & Bohrson, 1998).
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TOPABSTRACTINTRODUCTIONGEOLOGICAL SETTINGSAMPLESMETHODSRESULTSZIRCON MORPHOLOGICAL AND...DISCUSSIONCONCLUSIONSSUPPLEMENTARY DATAREFERENCES |
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Supplementary data for this paper are available from Journal of Petrology online.
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ACKNOWLEDGEMENTS |
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This paper was greatly improved by the careful reviews by G. Bergantz, U. Klötzli, U. Schaltegger, R. Parrish, J. Epstein and C. W. Naeser. This work has been made possible by the supportive assistance in all phases of the research provided by Ulrike Poller and Wolfgang Todt at MPI-Geochemie, Mainz, who are gratefully acknowledged. Joe Wooden at SUMAC is gratefully acknowledeged for support and assistance at SHRIMP II RG. Massimo Tiepolo and Alberto Zanetti are thanked for technical assistance and helpful discussions on LA-ICP-MS at CNR–University of Pavia. Philippe Turpaud provided advice and support for the editing of text and figures. Funding for this work was provided by MURST-Project PRIN 2005, and also by MIUR-‘Progetto Giovani Ricercatori-2001’ to G.P. Expenses were partly covered by a USGS Cooperative project with the University of Trieste.
*Corresponding author. Telephone: 00390405582222. Fax: 390405582213. E-mail: peressin{at}units.it
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