Radiometric dating of sedimentary rocks: the application of diagenetic xenotime geochronology

Earth-Science Reviews 68 (2005) 197 – 243 www.elsevier.com/locate/earscirev

Radiometric dating of sedimentary rocks: the application of diagenetic xenotime geochronology Birger Rasmussen * School of Earth and Geographical Sciences, University of Western Australia, Crawley, WA 6009, Australia Received 9 December 2003; accepted 24 May 2004

Abstract Recent advances in the field of geochronology have led to a greater understanding of the scale and duration of geological processes. It is currently possible to date igneous and metamorphic rocks by a variety of radiometric methods to within a million years, but establishing the depositional age of sedimentary rocks has remained exceedingly difficult. The problem is most pronounced for Precambrian rocks, where the low diversity and abundance of organisms have prevented the establishment of any meaningful biostratigraphic framework for correlating strata. Also, most Precambrian successions have been metamorphosed, rendering original minerals and textures difficult to interpret, and resetting diagenetic minerals. Xenotime (YPO4) is an isotopically robust chronometer, which is increasingly being recognized as a trace constituent in siliciclastic sedimentary rocks. It may start to grow during early diagenesis, typically forming syntaxial outgrowths on detrital zircon grains. Diagenetic xenotime occurs in a wide variety of rock types, including conglomerate, sandstone, siltstone, shale, phosphorite and volcaniclastic rocks, varying from early Archaean to Mesozoic in age. The formation of diagenetic xenotime is principally related to redox cycling of Fe-oxyhydroxides and microbial decomposition of organic matter, leading to elevated concentrations of dissolved phosphate and rare earth elements (REE) in sediment pore-waters. Xenotime has the properties of an ideal U – Pb chronometer, containing elevated levels of U (generally >1000 ppm) and very low concentrations of initial common Pb. In addition, it has an exceptional ability to remain closed to element mobility during later thermal events, and commonly yields concordant and precise dates. Because of the small size of diagenetic xenotime crystals and common textural complexities, an in situ isotopic technique with a spatial resolution of < 10 Am is required to successfully date xenotime; to date, this has only been achieved by ion microprobe. In metamorphosed sedimentary rocks, diagenetic xenotime retains its age information up to lower amphibolite facies in sandstone, and up to mid-upper greenschist facies in pelitic rocks. In many Precambrian basins (e.g., Witwatersrand Basin, South Africa), diagenetic xenotime is overgrown by chemically distinct and texturally younger xenotime related to burial diagenesis, contact metamorphism, hydrothermal alteration or regional metamorphism. With the aid of petrography, geochemical microanalysis and the use of isotopic techniques with fine spatial resolution, it may be possible to use xenotime to date early diagenesis, and potentially every major fluid and thermal event to have affected a depositional basin. D 2004 Elsevier B.V. All rights reserved. Keywords: Geochronology; Radiometric dating; Xenotime; Diagenesis; Sedimentary rocks; Precambrian

* Tel.: +61-8-6488-2666; fax: +61-8-6488-1037. E-mail address: [email protected] (B. Rasmussen). 0012-8252/$ - see front matter D 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.earscirev.2004.05.004

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1. Introduction Sedimentary rocks are the most abundant lithology in the geological record and sediments have been continuously deposited since at least 3.7 –3.8 billion years ago (Nutman et al., 1997; Appel et al., 1998). They record information on the composition of the continental crust, the chemical evolution of the hydrosphere and atmosphere, changes in climate and the evolution of life (Holland, 1984; Taylor and McLennan, 1985; Nisbet, 1987; Schopf and Klein, 1992; Bengtson, 1994; Windley, 1995). Although sedimentary rocks contain such a wealth of information, uncertainty about the age of many Precambrian sequences has prevented the establishment of an accurate temporal framework through which to understand the scale and duration of geological and biological events. Currently, the age of sedimentary rocks is usually constrained using a number of indirect dating techniques, including the dating of contemporaneous volcanic rocks, bracketing relationships of igneous and metamorphic rocks, and the dating of detrital and diagenetic minerals. However, many basins lack suitable volcanic rocks and bracketing relationships rarely constrain the age of successions to better than hundreds of millions of years. Dating detrital minerals only yields maximum ages that in most cases are of little value in determining the depositional age. Most diagenetic minerals are also unsuitable for geochronology because they form in low-temperature environments over prolonged intervals and have low closure temperatures, making them prone to later thermal resetting. The problem is particularly acute in Precambrian basins, where most successions are devoid of fossils suitable for stratigraphic correlation and have undergone some level of deformation and metamorphism, commonly resetting diagenetic minerals and seriously compromising their usefulness for radiometric dating. A recent development in sedimentary geochronology has been the identification of the U – Pb chronometer, xenotime, as a diagenetic phase in sedimentary rocks (Rasmussen, 1996; Rasmussen et al., 1998). Using an ion microprobe with 5– 10 Am spatial resolution, it is now possible to date xenotime, yielding precise dates that in some instances approximate the age of deposition (McNaughton et al., 1999;

Fletcher et al., 2000; Stern and Rainbird, 2001; Vallini et al., 2002; Rasmussen et al., 2004). The technique is still in its early stages of development, and its successful application requires a thorough understanding of the mechanism of xenotime formation, and its behaviour during diagenesis, hydrothermal alteration and metamorphism. The following section will present a brief review of the various radiometric methods available for dating sedimentary rocks (separated into traditional methods and dating diagenesis) and detail some of the difficulties of dating ancient basins. This will be followed by an account of the current knowledge regarding the occurrence, formation and geochronology of diagenetic xenotime. Because most case studies have focussed on Proterozoic successions, this discussion will primarily deal with the application of diagenetic xenotime geochronology to Precambrian strata.

2. Traditional methods 2.1. Dating interbedded volcanic rocks Lava flows and pyroclastic deposits (e.g., ashfall tuffs) can rapidly become incorporated into a depositional sequence, resulting in a succession of interbedded sedimentary and volcanic rocks that may be regarded as contemporaneous in age. Unaltered volcanic rocks are ideal for radiometric dating and consistently provide precise and stratigraphically meaningful dates (Dalrymple and Lanphere, 1969; Faure, 1986; Harland et al., 1990; Hanes, 1991; Dickin, 1995). With the K – Ar method, interbedded lavas and pyroclastic rocks are dated using mainly mica, hornblende and feldspar, while the Rb – Sr method mostly relies on mica or whole-rock samples. Conventional dating (e.g., the K – Ar method) has in recent years been superceded by 40Ar/39Ar analysis, which has become the prime method for dating Phanerozoic rocks. The precision achievable by 40 Ar/39Ar analysis has proven particularly useful in calibrating the Phanerozoic time-scale (Hess and Lippolt, 1986), and in recent years, demonstrating the synchrony of major flood basalt volcanism with periods of mass extinctions (Renne et al., 1995; Marzoli et al., 1999; Reichow et al., 2002). However, if volcanic rocks have been subjected to alteration or

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metamorphism, all methods may yield anomalous dates because of the susceptibility of mica, hornblende and K-feldspar to thermal resetting. This represents a particular problem in Precambrian rocks. Another powerful approach is U – Th – Pb geochronology of zircon from intercalated volcanic ash layers. Zircon is a common U-bearing trace phase in many volcanic rocks and is physically and chemically robust, providing concordant and very precise crystallisation ages despite metamorphism and alteration. Where the contemporaneity of the ashfall tuff can be demonstrated, the eruption age can be used as a proxy for the depositional age. Such an interpretation can be enhanced by the dating of several discrete horizons that yield consistent and meaningful dates. Two isotopic techniques are commonly used, of which isotope dilution thermal ionisation mass spectrometry (TIMS) is the most precise (Samson et al., 1989; Tucker et al., 1990, 1998; Mundil et al., 1996; Bowring et al., 1998; Ray et al., 2002; Parrish and Noble, 2003). In contrast, secondary ion mass spectrometry (SIMS), which measures only a small volume of zircon (typically 20 by 25 by 1 Am), gives dates with lower precision but with greater spatial resolution within grains (Compston and Williams, 1992; Barton et al., 1994; Byerly et al., 1996; Martin et al., 1998; Trendall et al., 1998; Rasmussen et al., 2002b; Ireland and Williams, 2003). In samples where zircon crystals are uniform (Fig. 1a) and crystallised during a single magmatic event, TIMS is preferable. However, if zircon grains are heterogeneous, for example comprising an older core and younger rim (Fig. 1b), then SIMS is better suited to resolving the age of the internal structures. The geochronology of ash beds is probably the most reliable and precise method for dating ancient sedimentary rocks, and is increasingly being applied to Precambrian sequences (Bowring and Grotzinger, 1992; Bleeker et al., 1999; Page et al., 2000; Blake et al., 2004). The approach has proven particularly useful in refining the age of the Precambrian – Cambrian boundary, and constraining key events in metazoan evolution (Compston et al., 1992, 1995; Cooper et al., 1992; Bowring et al., 1993; Grotzinger et al., 1995; Landing et al., 1998; Bowring and Erwin, 1998; Martin et al., 2000; Bowring and Schmitz, 2003). A major limitation of the approach is the absence of intercalated volcanic rocks in many sedimentary basins. In addition, many ultramafic, mafic and inter-

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mediate volcanic rocks either do not contain magmatic zircon, or contain zircon that is too small for mineral separation or dating (typically < 10 Am in size). Where suitable volcanic rocks are present, potential complications can arise due to: (1) the incorporation of older inherited crystals during eruption (xenoliths or xenocrysts) or sediment reworking (exotic detritus) (Odin et al., 1991; Baadgaard et al., 1988; Roden et al., 1990; Trendall et al., 1998; Pickard, 2003); (2) post-eruptive Pb-loss associated with metamorphism (Fig. 1c and d) or hydrothermal alteration (Pickard, 2002, 2003); and (3) metamictisation of high-U zircons (Fig. 1e and f), leading to crystal damage and loss of radiogenic Pb. Some of the complexities of interpreting U – Pb zircon age data from tuffs are perhaps most apparent when attempting to use dates to calibrate the geological time scale, where precision and accuracy are paramount. In such studies, it is often necessary to reconcile a range of ages generated by different decay schemes (U – Pb, K –Ar, Rb – Sr) and analytical techniques (e.g., SIMS and TIMS) (Harland et al., 1990; Tucker et al., 1990; Claoue´-Long et al., 1995; Compston, 2000a,b). In most instances, U – Pb ages are accurate and have precision comparable to other isotopic methods. In cases where U – Pb ages from SIMS and TIMS are inconsistent, it is often possible to distinguish between dates that are representative of eruption, from dates that are mixtures of inherited older zircon, true magmatic zircon and zircon that give young dates due to Pb-loss (Compston, 2000b). 2.2. Bracket ages In some sedimentary basins, the age of successions is confined by the radiometric dates of associated igneous and metamorphic rocks (Stockwell, 1968; Nisbet, 1987). The depositional age of a sedimentary rock must be greater than that of the igneous rock that intrudes it, but less than the age of the igneous rock upon which it was deposited. With this indirect method, the age of sediment deposition is bracketed between maximum and minimum ages. Although generally unsatisfactory for high-accuracy geochronology, such an approach may provide valuable constraints in the absence of more suitable material to date. The use of bracketing relationships is particularly common in unfossiliferous Precambrian sedimen-

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Fig. 1. (a) BSEM image of an euhedral zircon crystal from an ashfall tuff displaying simple oscillatory zoning. Rampur Shale, Semri Group, Lower Vindhyan. (b) An SEM image taken using cathodoluminescence (SEM-CL) showing zircon crystal consisting of an older euhedral core (315 F 7 Ma, 1j) and a bright rim (1.21 F 0.07 Ma, 1j; Brown and Smith, 2004). Ongatiti Ignimbrite, Taupo Volcanic Zone, New Zealand (photo courtesy of Stuart J.A. Brown). (c, d) SEM-CL images of detrital zircon grains from an upper greenschist facies quartzite displaying a bright luminescent rim (‘‘white pest’’) and a dark core. Palaeoproterozoic Mount Barren Group, Western Australia. (e) BSEM image of a highly metamict zircon crystal surrounded by magnetite (mt) and quartz (qtz). Mt Goyder Syenite, Pine Creek Inlier, northern Australia. (f) BSEM image of a high-U zircon that has undergone intense radiation-induced crystal damage. White specks are inclusions of thorite (arrows). Lamprophyre dyke, Pine Creek Inlier, northern Australia.

