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Journal of Petrology Volume 42 Number 9 Pages 1729-1749 2001
© Oxford University Press 2001
Comparison of Thermochronometers in a Slowly Cooled Granulite Terrain: Nagssugtoqidian Orogen, West Greenland
1DANISH LITHOSPHERE CENTRE, ØSTER VOLDGADE 10, DK-1350 COPENHAGEN K, DENMARK
2VRIJE UNIVERSITEIT, FACULTY OF GEOLOGY, DE BOELELAAN 1085, 1081 HV AMSTERDAM, THE NETHERLANDS
Received April 14, 2000; Revised typescript accepted February 3, 2001
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONUranium–Pb sphene and apatite, and 40Ar/39Ar hornblende, muscovite and K-feldspar ages from the core of the Proterozoic Nagssugtoqidian orogen, West Greenland, are used to constrain the timing of granulite-facies metamorphism and the subsequent cooling history. Metamorphic monazite growth occurred at 1858 ± 2, 1830 ± 1 and 1807 ± 2 Ma and defines the peak of metamorphism. The uncertainty in the cooling rates has to include the error in the decay constants of the systems used. This source of uncertainty is, however, negligible if a single decay scheme is used or when the age difference between the chronometers is large (>100 m.y.). Over the last two decades increasingly higher closure temperatures have been proposed. This trend reflects the difficulty of determining ‘absolute’ closure temperatures and in using a limited number of closure temperature estimates to infer closure temperatures of other geochronometers. Cooling rates at Ussuit were 2·9 ± 1·7°C/m.y. from 1762 Ma (
670°C) to 1705 Ma (500°C), 1·5 ± 1·1°C/m.y. from 1705 Ma to 1640 Ma (410°C), and 0·9 ± 0·4°C/m.y. between 1640 and 1416 Ma (200°C). Between 1720 and 1645 Ma cooling rates in Lersletten, 60 km north of Ussuit, are indistinguishable from those at Ussuit. After 1645 Ma, however, the area cooled to 200°C at a slightly faster rate of 2·6 ± 1·2°C/m.y. KEY WORDS: 40Ar/39Ar and U–Pb geochronometers; granulite metamorphism; slow cooling; T–t path
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INTRODUCTION |
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One of the central aims of current research in the geological sciences is to constrain rates of geological processes. Age control can be provided by isotopic age determinations, whereby a broad distinction can be made between ages that record mineral (re)crystallization, and ages that date cooling of a mineral through a certain temperature range. Although it is accepted that isotopic ages can provide time constraints on (re)crystallization events, the interpretation of isotope data as cooling ages has been the focus of much more debate.
To interpret isotope data from minerals as cooling ages, the diffusion behaviour of parent and daughter isotopes through the analysed material, and the temperature both need to be quantified. Diffusion rates, to a first-order approximation, can be calculated using equations for diffusion from a few simple geometries, such as a sphere, a cylinder, and a slab. From diffusion equations, using a convenient definition of isotopic closure, Dodson (1973)
derived a simple formula for isotopic closure, which is widely used within the geological sciences (e.g. Harrison & McDougall, 1982; Wijbrans & McDougall, 1988; Hames & Hodges, 1993). This approach is expected to give consistent results provided that some simple boundary conditions that govern diffusion are fulfilled. Deviations from theoretical behaviour are expected to occur when the concentration of the daughter isotope at the grain boundary is significantly higher than zero or second-order effects such as defect-controlled diffusion or sub-grain boundary diffusion have occurred (Villa, 1994, 1998; Lee, 1995).
Discussions as to whether to interpret dates in terms of cooling (Von Blanckenburg et al., 1989; Mezger et al., 1991) or (re)crystallization ages (Corfu, 1988; DeWolf et al., 1993) have concentrated on whether or not loss of daughter isotopes through volume diffusion, the prerequisite for interpreting isotope data in terms of cooling ages, is sufficiently fast to be relevant for geological processes. Loss of radiogenic daughters by recrystallization, enhanced by fluid circulation and strain, is a more efficient process than volume diffusion. Recrystallization is, therefore, expected to be the rate-controlling process in some geological instances (e.g. Villa, 1998). An additional complication is that, even assuming that the measured dates are cooling ages, the predicted values for closure temperature for many geochronometers are still debated. Therefore work is required to evaluate the nature of diffusion in geological settings in order to constrain closure temperatures (see Table 1).
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Evaluation of the role of volume diffusion and of defining closure temperatures of various geochronometers can best be carried out in a high-grade terrain that experienced slow cooling (e.g. Mezger et al., 1991), as under such conditions one can make the case that most isotope systems were fully reset at peak temperatures. In slowly cooled terrains the uncertainty in the age determination becomes a less critical factor with respect to the whole duration of the cooling episode, which allows for a better estimation, and comparison, of the closure temperatures.
In this study, the geochronology of a slowly cooled high-grade terrain is investigated to determine if: (1) the loss of daughter isotopes before mineral closure occurred through volume diffusion; (2) the relative closure temperatures of different geochronometers are precisely known; (3) the resultant time–temperature path is a realistic reflection of the ‘true’ thermal history. The isotope systematics of different geochronometers based on the K–Ar (40Ar/39Ar) and U–Pb systems are utilized. The tectonic implications of the data will be discussed elsewhere.
