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Diamondiferous Archean rocks of the Olondo greenstone belt (western Aldan–Stanovoy shield)
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Russian Geology and Geophysics 53 (2012) 1012–1022 www.elsevier.com/locate/rgg
Diamondiferous Archean rocks of the Olondo greenstone belt (western Aldan–Stanovoy shield) A.P. Smelov a,*, V.S. Shatsky b, A.L. Ragozin b, V.N. Reutskii b, A.E. Molotkov a a
Diamond and Precious Metal Geology Institute, Siberian Branch of the Russian Academy of Sciences, ul. Lenina 39, Yakutsk, 677980, Russia b V.S. Sobolev Institute of Geology and Mineralogy, Siberian Branch of the Russian Academy of Sciences, pr. Akademika Koptyuga 3, Novosibirsk, 630090, Russia Received 31 October 2011; accepted 23 March 2012
Abstract Diamond from metaultramafic rocks of the Mesoarchean (2.96–3.0 Ga) Olondo greenstone belt, located in the western Aldan–Stanovoy shield, has been studied. Diamonds occur in lenses of olivine–serpentine–talc rocks within metaultramafic rocks of intrusive habit, whose composition corresponds to peridotite komatiites. All diamonds from the metaultramafic rocks are crystal fragments 0.3 to 0.5 mm in size. Morphological examination has revealed laminar octahedra, their transitional forms to dodecahedroids, crystals with polycentric faces, and spinel twins. The crystals vary in photoluminescence color: dark blue, green, yellow, red, or albescent. Characteristic absorption bands in crystals point to nitrogen impurities in the form of A and B1 defects and tabular B2 defects. The crystals studied belong to the IaA/B type, common among natural diamonds. The overall nitrogen content varies from <100 to 3800 ppm. The relative content of nitrogen in B1 centers varies from 0 to 94%, pointing to long stay in the mantle. The carbon isotope ratio in the diamonds, 13C = –26‰, is indicative of involvement of subducted crust matter in diamond formation in the Archean. © 2012, V.S. Sobolev IGM, Siberian Branch of the RAS. Published by Elsevier B.V. All rights reserved. Keywords: diamond; Archean; komatiites; greenstone belt; Aldan–Stanovoy shield
Introduction The mining of primary diamond deposits and placers in the Siberian craton yields 99% of the overall diamond winning in Russia (Dobretsov and Pokhilenko, 2010). Diamonds from these deposits have been studied by a wide range of up-to-date methods (Logvinova et al., 2011; Skuzovatov et al., 2011; Zedgenizov et al., 2011). However, it is expected that the exhaustion of the largest deposits, Mir and Udachnaya kimberlite pipes, will cause a decline in diamond output. Therefore, search for new diamond deposits is an urgent task (Dobretsov and Pokhilenko, 2010). As reported in (Khaidarov and Chechetkin, 1990), R.A. Khaidarov (Central Research Institute of the Geology of Non-Ore Minerals, Kazan, Russia) found two diamond crystals, 0.1 and 0.8 mm, in chip samples of metaultramafic rocks sampled in the Olondo Mesoarchean greenstone belt (Drugova et al., 1983, 1988; Popov et al., 1990, 1995; Dobretsov et al., 1997). In 1987–1989, a test survey crew of the Siberian
* Corresponding author. E-mail address: [email protected] (A.P. Smelov)
Research Institute of Geology, Geophysics and Mineral Resources, working under the auspices of the All-Russia Lamproite Program, found diamonds in alluvia of the Olondo and Taryn-Yuryakh Rivers, draining volcanic deposits of the belt: 3 grains 0.5–1.0 mm in size; 31 grains 0.2–0.5 mm; 1 crystal fragment 1.1 × 1.1 mm; and 1 rhombododecahedroid crystal 1.1 × 1.0 mm. At the same time, V.V. Levitskii took stream sediment samples from the Olondo River and its tributaries. These samples were bulked after washing and treated at the Irkutsk Institute of Rare and Precious Metals and Diamonds (Irkutsk, Russia). They yielded 18 diamond crystals. Later, the Olondo