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tary rocks devoid of intercalated ashfall tuffs (e.g., Nisbet, 1987; Armstrong et al., 1990; Rainbird et al., 1996, 1998; Lowe, 1999). A typical example of such an approach comes from the Mesoproterozoic Bangemall Supergroup in Western Australia, which comprises a ca. 7-km-thick sequence of siliciclastic and carbonate rocks, devoid of intercalated ashfall tuffs (Muhling and Brakel, 1985; Martin and Thorne, 2002). The succession comprises the lower Edmund Group and overlying Collier Group (Martin and Thorne, 2002). The maximum age for sedimentation is derived from zircon in granites (1619 F 15 Ma, Nelson, 1998; Sheppard and Swager, 1999, and 1679 F 6 Ma, Pearson et al., 1996) in the underlying igneous basement (Gascoyne Complex). The lower Edmund Group is intruded by two generations of dolerite sills that yield U – Pb zircon and baddeleyite dates of 1465 F 3 and 1070 F 6 Ma, respectively (Wingate, 2002), whereas the overlying Collier Group is intruded only by the younger sills. These dates provide minimum ages for deposition, indicating that the lower succession (Edmund Group) is more than 1465 million years old and that the upper succession (Collier Group) is more than 1070 million years old. The absence of 1465-million-year-old sills in the upper succession provides a possible maximum age for the Collier Group. Together, the dating of igneous basement and intrusive sills, bracket deposition to within a ca. 150and 400-million-year interval for the Edmund and Collier Groups, respectively. The age of a sedimentary rock can also be constrained by dating metamorphic minerals in rocks unconformably beneath the sequence, thus providing a maximum age for the strata, or metamorphic minerals that have grown within the succession, thus yielding a minimum age. It is sometimes difficult to be certain whether a mineral formed during metamorphism or is detrital; this is a particular problem with mica in metapelites. A recent example comes from the Proterozoic Stirling Range Formation in Western Australia, where Rb – Sr whole-rock analyses of slate samples yielded an interpreted metamorphic age of 1126 F 40 Ma (Turek and Stephenson, 1966). However, with the discovery of what were thought to be Ediacaran-type fossils (ca. 600– 540 Ma) in the succession (Cruse et al., 1993), the Rb –Sr date was reinterpreted as the age of detrital mica, and hence a

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maximum age for deposition. Recent U – Pb dating of metamorphic monazite overgrowths (Fig. 2c –d) by SHRIMP (sensitive high-resolution ion microprobe) (1215 F 20 Ma; Rasmussen et al., 2002a), suggests that the Stirling Range Formation was indeed metamorphosed at ca. 1200 Ma, consistent with the earlier Rb– Sr dating. Uranium – lead geochronology of detrital zircon grains in the same succession provides a maximum age of ca. 2000 Ma, constraining deposition only to an 800-million-year interval (Fig. 3; Rasmussen et al., 2002a). A significant limitation of this approach is that depositional ages are rarely constrained to better than hundreds of millions of years. Also, there may be uncertainties about the stratigraphic relationship between the bracketed sedimentary succession and the dated rock; for instance, it can sometimes be difficult to determine whether a sedimentary succession was intruded by an igneous body or whether it was deposited upon older basement (e.g., Moorbath et al., 1987). Other drawbacks include the presence of zircon xenocrysts and xenoliths inherited at depth from the contributing source rock, or incorporated from the wall rock during late-stage emplacement. Inherited zircon can be quite common in intrusive igneous rocks, particularly in low-temperature granites (Chappell et al., 1998; Miller et al., 2003), which if dated, are likely to give erroneously ‘‘old’’ minimum ages. Another potential problem is post-emplacement thermal resetting of chronometers in igneous rocks, producing ‘‘young’’ maximum ages that may postdate the depositional age. 2.3. Dating detrital minerals A significant development in recent years has been the ability to date single detrital grains in a sedimentary rock or sediment. Such studies produce a spectrum of age populations which can provide information on the age of the provenance of a sediment, the orogenic history of ancient terranes, and palaeogeography and sediment-dispersal patterns (Kro¨ner and Todt, 1988; Davis et al., 1989; Ross and Parrish, 1991; Rainbird et al., 1992, 1997; Zhao et al., 1992; Wysoczanski et al., 1997; Adams et al., 1998; Hutson et al., 1998; Carter and Moss, 1999; Geslin et al., 1999; DeCelles et al., 2000; Fedo et al., 2003). Several techniques are used, including

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Fig. 2. (a) BSEM image of a monazite overgrowth (white) that has engulfed metamorphic Fe-oxide laths (see arrow) that are aligned with the rock fabric. Palaeoproterozoic Stirling Range Formation, Western Australia. (b) High-contrast, BSEM image of entire grain featured in (a) showing a rounded core partly surrounded by an inclusion-rich overgrowth. (c) BSEM image of monazite crystal (white) with three SHRIMP analytical pits. Palaeoproterozoic Stirling Range Formation, Western Australia. (d) High-contrast BSEM image of (c) showing a rounded, detrital core and an inclusion-rich overgrowth. The age of the core is a 207Pb/206Pb date, whereas the ages from the rim are 206Pb/235U and 208 Pb/232Th (italics) dates, given with F 1j precision. 40

Ar/39Ar dating of detrital mica, amphibole and Kfeldspar (Harrison and Be, 1983; Kelley and Bluck, 1989; Renne et al., 1990; Cohen et al., 1995; Najman et al., 1997, 2001), and U – Pb dating of detrital zircon (Drewery et al., 1987; Zhao et al., 1992; Avigad et al., 2003) and less commonly monazite (Cliff et al., 1991; Ross et al., 1991; Evans et al., 2001; White et al., 2001) and titanite (Krogh and Keppie, 1990). Apart from providing information about the source terrain, these techniques may also be used to constrain the maximum depositional age of a sedimentary rock

(Kro¨ner and Compston, 1988; Kinny et al., 1990; Rainbird et al., 1992, 1997, 1998; Maas and McCulloch, 1991; Cohen et al., 1995; Krapez et al., 2000; Nelson, 2001; Dawson et al., 2002; Rasmussen et al., 2002b; Fedo et al., 2003). For the study of Precambrian strata, detrital zircon is particularly useful because of its widespread distribution in sedimentary rocks and its physical and chemical stability. The robustness of zircon has been demonstrated by its ability to preserve age information despite later thermal events (>800 jC), so much so that it has been

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Fig. 3. Probability plot showing SHRIMP U – Th – Pb age data for detrital zircon grains (207Pb/206Pb dates, n = 82), detrital monazite cores (207Pb/206Pb, n = 20) and metamorphic monazite overgrowths (208Pb/232Th dates, n = 29).

possible to date the oldest crustal material (about 4100 – 4400 Ma) using detrital zircon grains from strongly metamorphosed, polycyclic sedimentary rocks (Froude et al., 1983; Compston and Pidgeon, 1986; Maas et al., 1992; Wilde et al., 2001; Mojzsis et al., 2001). Different strategies may be employed to date detrital zircon depending on the type of information sought. For a study of provenance, a variety of rock types (from proximal to distal deposits) may be sampled. For constraining the age of a succession, immature volcaniclastic rocks from regions of active volcanism and magmatism are the best candidates for providing detrital zircon ages that approximate deposition (Krapez et al., 2000; Nelson, 2001). However, in many Precambrian depositional basins, mature, siliciclastic sedimentary rocks tend to predominate. During SHRIMP dating, where the youngest zircon grain is being sought, those showing the least amount of abrasion are often preferentially analysed. However, the presence of euhedral zircon crystals in the mineral separates of a mature sandstone need not imply a local, young source, but may represent zircon inclusions in detrital grains that were released during sample crushing. Another strategy for obtaining the youngest detrital zircon grains is to analyse overgrowths rather than cores, to maximize the possibility of analysing the youngest detrital components (Fig. 1b). Once a ‘‘young’’ zircon grain has been identified, similar grains from the same population can be

targetted to generate a more precise maximum age for deposition. Although a useful technique in Precambrian basins, where rocks suitable for radiometric dating are scarce, U –Pb dating of detrital zircon has several limitations. For instance, the youngest detrital zircon grain or population may be significantly older than the depositional age of the sedimentary rock (Sircombe and Freeman, 1999; Nelson, 2001). Also, detrital zircon grains may not faithfully represent the age of the provenance region, as zircon is absent from many silica undersaturated igneous rocks (Krapez et al., 2000). In addition, high-U zircon grains rarely survive prolonged transportation and reworking, and are generally absent in mature siliciclastic sedimentary rocks (Heaman and Parrish, 1991). If mafic/ultramafic rocks or high-U igneous rocks are the youngest sources of detritus in a basin, they will not be represented in the detrital zircon population, in which case the age of the youngest zircon is unlikely to approximate the depositional age. Another significant drawback is the potential of zircon grains to undergo partial Pb-loss during metamorphism or hydrothermal alteration (Pidgeon, 1992; Mezger and Krogstad, 1997; Vavra et al., 1999; Zeck and Whitehouse, 2002), yielding dates that postdate the age of deposition (Krapez et al., 2000; Nelson, 2001). The ‘‘young’’ dates should be either disregarded or used as minimum ages, but in many instances, they cannot be distinguished from detrital

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ages and are incorporated into the general population, giving erroneous maximum ages. Many of the uncertainties regarding detrital zircon ages can be overcome if grains are selected following examination by scanning electron microscopy (SEM) using back-scattered electron (BSE) and cathodoluminescence (SEM-CL) imaging techniques (see Hanchar and Miller, 1993; Corfu et al., 2003). Imaging of individual grains can help distinguish internal heterogeneity such as the presence of cores and rims, areas of radiation-induced crystal damage and zones of metamorphic recrystallisation. The latter can sometimes be recognised using SEM-CL as bright rims or patchy domains within cores referred to as ‘‘white pest’’ (Fig. 1c and d) (Vavra et al., 1999; Zeck and Whitehouse, 2002). Dates from ‘‘white pest’’ domains can be highly variable and may yield primary zircon ages, ages that correspond with the timing of recrystallisation, or ages that are only partially reset.

3. Dating diagenesis 3.1. Glauconite Of all the diagenetic minerals, glauconite is perhaps the most widely used to date sedimentary rocks (Odin et al., 1978; Odin, 1982; Smith et al., 1993, 1998). The mineral occurs as mm-sized grains (Fig. 4a and b) composed of fine clay particles that form at the sediment-water interface of marine sediments (Cloud, 1955; Burst, 1958; McRae, 1972; Odin and Matter, 1981). The widespread distribution of the mineral and its demonstrated formation during early diagenesis have made glauconite invaluable for estimating the age of many Phanerozoic successions (Odin, 1982; Harland et al., 1990). However, the reliability of glauconite dating using the K –Ar and Rb –Sr methods has been questioned because it commonly yields dates that are inconsistent with the depositional age of the sediment (Hurley et al., 1960; McRae, 1972; Grant et al., 1984; Obradovich, 1988). Anomalous ‘‘young’’ ages have been attributed to the loss of radiogenic 40Ar and 87 Sr, or modifications in K and Rb contents, in response to burial diagenesis, metamorphism and weathering (Evernden et al., 1960; Thompson and Hower, 1973; Odin, 1982). In contrast, ‘‘older’’ ages have been attributed to inheritance from remnant detrital phases

in the glauconite (Cooper et al., 1971; Morton and Long, 1980; Odin and Dodson, 1982; Fischer, 1987). However, studies of glauconite from modern shelves have shown that inheritance of radiogenic 40Ar can largely be overcome by analysing evolved grains with K2O contents >6– 7 wt.% (Odin and Matter, 1981; Odin and Hunziker, 1982; Harris and Fullagar, 1989). Glauconite dating has been most effective for Cretaceous and younger sedimentary rocks, where the mineral provides about 40% of the age data for calibrating the geological time scale (Smith et al., 1998). Although some Precambrian sedimentary rocks have been dated using glauconite (Gulbrandsen et al., 1963; Tugarinov et al., 1965; Goode and Hall, 1981; Clauer, 1981), its application to dating older strata is hampered by the greater likelihood of isotopic disturbance during weathering, diagenesis or metamorphism. As the estimated closure temperatures for glauconite are relatively low (ca. 200 jC for K – Ar and ca. 230 jC for Rb – Sr; Odin et al., 1982), the possibility of thermal resetting is very high in Precambrian successions, where few rocks have escaped metamorphism. Another significant limitation is the apparent absence of glauconite in sedimentary rocks older than 2000 Ma (Odin and Matter, 1981). 3.2. Illite Diagenetic clay minerals, such as illite, have also been used to date sedimentary rocks (Clauer and Chaudhuri, 1995; Dickin, 1995). Illite is a potassic silicate mineral, which is abundant in sedimentary rocks as a detrital and diagenetic component (Wilson and Pittman, 1977; Schieber et al., 1998). Several isotopic systems (Rb– Sr, K – Ar, Sm – Nd, Pb –Pb) have been applied to the geochronology of illite in sedimentary rocks, providing information on provenance (Morton, 1985b), the age of sediment deposition (Compston and Pidgeon, 1962; Bofinger et al., 1968; Clauer, 1979, 1981, 1982; Bonhomme, 1982; Gauthier-Lafaye et al., 1996), the timing of discrete diagenetic events (Morton, 1985a,b; Hamilton et al., 1989; Hogg et al., 1993; Robinson et al., 1993; Dong et al., 1995; Rousset and Clauer, 2003), the age of hydrocarbon emplacement (Lee et al., 1985; Thomas, 1986; Liewig et al., 1987) and the timing of thermal events (Hower et al., 1963; Lee et al., 1989; Clauer et al., 1995; Evans, 1996; Zwingmann et al., 1998).

