Ion Microprobe Zircon U-Pb Age and Geochemistry of the Myanmar Jadeitite

By Shi, Guanghai Cui, Wenyuan; Cao, Shumin; Jiang, Neng; Jian, Ping; Liu, Dunyi; Miao, Laicheng; Chu, Bingbing

Abstract: Combined geochemistry and geochronology of the Myanmar jadeitite were determined. Bulk-rock trace element compositions display U-shaped REE patterns with pronounced positive Eu anomalies. The total REE abundances are very low, less than half chondritic, and the high field strength elements and some large ion lithophile elements are moderately enriched. These features indicate a metasomatic origin. There are three groups of zircons with different interior characteristics, cathodoluminescence, mineral inclusions, chemical compositions and sensitive high-resolution ion microprobe U- Pb ages. Group-I zircons, with a mean age of 163.2 +- 3.3 Ma, mostly have distinct oscillatory zoning, highest U and Th contents, and Na- free, Mg-rich mineral inclusions, and thus indicate an igneous (formation of oceanic crust) or hydrothermal (serpentinization and/ or rodingitization) event in the Middle Jurassic. Group-II zircons, with a mean age of 146.5 +- 3.4 Ma, have bright luminescence without oscillatory zoning and include jadeite and jadeitic pyroxene inclusions, suggesting that formation of the Myanmar jadeitites, as well as subduction of the eastern Indian oceanic plate, occurred in the Late Jurassic. Group-III zircons have an age of 122.2 +- 4.8 Ma, which represents a later unknown thermal event. Discovery of the Middle Jurassic zircons provides geochronological constraint on the tectonic evolution of the eastern Indo-Burman Range.

Jadeitite, a rock made up almost entirely of jadeitic pyroxene, is found in fewer than 10 locations worldwide. The largest and most important jadeitite (jade) deposit occurs in the Hpakan-Tawmaw serpentinite, Kachin State, northern Myanmar (former Burma), conglomerates derived from it (Chhibber 1934; Bender 1983), and similar smaller deposits in the region (Hughes et al. 2000). Next in economic importance, and the source of archaeological New World jade, are the deposits in the middle Motagua Valley, Guatemala (McBirney et al. 1967; Harlow 1994; Seitz et al. 2001). These deposits straddle the Motagua Fault, which connects to the Cayman Trough. A third important jadeitite source is the alluvial deposits of the Itoigawa-Omi area in the Hida Mountains, Japan (Komatsu 1990; Miyajima et al. 1999, 2001); these deposits are located in the so- called ‘Hida marginal zone’ (e.g. Tsujimori 2002), an east-west- trending unit that might be a pre-Jurassic fault zone. There are at least three further jadeitite localities in the Chugoku Mountains of SW Japan, in the Wakasa, Oya and Osayama areas. All these jadeitites occur in a serpentinite melange unit of the Early Palaeozoic Oeyama belt (e.g. Kobayashi et al. 1987; Tsujimori & Itaya 1999; Tsujimori et al. 2005), and these units are the eastern extension of the serpentinite melange of the Itoigawa-Omi area in the Hida Mountains. There is also a Mesozoic jadeitite locality in the Nishisonogi Peninsula, Kyushu, Japan (Nishiyama 1978; Shigeno et al. 2005). Relatively poorly studied deposits exist along the Ketchpel River, in the Pay-Yer massif, Polar Urals (Morkovkina 1960), and along the Yenisey and Kantegir rivers, in the Bonis Mtns., West Sayan- Khakassia (Dobretsov 1963) in Russia, and Itmurundy, near Lake Balkhash, Kazakhstan (Dobretsov & Ponomareva 1965). A small but well- described deposit occurs along Clear Creek in the New Idria serpentinite body, San Benito Co., California (Coleman 1961); this serpentinite is near the San Andreas Fault, also a strike-slip fault. In summary, jadeitite always occurs in serpentinites, and is typically associated with eclogite and blueschist (e.g. Harlow 1994; Shi et al. 2001), within sutures that have evolved from subduction to transpression (Harlow & Sorensen 2005). Apparently, jadeitic pyroxenes in the jadeitite are different from those from typical high-pressure or ultrahigh-pressure metamorphic belts in both P-T conditions and chemical environment (e.g. Essene & Fyfe 1967; Liou et al. 1997; Tropper et al. 1999; Trepmann & Stockhert 2001; Zhang et al. 2002; Liu & Ye 2004), although in some cases it is hard to distinguish jadeitic pyroxenes in blueschist-facies metagreywackes from those in jadeitite (e.g. Terabayashi et al. 1996). Thus geochemical and geochronological studies on jadeitites will yield useful information about their petrogenesis and tectonic relationship to subduction zones.