team of the Udokan mining enterprise conducted prospecting for diamonds in the catchment areas of the Olondo and Taryn-Yuryakh Rivers. Eight diamonds were found near the southwestern shore of Lake Tokko: two colorless transparent octahedroid crystals 0.7 × 0.8 mm in size and six diamond fragments (fraction –0.5 to +0.525 mm) with characteristic pale greenish-yellow color. Eleven diamonds of class 0.25 mm and two crystals 0.8 mm were found on the southern slope of the Krasnaya Gorka metaultramafic block in the upper reaches of Olondo. One of the larger diamonds was a flat-faced octahedron and the other, an irregularly
1068-7971/$ - see front matter D 201 2, V . S. S o bolev IGM, Siberian Branch of the RAS. Published by Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.rgg.2012.0 + 8.005
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Fig. 1. Map of terranes of the Aldan–Stanovoy shield according to (Smelov and Timofeev, 2007; Smelov et al., 2001), as subsequently amended: 1, deposits of the Siberian craton cover; 2–5, terranes: 2, granite–greenstone (WA—Western Aldan—AR, EBT—Batomga—Pt1), 3, tonalite–trondhjemite gneiss (TN—Tynda—ARPt1), 4, granulite–orthogneiss (ANM—Nimnyr—AR-Pt1, CG—Chogar—AR), 5, granulite–paragneiss (AST—Sutam—AR Archean, EUC—Uchur—Pt1); 6, tectonic mélange zones (am, Amga; kl, Kalar; tr, Tyrkanda); 7, faults (dj, Dzheltulak; ts, Taksakanda); 8, thrusts. The Batomga (EBT) and Uchur (EUC) terranes form the Eastern Aldan superterrane, and the Nimnyr (ANM) and Sutam (AST) ones form the Central Aldan superterrane.
shaped fragment. The smaller diamonds were transparent pale greenish-yellow irregularly shaped diamond fragments with rare preserved octahedron faces. A small sample taken from the talus of a small ultramafic block on the right bank of the Skvoznoi Creek (large right tributary of Taryn-Yuryakh) contained four pale greenish-yellow octahedral diamond crystals up to 0.25 mm in size (Yurgenson et al., 1999). Thus, about 80 diamond crystals were found within five years of prospecting works in the alluvial deposits. V.G. Gadiyatov inferred from analysis of available prospecting data and comparison of the geologic structure of the Olondo greenstone belt (Dobretsov et al., 1997; Drugova et al., 1983; Popov et al., 1990, 1995) with Dachine diamondiferous volcanic rocks in French Guiana (Capdevila et al., 1999) that komatiitic tuffs of the belt, dated to 2.96–3.01 Ga, were the diamond ore body. In 2002, he took samples from volcanic rocks of different compositions. Yu.G. Tyllar investigated rock samples from upper reaches of the Tokko River and found over 100 diamond grains in a single 150-g sample recognized in the field as komatiitic tuff (Gadiyatov, 2005; Gadiyatov et al., 2003a,b). Nearly all of them were irregular fragments 0.3–0.5 mm in size. Most grains were used up in diagnostics and external expertise. The physical parameters of the diamonds, as well as the mineralogical and petrochemical compositions of the host rocks remained unexplored. The very fact of the presence of ancient diamonds in rocks of the Aldan–Stanovoy shield is of great importance for understanding diamond formation and geodynamics in the early history of the Earth, because such findings are rare. The
oldest diamonds (3.0–2.8 Ga) are found in gold ore conglomerates of the Witwatersrand basin, South Africa. Relatively younger diamonds have been recorded in lamprophyres of the Abitibi–Wawa greenstone belt, Superior Craton (2.7 Ga; Stachel et al., 2006) and in metamorphosed kimberlites of Gabon, Central Africa (2.85 Ga; Henning et al., 2003). V.G. Gadiyatov kindly provided 25 diamond grains and crushed host rock (fraction 3–5 mm) from the Olondo greenstone belt for our study. Here we report the results of their investigation and analysis of the geologic location of the diamondiferous rocks.