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Fig. 4. (a) BSEM image of a single glauconite grain showing compositional zonation and regularly spaced shrinkage cracks (arrows). Lower middle Eocene Tallahatta Formation, Mississippi. (b) Glauconite grain with a rounded ilmenite grain as nucleus. Lower middle Eocene Tallahatta Formation, Mississippi. (c) Photomicrograph of a K-feldspar grain (light grey) surrounded by quartz grains (grey), kaolinite booklets (bottom, dark grey) and open pore space (black). Lower Cretaceous Barrow Group, Carnarvon Basin, Western Australia. (d) An SEM-CL photo of (c) showing the presence of diagenetic K-feldspar (non-luminescent rim) around a brightly luminescent detrital core. (e) Elongate monazite nodule (white) in a low-grade metamorphic mudrock. Lower Silurian Caerau Formation, central Wales. (f) Close-up of (e) showing matrix particles (chlorite, quartz and sericite/illite) engulfed by monazite cement (white).

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The ability of clay minerals such as illite to provide an estimate for the age of deposition has been of limited success, with many studies yielding equivocal results that either predate or significantly postdate the stratigraphic age (Hurley et al., 1963; Bonhomme et al., 1982; Bonhomme, 1987). The anomalous dates may have many causes, including: (1) inheritance from detrital components; (2) uncertainties about initial isotopic ratios (Rb –Sr and Sm – Nd); (3) extended period of mineral growth (Gorokhov et al., 1993); (4) heterogeneity within analysed samples (Burley and Flisch, 1989); (5) incomplete retentivity of daughter elements; (6) uncertainties regarding decay constraints; and (7) susceptibility to thermal resetting (Whitney and Hurley, 1964; Ohr et al., 1991). The latter is a significant problem for Precambrian sedimentary rocks, in which hydrothermal alteration and metamorphism have disturbed most isotopic systems (Bonhomme et al., 1982; Toulkeridis et al., 1998). Many early studies used whole-rock shale samples for dating, however, such material consists of a mixture of detrital and diagenetic components with heterogeneous initial isotopic compositions, and is therefore unlikely to produce a meaningful isochron (Clauer, 1982; Clauer and Chaudhuri, 1995). More recent studies have emphasized the need to carefully disaggregate and separate bulk samples into differently sized fractions (Aronson and Hower, 1976; Hunziker et al., 1986; Bonhomme, 1987; Liewig et al., 1987; Reuter, 1987; Clauer et al., 1993). Typically, the finer-grained fractions of shales yield the youngest radiometric dates; thus, the finest fractions ( < 0.02 to < 0.01 Am) are considered to contain material with the highest diagenetic-to-detrital clay ratios. Nevertheless, even if separation procedures are successful in removing the detrital component, it may still be difficult to determine whether the authigenic clay fraction formed during early diagenesis, burial or metamorphism, and as such the resultant isotopic dates can only be considered as minimum ages for the time of deposition. 3.3. K-feldspar K-feldspar is a common constituent of sedimentary rocks (Walker et al., 1978; Kastner and Siever, 1979; Milliken, 1989; Worden and Rushton, 1992; Lee and

Parsons, 2003), where it can occur as syntaxial overgrowths on detrital cores (Fig. 4c and d), or less commonly, as discrete authigenic crystals in carbonate rocks (Hearn and Sutter, 1985; Spo¨tl et al., 1998). Several techniques have been used in attempts to date the diagenetic overgrowths; one approach has been to physically separate overgrowths from detrital cores and to date the two components separately (Hearn et al., 1987; Girard et al., 1988, 1989; Warnock and van de Kamp, 1999). A major problem with this procedure is the difficulty of obtaining pure separates consisting entirely of authigenic overgrowths. However, where diagenetic K-feldspar has been successfully separated, K – Ar isotopic analysis has yielded dates that appear to be consistent with diagenesis (Girard et al., 1988, 1989), or hydrothermal fluid alteration (Hearn et al., 1987; Spo¨tl et al., 1998; Warnock and van de Kamp, 1999). An alternative approach is laser probe analysis, which allows overgrowths as small as 50 Am to be dated, avoiding potential complications of incomplete separation (Walgenwitz et al., 1990; Girard and Onstott, 1991). However, despite placement of the laser beam entirely on diagenetic overgrowths, Girard and Onstott (1991) obtained mixture ages, attributed to the release of Ar from the adjacent detrital K-feldspar grain during heating. In contrast, UV lasers are strongly absorbed by K-feldspar and cause minimal heating beyond the laser pit (Kelley et al., 1994; Hagen et al., 2001). Early studies suggest that the UV laser can achieve a spatial resolution of ca. 20 Am, with little or no contribution from the adjacent detrital core, yielding 40Ar/39Ar dates that correspond with peak burial (Hagen et al., 2001). A major limitation of this approach is the possibility of Ar loss during diagenesis, metamorphism and hydrothermal alteration. Recent petrographic studies of feldspar overgrowths by SEM and transmission electron microscope (TEM) reveal the presence of micron-sized subgrains with dislocations and micropores (Worden and Rushton, 1992; Lee and Parsons, 2003). The enhanced permeability of the subgrain microstructures is likely to aid Ar diffusion (with potential closure temperatures of V 150 jC), and therefore significantly hamper the use of feldspar overgrowths for dating diagenesis.

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3.4. Monazite Monazite is a common accessory mineral in granite, pegmatite and medium- to high-grade metamorphic rocks (Palache et al., 1951; Overstreet, 1967; Chang et al., 1996), and is widely used for U –Th – Pb geochronology (Parrish, 1990; Heaman and Parrish, 1991; DeWolf et al., 1993; Williams et al., 1996; Catlos et al., 2002; Foster et al., 2002; Harrison et al., 2002). Monazite has also been documented as a diagenetic phase (Burnotte et al., 1989; Milodowski and Zalasiewicz, 1991; Lev et al., 1998; Evans et al., 2002), and a very low- to low-grade metamorphic phase (Franz et al., 1996; Rasmussen et al., 2001, 2002a; Rasmussen and Fletcher, 2002). In organic-rich hemipelagites from the Welsh Basin, monazite is a relatively widespread, albeit volumetrically minor, authigenic phase, occurring as elongate nodules up to several mm long (Fig. 4e; Milodowski and Zalasiewicz, 1991; Evans and Zalasiewicz, 1996; Lev et al., 1998). The monazite nodules range in shape from minute randomly oriented crystals with prominent tabular crystal faces to large elongate nodules. The monazite engulfs, and partly replaces, detrital and diagenetic components in the host shale (Fig. 4f), indicating that the phosphate forms by replacement of the shale matrix. The absence of monazite nodules in early diagenetic apatite or calcite concretions that preserve pre-compaction sedimentary structures, suggests that monazite growth postdated the formation of both these minerals. Also, the lack of differential compaction around the nodules indicates that the monazite formed after sediment compaction (Milodowski and Zalasiewicz, 1991). Based on petrography and geochemistry, Milodowski and Zalasiewicz (1991) argue that the monazite formed as REE migrated from compacting turbidite muds into overlying, organic-rich hemipelagites, where they precipitated as bedding-parallel nodules. The authors speculate that the initial mineral phase may have been the hydrous LREE-phosphate, rhabdophane, which was transformed to monazite during burial diagenesis or metamorphism. Later deformation and metamorphism caused the rotation of the nodules and the development of cleavage within the shale (Milodowski and Zalasiewicz, 1991).

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Analysis of the monazite nodules yielded a Pb– Pb isochron age of 422 F 24 Ma (2j, MSWD = 0.9) (Evans and Zalasiewicz, 1996), and recently, a more-refined age of 419 F 14 Ma (2j, MSWD = 1.2) (Evans et al., 2002). A Sm –Nd isochron obtained from monazite and apatite yielded an apparent age of 553 F 22 Ma (2j, MSWD = 1.6), some 100 million years older than the time of deposition (Evans and Zalasiewicz, 1996). The erroneous Sm –Nd date was attributed to changes in the Nd isotopic composition of the diagenetic fluid, from which apatite initially precipitated, followed by monazite. The resultant data produced a mixing line rather than a true isochron. In contrast, the Pb –Pb isochron date is within error of the depositional age of the mudrocks, between ca. 430 and 435 Ma, based on biostratigraphic correlation (Gradstein and Ogg, 1996). The Pb – Pb monazite date coincides with Rb –Sr whole-rock ages from nearby Ordovician and Silurian mudrocks, 414 F 6 Ma (2j, MSWD = 2.5) and 431 F 10 Ma (2j, MSWD = 2.0), interpreted to record the timing of the burial-related smectite-illite transition (Evans, 1996). Evans and Zalasiewicz (1996) conclude that monazite can be used to date the onset of compactional dewatering, and that in rapidly deposited sequences, the technique may provide useful minimum ages for deposition. 3.5. Apatite Apatite is a common uranium-bearing accessory mineral in igneous and metamorphic rocks, and has been used for U –Pb geochronology by isotope dilution mass spectrometry (Oosthuyzen and Burger, 1973; Romer, 1996; Corfu and Stone, 1998; Chamberlain and Bowring, 2000), ion microprobe (Kennedy and Dante, 1997; Sano et al., 1999a,b) and laser ablation-inductively coupled plasma-mass spectrometry (Willigers et al., 2002). Apatite is also a relatively common diagenetic mineral, and is the main phase in phosphorite deposits where it occurs as cryptocrystalline aggregates (Fo¨llmi, 1996; Lucas and Pre´votLucas, 1997). In sandstones, apatite may be present as overgrowths on detrital apatite grains (Fig. 5a) (Bouch et al., 1995), or as pore-filling and replacive cements (Miller et al., 1989), and in shales, as earlydiagenetic concretions or nodules (Fig. 5b) (Smith, 1987; Milodowski and Zalasiewicz, 1991; Morad and Al-Aasm, 1994; Lev et al., 1998). The mineral forms

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very low U concentrations and high 204Pb contents commonly leading to imprecise dates. 3.6. Carbonates

Fig. 5. (a) BSEM image of a rounded apatite grain with a syntaxial overgrowth in a greenschist-facies quartz sandstone. Palaeoproterozoic Stirling Range Formation, Western Australia. (b) BSEM image showing a round ‘‘fossil’’ fragment comprising early diagenetic apatite, in a matrix of fine quartz and clay, along with diagenetic pyrite and carbonate. Ordovician ‘‘Sleepy Hollow Member’’, Utica Shale, New York State.

at various stages of basin evolution, beginning within metres of the sediment –water interface (Ruttenberg and Berner, 1993), and continuing during burial diagenesis and metamorphism. Apatite has been used to date diagenesis by TIMS U –Pb geochronology (Cumming et al., 1987; Chandler and Parrish, 1989; Miller et al., 1989), and more recently by using a multiple collector –inductively coupled plasma– mass spectrometer (MC –ICP– MS) to generate Lu –Hf and Pb/Pb mineral isochron dates (Barfod et al., 2002). A major limitation of U – Pb apatite geochronology is the

The use of U – Pb and Pb –Pb methods for obtaining the depositional age of sedimentary carbonate rocks is a relatively recent development (Moorbath et al., 1987; Smith and Farquhar, 1989; Jahn et al., 1990; Cuvellier, 1992; Russell, 1992; Babinski et al., 1995). In the first successful application of the technique, Moorbath et al. (1987) dated a stromatolitic limestone from the Masvingo greenstone belt of southern Zimbabwe, producing a Pb/Pb isochron age of 2839 F 33 Ma (2j, MSWD = 75). The limestone was considered previously to be intruded by granite with a Rb – Sr age of 3445 F 260 Ma (MSWD = 15; Hickman, 1974), but further Rb –Sr and Pb/Pb age data, combined with a re-evaluation of the field evidence, indicated that the limestone was deposited upon ca. 2900-million-year-old granite basement (Moorbath et al., 1987). However, the limestone occurs in a region of low-pressure metamorphism, estimated to have reached 400 – 500 jC, and the possibility of resetting could not be excluded. Nevertheless, following these promising results, a number of studies on Precambrian limestones have yielded Pb/ Pb isochron dates, variously interpreted to be early diagenetic, and a close approximation of the depositional age (Jahn et al., 1990; Russell et al., 1994; Babinski et al., 1995; Woodhead et al., 1998), late diagenetic (Jahn and Simonson, 1995; Woodhead and Hergt, 1997; Bau et al., 1999) and metamorphic (Russell et al., 1996). The application of the Pb/Pb isochron method to Phanerozoic corals and carbonate rocks is hampered by the extremely small quantities of 207Pb produced from the decay of 235U. Instead, the 238U – 206Pb isochron method is used, providing ages with relatively good precision that may approximate the age of deposition (Smith and Farquhar, 1989; DeWolf and Halliday, 1991; Smith et al., 1991, 1994; Russell, 1995; Jones et al., 1995). In the first successful demonstration of the method in younger rocks, Smith and Farquhar (1989) obtained a U – Pb isochron age of 374 F 22 Ma for corals from the Devonian Arkona Limestone in Ontario, Canada, that is in agreement with the known depositional age.