We have studied the thermal histories of the Ussuit and Lersletten areas, located in the central part of the Nagssugtoqidian orogen, West Greenland. The areas were selected because structural and metamorphic studies (see Van Gool et al., 1998, 1999) suggest that the areas were not tectonized after termination of high-grade, amphibolite- to granulite-facies, Nagssugtoqidian metamorphism, suggesting that the different parts of the two areas shared a common post-metamorphic thermal history. Furthermore, the exposure of various lithologies permits the investigation of a number of geochronometers that record isotopic closure over a wide temperature range. Published geological evidence and geochronological results from this study are used to argue that volume diffusion was the rate-controlling mechanism responsible for the loss of radiogenic daughter isotopes from the geochronometers before isotopic closure. Isotope data from the Nagssugtoqidian orogen are used in a general discussion of the closure temperatures of the analysed geochronometers.
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GEOLOGICAL SETTING |
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The two areas of interest are located in the central part of the Nagssugtoqidian orogen (NO; Fig. 1) and are referred to as the Ussuit and Lersletten areas. The NO is an ENE-striking,
250 km wide belt of medium- to high-grade gneisses consisting predominantly of variably reworked equivalents of Archaean dioritic to granitic banded gneisses exposed in the southern foreland of the orogen (McGregor et al., 1991; Friend et al., 1993; Nutman et al., 1993; Kalsbeek & Nutman, 1996; Connelly & Mengel, 2000). The various suites of gneisses are intercalated with Archaean and Proterozoic metasediments, metavolcanics and ultramafic rocks (Kalsbeek & Nutman, 1996; Nutman et al., 1999). The metasediments include metapelites, psammitic gneisses, and calc-silicate-bearing rocks. Proterozoic orthogneisses referred to as the Arfersiorfik suite (with protolith ages varying from 1920 to 1870 Ma) are volumetrically subordinate and mainly restricted to the central part of the NO (Kalsbeek et al., 1987; Kalsbeek & Nutman, 1996; Connelly et al., 2000).
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Kalsbeek and co-workers (Kalsbeek et al., 1987; Kalsbeek & Nutman, 1996) suggested, on the basis of elemental and radiogenic isotopic data obtained from the Proterozoic Arfersiorfik suite, that these rocks were formed in a Proterozoic arc environment (Kalsbeek et al., 1987; Whitehouse et al., 1998). Regional Nagssugtoqidian thrusting, associated with subsequent continental collision in the NO (Bridgwater et al., 1973) started after 1873 +7/-4 Ma, the age of the youngest identified arc rocks (Connelly et al., 2000). A 1825 ± 1 Ma U–Pb zircon age from an undeformed pegmatite that cuts across the foliation associated with the regional thrusting provides a minimum age for the main thrusting episode associated with the Nagssugtoqidian orogeny (Van Gool et al., 1998, 1999; Connelly & Mengel, 2000). The preservation of melts in the high-strain zones indicates that shearing occurred at high temperatures. While high temperatures prevailed and partial melting continued to occur, the thrusts or shear zones were deformed in kilometre-scale, open, upright folds (Hamner et al., 1997; Passchier et al., 1997; Manatschal et al., 1998; Van Gool et al., 1998, 1999). This later deformation resulted in the currently observed ENE-trending structural grain of the orogen. Static recrystallization caused an obliteration of mylonitic fabrics (Van Gool et al., 1997), which indicates that the observed high-grade equilibrium textures developed during this period of metamorphism. Nagssugtoqidian metamorphism reached amphibolite- to granulite-facies conditions with metamorphic temperatures in excess of 800°C (Davidson, 1979; Mengel, 1983). A final stage of strike-slip deformation occurred in isolated, narrow, steep shear zones. Monazites from an undeformed pegmatite cutting across one of these late shear zones yield a U–Pb age of 1772 ± 2 Ma (Connelly et al., 2000). The age is interpreted as an intrusion age and suggests that Nagssugtoqidian deformation ceased before 1772 Ma.
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ISOTOPIC CLOSURE |
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Relative estimates of closure temperatures can be derived empirically by either sequencing isotope ages of different minerals with a common thermal history and assigning the highest closure temperature to the mineral with the oldest age (e.g. Cliff & Cohen, 1980
; Von Blanckenburg et al., 1989) or by modelling the crystal-chemical framework of different minerals (Dahl, 1997). Once the absolute closure temperatures of two or more geochronometers are known, an unknown closure temperature can be derived once the ages of the thermochronometers involved have been determined. Unfortunately, this approach does require initial ‘absolute’ closure temperature estimates, and different approaches have been employed to constrain closure temperatures. Careful petrographic and thermobarometric studies can potentially provide temperature estimates. There are, however, several difficulties with this approach: (1) relating these temperatures to isotope ages can be problematic; (2) many of the dated minerals are minor phases, whose mineral-forming reactions are poorly understood (e.g. not well constrained in P–T space) and not used in thermobarometric studies; (3) the temperature estimates are based on the (re)crystallization of minerals whereas stable phases are a prerequisite of the volume diffusion theory. In an alternative approach two diffusion parameters, diffusivity (Do) and activation energy (Ea), are determined by plotting the data from diffusion experiments in an Arrhenius plot (e.g. Harrison, 1981; Harrison et al., 1985; Cherniak, 1993). The slope and y-intercept of a line through the diffusion data represent the Ea and Do, respectively. Subsequently, a closure temperature can be calculated using the Dodson formula [Dodson, 1973; see Ganguly & Tirone (1999) for a more recent discussion] through a combination of the Ea and Do (Table 1), the effective diffusion radius of the analysed material, and the cooling rate during closure of the isotope system. When assessing the experimental uncertainty (Table 1) of this approach the assumption is made that no error correlation exists between the errors determined for Ea and those for Do (i.e. 