The Myanmar jadeitite, the largest and most important deposit, is found as veins or blocks connected to veins. The discoveries of metasomatic ureyite (kosmochlor) (Yang 1984; Harlow & Olds 1987; Shi et al. 2005a), metasomatized glaucophane and five other sodic- or sodic-calcic amphibole associations (Shi et al. 2003), as well as methane-bearing fluid inclusions (Shi et al. 2005b) in the Myanmar jade deposits suggest that widespread metasomatism must have occurred during the formation of the Myanmar jadeitite. Sorensen et al. (2006) reported trace element and oxygen-isotope geochemistry of the Myanmar jadeitite and suggested that jadeitite-depositing fluids had multiple sources or evolved in composition along their flow path (or both). Indicative textures are lacking, however, so more geochemical data is needed to determine the jadeitite’s origin. Furthermore, the lack of geochronological data for the Myanmar jadeitite, its host serpentinite and adjacent rock associations restricts integrated and further study in this area. Here we present geochemical data and sensitive high-resolution ion microprobe (SHRIMP) zircon U-Pb dating of jadeitites from the Myanmar jadeitite area. The ages have been interpreted using petrographic analysis and zircon textures. Implications for jadeitite petrogenesis and the tectonics of the eastern Indo-Burman Range are discussed.

Geological setting

The Myanmar jadeitite, located in the western part of the Sagaing fault belt in the Hpakan area of the Kachin State, belongs to the Indo-Burman Range, part of the eastern subduction zone where oceanic crust of the Indian plate was overridden by the Burmese platelet (Fig. 1). The Sagaing fault is a major strike-slip right-lateral continental fault that extends over 1200 km and connects to the Andaman spreading centre at its southern end. It has been interpreted by some workers as a plate boundary between India and Indochina (LeDain et al. 1984; Guzman-Speziale & Ni 1996). Other workers have pointed out, however, that a significant part of the relative motion between the two plates could be accommodated by other tectonic structures in the Myanmar Central Basin and the Shan plateau scarps (Hla Maung 1987; Holt et al. 1991; Bertrand et al. 1999, Bertrand & Rangin 2003). The Indo-Burman Range covers the area of Myanmar between the Myanmar Central Basin and the western border, including the offshore regions and islands of the Bay of Bengal. The eastern boundary of the range is generally defined by a discontinuous line of ophiolite and ophiolite-derived blocks. The accumulation and deformation of rocks in this terrane have taken place within a zone where oceanic crust of the Indian plate was subducted beneath the Burmese platelet. The rocks in the Indo- Burman Range are progressively younger from east to west.