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This method has potential for the dating of sedimentary sequences. However, several problems exist, including: (1) the redistribution of U and Pb during diagenesis, hydrothermal alteration and metamorphism, yielding potentially meaningless dates; (2) possible incorporation of ‘‘older’’ detrital carbonate; (3) late-diagenetic or metamorphic recrystallisation of carbonate cement, resulting in ‘‘young’’ depositional ages; and (4) relatively poor precision with analytical errors commonly z 20 Ma. In some instances, the scatter of data is so great that isochrons cannot be obtained (Jahn and Cuvellier, 1994). Despite the potential limitations, the technique may provide dates that approximate the age of deposition (Woodhead et al., 1998), and in many cases, yield diagenetic dates that can be treated as useful minimum ages. 3.7. Sm – Nd in mudrocks In recent years, the Sm – Nd method has been used in an attempt to date sediment deposition and diagenesis of shales (Stille and Clauer, 1986; Ohr et al., 1991, 1994; Bros et al., 1992; Schaltegger et al., 1994). Using this method, a series of fine-grained fractions (typically between < 0.2 and 1.0 Am) are analysed because they are considered to consist entirely of diagenetic clays, with little or no contribution from detrital components. Samples are initially disaggregated by mortar and pestle, and then the finegrained fractions (e.g., < 0.2, 0.2– 0.4 Am, etc.) are separated from coarser material by gravity settling and centrifuge. The finer-sized fractions are then leached in acid, and both the leachates and insoluble residues are analysed by mass spectrometer (see Halliday et al., 1989 for analytical details). The method relies on: (1) pronounced Sm –Nd fractionation between leachates and residues during diagenesis; (2) that the different REE hosts formed during a single event from pore fluids of the same isotopic composition; and (3) that the Sm – Nd system has remained closed subsequent to diagenesis. In one of the few Precambrian case studies, Bros et al. (1992) sampled carbonaceous shales (total organic carbon contents of 5.79 –12.6%) from the Palaeoproterozoic Francevillian Basin in Gabon. Leachates and their insoluble residues from a series of fine-grained fractions ( < 0.2, 0.2 – 0.4, 0.4 – 0.8, 0.8– 2.0 and 6.0– 50 Am) were analysed from two samples, yielding

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Sm – Nd isochron ages of 2099 F 115 Ma (2j, M S W D = 0 . 5 3 1 ) a n d 2 0 3 6 F 7 9 M a ( 2 j, MSWD = 0.519). Data from the coarser fractions ( z 0.4 Am) and whole rock samples plotted beneath both isochrons reflecting heterogeneity within the samples due to the presence of detrital components. The isochron dates, derived from the acid-soluble and insoluble pairs of the < 0.2- and 0.2– 0.4-Am fractions, are interpreted to record multiple phases of illite growth during early diagenesis (Bros et al., 1992). The Sm – Nd ages are within error of interbedded volcanic rocks (Rb– Sr isochron age of 2143 F 143 Ma; Bonhomme et al., 1982), and are older than Rb – Sr (1875 F 29 and 1867 F 78 Ma) and K – Ar ages (1500 – 1900 Ma) from clay minerals ( < 2 Am) in the same succession (Bonhomme et al., 1982). The Sm –Nd method has the advantage over the K –Ar and Rb – Sr methods of greater resistance to thermal resetting (Stille and Clauer, 1986; Bros et al., 1992), although there is evidence for isotopic resetting during metamorphism and hydrothermal alteration (Toulkeridis et al., 1994, 1998). Significant limitations of the technique include the relatively low precision that is often attained, as well as the possibility that Sm – Nd isochron ages are erroneously old because of contributions from detrital components. For instance, fine-grained sedimentary rocks from the Central Iberian Zone, Spain, yielded a whole-rock isochron age of 1524 F 173 Ma, more than 1000 million years older than the depositional age (ca. 490 Ma) (Na¨gler et al., 1992). Another potential drawback is the possibility that diagenetic mineral growth occurred over a prolonged period, and that the various authigenic phases precipitated at different stages from a fluid with evolving Nd isotopic composition (cf. Evans and Zalasiewicz, 1996). A final consideration is that the technique is relatively expensive and timeconsuming, requiring the separation and acid-leaching of numerous size fractions, prior to analysis by mass spectrometer. 3.8. Re –Os of organic-rich mudrocks The decay of 187Re to 187Os by beta-particle emission provides a potential tool for determining the age of sedimentary rocks. As with many other isotopic techniques, Re – Os dating requires that: (1) initial ratios (187Os/188Os) of contemporaneous sam-

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ples are homogeneous; (2) the Re –Os isotopic system remains closed after deposition; and (3) the samples display a range of Re/Os ratios to define the isochron. A number of recent studies have generated Re – Os whole-rock ages from organic-rich shales, yielding dates that are consistent with the age of deposition (Ravizza and Turekian, 1989; Singh et al., 1999; Cohen et al., 1999; Ripley et al., 2001; Creaser et al., 2002; Selby and Creaser, 2003). In an early application of the technique, Ravizza and Turekian (1989) generated a whole-rock isochron with an age of 354 F 49 Ma, within error of the stratigraphic age ( f 360 Ma). However, the low precision and evidence for element mobility was not encouraging. Recent studies have achieved better precision ( F 2 – 3% uncertainties) and results indicate that the Re – Os system remains essentially closed after deposition (Singh et al., 1999; Cohen et al., 1999) despite increasing levels of thermal maturation (Creaser et al., 2002). The technique is still in its infancy, but early results suggest that Re – Os dating of carbonaceous shales may provide valuable constraints on the age of deposition.

4. Diagenetic xenotime—petrography and genesis As illustrated in the previous section, there are a variety of approaches for dating diagenesis, however, all are hampered by limitations, especially when applied to Precambrian strata. The following section summarises a new method of dating sedimentary rocks; U –Th – Pb geochronology of diagenetic xenotime. In the absence of suitable volcanic rocks, this method may prove to be the most useful technique for dating Precambrian sedimentary rocks. 4.1. Background Xenotime is an yttrium phosphate (YPO4) that occurs as an accessory mineral in acidic and alkaline igneous rocks, pegmatites, carbonatites, gneisses, schists, Alpine-type vein deposits and in placer deposits (Palache et al., 1951; Milner, 1962; Mariano, 1989; Wark and Miller, 1993; Casillas et al., 1995; Bea, 1996a,b). It has a tetragonal crystal structure, and is isostructural with zircon (ZrSiO 4 ) and thorite (ThSiO4) (Burt, 1989). The optical properties of

xenotime are very similar to those of zircon, making it exceedingly difficult to positively identify using a standard microscope (Milner, 1962). Both minerals have very high refractive indices, strong birefringence, weak pleochroism and are uniaxial positive. Like zircon, xenotime is physically and chemically stable, and is ‘‘almost insoluble in acids’’ (Palache et al., 1951). The mineral was first described by the great Swedish chemist Berzelius in 1824, from a site in Ytterby, Sweden, and later named by Beudant in 1832 apparently after the Greek for ‘‘stranger to’’ and ‘‘honour’’. However, the listed derivation is ‘‘vain’’ and ‘‘honour’’, as if the mineral name were kenotime, to signify that the mineral was mistakenly thought by Berzelius to contain a new element (Palache et al., 1951). In its accepted form, the name is highly appropriate as it alludes to the fact that xenotime is small, rare and has long gone unnoticed. 4.2. Discovery of diagenetic xenotime Diagenetic outgrowths on zircon were first reported by Butterfield (1936), who described pyramidal crystals (Fig. 6) in complete optical continuity with their detrital zircon substrate. The mineral was assumed to be zircon, but a positive identification could not be made. Some speculated that the mineral ‘‘may be a rare earth mineral isomorphous with zircon’’ (Dr. A. Brammall, pers. comm., in Smithson, 1941). The implication that zircon could dissolve and reprecipitate in a diagenetic environment was considered highly improbable, and prompted the suggestion that the outgrowths were ‘‘relics of the outer layers of strongly zoned zircon crystals, chipped off along cleavageplanes during attrition and transport’’ (see discussion and reply in Smithson, 1940). Despite the uncertainty about the nature of the mineral outgrowth, zircon was soon listed as a rare authigenic phase in sedimentary rocks (Pettijohn, 1949; Twenhofel, 1950; Packam and Crook, 1960). After the initial paper by Butterfield, a steady number of reports appeared documenting outgrowths on detrital zircon (Smithson, 1937, 1940; Tyler et al., 1940; Bond, 1948; Hutton, 1950; Kilpady and Deshpande, 1955; Awasthi, 1961; Milner, 1962; Rajulu and Nagaraja, 1966). Again, the mineral was not conclusively identified, although in one study, refractive

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Fig. 6. Drawings of large pyramidal outgrowths on zircon grains as originally observed by Butterfield in Palaeozoic sandstones from the Pennine region. Length of zircon grain 1 is 0.35 mm; grain 2 is 0.24 mm; grains 3 – 9 and 11 are f 0.15 mm; grain 10 is 0.19 mm. From Butterfield (1936) ‘‘Outgrowths on zircon’’, Geological Magazine LXXIII, XI, p. 512; republished with permission from Cambridge University Press.

index determinations on a number of secondary outgrowths indicated that the mineral was zircon, leading the author to conclude that ‘‘Brammall’s suggestion that the secondary material might be some rare earth mineral isomorphous with zircon is not substantiated’’ (Hutton, 1950). Although there is no reason to doubt the mineral identification (see also Hutton, 1947), the outgrowth featured in the paper (see Hutton, 1950, plate 1, fig. 14), is probably not diagenetic, but rather a detrital, composite zircon grain. In a study of accessory minerals in granite and other igneous rocks, the occurrence of zircon-xenotime intergrowths was thought to be significant (Hoppe, 1951), leading some to caution that ‘‘confir-

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mation is needed that the outgrowths are in all instances zircon and not xenotime’’ (Poldervaart, 1955). Nevertheless, in a study of Neoproterozoic sedimentary rocks in Scandinavia, Saxena (1966a,b, 1968), suggested that not only were the syntaxial outgrowths composed of zircon, but that entire zircon grains formed during authigenic processes within sediments. While the authigenic origin of rounded zircon grains was challenged (Marshall, 1967, 1968; Kalsbeek, 1967), it was conceded that outgrowths and overgrowths might be diagenetic zircon. The debate highlights the difficulty of identifying minute crystals by optical microscopy. With the advent of the SEM and energy dispersive spectrometer (EDS), it has become a relatively straightforward process to identify minerals as small as 1 –2 Am and determine their elemental composition (Welton, 1984; Krinsley et al., 1998). In a recent study investigating the diagenetic history of a Permian sandstone from Western Australia, minute ( < 5 Am) irregular to pyramidal outgrowths were identified on detrital zircon grains (Rasmussen and Glover, 1994). Analysis of the outgrowths by EDS showed that the mineral was an yttrium phosphate, containing minor heavy REE (mainly Gd, Dy, Er and Yb), identifying the mineral as xenotime. Subsequent work on another Palaeozoic sandstone (Rasmussen and Glover, 1996), and later, Precambrian sedimentary rocks (Rasmussen, 1996; Rasmussen et al., 1998), indicated that xenotime might be a relatively common, albeit volumetrically minor, constituent of sandstones. Among the many samples examined were sandstones from the same succession as studied by Butterfield (1936); the rocks contain a diverse range of detrital zircon grains, many of which have pyramidal outgrowths (Fig. 7), identified by EDS as xenotime (Fig. 8). Since 1994, and after thousands of analyses of outgrowths in hundreds of samples, the outgrowth mineral is almost without exception xenotime. In some cases, minute, irregular outgrowths of zircon have been observed in low-grade metamorphic rocks (Fig. 9) (see also Dempster et al., 2004). 4.3. Distribution and habit Diagenetic xenotime overgrowths occur mainly in medium- to coarse-grained sandstones, but are also present in conglomerate (England et al., 2001a,b),

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Fig. 7. Detrital zircon grains with minute, pyramidal outgrowths. The zircon grains occur in feldspathic sandstones collected from the same region as material examined by Butterfield. Plane polarised light from a mount of heavy mineral separates. Carboniferous Lower Coal Measures, United Kingdom.