Geology of the Olondo greenstone belt The Aldan–Stanovoy shield is located on the southern margin of the Siberian craton. It is built mainly by deeply metamorphosed rocks of the granulite facies and, to a lesser extent, by rocks of the amphibolite and greenschist facies. As shown in Fig. 1, five terranes are recognized in the shield on the base of rock composition, metamorphism degree, and age: the Western Aldan granite–greenstone composite terrane, Central Aldan and Eastern Aldan superterranes in the north; Tynda tonalite–trondhjemite-gneiss composite terrane and Chogar granulite–orthogneiss terrane (Smelov et al., 2001) in the south. In a collision, which occurred 1.9 Ga BP, the composite terranes and superterranes formed the base of the North Asian craton (Smelov and Timofeev, 2007). The Western Aldan granite–greenstone terrane is the oldest com-
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Fig. 2. The center and east of the Western Aldan granite–greenstone terrane (Smelov, 1996). 1, Neoproterozoic–Cenozoic deposits of the platform cover and depressions of the Mesozoic and Cenozoic activity stages; 2, Paleoproterozoic metamorphosed subplatform deposits, the Udokan complex (graben–synclines: LKh, Lower Khani; Ol, Oldongso; Ug, Ugui); 3, outcrops of greenstone belt rocks (Subgan greenstone complex): I, Itchilyak fragment; Ya, Yaelakh fragment; Old, Olondo structure; Tm, Tasmiela structure; Sl, Syrylyr structure; Tn, Tungurcha structure; 4, tonalite–trondhjemite gneisses and migmatites of the Olekma complex; 5, granulite (Kurulta) complex (a, paragneisses; b, orthogneisses); 6, metagabbro–diorite–tonalite–trondhjemite complex (Amn, Amnunnakta block; Tu, Tungurchakan block); 7, tonalites and trondhjemites (UO, Ust’-Oldongso block); 8, granites and bastard granites, poorly defined; 9, geologic boundaries (a); faults with steep displacement surfaces: A, Amga; Kh, Khani; Ch, Chara (b); c, thrusts: S, South-Chulman thrust.
posite one, which was not reworked at granulite facies conditions in the course of the Paleoproterozoic collision. The Western Aldan composite terrane is 400 × 350 km in size. It borders the Baikal–Patom fold-and-thrust belt in the west. In the east and south, it borders the Amga and Kalar tectonic mélange zones, respectively. In the north, the terrane is overlain by Upper Riphean and Vendian deposits of the Siberian Craton sedimentary cover (Fig. 2). The terrane is built by Archean bodies of various types, which were metamorphosed in broad temperature and pressure ranges (Mironyuk et al., 1971; Smelov, 1989). Tonalite–trondhjemite orthogneisses are predominant there. They are bulked into the Olekma complex. They form several large linear blocks separated by four meridional belts 300 km in length and up to 30 km in width, where tectonic flakes of greenstone rocks of the Subgan complex are concentrated. The greenstone belts are typically confined by blastomylonites. The Kurulta granu-
lite complex forms several blocks separated by faults and tectonic flakes (Fig. 2). All these blocks and zones can be interpreted as individual terranes. Therefore, the Western Aldan terrane is determined as a composite one, consisting of several terranes. Their main feature is the presence of Archean greenstone bodies and tonalite–trondhjemite orthogneisses. The Olekma tonalite–trondhjemite complex includes compositionally uniform biotite, biotite–amphibole, and amphibole plagiogneisses and gneisses. They have Na : K ratios within 2.5–5.0 and belong mainly to the high-alumina type. They are enriched in LREE and Sr and depauperated in U (Dobretsov, 1986). The associated mafic schists and amphibolites constitute up to 10% of the complex (Cherkasov, 1979). The age of the tonalite–trondhjemite gneisses is 3.2–2.8 Ga (Baadsgaard et al., 1990; Neymark et al., 1993; Nutman et al., 1992). The progressive metamorphism of the rocks corresponds to conditions of the moderate pressure