Primary jadeitite deposits occur as veins (called ‘dykes’ by Chhibber 1934) crosscutting serpentinized peridotite bodies that belong to the so-called Hpakan-Tammaw ultramafic body (Fig. 1b) or ophiolite (Shi et al. 2001). The jadeitite veins are almost vertical, strike north-south, and are 1.5-5 m wide and 5-100 m long. The boundary zone between the serpentinized peridotite body and the jadeitite vein is metasomatic amphibolite (Fig. 2) consisting of six types of sodic to sodic-calcic amphiboles (eckermannite, magnesiokatophorite, nyboite, glaucophane, rich-terite, winchite) (Shi et al. 2003). The amphibolite zone ranges from 1 to 50 cm wide. As a result of several stages of later deformation, fragments of the boundary zone can also be found intermingled with blocks of jadeitite that often have a preferred orientation. Within or associated with the amphibolite zones, kosmochlor, Cr-bearing jadeite and some Cr-bearing omphacite (Omp) commonly occur as corona aggregates or small blocks (Shi et al. 2005a). Jadeitized omphacitites are also present (Yi et al. 2006). The jadeitite veins are crosscut in some places by fine veins of late-stage albitite, which are commonly less than 5 mm wide. Outside the ultramafic bodies, high-pressure rocks such as phengite-bearing glaucophane schists and stilpnomelane-bearing quartzites, and amphibolite- facies rocks such as garnet-bearing amphibolites and diopside- bearing marbles occur as tectonic intercalations (Shi et al. 2001).

Petrography of the jadeitite

A few jadeitites retain an undeformed massive structure, with euhedral to subhedral jadeite prisms up to >20 mm long. Most, however, are deformed and fine-grained, with occasional domains of coarse-grained, less deformed or undeformed jadeite (Fig. 3). Cathodoluminescence (CL) images reveal a rhythmic zoning pattern within the coarse-grained jadeite crystals (Fig. 4), similar to observations by Harlow (1994) and Sorensen et al. (2006), who provided CL images of jadeitites from various localities. Most of the deformed jadeitites are fine-grained aggregates of oriented crystals (Fig. 5). Mechanical jadeite twinning as described by Trepmann & Stockhert (2001) also occurs in the jadeitite. Crystallographic orientation is obvious in small domains, indicating rotation recrystallization within ductile deformation. Most grains have undulose extinction. A small-scale shear zone crosscutting the foliation defined by the oriented grains reflects multiple stages of deformation with progressive strain localization. Most of the zircon grains in the Myanmar jadeitite show signs of brittle deformation. Zircon fragments are distributed along ductile-deformed jadeite foliations (Fig. 6). The fragments are irregularly shaped, ranging in size from less than 5 [mu]m to more than 100 [mu]m, and are distributed discontinuously within ductile-deformed jadeite aggregates.

Most of the jadeitites are white, consisting of monomineralic jadeite, with minor accessory minerals such as omphacite, amphibole and zircon, and albite and analcite occur as rare secondary alteration products. The jadeites in white jadeitite are very pure, with jadeite (XJd) contents of more than 98 mol.% (Shi et al. 2003, 2005b). Jadeites from the green jadeitites contain a component of kosmochlor (Shi et al. 2005a), whereas jadeites from the grey jadeitites consist of diopside (Yi et al. 2006).

Internal textures and mineral inclusions of analysed zircons

The jadeitite sample selected for zircon dating was 5.8 kg of white to pale grey jadeitite from the Myanmar tract (sample location is shown in Fig. 1b). After the separation procedure (crushing, sieving, gravity separation, electric-magnetic separation and microscopic selection) at Langfang laboatory, Hebei Geology and Resource Bureau, about 200 zircon grains or fragments were extracted. The zircons are slight yellow-pink. Most of the grains have cracks or appear to be fragments; only a few are intact crystals. The intact grains are euhedral, 50-300 [mu]m long and have an aspect ratio of 1.5-3.0. These features, in combination with the internal texture described below, are a result of natural brittle deformation, not the mineral separation process. Small amounts of pyrite and galena were also recovered.