siltstone and shale (Milodowski and Zalasiewicz, 1991; Rasmussen, 1996; Rasmussen et al., 1998). In one sample, exceptionally large overgrowths were found in a sandy phosphorite (Vallini et al., 2002). Xenotime is mostly associated with sedimentary rocks deposited in deltaic or coastal environments (Butterfield, 1936; Smithson, 1937; Rasmussen, 1996), but is also known in sandstone deposited in fresh water environments (Bond, 1948; Rasmussen, 1996). To date, xenotime has not been found in carbonatedominated sequences or aeolian deposits. A major control on the likely presence of diagenetic xenotime in a rock is zircon, although sedimentary rocks with high concentrations of zircon (>200 –300 grains per polished thin section) rarely contain large diagenetic xenotime crystals. Where visible in transmitted light, xenotime crystals are clear, colourless, and in optical continuity with

their detrital zircon substrate. They form on many different zircon grain types and show no obvious preference for composition or age of the substrate. In sandstones, diagenetic xenotime is mostly present as minute pyramidal overgrowths that line the surfaces of detrital zircon grains (Fig. 10a and b), or partially infill voids and cavities (Fig. 10c and d). Most xenotime overgrowths are less than a few microns in size, but in some samples, larger single crystals (up to 20 Am) are present (Fig. 10e). The xenotime usually occurs as isolated crystals attached to the rounded zircon substrate (outgrowths), but in exceptional circumstances, the xenotime may completely surround the detrital core (overgrowths) (Fig. 10f). In places, the development of euhedral xenotime crystals has been hindered by the proximity of adjacent detrital particles leading to the formation of irregularly shaped outgrowths (Fig. 11a– c). In some samples, early diagenetic clay coatings or collophane crusts have prevented xenotime from nucleating on zircon substrates (Fig. 11d –f). In contrast to Phanerozoic examples, xenotime outgrowths in Proterozoic and Archaean sedimentary rocks are typically irregular and discontinuous, although distinct pyramidal crystals can be recognised in some samples (Fig. 12a). Many outgrowths comprise several compositionally distinct zones (Fig. 12b – f). In one locality, xenotime overgrowths occur as pore-filling cement, up to 300 Am in size (Vallini et al., 2002) (Fig. 12c). 4.4. Relative timing of xenotime precipitation The pyramidal shape of many xenotime crystals (Figs. 7 and 8), as well as the relative softness of xenotime (4– 5 on Mohs’ hardness scale in contrast to 7.5 for zircon) (Palache et al., 1951), supports a postdepositional origin, as the euhedral overgrowths are unlikely to have survived mechanical abrasion during transportation. In addition, the bond between the detrital zircon substrate and syntaxial xenotime overgrowth is not particularly strong, as demonstrated during polished thin-section preparation, where zircon grains may be plucked out of the section, leaving behind a cavity surrounded by a rim of diagenetic xenotime (Fig. 13a). During heavy mineral separation and SHRIMP mount preparation, xenotime overgrowths may be dislodged from their detrital zircon

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Fig. 8. (a) BSEM image of a detrital zircon grain with two pyramidal xenotime outgrowths. Carboniferous Lower Coal Measures, United Kingdom. (b) Close-up of the xenotime pyramids showing the location of two EDS analysis spots. (c) EDS plot showing a combined phosphorus (K-alpha) and yttrium (L-alpha) peak, and several smaller peaks coinciding with the heavy REE. (d) EDS plot with two main peaks coinciding with the silicon (K-alpha) peak and zirconium (L-alpha) peak.

host and ‘‘float’’ in the epoxy resin (Fig. 13b). The boundary may also act as a line of weakness during deformation (Fig. 13c), possibly due to metamictisation of the zircon surface by the adjacent xenotime. A post-depositional origin for xenotime outgrowths is supported by textures showing xenotime crystals partly surrounding and engulfing detrital and diagenetic minerals (Fig. 14a – d). Establishing the timing of xenotime growth relative to other diagenetic minerals can be difficult because xenotime is scarce and crystals are small. In a few cases, petrographic textures show xenotime overgrowths engulfed by syntaxial quartz (Fig. 14a), K-feldspar cement, opaline silica and pore-filling carbonate cement (Rasmussen and Glover, 1994, 1996), suggesting that xenotime is one of the earliest diagenetic minerals to form. In one sample, a framboidal pyrite crystal

appears to have grown around a pyramidal xenotime overgrowth (Fig. 14b), suggesting that xenotime growth preceded pyrite formation. Based on petrographic observations of mostly Palaeozoic and Mesozoic sandstones, xenotime appears to be almost exclusively a diagenetic precipitate (Rasmussen and Glover, 1994, 1996; Rasmussen, 1996). However, in many Precambrian basins, xenotime overgrowths comprise several compositionally distinct zones, some of which enclose diagenetic and metamorphic minerals, indicating a post-diagenetic origin for at least some of the xenotime. 4.5. Conditions of formation Diagenetic phosphate minerals are concentrated in regions of oceanic upwelling where bottom

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Fig. 9. (a) Elongate detrital zircon grain in a slate lined with irregular zircon outgrowths (arrows). Surrounding minerals are quartz and sericite. Proterozoic Cardup Group, Western Australia. (b) Sub-rounded to sub-angular zircon grain showing truncation of internal zoning with small, irregular zircon outgrowths (arrows). The sample is a shale composed primarily of quartz and sericite/illite. Hamersley Group, Pilbara Craton, Western Australia. (c) Rounded zircon grain in a quartzite with irregular zircon outgrowths that extend outwards along grain boundaries (arrows). Palaeoproterozoic Mount Barren Group, Western Australia.

currents supply nutrients leading to high biological productivity in the photic zone. Shortly after the organic debris settles on the seafloor it is decomposed releasing dissolved phosphate into the sediment pore-waters (Burnett, 1977; Froelich et al., 1988; McArthur et al., 1988; Glenn, 1990; Fo¨llmi, 1996). Seafloor phosphate concentrations may also be enhanced by the reduction of Fe-oxyhydroxides that scavenge dissolved phosphate from the water column. The combination of physical and biological processes cause a build-up of dissolved phosphorus, leading to the precipitation of diagenetic phosphate minerals (primarily francolite or carbonate fluorapatite) at or near the sediment–water interface. Marine phosphatization is geologically instantaneous and is responsible for some truly spectacular fossil preservation (Bengtson and Yue, 1997; Xiao et al., 1998). Many phosphorite deposits contain elevated concentrations of Y and REE (Goldberg et al., 1963; Schofield and Haskin, 1964; Altschuler, 1980). A recent geochemical study of a Proterozoic phosphorite in Western Australia revealed yttrium contents of 400– 500 ppm over a 10-m interval (Vallini, 2000; Vallini et al., 2002). Detailed petrography and SEM

imaging of the interval revealed that the main mineralogical host for Y and heavy REE is xenotime, which is present as unusually large overgrowths on detrital zircon grains (Vallini et al., 2002) (Fig. 14c). The main source of the dissolved phosphate is, as with carbonate fluorapatite, the organic matter that undergoes decomposition shortly after deposition. The source of the Y and REE is less certain, but may have been derived in part from seawater, decomposing organic complexes, clay particles and reduced Feoxyhydroxies in the matrix and interbedded shale laminae. Pore-water data and sequential leaching techniques suggest that phosphates also form at the sediment – water interface on continental shelves away from the influence of upwelling ocean currents (Berner, 1990; Ruttenberg and Berner, 1993). However, the concentration of authigenic phosphates is generally much lower because of dilution due to higher rates of sediment accumulation, and moderate biological productivity. The pore-water from which diagenetic xenotime crystals precipitated must have been locally enriched in HREE, Y and dissolved phosphate. Although the concentration of these chemical species is very low in sea water (Turekian, 1969; Krauskopf,

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Fig. 10. BSEM images showing association between detrital zircon and diagenetic xenotime. (a) An elongate zircon grain lined by numerous minute ( < 2 Am) xenotime outgrowths (arrows). Carboniferous Millstone Grit, United Kingdom. (b) Close-up of zircon grain showing a line of xenotime outgrowths (white) along margin. Several larger outgrowths have engulfed the clay matrix adjacent to the zircon. Carboniferous Millstone Grit, United Kingdom. (c, d) Zircon crystals with masses of xenotime crystals partly infilling vacuoles formerly occupied by glass. Proterozoic Capricorn Formation, Western Australia. (e) A large, pyramidal xenotime outgrowth on a rounded zircon grain. Carboniferous Millstone Grit, United Kingdom. (f) A detrital zircon grain entirely surrounded by xenotime, forming a complete overgrowth. Palaeoproterozoic Mount Barren Group, Western Australia.

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Fig. 11. (a) BSEM image of a fractured zircon grain surrounded by xenotime overgrowth, that has grown around the adjacent detrital apatite clast (ap). Palaeoproterozoic Mount Barren Group, Western Australia. (b) BSEM image of an euhedral zircon with a diagenetic xenotime outgrowth that has partially engulfed kaolinite booklets in the adjacent pore. Carboniferous Millstone Grit, United Kingdom. (c) BSEM image of diagenetic xenotime that has partially engulfed detrital clay matrix and kaolinite booklets. Carboniferous Millstone Grit, United Kingdom. (d) BSEM image of a detrital zircon from a heavy mineral separate lined by a thin coating of authigenic clay. Carboniferous Millstone Grit, United Kingdom. (e) BSEM image of two detrital zircon grains; one surrounded by authigenic xenotime and the other with minute irregular xenotime outgrowths. (f) Close-up of zircon in (e) showing the presence of a coating of microcrystalline apatite that appears to have inhibited xenotime growth. Palaeoproterozoic Mount Barren Group, Western Australia.

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Fig. 12. BSEM images of zoned diagenetic xenotime from Precambrian sedimentary rocks. (a) Xenotime outgrowth on zircon comprising a broad pyramid, surrounded by a possible metamorphic rim. Palaeoproterozoic Stirling Range Formation, Western Australia. (b) Close-up of part of a large xenotime overgrowth comprising an inner, inclusion-rich zone (xt1) surrounded by a massive, compositionally zoned outer rim (xt2) with euhedral inclusions of oxidised pyrite (black, top right). Palaeoproterozoic Pentecost Sandstone, Kimberley Group, Western Australia. (c) A cluster of detrital zircon grains with complete xenotime overgrowths displaying complex compositional zonation. Ion microprobe dating indicates that the xenotime precipitated during at least three temporally distinct periods. Palaeoproterozoic Mount Barren Group, Western Australia. (d) Euhedral zircon with an irregular xenotime outgrowth, partly intergrown with thorite (th). Archaean Fortescue Group, Western Australia. (e) A small zircon core surrounded by two compositionally distinct zones of xenotime. Palaeoproterozoic Maraloou Formation, Western Australia. (f) Rounded zircon core surrounded by authigenic xenotime comprising three compositionally distinct growth zones, as well as fine crystals of thorite (th). Palaeoproterozoic Maraloou Formation, Western Australia.

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Fig. 13. (a) BSEM image of a void (black) surrounded by xenotime overgrowth produced after the detrital zircon core was plucked out of the polished thin section during polishing. Palaeoproterozoic Mount Barren Group, Western Australia. (b) BSEM image of a line of pyramidal xenotime crystals (white) surrounded by detrital zircon grains. The xenotime has been displaced from the zircon substrate during mount preparation. Carboniferous Millstone Grit, United Kingdom. (c) BSEM image of an elongate detrital zircon grain partly surrounded by displaced xenotime outgrowths. The gap between the zircon and base of xenotime is filled by quartz. Palaeoproterozoic Mount Barren Group, Western Australia.

1977), their concentration in coastal sediments may be significantly increased during suboxic– anoxic reduction of Fe- and Mn-(hydr-)oxides (Elderfield and Sholkovitz, 1987; Sholkovitz et al., 1989). The REE concentration may also be raised by the release of REE adsorbed onto clay mineral surfaces (Balashov and Girin, 1969; Roaldset and Rosenqvist, 1971; Fleet, 1984), and from the partial dissolution and surface reactions of REE-bearing minerals. Similarly, the concentration of dissolved phosphate may be raised in sediment pore waters as phosphate is released from organic matter during microbial decomposition (Krom and Berner, 1981; Berner et al., 1993; Fo¨llmi, 1996). Organic matter is a major sink for reactive phosphate in the ocean and its remineralisation within sediments exerts a major control on porewater phosphate concentrations. Other important processes include the desorption of phosphate from the surfaces of Fe- and Mn-(hydr-)oxide particles during suboxic to anoxic reduction (Berner, 1973; Froelich et al., 1977; Krom and Berner, 1981; Lucotte et al., 1994), and the dissolution of phosphatic skeletal debris (Suess, 1981).

It is from such REE- and phosphate-enriched sediment pore-waters that xenotime probably precipitated. Petrographic textures showing xenotime crystals engulfed by siderite rhombs and partly embedded in framboidal pyrite (Rasmussen and Glover, 1994, 1996) (Fig. 14b) support an early diagenetic origin for xenotime, with growth probably starting within the zone of bacterial sulphate reduction. The close association between diagenetic xenotime and detrital zircon may be explained by the two minerals sharing the same crystal structure. Consequently, detrital zircon grains acted as ideal nucleation sites, where porewater REE and phosphate combined to precipitate as xenotime crystals.