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amphibolite facies (T = 650–700 °C, P = 5.5–7.0 kbar) or, less frequently, the epidote–amphibolite ones. Nearby the greenstone belts, orthogneisses have been reworked under the conditions of the epidote–amphibolite facies (T = 475– 535 °C, P = 5.0–4.0 kbar) to form heteroblastic, glomeroblastic, and blastomortar textures (Smelov, 1989). The Kurulta granulite complex, which builds up the Olomokit, Chara, and Kalar blocks, consists mainly of orthorocks: enderbite and charnokite gneisses. There are also paragneiss bodies, which include amphibole- and pyroxenecontaining gneisses and schists, garnet–biotitic gneisses, sometimes with cordierite and sillimanite. The blocks themselves are relatively thin tectonic flakes thrust over rocks of the trondhjemite–gneiss and greenstone complexes (Stognii et al., 1996). Little is known about the age of the rocks. It has been found that rocks with preliminary Nd age 3.5 Ga gave rise to protoliths of the garnet–biotite gneisses of the Olomokit block (Kovach et al., 1995), and protoliths of two-pyroxene crystalline schists formed 3.15 Ga BP (Levchenkov et al., 1987). The Subgan greenstone complex was recognized by E.P. Mironyuk and V.S. Fedorovskii in 1963–1968. Bodies of greenstone deposits vary in section patterns and metamorphism conditions (Smelov, 1989). The formation of the sedimentary and volcanogenic rocks of the greenstone belts is dated to two time spans: 3.2–3.0 and 3.0–2.7 Ga BP (Smelov et al., 2001). These time spans were separated by a stage of thrusting and metamorphism of young greenstone belts about 3.0 Ga BP (Nutman et al., 1992; Smelov, 1996). Stratigraphic relationships between the greenstone belts and the neighboring rocks of the tonalite–trondhjemite complex are unknown. The contacts are mainly tectonic or intrusive. The largest volumes of mafic and ultramafic volcanic rocks in combination with felsic and intermediate ones occur in the Tokko–Khani greenstone belt, which is located in the middle of a composite terrane (Fig. 2). The Chara–Tokko and Temulyakit–Tungurcha greenstone belts, located west and east of Tokko–Khani, are characterized mainly by tholeiitic volcanism, notable amounts of ferruginous quartzites, and appearance of carbonate–terrigenous rock assemblages (Temulyakit–Tungurcha and Saimagan) (Smelov, 1989). The southern portion of the Tokko–Khani belt has been best studied. It is referred to in the literature as the Olondo greenstone belt. The Olondo greenstone belt is a V-shaped body of submeridional strike in plan (Fig. 3). It is built by mafic varieties of volcanogenic and volcanosedimentary deposits, characteristic of Archean greenstone belts (Popov et al., 1990). Its section includes three metavolcanic rock associations differing in rock composition and spatial location. Comprehensive lithostratigraphical studies were insufficient for proper understanding the succession of primary superposition of rocks, which had experienced intense metamorphism under the conditions of the epidote–amphibolite facies of andalusite–sillimanite type (T = 550–570 °C and P = 3.5–4.5 kbar) and repeated deformations. Moreover, the belt contains numerous intrusion bodies of various ages and compositions, which also complicate the reconstruction of the stratigraphy of vol-