The selected zircons, together with standard zircons, were mounted in epoxy resin and polished for cathodoluminescence (CL), backscattered electron (BSE) and secondary electron (SE) imaging, and SHRIMP analyses. CL and BSE images of zircons, as well as chemical compositions of mineral inclusions in zircons, were obtained by electron microprobe analysis (EMPA) using a CAMECA-SX- 51 instrument at the Institute of Geology and Geophysics, Chinese Academy of Sciences (IGGCAS), with the same analytical conditions as described by Shi et al. (2005a).

Because of small sizes and possibly not flat surfaces of mineral inclusions within zircons, the data obtained by EMPA (Table 1) approximate their inherent compositions. According to these compositions, two distinct types of inclusions were identified in the Group-I and Group-II zircons, respectively. Mineral inclusions in Group-I zircons are Ni-bearing Mg-rich silicates (5.98-20.70 wt % MgO, 0.04-1.05 wt % NiO); they have variable FeO and Al^sub 2^O^sub 3^ with minor CaO, but no Na^sub 2^O was detected (Table 1). Although it is hard to distinguish the mineral varieties of the inclusions, their chemical features suggest that the inclusions formed in a different environment from that of the jadeitite, which is enriched in Na and Al and thus may represent an earlier formed stage. In contrast, inclusions in Group-II zircons are Na-rich minerals with compositions close to those of jadeite and jadeitic pyroxene (omphacite) (Table 1).

There are at least three groups of zircons according to their internal textures, luminescence (Fig. 7), mineral inclusions (Table 1) and chemical compositions (U and Th contents and Th/U ratios; Table 2). Group-I zircons are typically zoned and have Na-free, Mg- rich silicate mineral inclusions and the highest U and Th contents and Th/U ratios (U 230-4352 ppm, 947 ppm on average; Th 22-1305 ppm, 221 ppm on average; Th/U 0.049-0.31, 0.20 on average) (Table 2). They commonly occur as cores surrounded by Group-II zircon (Fig. 8), and thus formed prior to formation of the Group-II zircons. Group- II zircons are much brighter than Group-I zircons and have no oscillatory zoning. They have mineral inclusions of jadeite and jadeite-rich pyroxene without Na-free mineral inclusions (Fig. 7, Table 1), and have lower U and Th contents and Th/U ratios (U 73- 413 ppm, 204 ppm on average; Th 5-41 ppm, 21 ppm on average, Th/U 0.075-0.178, 0.11 on average) (Table 2). The sizes of Group-II zircon rims vary greatly from 1 [mu]m to 50 [mu]m. They occur mainly as concordant or near-concordant rims on Group-I zircon, and are interpreted as coeval with the formation of the jadeitite. Group- III zircon has the lowest U and Th contents and Th/U ratios. This zircon occurs as small late-stage veins crosscutting Group-I and Group-II zircon, and possibly formed as a result of local recrystallization. Zircon grains commonly consist of Group-I and Group-II zircons; only a few grains have all three groups in one crystal. Most of the Group-III zircons are too narrow for SHRIMP analysis. Some zircons were difficult to classify on the basis of zoning, and were classified mainly by their chemical compositions and ages. Small jadeite or jadeite-rich pyroxene inclusions in Group- II zircons indicate that Group-II zircon is coeval with jadeitite formation. Group-I zircon, which was formed before Group-II and has Na-free Mg-rich silicate inclusions, is inferred to have crystallized during an earlier igneous (formation of oceanic crust) or hydrothermal (serpentinization and/or rodingitization of the oceanic crust) event.

Zircon SHRIMP U-Pb dating

Zircon U-Th-Pb analyses were performed on SHRIMP II at the Beijing SHRIMP Laboratory, Institute of Geology, Chinese Academy of Geological Sciences. The standard SL13 (572 Ma, 238 ppm U) from the Research School of Earth Sciences, Australian National University, was used to calibrate the U, Th and Pb contents of the standard TEM and the analysed samples. Inter-elemental fractionation was calibrated using Pb/U, UO/U and TEM (417 Ma) (Williams 1998). Data were processed using PRAWN software. Because of small amounts of ^sup 207^Pb formed in young (i.e. <800 Ma) zircons, which results in low count rates and high analytical uncertainties, the determination of the ages for young zircons has to be based primarily on their ^sup 206^Pb/^sup 238^U ratios. Common Pb corrections were made using the measured ^sup 207^Pb and Th/U (Compston et al. 1984). Analytical results are listed in Table 2, in which errors are +-1sigma, whereas error for weighted mean ages of samples is quoted at 95% confidence limits.