5. Geochemistry Despite the growing importance of xenotime, only a limited amount of geochemical data is available, in contrast to other U-bearing accessory minerals such as zircon and monazite. The composition of xenotime has been determined from granite (Jefford, 1962; Suzuki et

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Fig. 14. (a) BSEM image showing tiny pyramidal xenotime outgrowths (white) on zircon engulfed by diagenetic quartz cement (grey). Carboniferous Lower Coal Measures, United Kingdom. (b) BSEM image of a tiny zircon crystal surrounded by two pyramidal overgrowths that are partly engulfed by an oxidised pyrite framboid. Carboniferous Lower Coal Measures, United Kingdom. (c) BSEM image of xenotime overgrowth containing authigenic pyrite crystals. Palaeoproterozoic Mount Barren Group, Western Australia. (d) Pore-filling xenotime cement that has partly surrounded diagenetic aluminophosphate crystals (ALPO), kaolinite booklets and microcrystalline quartz. Palaeoproterozoic Warton Sandstone, Kimberley Group, Western Australia.

al., 1992; Wark and Miller, 1993; Casillas et al., 1995; ˚ mli, 1975; Demartin et al., Fo¨rster, 1998), pegmatite (A 1991; Petersen and Gault, 1993), carbonatite (Wall and Mariano, 1996), metamorphic rocks (Suzuki and Adachi, 1991; Franz et al., 1996; Pan, 1997), hydrothermal deposits (Demartin et al., 1991; Kerrich and King, 1993), sedimentary rocks (Kositcin et al., 2003) and heavy mineral deposits (van Emden et al., 1997). Xenotime comprises mostly Y2O3 (26.8 – 53.5 wt.%), P2O5 (32.0 –37.4 wt.%) and a range of the heavier REE (15 – 45 wt.%); Nd2O3 ( < 0.02 – 1.5 wt.%), Sm2O3 ( < 0.02 – 2.0 wt.%), Gd2O3 (0.5 – 9.71 wt.%), Tb2O3 (0.3 – 1.51 wt.%), Dy2O3 (2.1 – 8.6

wt.%), Ho2O3 (0.5 – 2.0 wt.%), Er2O3 (1.67 – 7.4 wt.%), Yb2O3 (0.7 – 15.1 wt.%), Lu2O3 (0.1 – 2.3 wt.%). The high concentration of heavier REE in xenotime can be explained by the similar ionic radius ˚ for Gd3 + to of the lanthanide RE3 + ion series (1.05 A 3+ ˚ 0.98 A for Lu ; octahedral coordination) and the ˚ ) (Shannon, 1976; Burt, 1989; Ni et Y3 + ion (1.02 A al., 1995). The actinide elements, U and Th, range in abundance from < 0.01 to 6.7 wt.% and < 0.01 to 6.0 wt.%, respectively. In contrast to monazite, xenotime typically has a U/Th ratio >1, reflecting the preferred ˚ ) over substitution of Y3 + by the smaller U4 + (1.05 A

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Table 1 Electron microprobe analyses of igneous, metamorphic, hydrothermal and diagenetic xenotime 1 Oxide (wt.%) SiO2 0.23 P2O5 35.38 CaO nd Y2O3 46.13 ThO2 nd UO2 nd nd La2O3 Ce2O3 nd Pr2O3 nd 0.22 Nd2O3 Sm2O3 0.49 Eu2O3 0.13 Gd2O3 1.91 0.42 Tb2O3 Dy2O3 3.84 Ho2O3 1.01 Er2O3 3.78 0.68 Tm2O3 Yb2O3 4.25 Lu2O3 0.86 Total 99.33

2

3

4

5

6

7

8

9

10

0.95 34.48 0.04 40.99 1.30 1.76 0.14 0.10 0.08 0.36 0.68 0.07 2.22 0.58 5.05 1.07 3.99 0.70 3.89 0.52 100.49

2.64 30.45 0.54 38.34 1.27 nd 0.30 bd bd 0.34 1.34 1.32 6.11 1.93 8.58 1.06 2.03 nd 0.65 nd 99.66

0.42 33.97 0.02 45.35 0.01 0.17 0.03 0.00 0.04 0.04 0.00 nd 3.37 0.52 5.36 1.28 4.45 nd 3.57 0.50 99.10

0.42 33.84 0.02 43.43 0.20 0.41 0.00 0.03 0.00 0.15 0.21 nd 4.75 0.72 5.13 1.21 4.65 nd 4.03 0.87 100.05

0.09 34.67 0.08 43.31 0.02 0.35 0.00 0.09 0.01 0.46 0.34 nd 3.92 0.57 4.75 1.31 4.79 nd 4.17 0.79 99.72

0.05 35.09 0.07 39.53 0.03 0.08 0.00 0.00 0.00 0.13 1.36 0.82 6.96 1.29 8.45 1.39 3.05 0.27 1.75 0.00 100.33

0.14 35.90 0.07 42.60 0.16 0.66 0.00 0.00 0.01 0.06 0.36 0.21 3.49 1.12 8.40 1.61 3.34 0.41 2.62 0.15 101.35

0.40 33.80 0.58 42.61 0.79 0.28 nd 0.05 nd 0.14 0.37 0.56 3.00 0.82 6.77 1.22 3.87 0.62 3.18 0.35 99.52

2.15 32.49 0.17 46.11 0.22 0.65 nd 0.03 nd 0.15 0.28 0.22 1.40 0.51 5.43 1.22 4.43 0.51 3.96 0.48 100.46

0.896 0.092 0.020 0.709 0.010

0.973 0.014 0.001 0.816 0.000 0.001 0.000 0.000 0.001 0.001 0.000 0.016 0.038 0.006 0.058 0.014 0.047

0.977 0.014 0.000 0.784 0.001 0.003 0.000 0.000 0.000 0.002 0.003

0.991 0.003 0.003 0.777 0.000 0.002 0.000 0.001 0.000 0.005 0.004

0.974 0.014 0.021 0.772 0.006 0.002

0.922 0.072 0.006 0.823 0.001 0.004

0.001

0.000

0.053 0.008 0.056 0.005 0.050

0.044 0.006 0.052 0.014 0.051

0.037 0.009 2.016

0.042 0.008 2.006

0.043 0.008 2.004

0.002 0.004 0.007 0.034 0.009 0.074 0.013 0.041 0.007 0.033 0.004 2.018

0.002 0.003 0.003 0.016 0.006 0.059 0.013 0.047 0.005 0.040 0.005 2.026

Formula calculated to 4(O) P 0.992 0.977 Si 0.008 0.032 Ca 0.002 Y 0.813 0.730 Th 0.012 U 0.006 La 0.002 Ce 0.001 Pr 0.001 Nd 0.003 0.004 Sm 0.006 0.008 Eu 0.002 Gd 0.021 0.025 Tb 0.005 0.054 Dy 0.041 0.011 Ho 0.011 0.042 Er 0.039 0.007 Tm 0.007 Yb 0.043 0.005 Lu 0.009 0.006 Total 2.000 1.990

0.004

0.004 0.016 0.001 0.070 0.022 0.096 0.012 0.022 0.040 0.007 0.021 2.017

˚ ). The amount of U and Th is the larger Th4 + (1.09 A largely controlled by two charge-balanced coupled substitution mechanisms with Si and Ca (van Emden et al., 1997) (Eqs. (1) and (2)): ðY; REEÞ3þ þ P5þ ¼ ðTh; UÞ4þ þ Si4þ

ð1Þ

1.003 0.002 0.003 0.710 0.000 0.001 0.000 0.000 0.000 0.002 0.016 0.010 0.078 0.014 0.092 0.015 0.032 0.003 0.018 0.000 1.998

1.004 0.005 0.002 0.749 0.001 0.005 0.000 0.000 0.000 0.001 0.004 0.002 0.038 0.012 0.089 0.017 0.035 0.004 0.026 0.001 1.996

and 2ðY; REEÞ3þ ¼ ðTh; UÞ4þ þ Ca2þ :

ð2Þ

However, the presence of Ca and Si abundances greater than those predicted through the incorporation

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of Th and U suggest that additional mechanisms may be involved (van Emden et al., 1997; Forbes, 1999). Diagenetic xenotime generally has lower U and Th contents than igneous varieties (Table 1) (Forbes, 1999; Kositcin et al., 2003). In the Witwatersrand Basin, South Africa, diagenetic xenotime can be differentiated from hydrothermal varieties by its higher U and Th contents, and from igneous (detrital) xenotime by its higher Eu, Dy and Gd concentrations and lower Gd/Yb ratio (Kositcin et al., 2003). The chondrite-normalised REE pattern for diagenetic xenotime is very similar across different sedimentary units, displaying smooth broad peaks from Tb to Yb (Fig. 15a –c). The patterns differ from igneous xenotime by the absence of an Eu anomaly (Fig. 15d), although diagenetic xenotime from the Witwatersrand Basin displays a small negative anomaly (Fig. 15c). Metamorphic and hydrothermal xenotime crystals (Fig. 15e and f) display very similar REE patterns to diagenetic xenotime, and cannot be distinguished solely on the basis of REE chemistry. In basins with complex post-depositional thermal histories and multiple stages of xenotime growth, the application of in situ microanalysis may help identify and characterise different growth zones. Determination of the origin of xenotime crystals via thorough petrographic studies prior to dating by in situ methods is critical to avoid obtaining mixed data from adjacent, temporally distinct zones.

6. U – Pb geochronology Xenotime is an ideal U –Pb geochronometer because of its high levels of uranium, low initial lead

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concentration, and ability to remain isotopically closed after formation (Heaman and Parrish, 1991; McNaughton et al., 1999; Fletcher et al., 2000). Xenotime is also highly resistant to metamictisation despite U concentrations (>10,000s ppm) that would render other minerals (e.g., zircon) useless for geochronology. Despite the exceptional qualities of xenotime, it has only been used in a small number of studies; primarily to date high-grade metamorphism ((Ko¨ppel and Gru¨nenfelder, 1975; Aleinikoff and Grauch, 1990; Zhu et al., 1997; Kamber et al., 1998; Hawkins and Bowring, 1999; Crowley and Parrish, 1999; Simpson et al., 2000; Viskupic and Hodges, 2001), but also crystallisation (Scha¨rer, 1984; Heaman and Parrish, 1991; Hodges et al., 1992; Hawkins and Bowring, 1997) and hydrothermal processes (Compston and Matthai, 1994; Scha¨rer et al., 1999; Petersson et al., 2001). In most of these studies, xenotime yielded concordant and precise U –Pb dates. The potential use of xenotime to date sedimentary rocks stems from petrographical studies of diagenetic rare-earth phases in sandstones (Rasmussen and Glover, 1994; 1996). Xenotime outgrowths were initially identified in a Permian sandstone in the Carnarvon Basin, Western Australia, but as more successions were examined, it soon became apparent that xenotime was a widespread diagenetic trace phase that could potentially be used to constrain the age of sediment deposition. At about the same time, a sensitive, high-resolution ion microprobe (SHRIMP) was brought to Perth, and soon became available to researchers. Early work focused on zircon geochronology using a standard spot size of ca. 20 –25 Am. Because most xenotime overgrowths were < 10 Am in size, diagenetic xen-

Notes to Table 1: ˚ mli, 1975). 1. Xenotime associated with muscovite, euxenite and calcite, Gloserheia granite pegmatite, Froland, southern Norway (A 2. Xenotime from an Archaean pegmatite, Yilgarn Craton, Western Australia (Kositcin et al., 2003). 3. Xenotime overgrowth on niobian rutile in quartz-apatite rock, Kangankunde carbonatite complex, southern Malawi (Wall and Mariano, 1996) (includes Sc2O3 = 0.75, TiO2 = 0.43, FeO(T) = 1.58). 4 – 6. Xenotime from greenschist facies (chlorite) metapelite (2), amphibolite facies (andalusite – garnet – sillimanite – staurolite) metapelite (3) and upper amphibolite/granulite facies (cordierite – garnet – biotite – sillimanite – K-feldspar) metapelite (4), Variscan mountain belt, NE Bavaria, Germany (Franz et al., 1996). 7. Hydrothermal xenotime, Parktown Formation, Witwatersrand Basin, South Africa (Kositcin et al., 2003). 8. Diagenetic xenotime in quartzite from Krugersdorp Formation, Witwatersrand Basin, South Africa (Kositcin et al., 2003) (includes PbO2 = 0.06). 9. Diagenetic xenotime overgrowth in quartz sandstone, Warton Sandstone, Kimberley Group, Western Australia (Forbes, 1999) (includes FeO(T) = 0.03, PbO = 0.10). 10. Diagenetic xenotime outgrowth on zircon in Palaeozoic sandstone (Forbes, 1999) (includes PbO = 0.04).

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Fig. 15. Chondrite-normalised plots of diagenetic, igneous, metamorphic and hydrothermal xenotime. (a) Diagenetic xenotime in feldspathic sandstone from the Carboniferous Millstone Grit (data from Forbes, 1999). (b) Diagenetic xenotime outgrowths in quartz sandstone from the Warton and Pentecost Sandstones, Kimberley Group (data from Forbes, 1999). (c) Diagenetic xenotime from the Archaean Witwatersrand Basin, South Africa (data from Kositcin et al., 2003). (d) Xenotime from an Archaean pegmatite in the Yilgarn Craton, Western Australia; a granite pegmatite in the Grenville Province, Canada (XENO1 of Stern and Rainbird, 2001); a granite pegmatite from Aust-Agder province, Norway (XENO2 of Stern and Rainbird, 2001). Electron microprobe data from Kositcin et al. (2003). (e) Metamorphic xenotime from the greenschist and amphibolite facies metapelites, Germany (data from Franz et al., 1996). (f) Hydrothermal xenotime in quartzite and conglomerates from the Archaean Witwatersrand Basin, South Africa (data from Kositcin et al., 2003). Chondrite composition after Nakamura (1974); Pr, Tb, Ho and Tm values are from Evensen et al. (1978).