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canosedimentary strata and their facies transitions (Popov et al., 1995). The first rock association occurs in the margin of the V-shaped body (Fig. 3). It is characterized by alternating mafic and ultramafic metavolcanic rocks. Rocks of same lithological groups differ in geochemistry, depending on their spatial location. In the eastern part of the body, the planar structures and boundaries of rock bodies generally dip to the east. The following sequence is observed from the axis to margins of the body: intermediate and felsic metavolcanic rocks; metatholeiites; ultramafic metavolcanic rocks (chlorite–actinolite schists with interbeds of anthophyllite–carbonate rocks); and metatholeiites with interbeds of komatiites in the form of anthophyllite–carbonate rocks and intermediate and felsic volcanic rocks, as well as ferruginous quartzites and metapelites. The thickness of the eastern branch of the body is estimated to be about 400 m with regard to large folds notable in the map scale. The Sm–Nd isochronous age was established from bulk samples of ultramafic and mafic volcanic rocks to be in the range 2973 ± 48 to 2959 ± 17 Ma (εNd(T) = 2.2 ± 0.1) (Pukhtel and Zhuravlev, 1993), and the U–Pb age (SHRIMP) of zircons from interbeds of intermediate and felsic volcanic rocks, 2986 ± 6 to 3005 ± 5 Ma (Baadsgaard et al., 1990; Nutman et al., 1992). In the western branch of the body, the first association is dominated by amphibole–plagioclase schists (metatholeiites) with frequent thin interbeds of metavolcanic rocks of intermediate composition. Ultramafic metavolcanic rocks are absent. The thickness of this part of the body without regard to folding is 250–300 m. The second rock association builds the central part of the belt. It consists of amphibole and biotite–amphibole microgneisses of intermediate and felsic compositions. Their effusive origin is proven by well-preserved remains of primary structures and textures. Typically, the proportion of felsic rocks increases towards the eastern part of the body. The microgneisses host rare thin (5–10 m) interbeds of amphibole schists (metabasalts), sulfidized garnet amphibolites, and biotite–garnet (sometimes with staurolite) microgneisses. The association includes rocks formed from lavas, tuffs, and tuffaceous sandstones, as well as vein stones. The thickness of the body in the eastern part of the belt is estimated to be about 500 m. The U–Pb isochronous age of zircons from metadacites of this association is 2950 ± 50 Ma (Bibikova et al., 1984). The third rock association is confined to the axial part of the eastern branch of the belt. It consists of frequently alternating interbeds of metatholeiites, metatuffs, metasandstones, and metapelites. Less frequent are metaandesites, which presently include amphibole, amphibole–biotite, and biotite schists; microgneisses; and high-silica rocks. The metatuffaceous sandstones and metasandstones display a clear rhythmicity, emphasized by interbeds with carbonate cement. Rocks with well-preserved structures of tuff breccias containing volcanic rocks of various compositions and agglomerate
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Fig. 3. Geological outline of the Olondo greenstone belt according to (Popov et al., 1995). 1, Quaternary fluvioglacial deposits; 2, bifeldspathic granites; 3, terrigenous deposits of the Udokan complex; 4, amphibole, biotite–amphibole, and biotite microgneisses (metaandesites and metadacites, their tuffs, and metasandstones); 5, variegated deposits: intercalation of microgneisses and amphibole–plagioclase schists; 6, amphibole–plagioclase schists (metatholeiites); 7, actinolite–chlorite schists (komatiites and komatiite basalts); 8, carbonate–anthophyllite schists (komatiites); 9, gabbro–amphibolites and hornblendites; 10, olivinites, serpentinites, and talc rocks (blocks indicated with encircled numerals: 1, Krasnaya Gorka; 2, Tokko; 3, Taryn-Yuryakh); 11, differentiated bodies of gabbro–diorite–tonalite composition; 12, trondhjemites; 13, tonalite–trondhjemite gneisses and migmatites; 14, picrite dikes; 15, geologic boundaries; 16, (a) proven faults, (b) conjectured faults; 17, site of diamondiferous sample 18/4.