= 0). The closure temperature estimates reported in Table 1 are based on a variety of geological evidence and experimentally determined diffusion parameters. This is discussed more fully below. |
SAMPLE DESCRIPTION |
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Hornblende
Two hornblende-bearing granodiorite samples (449185 and 431279) were taken from a sheet-like body, up to 500 m thick, of Proterozoic medium-grained foliated Arfersiorfik gneiss. Several tens of metres from sample locality 431279, the gneiss is cross-cut by a coarse-grained undeformed pegmatite (sample 431283). Sample 431270 represents a similar type of pegmatite cutting across a foliated and folded Proterozoic orthogneiss. Samples 431284 and 449146 represent the mafic components of grey tonalitic to granodioritic orthogneisses. The mafic rocks form boudins in more felsic gneisses. All six samples have a granoblastic texture and contain 0·5–2 mm hornblende crystals that coexist with plagioclase, biotite, quartz and variable amounts of clinopyroxene.
Muscovite
Muscovite is rare in the orogen, which precluded its extensive application as a geochronometer (for the 40Ar/39Ar system). Sample 449186 represents a horizon of medium- to coarse-grained muscovite-bearing psammite (quartz constitutes 90% of the rock) in a biotite–sillimanite–garnet–cordierite-bearing metapelite unit of 30 m thickness. Muscovite occurs as undeformed crystals of up to 2 mm. The muscovite crystals occur in a narrow shear zone that is part of a system of post-granulite facies metamorphic shear zones that ceased to be active around 1772 Ma (Connelly et al., 2000). Sample 415528 is a muscovite–quartz–plagioclase schist from a unit of metasediments, mainly consisting of biotite-bearing schists, of several kilometres thickness.
K-feldspar
The analysed K-feldspars (413777 and 415925) are derived from fine- to medium-grained, grey biotite tonalites. The rocks contain quartz, plagioclase and biotite, and variable amounts of K-feldspar. The K-feldspar in all analysed samples occurs as anhedral 0·5–1 mm crystals. Although in places no obvious field or geochemical differences exist between Archaean and Proterozoic gneisses (Kalsbeek et al., 1987), the tonalitic gneisses analysed probably have Archaean protolith ages, because the samples were taken outside the area where Proterozoic rocks are known to occur. Geochemical and U–Pb, Rb–Sr data on the tonalitic gneisses have been published previously by Kalsbeek et al. (1987).
Sphene, apatite, plagioclase, calcite and phlogopite
Calc-silicate-bearing rocks occur as coarse-grained units, of decimetre to several tens of metres thickness, that are characterized by an alternation of pure marbles with centimetre- to metre-thick horizons of calcite–dolomite–forsterite–diopside- or calcite–dolomite–tremolite–diopside-bearing rocks. The latter have been analysed and consist of variable amounts of plagioclase, K-feldspar, phlogopite, spinel, sphene and apatite. Although the calc-silicate-bearing units are boudinaged and intensely foliated, the analysed samples (448962, 448963, 448978, 449018, 449074, 449187) show an undeformed granoblastic, polygonal texture, which suggests that deformation preceded (re)crystallization. Sphene occurs as euhedral crystals up to 2 cm long; apatite occurs as subhedral crystals up to 5 mm long. In the hand samples 448962, 448963, 448978 and 449187 sphene represents up to 2% of the total rock volume. Plagioclase, calcite and phlogopite occur in the matrix and coexist with apatite and sphene.
Monazite
The protoliths of the monazite-bearing metapelitic rocks are considered to have a Proterozoic sedimentation age, because the rocks contain zircons, regarded as detrital from their appearance, that yield Proterozoic U–Pb ages (Nutman et al., 1999). The monazite-bearing metapelites are fine- to medium-grained, well-foliated, garnet–biotite–quartz–plagioclase–K-feldspar schists containing up to 10% sillimanite, and locally significant amounts of graphite (up to 10%). The analysed pale yellow 0·25 mm monazite crystals are indistinguishable under an optical light microscope. The paragneisses occur as horizons of 10–100 m thickness intercalated with the Archaean and Proterozoic gneisses.
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ANALYTICAL PROCEDURES |
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Rock samples were crushed using a jaw crusher and a disc mill. A concentration of mineral separates was achieved by wash pan, Franz isodynamic magnetic separator and heavy liquids. Final purification was achieved by hand picking. Large (>2 mm) sphene crystals were removed from hand specimens using a scalpel.