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otime was initially considered too small for SHRIMP geochronology. However, the examination of additional samples led to the discovery of larger xenotime overgrowths (Rasmussen, 1996; Rasmussen et al., 1998), and increasing experience with SHRIMP methodologies led to the first successful attempt to date ca. 10 Am xenotime overgrowths in the late 1990s (McNaughton et al., 1999). Since then, a growing number of sedimentary successions have been dated using a combination of diagenetic xenotime and detrital zircon geochronology (England et al., 2001a,b; Vallini et al., 2002; Rainbird et al., 2003; Rasmussen et al., 2004). Following the initial work on diagenetic xenotime, the mineral has now become a chronometer of contact metamorphism (Rasmussen et al., 2001), low-grade regional metamorphism (Dawson et al., 2003) and hydrothermal mineralisation (England et al., 2001a,b; Brown et al., 2002; Rasmussen et al., 2001; Pigois et al., 2003; Tallarico et al., 2004). 6.1. Analytical dating techniques A variety of techniques are currently available for U – Th –Pb geochronology of common accessory minerals such as zircon, monazite and xenotime (Ireland, 1999; Poitrasson et al., 2002; Harrison et al., 2002; Parrish and Noble, 2003; Ireland and Williams, 2003) (Table 2). The two main techniques used are SIMS and TIMS. The techniques are quite different, with

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SIMS allowing in situ analysis of very small volumes of material, whereas TIMS involves the chemical digestion and analysis of entire crystals. Isotope dilution TIMS is more precise than SIMS, mainly because of the larger volume of material that is analysed. However, if minerals are heterogenous, comprising differently aged growth zones or areas of recrystallisation, then analysis by ion microprobe (using a 10 –25-Am spot) is more appropriate. Early attempts to separate diagenetic xenotime from zircon by acid dissolution were unsuccessful, as both minerals appear to be equally insoluble, and thus it has not yet been possible to date diagenetic xenotime by TIMS. Instead, the minute size of xenotime crystals and their growth during multiple events is more conducive to dating by in situ microbeam techniques. At the University of Western Australia, analysis by ion microprobe is favoured over other dating techniques because of the fine spatial resolution ( f 10 Am), shallow penetration depth ( f 1 – 2 Am; allowing repeat measurements on the same spot) and high precision ( F 1 – 2% at 2j; under optimum conditions). The fine spatial resolution offered by SIMS is particularly important because xenotime crystals larger than 10 Am in size are rare (Fig. 16a). A small spot size is also essential for avoiding overlap onto the detrital zircon host (Fig. 16b), temporally distinct growth zones (Fig. 16c), fine matrix inclusions, fractures, and specks of Th- or U- bearing mineral inclusions (e.g., thorite).

Table 2 Techniques for U – Th – Pb dating of xenotime (after Poitrasson et al., 2002) Technique

EMPA

SIMS

LA – ICP – MS

ID – TIMS

Name

Electron microprobe analysis (aka CHIME) Chemical No No F 30%; not very suitable for xenotimea 1 – 5 Am Circular < 5 Am Yes 5 – 10 min Mineral

Secondary ion mass spectrometry (e.g., SHRIMP) Isotopic Yes Yes F 1 – 2%b

Laser ablation – inductively coupled plasma – mass spectrometry Isotopic Possible Yes F 1 – 10%

Isotope dilution – thermal ionisation mass spectrometry Isotopic Yes Yes F 0.3%

5 – 25 Am Elliptical 1 – 5 Am Yesc 15 – 30 min Mineral

5 – 400 Am Variable 10 – 20 Am No < 2 – 5 min Mineral, glass or tracer solution

Entire grain or fragment na na No Days Tracer solution

Chemical or isotopic 204 Pb correction 207 Pb/206Pb dates Average 2j precision for Palaeozoic ages Spatial resolution Spot shape Penetration depth Repeat analyses Analysis time Standard type a

The precision achievable by EMPA dating of xenotime is significantly reduced by the overlap of the primary Pb peak onto the Y peak. Under optimum conditions is analogous to zircon (see Fletcher et al., in press). c It is possible to perform multiple analyses in the same spot but there are limitations (see Fletcher et al., 2000). b

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Fig. 16. (a) BSEM image of a large xenotime outgrowth with ca. 8 Am SHRIMP analytical pit. Carboniferous Millstone Grit, United Kingdom. (b) Rounded, elongate zircon with patchy xenotime outgrowths surrounded by quartz and metamorphic sericite. Arrow points to a SHRIMP pit that partially overlaps the zircon host. Palaeoproterozoic Stirling Range Formation, Western Australia. (c) A large xenotime overgrowth showing the location of numerous SHRIMP analytical pits, and the 207Pb/206Pb age obtained from each pit. Palaeoproterozoic Mount Barren Group, Western Australia (photo taken by Daniela Vallini).

Currently, the best spatial resolution (ca. 1 Am) is achieved by chemical Th – U –total Pb analysis using an electron microprobe. The technique has been used to date monazite (Suzuki and Adachi, 1991, 1994; Montel et al., 1996; Rhede et al., 1996; Cocherie et al., 1998; Williams et al., 1999), and to a lesser extent zircon (Geisler and Schleicher, 2000) and xenotime (Suzuki and Adachi, 1991, 1994; Asami et al., 2002; Grew et al., 2002). The greater spatial resolution (1 Am) offered by electron microprobe permits the study of fine growth zones within single crystals and could potentially double or triple the number of xenotime crystals that could be dated. However, the technique suffers from problems of low precision due to the relatively poor detection limit for Pb and the inability to assess concordance of the parallel U – Pb decay systems (Poitrasson et al., 2002; Catlos et al., 2002). A major limitation specific to xenotime is the overlap of the primary Pb (Ma) X-ray peak onto the Y (Lg) X-ray peak, forcing Pb counts to be measured from lower intensity Pb peaks, significantly reducing the precision of dates (Forbes, 1999). Another technique that holds great promise for U – Pb geochronology is laser ablation –inductively coupled plasma – mass spectrometry (LA – ICP – MS) (Parrish et al., 1999; Li et al., 2001; Machado and Simonetti, 2001; Kosler and Sylverster, 2003). Under ideal conditions and using a large spot size (between 25 and 125 Am), the technique can yield U – Pb dates that approach the accuracy and precision delivered by SIMS in a fraction of the time (Horn et al., 2000). However, LA – ICP –MS requires a much larger volume of material than SIMS because of its lower efficiency at detecting ions per total atoms released (0.04% compared to z 1.5% for SHRIMP II; Compston, 1996), leading to a reduction of the spatial resolution or ability for high-resolution depth profiling (Horn et al., 2000). Another drawback is isotopic and element bias during ablation. However, analytical strategies, including the use of fast, short wavelength lasers with multicollector arrays, should help to reduce the effect of chemical fractionation (Machado and Simonetti, 2001; Poitrasson et al., 2002). The main application of the technique has been to date detrital zircon grains (Machado et al., 1996; Scott and Gauthier, 1996; Fernandez-Suarez et al., 1999; Reiners et al., 2002). In such studies, the technique is highly suitable because high-precision dates are

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generally not critical, and because analysis times are very fast (1 –2 min), a large number of grains can be dated. Future technical advancements will no doubt lead to improvements in the precision and accuracy of the technique such that it may soon rival SIMS as the main source of U – Th – Pb age data (Kosler and Sylverster, 2003). An exciting new development is the potential use of the Cameca NanoSIMS 50 for U –Th – Pb geochronology of accessory minerals as small as 1 – 2 Am in size. Given the complexity of many xenotime overgrowths, the capability to date individual growth zones could be of immense value. Nevertheless, the ability to date ever-smaller volumes of material comes at the expense of diminishing precision. Also, U –Th – Pb geochronology by NanoSIMS is as yet untested, although initial trials at the University of Western Australia are very promising. 6.2. Analytical procedures Diagenetic xenotime geochronology requires exhaustive amounts of SEM examination. Although diagenetic xenotime is relatively common, most overgrowths are too small for U –Pb geochronology by ion microprobe. Consequently, a large suite of samples has to be carefully scrutinised before a sufficient number of suitably sized crystals (>10 Am) can be found; generally < 1% of the population. In most instances, it may require an initial collection of up to 20– 30 hand-specimens to identify the most favourable rock-type, and then the preparation of additional polished thin sections from those specimens. In early studies, once suitable samples had been identified, zircon grains (and attached xenotime overgrowths) were extracted from hand-specimens following standard crushing, heavy-liquid and magnetic separation techniques (McNaughton et al., 1999; Fletcher et al., 2000). From the zircon separates, individual grains with pyramidal or irregular overgrowths were hand-picked and mounted in an epoxy disk and examined by SEM. With this approach, many ‘‘overgrowths’’ were found to comprise fragments of quartz attached to zircon grains, or irregular fractured and cleaved zircon grains. Also, many samples with diagenetic xenotime overgrowths in polished thin-sections (PTS) failed to yield xenotime overgrowths in the heavy mineral separates; this was a

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particular problem with tightly cemented quartz sandstones or quartzites. It seems likely that the xenotime overgrowths became detached from their zircon substrate during crushing and mineral separation. It has since become standard procedure to drill out or cut large xenotime crystals from polished thin sections and embed 10 –12 plugs or fragments, along with a zircon-xenotime standard set, into an epoxy mount. Crystal recovery varies from between 75% and 100%, and is significantly less costly and time consuming than mineral separation. To avoid lengthy delays between each analysis, the epoxy mounts are photographed in reflected light at low magnification, and individual PTS plugs are photographed at various magnifications to help locate individual crystals and to maximise spot placement by avoiding mineral inclusions, cracks and pits. The application of diagenetic xenotime geochronology to Phanerozoic sedimentary rocks requires accurate determinations of Pb/U and Pb/Th, and is hampered by the effects of variable trace element compositions (Fletcher et al., 2000). However, recent improvements in calibration procedures, including matrix corrections for U, Th and REE, have led to greater accuracy; ca. 1% for Pb/U and < 2% for Pb/Th (Fletcher et al., in press). Also, the development of new xenotime standards has reduced the chemical differences between the standard and sample, giving greater control on matrix effects (Fletcher et al., in press). The methodologies for analysing Precambrian xenotime crystals using 207Pb/206Pb dates are well established, and have been discussed in earlier publications (McNaughton et al., 1999; Fletcher et al., 2000). In the following section, the application of diagenetic xenotime geochronology to several Precambrian sedimentary sequences is discussed. 6.3. Case studies 6.3.1. Kimberley Group, northwestern Australia In the marine sandstones of the Proterozoic Warton and Pentecost Sandstones, Kimberley Basin, northwestern Australia (Griffin and Grey, 1990), diagenetic xenotime occurs as syntaxial outgrowths up to 20 Am (Fig. 12b) and as pore-filling cement unassociated with detrital zircon grains (Fig. 14d). Some of the larger outgrowths display two distinct compositional zones

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(Fig. 14b). Locally, xenotime overgrowths are surrounded by syntaxial quartz overgrowths, which appear to have precipitated before deep burial compaction. Prior to xenotime geochronology, the age of the sandstones was constrained only to within a billionyear interval. A minimum age of < 750 Ma is interpreted from unconformably overlying glacial deposits, the oldest of which has been correlated with the Sturtian glaciation (Coats and Preiss, 1980) or alternatively to the younger Marinoan glaciation (Plumb, 1996; Grey and Corkeron, 1998). A maximum age is given by mafic intrusive rocks (1790 F 4 Ma; Page and Sun, 1994), which intrude up to, but not through, the sandstones. Xenotime U –Pb data from the Warton and Pentecost Sandstones yield a spectrum of ages (Fig. 17), which is probably related to two different episodes of growth. The oldest 207Pb/206Pb age population is

1704 F 7 Ma for the Warton Sandstone and 1704 F 14 Ma for the Pentecost Sandstone (Fig. 17) (McNaughton et al., 1999). It therefore seems likely that the onset of xenotime precipitation predated 1704 Ma, which can only be considered as a minimum age for deposition. Although the xenotime ages from the Warton and Pentecost Sandstones are indistinguishable, the chemistry of the xenotime crystals is distinct (300 –1100 ppm U, Th/U ratios of 5.0– 30 for the Pentecost Sandstone; 1500– 4000 ppm U, Th/U ratios of 0.25– 0.7 for the Warton Sandstone). This suggests that the xenotime formed from two different pore fluids, probably at different times, but within analytical uncertainties in the Pb/Pb dates. 6.3.2. Witwatersrand Supergroup, South Africa The Witwatersrand Basin in South Africa is a thick (up to 7.5 km) succession of siliciclastic sedimentary

Fig. 17. Probability plots of SHRIMP U – Pb age data for xenotime outgrowths and detrital zircon grains in the Warton and Pentecost Sandstones, Kimberley Group, Western Australia.