Twenty-nine analyses were obtained from 19 zircon grains or fragments: 17 analyses on Group-I zircon, nine on Group-II and one on Group-III zircon. Spot 1.2 on Group-III zircon and two other spots (8.1 and 18.1) were not very clear. Analyses on Group-I and Group-II zircons do not show significantly different U-Pb ages. However, if they are treated as a coherent group, the mean square weighted deviation (MSWD) is high (Fig. 9a and b). Given the differences between the two groups of zircons in interior characteristics, mineral inclusions and luminescence, the ages were calculated separately for each group. The 17 analyses on Group-I zircons yield a weighted mean age of 163.2 +- 3.3 Ma, whereas the nine analyses on Group-II zircons yield a significantly lower weighted mean age of 146.5 +- 3.4 Ma (analyses of spots 8.1 and 18.1 are not included) (Fig. 9). The calculations yield much lower MSWD (Fig. 9c-f) for both groups of zircons. This suggests that analyses should be carried out separately for each zircon group. Because most Group-III zircons are too narrow, only one analysis (1.2 in Table 2) on Group-III zircon was obtained.

Geochemistry

Eight samples were selected for chemical analysis, including seven fresh white jadeitites, and one sample (FR-1) containing minor albite and a hydrous sodic silicate (analcime?), possibly secondary minerals after jadeites. The samples were ground in an agate mortar to c. 200 mesh. Major oxides were analysed by XRF using a Phillips PW2400 system at the IGGCAS. Fused glass discs were used and the analytical precisions were better than 5%, estimated from repeat analyses of GSR-3 (basalt, Chinese standard reference material; see Fan et al. (2004)). Trace element abundances were obtained by inductively coupled plasma mass spectrometry (ICP-MS) using a VG- PQII system also at the IGGCAS. Samples were dissolved in distilled HF + HNO3 in 15 ml Savillex Teflon screwcap beakers at 120 [degrees]C for 6 days, dried and then diluted to 50 ml for analysis. A blank solution was prepared and the total procedural blank was <50ng for all trace elements. Precision for all trace elements is estimated to be 5% (some lower abundance elements as Tb, Ho, Tm and Lu are estimated to be about or more than 10%) and accuracy is better than 5% for most elements, as determined by analyses of the GSR-3 standard (see Tang et al. 2006). The results are given in Table 3.

Except for sample FR-1, which has a slightly higher SiO^sub 2^ and loss on ignition (LOI), the jadeitites have similar major chemical characteristics; their major element concentrations are closely equivalent to that of the nearly pure jadeite. All the samples have the same chondrite-normalized REE patterns (Fig. 10), with pronounced positive Eu anomalies (deltaEu= 1.31-4.39). The patterns show light REE (LREE; La-Nd) enrichments, flat middle REE (MREE; Sm-Ho), and heavy REE (HREE; Er-Lu) enrichments, which are similar to these of jadeitite from the Itoigawa-Ohmi district, Japan (Morishita et al. 2007). The total REE abundances of the samples are very low (0.278-1.279 ppm), less than half chondritic. All the studied samples also have nearly the same trace element contents (Fig. 11). Their high field strength elements are moderately enriched, and also some large ion lithophile elements, such as Ba, Sr and U. In general, the trace element contents of these rocks are very low. Discussion