U–Pb analyses
Fragments of 0·2–1 cm sphene and 0·02–0·5 cm apatite crystals, and single 250 µm monazite crystals, were analysed. The Pb from coexisting low U–Pb phases (plagioclase, phlogopite and calcite) was used for a common lead correction. The separates used for U–Pb analyses were washed in purified water, weighed, spiked with a mixed 205Pb/235U tracer and subsequently dissolved (see footnote to Table 2). It proved very difficult to dissolve the small fragments of sphene crystals. Although insoluble white grains evolved after the HF–HNO3 treatment, suggesting CaF2 armouring, the assumption was made that all the U and Pb was in solution. Plagioclase, phlogopite and calcite were dissolved overnight on a hotplate, using HF for plagioclase and phlogopite and HCl for calcite.
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U and Pb were separated with conventional HBr–HCl–HNO3 techniques on a AG1-X8 (100–200 mesh) resin bed in 300 µl columns. U and Pb were loaded with silica gel and H3PO4 on Re filaments. The samples were analysed on a Fisons VG Sector 54-30 instrument in static multicollector mode or in single-collector mode on a Daly detector. The total analytical Pb blank was 90 pg during the study. A mass fractionation correction of +0·070%/a.m.u., referenced to the recommended isotopic values of NBS Pb standard SRM-981 and SRM-983, was determined and applied to the measured ratios. SRM-981 analyses (n = 10) yielded 16·900 ± 0·012 for 206Pb/204Pb, 0·91400 ± 0·00060 for 207Pb/206Pb and 2·1632 ± 0·0016 for 208Pb/206Pb (all 2
of the population). SRM-983 yielded 206Pb/204Pb of 2722·4 ± 61 (n = 28). Isotope ratios, concentrations and error calculations (2 errors are reported) were calculated using the Isoplot software of Ludwig (1982). Concordia intercepts and isochron ages were calculated using Isoplot/Ex 1.0 (Ludwig, 1998).
40Ar/39Ar analyses
Hornblende and muscovite mineral separates were irradiated for 50 h in two batches in the Cd-shielded HFPIF pool-side facility, at the ECN/EU high-flux reactor in Petten, The Netherlands. The ANU laboratory standard biotite GA1550 (measured K–Ar age 97·9 Ma) was used to calculate the J factor, which could be estimated with an uncertainty of 0·5% in both irradiations. System blanks were measured between every 5–6 unknowns. Blank corrections were made based on an average of two blank runs measured before and after the unknowns. The 1 mg hornblende and 0·3 mg muscovite and K-feldspar were dated by 40Ar/39Ar laser step-heating experiments using the argon laserprobe facility at the Vrije Universiteit Amsterdam. The laserprobe consists of an 18 W continuous argon ion laser, a low-volume gas purification line with Zr–Al and Fe–V–Zr gettering, and a Mass Analyser Products Ltd 215-50 noble gas spectrometer (Wijbrans et al., 1995).
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RESULTS |
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40Ar/39Ar analyses of hornblende, muscovite and K-feldspar
Figure 1 shows sample locations, release spectra and plateau ages of all samples. Table 2 gives the analytical data for all the 40Ar/39Ar experiments, and Table 3 summarizes the analysed rocks and minerals, plateau and total fusion ages, and the percentage of 39Ar incorporated in the plateau. All hornblende and muscovite release spectra yield plateau ages, which are defined as a minimum of three contiguous steps (within 2
error) representing over 50% of the total released 39Ar with concordant ages. The argon release spectrum of K-feldspar 415925 forms a plateau, but the spectrum of sample 413777 is more complex. During the initial heating steps of this sample increasingly older ages were measured until 25–55% of the 39Ar was released. In the remaining part of the spectrum the apparent ages fluctuate, outside 2 experimental error, around a ‘plateau’.
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U–Pb analyses of monazite, sphene and apatite
Sample locations, concordia plots and ages are shown in Fig. 2, and analytical data are presented in Table 4. Low beam intensities as a result of small sample sizes caused a large relative error in some analyses. The discordance of U–Pb sphene and apatite analysis, as reported in this study, is not unique. Reverse and normally discordant analyses representing different fractions from an individual sample have previously been noted by a number of workers (e.g. Corfu et al., 1985; Krogstad & Walker, 1994; Corfu & Stone, 1998). Despite having different degrees of discordance, the different mineral fractions yield only minor variations in 207Pb/206Pb ages and have recent lower concordia intercept ages, which suggests that the analysed fractions were affected by varying, but small, degrees of recent Pb or U loss. Given that all samples analysed in this study were taken from the Earth’s surface, weathering could well be the explanation for the observed discordance.
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DISCUSSION |
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Volume diffusion or recrystallization
It has long been recognized (e.g. Jäger, 1967
) that minerals can lose radiogenic daughter isotopes during metamorphic events following the original crystallization of the rock. Before interpreting mineral isotopic data as a cooling age, and thus assigning temperature significance to the age, it must be ascertained that volume diffusion, not (re)crystallization, controlled the isotope systematics in the analysed phase.
The only recognized fluid-catalysed recrystallization in the NO is the widespread retrogression of granulite-facies gneisses to amphibolite-facies rocks. On the basis of the consistency of 40Ar/39Ar hornblende ages in both granulite- and amphibolite-facies gneisses (see also Willigers, 1999; Willigers et al., 1999), it is suggested here that the fluid-driven recrystallization of granulite-facies rocks occurred above the closure temperature of hornblende and therefore is unrelated to the age recorded by the hornblende.