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rocks that host major gold and uranium mineralisation associated with unconformity-bounded quartz-pebble conglomerates and quartzose sandstones (Liebenberg, 1955; Robb and Meyer, 1995). The timing of mineralisation is controversial, with debate divided between the ‘‘modified placer’’ model (Minter et al., 1993; Frimmel, 1997; England et al., 2001b; Kirk et al., 2002) and the ‘‘hydrothermal’’ model (Barnicoat et al., 1997; Phillips and Law, 2000). The age of the Witwatersand succession is poorly constrained. A maximum age of 3074 F 6 Ma is provided by SHRIMP U – Pb dating of volcanic rocks from the unconformably underlying Dominion Group (Armstrong et al., 1991). Detrital zircon U –Pb age data from horizons within the Witwatersrand Supergroup indicate a source age of between ca. 3.3 and 2.9 Ga (Barton et al., 1989; Robb et al., 1990; Poujol et al., 1999). A minimum age of 2714 F 8 Ma is derived from flood basalts in the overlying Ventersdorp Supergroup (Armstrong et al., 1991). Xenotime has been identified in several studies of mineralised strata from the Witwatersrand Basin (Ramdohr, 1958; Oberthu¨r, 1987). In a recent study, England et al. (2001b) found that xenotime was relatively common in horizons throughout the succession (about 75% of samples, 73 polished thin-sections), but only in trace quantities as minute crystals (mostly < 5 Am). Based on textural observations, several different types of xenotime were observed: (i) rounded to sub-rounded detrital grains in heavy mineral seams, (ii) outgrowths on detrital zircon grains, (iii) infill of fractured zircon and (iv) discrete, irregular crystals within the sediment matrix. Examination by back-scattered electron microscope (BSEM) reveals that much of the xenotime is characterised by complex compositional zoning, which, in many instances, is less than 10 Am wide. Also, some of the xenotime contains numerous fine specks of U- and Th-rich mineral inclusions. In the first attempt to date diagenetic xenotime from the Witwatersrand Basin, England et al. (2001a) obtained a spread of 207Pb/206Pb dates from ca. 2760 to 2040 Ma. Such a broad range (>700 million years) had not been encountered in previous studies of diagenetic xenotime (cf. McNaughton et al., 1999; Fletcher et al., 2000), and suggests that the history of xenotime growth in the Witwatersrand was multistaged and complex. The nature of the U –Pb xen-

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otime data is perhaps not surprising given that the Witwatersrand Supergroup was affected by several hydrothermal and metamorphic events: (i) syn-Central Rand compressive deformation (Barnicoat et al., 1997); (ii) flood-basalt volcanism during deposition of the Ventersdorp Supergroup at 2714 F 8 Ma (Armstrong et al., 1991); (iv) emplacement of the Bushveld Igneous Complex at 2061 F 4 Ma (Walraven, 1997); and (v) the Vredefort deformation event at 2023 F 4 Ma, interpreted as the result of a meteorite impact (Kamo et al., 1996). Concordant ages obtained from authigenic xenotime are interpreted to record diagenetic, hydrothermal and metamorphic events, although not all can be matched to known tectonothermal events. A single analysis yielded a 207Pb/206Pb age of 2764 F 5 Ma, which is the oldest unequivocal age for any authigenic mineral in the Central Rand Group, and confines sedimentation to between ca. 2764 and ca. 2886 Ma, the age of the youngest detrital zircon. A peak in xenotime ages at ca. 2717 Ma is interpreted to correlate with emplacement of the Ventersdorp flood basalts. Matrix-filling xenotime, which is associated with hydrothermal brannerite, gersdorffite and gold, yields a spread of ages from 2382 to 2039 Ma. The younger ages may relate to the emplacement of the Bushveld Igneous Complex ( f 2060 Ma) and possibly associated with the Vredefort event ( f 2020 Ma). The relatively poor definition of some age peaks partly reflects the inability to resolve the ages of complexly zoned xenotime. Most xenotime outgrowths are quite small ( < 10 Am), and comprise zones that grew during diagenesis, hydrothermal alteration and metamorphism; thus, many analyses will inevitably represent mixture ages when the maximum spatial resolution is 10 Am. Another result not observed in previous studies, is the presence of highly discordant dates, suggesting that the xenotime may have undergone Pb loss. In zircon, Pb-loss is common in grains with high U contents, leading to crystal damage and increased susceptibility to element diffusion and resetting (Mezger and Krogstad, 1997). However, the xenotime grains from the Witwatersrand Basin have U concentrations comparable to xtc (ca. 10,000 ppm), a UWA standard that yields concordant (within 1 – 2%) and precise dates (McNaughton et al., 1999; Fletcher et al., 2000). Rather than Pb-loss, the

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discordance may be due to the presence of U- and Th-rich inclusions. The most discordant dates correlate with grains containing numerous < 1 Am specks of uraninite and thorite (Kositcin et al., 2003). These minerals are prone to Pb-loss, and although they represent only a minute fraction of the total volume analysed, their contribution to the Pb counts may be significant. That minor U- and Th-rich phases are the cause of discordance is seemingly demonstrated by the generation of a concordant date from a ‘‘clean’’ area of a grain that also yields two highly discordant dates from regions with numerous uraninite inclusions. A major conclusion of the Witwatersrand geochronological studies (England et al., 2001a; Kositcin et al., 2003) is that xenotime growth is not restricted to diagenesis, as observed in earlier studies, but also proceeds during later tectonothermal and metamorphic events. The age data also show that the development of xenotime outgrowths on detrital zircon grains is not diagnostic of diagenetic growth, but may also occur during subsequent thermal and fluid events. The results highlight the potential usefulness of xenotime to date multiple events in the evolution of a sedimentary basin, as well as some of the complications. However, this has only been possible by combining detailed SEM petrography, with in situ microanalysis and microbeam U – Pb geochronology, allowing the multiple stages of xenotime growth to be differentiated, and to establish a succession of fluid events from diagenesis through to late-stage hydrothermal alteration. With new developments, it may soon be possible to derive U –Pb isotopic data from an area as small as 1 – 2 Am, allowing better discrimination between temporally distinct growth zones. 6.4. Effects of metamorphism on diagenetic xenotime Currently, little is known about the preservation of diagenetic xenotime during prograde metamorphism. Detrital xenotime in chloritoid-garnet quartzite from the Mount Barren Group in Western Australia persists up to lower amphibolite (between ca. 400 and 600 jC) and yields precise and concordant dates. Also, detrital xenotime grains in the Witwatersrand Basin that have undergone several deformational and metamorphic events (up to 350 F 50 jC; Phillips and Law, 2000) still yield concordant and precise dates (up to 3.1 Ga)

despite U concentrations of over 10,000 ppm (England et al., 2001a; Kositcin et al., 2003). Diagenetic xenotime outgrowths have been documented in lower to mid-upper greenschist facies metasandstone (England et al., 2001a; Vallini et al., 2002; Kositcin et al., 2003; Rasmussen et al., 2004), and in one case, from a lower amphibolite facies quartzite (Dawson et al., 2003). In the Witwatersrand Basin, xenotime overgrowths yield precise, and in most instances, concordant U – Pb ages (ca. 2780 – 2760 Ma) interpreted to be diagenetic (England et al., 2001a; Kositcin et al., 2003). The U – Pb isotopic system in diagenetic xenotime appears to have remained closed despite a long (>750 million years) and protracted history of hydrothermal alteration, deformation and metamorphism. Similarly, diagenetic xenotime in lower-mid greenschist facies quartzite from the Proterozoic Stirling Range Formation appears to have been unaffected by metamorphism (Rasmussen et al., 2004). The diagenetic xenotime yields an oldest age population of ca. 1800 Ma (Rasmussen et al., in 2004), some 600 million years older than metamorphism dated at 1200 Ma from metamorphic monazite (Rasmussen et al., 2002a). In metapelites from the Mount Barren Group in southwestern Australia, diagenetic xenotime is common up to lower greenschist facies, typically occurring as irregular to pyramidal outgrowths on detrital zircon or xenotime grains (Fig. 18a). Under greenschist facies conditions, diagenetic xenotime becomes less common relative to metamorphic varieties, which typically occur as isolated equant crystals with a tendency to form euhedral faces aligned with the fabric (Fig. 18). However, metamorphic xenotime may still precipitate as outgrowths on detrital zircon and xenotime grains (Fig. 18). At mid-upper amphibolite facies (garnetstaurolite-kyanite), xenotime comprises entirely metamorphic crystals that yield a U – Pb age of 1206 F 8 Ma (MSWD = 0.73), interpreted as peak metamorphism (Dawson et al., 2003). A single old (concordant) outlier yielded an age of 1297 F 16 Ma and may represent a mixing age with minor overlap onto diagenetic xenotime, or possibly detrital xenotime. Once medium- to high-grade metamorphic conditions have been reached, xenotime occurs as a relatively widespread accessory mineral (Hawkins and Bowring, 1999; Bea and Montero, 1999; Pyle and Spear, 2000; Spear and Pyle, 2002), with a postulated

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Fig. 18. BSEM images of xenotime-zircon associations in mid greenschist to lower amphibolite facies metapelites from the Palaeoproterozoic Mount Barren Group in Western Australia. (a) Detrital zircon with large xenotime outgrowths in a crenulated quartz-mica schist. (b) An irregular zircon core surrounded by a rim of xenotime in a quartz-mica schist. (c) Outline of former zircon grain in an equant xenotime crystal. (d) Elongate xenotime crystal aligned with fabric. Note that muscovite plate is partly engulfed by xenotime (see arrow).

closure temperature of >650 jC (Heaman and Parrish, 1991) or even z 750 jC (Dahl, 1997). During prograde metamorphism, partitioning of REE between coexisting xenotime and monazite has been shown to relate to temperature, rendering xenotime a possible geothermometer (Heinrich et al., 1997; Gratz and Heinrich, 1997; Andrehs and Heinrich, 1998; Gratz and Heinrich, 1998; Viskupic and Hodges, 2001). A garnet-xenotime thermometer has also been developed, based on the observation that Y concentrations in garnet decrease with increasing temperature in metapelites (Pyle and Spear, 2000).

7. Summary Xenotime is a widely distributed diagenetic phase in siliciclastic sedimentary rocks, varying in age from mid-Archaean to Mesozoic. Diagenetic xenotime has been overlooked in the past because of its extremely small size (typically < 10 Am), and its similar optical properties to zircon, making positive identification exceedingly difficult. It starts to precipitate during early diagenesis, but may also form during burial diagenesis, contact metamorphism, very-low to low-grade regional metamorphism and hydrothermal alteration. Xenotime is an ideal U –Pb chronometer because it has high U contents (typically 1000s ppm) and low

initial Pb concentrations, and an exceptional ability to remain closed to U and Pb mobility, allowing for precise and concordant age determination. In addition, the U – Pb isotopic system has the advantage of two independent radioactive decay schemes, providing two sets of age data that allow an internal assessment of whether the isotopic system has remained closed after mineral growth. Once formed, xenotime does not appear to undergo Pb-loss (despite very high U concentrations) nor is it highly susceptible to later thermal resetting. Precise and concordant diagenetic ages have been obtained from mid-upper greenschist and lower amphibolite facies metasedimentary rocks. In sedimentary basins unaffected by tectonothermal events, xenotime may provide several diagenetic ages, the oldest of which can be regarded as a robust minimum age for deposition, and in some instances, a close approximation of the age of deposition. In more complex basins, the growth history of xenotime provides challenges, but with them, opportunities to establish the timing of multiple post-depositional events at an unprecedented level of precision. The challenge is to unravel the number of growth stages and determine whether individual dates are meaningful. This can be accomplished by applying a sample strategy cognisant of the wider geological context (e.g., proximity to intrusions or hydrothermal alteration zones), in conjunction with SEM petrography, in situ microanalysis, X-ray element maps and micro-

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beam U – Pb geochronology. By applying a combination of techniques, it may be possible to elucidate the timing of diagenesis and potentially every major tectonothermal event that has affected a basin. Never before has it been possible to date such an array of low-temperature processes and attain such a level of accuracy and precision. The successful application of U –Pb xenotime geochronology should lead to a better understanding of the timing and interaction of Earth processes. Acknowledgements I would like to thank Ian Fletcher and Neal McNaughton, without whose assistance diagenetic xenotime dating would not have been possible, as well as the many researchers who have participated in work on xenotime at UWA, including Galvin Dawson, Gavin England, Duncan Forbes and Daniella Vallini. I wish to thank Ian Fletcher, Bryan Krapez, Neal McNaughton and Steve Sheppard for discussion and comments, and the staff of UWA’s Centre for Microscopy and Microanalysis for their assistance. I extend my gratitude to Jane Evans, Paul Hill, Steve Sheppard and Jan Zalasiewicz for providing access to rock samples. George Gehrels and Robert Rainbird are thanked for helpful reviews. This work was supported by an Australian Research Council (ARC) fellowship and grant. References Adams, C.J., Barley, M.E., Fletcher, I.R., Pickard, A.L., 1998. Evidence from U – Pb zircon and 40Ar/39Ar muscovite mineral ages in metasandstones for movement of the Torlesse suspect terrane around the eastern margin of Gondwanaland. Terra Nova 10, 183 – 189. Aleinikoff, J.N., Grauch, R.I., 1990. U – Pb geochronologic constraints on the origin of a unique monazite-xenotime gneiss, New York. American Journal of Science 290, 522 – 546. Altschuler, Z.S., 1980. The geochemistry of trace elements in marine phosphorites: Part I. Characteristic abundances and enrichment. In: Bentor, Y.K. (Ed.), Marine Phosphorites. Special Publication-SEPM, vol. 29, pp. 9 – 30. ˚ mli, R., 1975. Mineralogy and rare earth geochemistry of apatite A and xenotime from the Gloserheia Pegmatite, Froland, southern Norway. American Mineralogist 60, 607 – 620. Andrehs, G., Heinrich, W., 1998. Experimental determination of REE distributions between monazite and xenotime: potential

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