Petrogenesis of jadeitite and derivation of jadeitite-forming materials

Previous genetic models for the origin of pure jadeitite include metasomatism and metamorphism (Coleman 1961, 1980; Mevel & Kienast 1986; Goffe et al. 2000; Harlow & Sorensen 2005), pressure solution and redeposition in fractures (Harlow 1994), and crystallization from a fluid phase (Shi et al. 2000). The systematic absence of quartz coexisting with jadeite precludes formation of jadeitite by reaction in a closed system from preexisting albite. In the Myanmar jadeite area, a distinct amphibolite boundary zone between the jadeitite bodies and their serpentinized host peridotite, as well as textures and compositions of kosmochlor and chromian jadeite, indicate that metasomatic reactions took place (Shi et al. 2003, 2005a). There is a striking similarity between the jadeite occurrences in Myanmar and Guatemala. In Guatemala, jadeite occurs as veins in serpentinites along the left-lateral Motagua strike- slip fault zone that forms the boundary between the North American and Caribbean plates (e.g. Johnson & Harlow 1999). For the pure jadeitites, however, except for the presence of a hydrous fluid during formation of jadeitite as revealed by primary fluid inclusions (Johnson & Harlow 1999; Shi et al. 2005b), there is no clear petrographic evidence for their metasomatic origin, especially for those jadeitites from large veins or blocks in the Myanmar jadeitite.

Geochemical features of the studied jadeitites preclude the Myanmar jadeitite from being a rock of igneous origin (Chhibber 1934), or a metamorphosed igneous rock, but point to a metasomatic origin as suggested by Coleman (1961), Harlow (1994), Shi et al. (2005b), Sorensen et al. (2006) and Morishita et al. (2007). The CL image of Figure 4 suggests that jadeite has precipitated directly from fluid (see Sorensen et al. (2006)). The total REE abundance and REE patterns, which are strikingly similar in samples, can be used to infer the original rock types. The eight jadeitites investigated yield a W-like normalized pattern; that is, a U-shaped REE pattern with strong positive Eu anomalies (Fig. 10), which is very rare in all types of igneous rocks. Of all igneous rocks, only some anorthosites exhibit seemingly similar patterns with strong positive Eu anomalies and LREE enrichments (e.g. Bhattacharya et al. 1998; James et al. 2002; Chartier et al. 2005). These anorthosites, however, have total REE contents nearly 10 times higher than those of the Myanmar jadeitites, and their HREE are depleted, the inverse of the pattern in the jadeitite. Recent experimental studies have demonstrated that the chlorine content and redox potential of a fluid is a major factor controlling LREE and Eu complexing and transportation (Allen & Seyfried 2005), and that the presence of plagioclase is not required for the generation of LREE-enriched fluid compositions with positive Eu anomalies.

The normalized REE patterns of the Myanmar jadeitites are similar to those of fluid-dominated serpentinized abyssal peridotites at the Mid-Atlantic Ridge, which are characterized by high sulphur and development of U-shaped REE patterns with strong positive Eu anomalies. These are also the characteristic of hot (350-400 [degrees]C) vent-type fluids discharging from black smoker fields (Paulick et al. 2006). Paulick and others considered that these REE patterns are similar to the REE characteristics of hot (350-400 [degrees]C) hydrothermal fluids discharging at black smoker sites above ultramafic-hosted hydrothermal systems such as Rainbow and Logatchev (Douville et al. 2002), and inferred that serpentinization involving vent-type fluids and high fluid/rock ratios imposed the REE signature of the fluid onto the serpentinites. Recent study has shown, however, that the REE patterns of the serpentinized peridotite that hosts the Myanmar jadeitite (Fig. 10) are roughly flat and show a remarkable similarity to those of mafic rocks, do not have any characteristics of the jadeitites, and are very different from those of the fluid-dominated serpentinized abyssal peridotites at the Mid-Atlantic Ridge (Paulick et al. 2006). Furthermore, their total REE content is about 10 times than that of the jadeitites (Shi et al. 2001). Apparently, the REE systematics of the serpentinized peridotites did not dominate those of the jadeitites. Combined with obvious enrichments in some incompatible elements (e.g. Ba, Sr), along with high Pb abundances, which suggest that a relatively recent addition of subduction-related fluids is possible because Pb is highly soluble in such fluids, it is inferred that the jadeitite-forming medium was hydrothermal fluids enriched in Na, Al and Si. These fluids or melts were possibly associated with subduction and/or possibly with discharge at black smoker sites from ultramafic-hosted hydrothermal systems (e.g. Douville et al. 2002), explaining the existence of some sulphide minerals in the Myanmar jadeitite.