The latest Nagssugtoqidian deformation ended before 1772 Ma according to Connelly et al. (2000). Although some sphene U–Pb data from this study yield ages slightly older than 1772 Ma, the crystals analysed were sampled outside the late deformation zones and are therefore unlikely to represent recrystallization ages. The other analysed geochronometers yield ages that are much younger than 1772 Ma, and thus unrelated to any recognized deformation event. Additional arguments against post-high-grade metamorphic recrystallization include: (1) the observed high-temperature equilibrium textures, which indicate a lack of recrystallization post-dating high-grade Nagssugtoqidian metamorphism; (2) the observed spread in the ages of different chronometers from a single hand specimen or a restricted area. Similar ages would be expected if the isotopic systematics were controlled by a discrete recrystallization event, either fluid or deformation induced.
Hornblende and muscovite
The progressive closing of an isotope system, during an event of slow cooling, will result in the development of diffusion gradients of radiogenic daughters across the effective diffusion domain of a given geochronometer. Diffusion profiles, however, are not necessarily reflected in the argon release spectra. First, as a result of disintegration of hydrous phases during the step-heating experiment, argon release will probably not occur through volume diffusion (e.g. Lee et al., 1991). Second, radiogenic argon in-growth for many hundreds of million years, at temperatures far below the closure temperature, will largely obliterate any diffusion profile that may span apparent ages with a range of only tens of million years. Following the diffusion theory (Dodson, 1973) the total fusion age date represents an integrated age from the entire release spectrum. This age indicates when the geochronometer passed through its ‘blocking temperature’. Given that all reported plateau ages incorporate >70% of the released 39Ar, which probably represents a mixture of core- and rim-derived argon, the plateau age will closely approximate the time passed since isotopic closure of the mineral occurred. The 1733 ± 20 Ma (n = 6) hornblende, and the 1640 ± 12 Ma (Ussuit) and 1649 ± 13 Ma (Lersletten) muscovite 40Ar/39Ar ages are interpreted as recording the time at which lithologic units in the central part of the orogen cooled through 600°C and 410°C (Table 1), respectively.
K-feldspar
The argon release spectrum from K-feldspar 415925 (Lersletten area) forms a plateau at 1569 ± 13 Ma. Although the ages of individual steps vary outside 95% experimental error, a ‘plateau’ of 1416 ± 14 Ma can be recognized in the argon release spectrum of the second K-feldspar (413777, Ussuit). The complexities reflected in the argon release spectrum could be the result of: (1) a combination of microstructures in the K-feldspar crystals and the slow cooling it experienced (e.g. Zeitler, 1987; Parsons et al., 1999); (2) intracrystalline temperature gradients, as indicated by the uneven glow of the mineral separate during the incremental heating experiments; (3) recoil of 37Ar and 39Ar from Ca-rich and K-rich domains. These complications may have resulted in an erratic release of argon. The complexities observed in the release spectrum of sample 413777 are also seen in five additional K-feldspar argon release spectrum from the NO (Willigers, 1999). On the basis of the trend in the K-feldspar data, whereby progressively older ages are measured from south to north, we assign an age significance to both K-feldspar analyses.
Monazite
The high estimates of the monazite closure temperature (>800°C; Kamber et al., 1998) and the estimates of temperatures during Nagssugtoqidian metamorphism up to 800°C (Davidson, 1979; Mengel, 1983) prevent a straightforward interpretation of the monazite ages as either cooling or growth ages. The conditions leading to monazite growth in metamorphic rocks are poorly understood (Lanzirotti & Hanson, 1996). If the monazite ages record mineral growth, it is not obvious what reactions or events are being dated. Sawka et al. (1986) argued that monazite growth was driven by dehydration reactions during prograde metamorphism and anatexis. Other studies have demonstrated that monazite growth can occur under greenschist-facies conditions and even during diagenesis (Bonn, 1988; Milodowski & Hurst, 1989). These observations illustrate the difficulty of linking monazite growth to a particular part of a metamorphic cycle.
Monazite from three samples yields distinct 207Pb/206Pb ages: 1858 ± 2 Ma (449108), 1830 ± 1 Ma (449105) and 1808 ± 2 Ma (448997). On the basis of the data at present available, it is not possible to determine whether the 1858 ± 2 Ma monazite age dates closure of the U–Pb system during cooling following a thermal event or monazite growth during prograde or peak metamorphism. The 1858 ± 2 Ma monazite age implies that post-1860 Ma temperatures, at least at a regional scale, remained below the closure temperature of the U–Pb system in monazite. It is therefore suggested here that the two youngest (1830 ± 1 Ma and 1808 ± 2 Ma) and possibly the oldest (1858 ± 2 Ma) generation of monazites ages date discrete phases of monazite-forming reactions, reflecting continuing metamorphism during the Nagssugtoqidian orogenesis.
Sphene
Uranium–Pb sphene ages record significant variation (Table 4) even for different fragments from an individual hand sample. These dates cannot be explained by differential cooling and are therefore probably related to differences in closure temperature of the analysed fragments. Variations in the closure temperature could be related to variations in the effective diffusion radius rather than compositional control. The assumption that the crystal radius equals the effective diffusion radius might be incorrect in these sphene crystals. In imperfect crystals, cracks or crystal defects can significantly reduce the effective diffusion radius (e.g. Harrison, 1981; Cherniak et al., 1991), which lowers the closure temperature of the mineral grain relative to gem quality material. Those crystals with the largest effective diffusion radius, and thus the highest closure temperature, will close first and yield the oldest ages. We make the assumption that the effective diffusion radius equals the crystal radius and this assumption is best approximated by those sphene crystals yielding the oldest ages. Unfortunately, in this study, the oldest sphene ages also yielded the largest age errors, therefore an average age of 1758 ± 12 Ma (n = 7) is used as the age dating the cooling of Ussuit through 670°C (Fig. 2).