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tuffs occur in some exposures. The thickness of the body is tentatively estimated to be 200 m. Intrusions constitute no less than 30% of the volume of the Olondo greenstone belt. Ultramafic, mafic, and felsic rocks are recognized there. They belong to different age stages. The oldest intrusions include numerous bodies of ultramafic rocks of intrusive habit, sill-like bodies of gabbro–amphibolites, and differentiated blocks of gabbro–diorite–tonalites. The ultramafic rocks are metamorphosed dunites and peridotites. They have experienced profound serpentinization and, in places, secondary talc development and carbonatization. Secondary rock transformation destroyed the initial structures, textures, and minerals almost everywhere. The Sm–Nd isochronous age in bulk samples of peridotites of the Krasnaya Gorka block is 3003 ± 117 Ma (εNd(T) = 1.12 ± 0.1) (Pukhtel and Zhuravlev, 1993). Gabbro–amphibolites are predominant in the belt. The largest bodies are differentiated from actinolite–chlorite ultramafic rocks and monomineral amphibolites (hornblendites) at the bottom to mesocratic gabbro–amphibolites at the roof. These bodies are sill-shaped, 50 to 350 m in thickness. They are confined to boundaries between rock types. The rocks differ from metabasalts in being more massive in structure and containing remains of igneous textures. They are sheeted along contacts and within zones. A complex of differentiated gabbro–diorite–tonalite bodies was reliably recognized in the central part of the belt in the branching region (Fig. 3). Intrusive associations are observed between rocks of the complex contrasting in composition. The U-Pb SHRIMP age of zircons from diorites of the complex is 3018 ± 10 Ma (Baadsgaard et al., 1990; Nutman et al., 1992). The youngest intrusions are bodies of granitoids of tonalite– trondhjemite composition, widespread in the margin of the Olondo belt and forming large blocks in the center. The tonalites and trondhjemites are represented by amphibole and biotite varieties. They are sometimes migmatized. They contain inclusions of all types of volcanic and metaultramafic rocks and gabbro–amphibolites, varying in size. The U–Pb SHRIMP age of zircons from tonalites of the eastern contact of the belt is 2862 ± 14 Ma (Baadsgaard et al., 1990; Nutman et al., 1992). Dike-like bodies of ultramafic and mafic rocks, assigned to the picrite series (Dobretsov et al., 1986), vary in thickness from tens of centimeters to meters. They are confined to borders between rock associations and contacts with early intrusions. They are boudined and sheeted along contacts. The Sm–Nd isochronous age from bulk samples of magnesium-rich picrites is 2202 ± 41 Ma (εNd(T) = 2.2 ± 0.2) (Pukhtel and Zhuravlev, 1993).
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350 microanalysis system (Oxford Instruments) at accelerating voltage 20 kV and current 1 nA. The accuracies of determination of elements (wt.%) were: Si Kα, 0.51; Al Kα, 0.13; Cr Kα, 0.18; Fe Kα, 0.28; Mn Kα, 0.14; Mg Kα, 0.42; Ni Kα, 0.37; O, 0.69. All analyses were done at the Diamond and Precious Metal Geology Institute, Yakutsk, by N.V. Leskova. The analyses of rock-forming, rare, and rare-earth elements in the rocks were done by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) by I.V. Nikolaeva, Institute of Geology and Mineralogy, Novosibirsk. Carbon isotope ratios in diamonds were analyzed at the Analytical Center, Institute of Geology and Mineralogy, Novosibirsk. As the diamonds were small, their carbon isotope ratios were determined in a 0.3-mg weight of five fragments. The weight was loaded into a platinum capsule and placed into a quartz reactor with purified copper oxide. The weight was combusted at 950 °C. The resulting carbon dioxide was purified and collected into a dismountable vacuum trap. Diamonds were prepared for isotopic studies as described in (Reutskii et al., 1999). Carbon