The jadeitite-forming medium is highly likely to have been derived from recycled subducted oceanic slabs. The required aqueous solution rich in Na, Al and Si is reported to be characteristic of the descending slab near the transition from blueschist to eclogite (Manning 1998). In addition, Ba-rich minerals have been reported in other jadeitites or albitite-metasomatized rocks in serpentinite- matrix melanges (Kobayashi et al. 1987; Harlow 1995). Very recently, barian feldspars were reported by Morishita (2005) in a jadeitite from the Itoigawa-Ohmi district in the Renge high-P-T metamorphic belt, Japan. Morishita implied that Ba-bearing fluids forming the jadeitite were possibly produced by dehydration of Barich sediments in the early stages of subduction of oceanic crusts and could be related to the serpentinization of the mantle wedge. In the Myanmar jadeitite, the enrichment in Ba (Fig. 11) relative to other trace elements suggests the same interpretation for the derivation of jadeitite-forming materials.

Interpretation of SHRIMP U-Pb dating

The weighted mean age of 163.2 +- 3.3 Ma (Fig. 9) for Group-I zircons is interpreted as the age of igneous or hydrothermal (serpentinization) rocks prior to the Myanmar jadeitite formation. Because oscillatory-zoned zircons are also common in hydrothermal ones (e.g. Dubiska et al. 2004; Tsujimori et al. 2005), the oscillatory zoning, as well as jadeite-absent Mg-rich mineral inclusions in most Group-I zircons suggest that these zircons crystallized in either a magmatic or a hydrothermal environment. As the Myanmar jadeitite has no indicator of an igneous origin or of a Na-free, Mg-rich environment, Group-I zircons are thus inherited relative to the jadeitite, and are inferred to be associated with the formation of the jadeitite-host ultramafic rocks, known as the Tamaw-Hpakan ultramafic bodies. These ultramafic bodies belong to the ophiolite (Shi et al. 2001) that primarily formed as an oceanic slab of the Indian plate and subsequently became serpentinized as serpentinite-matrix melange before or during subduction beneath the Burmese platelet (Fig. 1). Thus the oceanic crust of the Indian plate, which included the Myanmar jadeitite, formed or was subsequently serpentinized and/or rodingitized in the late Middle Jurassic.

The weighted mean age of 146.5 +- 3.4 Ma for Group-II zircons (Fig. 9) is interpreted as the formation age of the jadeitite. The Group-II zircons have lower Th/U ratios and no apparent growth zoning, indicating a metasomatic event related to jadeitite formation. This interpretation is supported by the presence of jadeite inclusions and the absence of any Mg-rich mineral inclusions in the Group-II zircon. Because most of these Group-II zircons crystallized as a concordant overgrowth surrounding pre-existing Group-I zircons, this coeval event is inferred to be under a long- term and relatively stable environment. Thus, the Myanmar jadeitites are Late Jurassic in age. Because jadeitites form only in subduction zones under high-P-T conditions, this Late Jurassic age also dates subduction of the oceanic slab at the eastern edge of the Indian plate.

One young age of 122.2 +- 4.8 Ma for Group-III zircon is interpreted as the age of a later thermal event after the jadeitite formation. The Th/U ratio of this zircon is the lowest. The Th/U ratio, lack of growth zoning, and the growth habit crosscutting Group-I and Group-II zircons indicate a very local and short-term recrystallization, possibly coeval with later ductile deformation of the jadeitite. If so, this deformation occurred in the Early Cretaceous.