Apatite
A closure temperature for the U–Pb system in apatite of 600°C can be calculated using diffusion data from Cherniak et al. (1991) and assuming that the crystal radii, of the analysed crystals, equal the effective diffusion radii. A closure temperature of 600°C equals the closure temperature used for the K–Ar system in hornblende. However, given that hornblende ages are consistently older and indications for discrete apatite forming reactions at low temperatures are lacking (see the Ussuit and Lersletten cooling curves) it seems that 600°C is an overestimation of the closure temperature. Cherniak et al. (1991) suggested that short circuits such as cleavage planes or fractures in apatite crystals limit the effective diffusion radii in apatite to <500 µm. On the basis of this effective diffusion radius and a cooling rate of 2°C/m.y., a closure temperature of 500°C can be calculated. This closure temperature has been used in our interpretation.
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UNCERTAINTIES IN THE DECAY CONSTANTS |
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Renne et al. (1998)
pointed out that the error of different decay constants significantly contributes to the uncertainty of an isotope age, e.g. considering solely the error on the decay constants. The uncertainties of Proterozoic isotope ages when solely considering the errors on the decay constants for U–Pb and K–Ar are 6 Ma and 20 Ma, respectively. On the basis of a comparison of U–Pb and 40Ar/39Ar ages, Min et al. (2000) suggested that the total 40K decay constant is 2% (2) lower than values currently used by geochronologists and that the expected discrepancy between U–Pb and 40Ar/39Ar ages is controlled by the uncertainty on the total 40K decay constant.
The uncertainties of the decay constants, and thus isotope ages, will propagate into the errors of timing and rates of cooling, which are based on a comparison of a series of thermochronometers. Figure 3 illustrates the effect of a 2% decrease of the 40K decay constant on the measured cooling rates in the Ussuit area. Although absolute ages will be affected by the uncertainty of the decay constant, the age differences, and thus the cooling rates, will be only slightly altered when the isotope ages are based on a single decay scheme. A cooling rate solely based on 40Ar/39Ar ages on hornblende and muscovite will decrease from 2·4 to 2·1°C/m.y., a variation that is much smaller than the uncertainty of the cooling rate. However, calculated cooling rates based on the comparison of U–Pb and 40Ar/39Ar ages will be affected by the error on the decay constants. A cooling rate based on a U–Pb sphene and a 40Ar/39Ar K-feldspar age, which are 350 m.y. different, is only slightly affected and increases from 1·7 to 1·9°C/m.y. (Fig. 3). However, the effect becomes pronounced when the age difference between the U–Pb and 40Ar/39Ar ages is relatively small. The cooling rate based on sphene and hornblende ages, which are 25 m.y. different, increases from 1·9 to 70°C/m.y. (Fig. 3). We do not consider an initial cooling rate of 70°C/m.y. very likely, because the fast cooling rate based on the comparison of U–Pb and 40Ar/39Ar ages contrasts strongly with the cooling rates calculated on the basis of thermochronometers using a single decay scheme. Alternatively, the sphene and hornblende ages could, given that after recalculation of the 40Ar/39Ar data the two ages are indistinguishable, represent a tectonic event. However, the Nagssugtoqidian orogenesis ended around 1770 Ma (Connelly et al., 2000) and there is no evidence for a regional tectonic event that resulted in a (re)crystallization of both the hornblende and sphene.
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The ages used in our interpretation do not incorporate the error on the decay constants in the age uncertainty reported. In not doing so we allow for (1) a better comparison of isotope ages based on a single parent–daughter scheme and (2) a future correction once the different decay constants have been better constrained. Errors incorporating the uncertainty of the decay constants are, however, listed in the tables. The decay constants used for the 40Ar/39Ar and U–Pb systems used by us are those recommended by Steiger & Jäger (1977).
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AN EVALUATION OF THE THEORETICAL CLOSURE TEMPERATURES |
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To assess the dependence of the closure temperature on frequency factor (Do), activation energy (Ea), size and shape of the diffusion domain and the cooling rate, the diffusion parameters are individually varied and the effect is monitored (Fig. 4). Sphene was taken as an example, but the relations shown in the figure also apply to other geochronometers. The reported 2
error of the calculated closure temperature solely reflects the errors of Do and Ea. The choice of the shape of the diffusion domain (sphere, cylinder or plate) has only a limited effect on the closure temperature, and it does not change the value of the closure temperature outside the reported 2 error. A cooling rate of 2°C/m.y. has been used in the closure temperature calculation. The ages presented in this study clearly indicate that the rocks in the Ussuit area cooled slowly, even if rather extreme theoretical closure temperatures for the different geochronometers are considered. The effect of the cooling rate on the calculated closure temperatures is therefore minor.
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The closure temperature is most sensitive to Do, Ea and the effective diffusion radius. In the error calculation the assumption has been made that no error correlation exists between Do and Ea. If, however, a negative error correlation exists between the two parameters, the reported errors on those closure temperatures estimates that are based on laboratory diffusion experiments are overestimated (Table 1). It should be noted that the effect of the error of the activation energy is more important than is the uncertainty of the diffusivity. A change of the diffusivity by its 2 error will change the closure temperature by 10°C, whereas the closure temperature changes by >30°C when the activation energy is changed by its 2 error. Although the theoretical closure temperatures might yield a higher precision than reported, inaccuracies could result from artefacts related to the diffusion experiments. Hames & Bowring (1994) pointed out that Robbins’ intensive grinding of the analysed muscovite material caused an extreme sheet-like aspect ratio, which might have increased the bulk dif-fusivity by enhancing loss from the 001 surfaces. On the basis of a field study, Kirschner et al. (1996a) argued for a higher closure temperature of the K–Ar system in muscovite than suggested by Robbins’ (1972) diffusion experiments. Villa et al. (1996) argued that Harrison (1981) overestimated the rate of argon diffusion in hornblende, as a result of mineral decomposition during the experiments and mineral impurities in the investigated material. In addition, alteration of the bulk diffusivity as a result of damage to the mineral lattice induced by the ion implantation techniques used for some Pb diffusion experiments (see Cherniak et al., 1991) or irradiation damage occurring during 40Ar/39Ar sample preparation (Hames & Bowring, 1994) could also result in an overestimation of diffusion rates.
In the calculation of the closure temperatures of the analysed minerals, the assumption has been made that the radii of the analysed crystals equal the effective diffusion radii (Table 1). Several workers (e.g. Harrison, 1981; Cherniak et al., 1991) suggested, however, that this assumption cannot be made a priori. A correct estimation of the size of the effective diffusion radii is problematic and an inappropriate choice of the effective diffusion radii can have a significant effect on the calculated value for the closure temperature. The consistency of hornblende ages throughout the NO, interpreted as an effect of homogeneous cooling across the orogen (Willigers, 1999; Willigers et al., 1999), suggests a consistent closure temperature, and thus a uniform effective diffusion radius for each of the analysed hornblendes. The apatite data from the Ussuit area illustrate the difficulty of assigning the correct effective diffusion radii to a thermochronometer.
Over the last two decades progressively higher closure temperature estimates have been proposed for many geochronometers (Table 1). These new estimates particularly affected those geochronometers with relative high closure temperatures (Table 1). The closure temperature of the U–Pb system in monazite was initially estimated at 550°C (Wagner et al., 1977), but was re-estimated at 650–700°C (Copeland et al., 1988; Parrish, 1988, 1990; Mezger et al., 1991). Recently a closure temperature of >800°C was proposed (Kamber et al., 1998). Figure 5 illustrates the different T–t paths obtained using current closure temperature estimates and those used in the early 1980s and early 1990s. The geochronometers define a curved line, in each individual graph, which steepens when the progressively more recent closure temperature estimates are used. These figures illustrate that, although the consensus on the relative closure temperatures has not changed, some of the earlier absolute closure temperature estimates are currently regarded as too low.
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On the basis of an interpolation between the closure temperatures for Sr in biotite and muscovite, Purdy & Jäger (1976) estimated the closure temperature of Ar in muscovite at 350°C. The closure temperature of the Rb–Sr system in biotite was based on an inferred temperature of 300°C for the stilpnomelane-out isograd. It is now thought that this isograd occurs at higher temperature. Therefore, the closure temperature of the K–Ar system in muscovite had to be raised (Kirschner et al., 1996b). Similarly, Von Blanckenburg et al. (1989) used the closure temperatures of the K–Ar system in hornblende and biotite (Harrison, 1981; Harrison et al., 1985) to estimate the closure temperature of the K–Ar system in muscovite. The inferred closure temperatures of muscovite will be too low, according to the current consensus, because more recent studies estimated the closure temperature of the K–Ar system in hornblende 120°C higher (Table 1).
Re-evaluation of the closure temperatures has two important implications when the T–t paths of the NO are considered: (1) according to the current general consensus the temperatures during the initial phase of cooling were previously underestimated by >250°C; (2) the overall cooling rate increases from 0·6°C/m.y. to 1·3°C/m.y. using the ‘1980s’ and ‘recent’ closure temperature estimates. Given the uncertainties of the individual closure temperature, there may be a large error in the inferred cooling histories. Well-constrained cooling histories can be obtained only by considering many geochronometers of which independent closure temperature estimates (based on, for example, experimental diffusion data) are available.
Although the predicted values of the closure temperatures of the employed geochronometers are debatable, the consistency between the determined ages and assigned closure temperatures as observed in this study and many others (e.g. Hanson & Gast, 1967; Berger & York, 1981; Mezger, 1990; Mezger et al., 1991; Krogstad & Walker, 1994) strongly suggests that there are consistent systematic differences in the rates at which volume diffusion occurs through different mineral phases and that the relative closure temperatures of different geochronometers are internally well constrained. The presence of inclusions and microstructures, and variations in diffusion as a function of the chemical composition as reported by several workers (e.g. Villa et al., 1996; Parsons et al., 1999) will almost certainly affect isotopic systematics and hence closure temperatures. Given the reproducible isotope ages of individual geochronometers, however (e.g. the hornblende ages in this study), it seems either that those effects are constant in all individual geochronometers or that there is no significant control on the closure temperature of the material analysed in this study.