Tectonic implications

The new SHRIMP zircon U-Pb ages provide geochronological constraints on the tectonics of the eastern Indo-Burman Range. Previous workers have suggested that the oldest rocks in the east Indo-Burman Ranges are a Late Cretaceous sedimentary section, including a dismembered ophiolite, unconformably overlying Carnian- age sediment, which in turn structurally overlies older schists, including blocks of serpentinite, gabbro, basalt, Triassic sandstone and schist (Mitchell 1993; Acharyya 2007). A string of ultramafic rocks of Late Cretaceous-Eocene age containing high-pressure metamorphic facies and including the famous jade mines of the Hpakan- Tawmaw jade tract is distributed along and adjacent to the trend of the Sagaing fault (Hughes et al. 2000). Hutchison (1975) and Charusiri et al. (1993) interpreted these rocks as Late Cretaceous ophiolites marking the collision zone of the West Burma plate and the Shan-Thai block. Our new SHRIMP U-Pb dating on zircons unambiguously demonstrates the presence of Middle Jurassic zircons of igneous or hydrothermal origin, which makes it necessary to reinterpret the tectonic structure and evolution of the Indo-Burman Range. Geochronological investigations of the adjacent eastern area have proposed the existence of Jurassic oceanic slabs beneath the southeastern margin of Eurasia. Very recent investigations by Barley et al. (2003) showed that strongly deformed granitic orthogneisses near Mandalay in the high-grade Mogok metamorphic belt contain Jurassic (c. 170Ma) zircons that partly recrystallized during c. 43 Ma high-grade metamorphism. A hornblende syenite from Mandalay Hill also contains Jurassic zircons. Barley et al. found these results to be consistent with the interpretation by Khin (1990) of arc to back- arc basin environments in western Myanmar at this time and also the suggestion of Searle et al. (1999) that the southern margin of Eurasia converted from a passive margin to an Andean-type convergent margin in the mid-Jurassic.

As the age of formation of the jadeitite can also be regarded as the subduction age of the host oceanic slabs, the narrow time span (c. 16 Ma) from oceanic slab generation-hydrothermal alteration at c. 163 Ma to its subduction at c. 147 Ma leads to two possible interpretations: one is that there had existed a very small ocean or ocean-like sea between the Indian plate and the Burmese (Myanmar) platelet; the other is that the ocean between the Indian plate and the Myanmar platelet was perhaps not narrow but was nearly enclosed by the Middle Jurassic.

We are indebted to R. X. Zhu and L. C. Chen for their kind support during the field trip and subsequent research, as well as X. D. Jin, H. Li, Q. Mao and Y.G. Ma, for help with the ICP-MS, XRF and EMP work. Extensive reviews by I. Williams, and constructive reviews by T. Tsujimori, greatly helped to improve the manuscript and are gratefully acknowledged. The research was supported by the National Science Foundation of China (40672046,40221402 and 40572049).

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Received 10 March 2007; revised typescript accepted 29 May 2007.

Scientific editing by Martin Whitehouse

GUANGHAI SHI1,2, WENYUAN CUI3, SHUMIN CAO4, NENG JIANG2, PING JIAN5, DUNYI LIU5, LAICHENG MIAO2 & BINGBING CHU1

1 China University of Geosciences, Beijing 100083, China (e- mail: shiguanghai@263.net.cn)

2 Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, China

3 School of Earth and Space Sciences, Peking University, Beijing 100871, China

4 Guangdong Province Material Testing Center, Guangzhou 510080, China

5 SHRIMP Laboratory, Institute of Geology, Chinese Academy of Geological Sciences, Beijing 100037, China

Copyright Geological Society Publishing House Jan 2008

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