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THE USSUIT AND LERSLETTEN COOLING CURVES: COOLING AFTER A SINGLE EVENT OR A RECORD OF MULTIPLE TECTONIC EVENTS |
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Before constructing a T–t path one must establish if the data reveal a single event or multiple thermal events. Repetitive crustal reheating is feasible if caused by a series of deformational events and/or emplacement of igneous bodies. Voluminous igneous bodies younger than the
1920–1870 Ma arc-related rocks are, however, unknown in the NO and the presence of post-peak metamorphic voluminous igneous bodies, emplaced below the present erosional surface, is considered unlikely. Thermal aureoles of large igneous bodies would reset locally the isotope systems of the neighbouring host rocks, which would, for example, lead to strongly variable hornblende ages. Such variation, however, is not observed in the data from the orogen (see also Willigers, 1999; Willigers et al., 1999). A series of reheating events decreasing in magnitude in a regular fashion over a 350 m.y. period is considered to be geologically unrealistic. |
THERMAL EVOLUTION OF THE USSUIT AND LERSLETTEN AREAS |
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Figure 6 shows cooling paths for the Ussuit and Lersletten areas, which combine measured 40Ar/39Ar and U–Pb ages with the appropriate closure temperatures. The starting point of the cooling curve, however, is based on U–Pb monazite and zircon ages (Kalsbeek et al., 1987
; Kalsbeek & Nutman, 1996; Connelly & Mengel, 2000; this study), and thermobarometric data (Davidson, 1979; Mengel, 1983). High-grade Nagssugtoqidian metamorphism between 1850 and 1800 Ma was associated with temperatures of 800°C (Davidson, 1979; Mengel, 1983). The Ussuit area cooled from 670°C to 500°C between 1762 and 1705 Ma, implying a cooling rate of 2·9 ± 1·7°C/m.y. Between 1705 and 1640 Ma the cooling rate decreased to 1·5 ± 1·1°C/m.y. The rate of cooling to 200°C (1416 Ma) diminished to 0·92 ± 0·4°C/m.y. Post-1645 Ma cooling in Lersletten occurred slightly faster at 2·6 ± 1·2°C/m.y.
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CONCLUSIONS |
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When erecting detailed cooling histories for an area the following points should be assessed:
- the errors in radioactive decay constants propagate into the uncertainty of cooling rates. This effect is, however, minor if the estimated cooling rate is based on thermochronometers using a single decay scheme or when the age difference between the chronometers is large (>100 m.y.).
- Over the last two decades progressively higher closure temperature estimates have been proposed, which in particular affected those geochronometers with the highest closure temperatures. The relative closure temperature estimates, by contrast, were only marginally affected. This trend reflects the difficulty in determining the ‘absolute’ closure temperatures.
- The systematic decrease in ages of a series of thermochronometers and the reproducibility of ages from individual geochronometers suggests that (micro-)inclusions, microstructures and/or chemical variations had either a limited effect on mineral diffusion behaviour or affected the individual geochronometers in a consistent fashion.
On the basis of the integration of isotope data presented in this study with structural and metamorphic field studies, it is reasoned that the relatively simple apparent cooling paths for the central part of the Nagssugtoqidian orogen are not an artefact of a more complex thermal history. Cooling in the Ussuit area between 1762 Ma (670°C) and 1705 Ma (500°C) occurred at 2·9 ± 1·7°C/m.y. Between 1705 and 1640 Ma the cooling rate decreased to 1·5 ± 1·1°C/m.y., and around 1416 Ma (200°C) the cooling rate further diminished to 0·92 ± 0·4°C/m.y. Post-1645 Ma cooling in Lersletten occurred slightly faster at 2·6 ± 1·2°C/m.y. U/Pb monazite ages of 1858 ± 2, 1830 ± 1 and 1807 ± 2 Ma are interpreted as growth ages in the Ussuit area during pro- or high-grade Nagssugtoqidian metamorphism.
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ACKNOWLEDGEMENTS |
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First and foremost we would like to acknowledge the help of the late David Bridgwater, without whose enthusiasm this work would never have been performed. J. van Gool and F. Mengel are thanked for many discussions on the regional geology and comments on earlier drafts of this paper. Helpful comments by R. Frei, K. Mezger, M. Rosing and L. M. Page led to significant improvements in this paper. J. van Gool and F. Mengel are thanked for providing additional samples. Constructive reviews by M. Cosca, G. Davies, M. Dodson, K. Mezger and P. Renne are gratefully acknowledged. This study was supported in the form of an EU-TMR Ph.D. scholarship grant (ERBFMBICT960843) to B.J.A.W. Fieldwork was funded by and carried out under the auspices of the Danish Lithosphere Centre (DLC), which receives its funding from the Danish National Research Foundation. This paper is NSG Publication 010608.
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FOOTNOTES |
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*Corresponding author. Present address: Geological Museum, Øster Voldgade 5–7, DK-1350 Copenhagen K, Denmark. Telephone: 0045-38-14-26-63. Fax: 0045-33-11-08-78. e-mail: bw{at}dlc.ku.dk
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