isotope ratios were determined in carbon dioxide with a Finnigan MAT Delta mass spectrometer in the dual inlet mode. To ensure the accuracy of δ13C determination, it was done against the international USGS-24 standard (graphite, δ13C = –15.9). The reproducibility of the data with reference to the standard, including sample preparation, did not exceed 0.1 ‰ (2σ). As the amount of the material was limited, the result was obtained in a single measurement: δ13C = –26.0 ‰ PDB. Infrared absorbance was measured with an IFS 130v Fourier transform IR spectrometer (Bruker). Spectra were recorded within the wavelength range 2.5–12.5 µm (wavenumbers 4000–800 cm–1), beam aperture from 30 × 30 to 80 × 80 µm. The absorbances at the maximum of the two-phonon lattice absorption (2030 cm–1) were used as an internal standard (I2030 = 12.8 cm–1) to determine the net thickness of the sample. The interpretation of the IR spectra involved attribution of characteristic bands to impurity defects and quantitative analysis: determination of concentrations of nitrogen atoms belonging to various optically active centers. The following relationships exist between intensities of the main impurity-induced A and B1 bands and absorbances at 7.8 µm (1282 cm–1) and 8.5 µm (1175 cm–1): αA = 1.2α1282– 0.49α1175, αB1 = 1.2α1175–0.51α1282 (Bokii et al., 1986). Nitrogen concentration was calculated from spectral properties by using the relationships NA (ppm) = 17.5 × αA (Evans, 1992) and, NB1 (ppm) = 43 × αB1 (Sobolev and Lisoivan, 1972).
Mineralogy and geochemistry of diamondiferous metaultramafic rocks Materials and methods Relationships of minerals in diamondiferous rocks were determined by light microscopy in thin sections. Chemical compositions of minerals were determined with a jeol JSM6480LV electron microscope equipped with an INCA Energy
Sample 18/4 had been taken by V.G. Gadiyatov on the left bank of the Tokko River at the boundary between a thin body of metaultramafic rocks of intrusive habit and actinolite–chlorite schists (komatiites and komatiite basalts), as shown in Fig. 3. The mineralogy of the diamondiferous rocks was
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studied in thin sections from fragments 3–5 mm in size. The rock consisted of a fine-crystalline aggregate with olivine, serpentine, talc, and carbonate grains no larger than 200 µm and dust-like chromium magnetite. Typical compositions of these minerals are shown in Table 1. The olivine compositions had large values of XMg = 0.91 and NiO concentrations: 0.27–0.76 wt.%. The contents of magnesium and nickel and relationships with metamorphic minerals point to the primary igneous nature of the olivine. The serpentines had XMg = 0.96, and Cr2O3 contents within 0.92—1.32 wt.%. An admixture of 0.37 wt.% NiO was present in the talc. Judging from chemical composition, the carbonate mineral was ferruginous hydromagnesite (Table 1). Chromium magnetite particles in the rocks were no larger than 20 µm. They had high Cr2O3 contents: 7.74–8.81 wt.%. Rocks with similar mineral associations and chemical compositions (olivine with XMg = 0.88– 0.89 and NiO = 0.36–0.78 wt.%, serpentine with XMg = 0.94–0.95 and NiO = 0–0.25 wt.%, and chromium-magnetite with Cr2O3 = 4.7–8.4 wt.%) occur in the form of olivine– talc–serpentinite lenses in compact metaultramafic rocks of intrusive habit, which build the Taryn-Yuryakh, Krasnaya Gorka, and Tokko blocks (Fig. 3) (Dobretsov et al., 1986; Popov et al., 1990). The chemical composition of the diamondiferous olivine– talc–serpentine rock is shown in Table 2. As evident from its magnesium oxide content and proportions of other oxides (CaO/Al2O3 = 0.27, MgO/FeO = 6.2, and Al2O3/TiO2 = 21.8), it was close to peridotitic komatiites. According to the CPIW norm from dry residue, the olivine content was 63.9%, and hypersthene, 30.3%. The rest included normative albite,
magnetite, ilmenite, and apatite. The total REE content in the sample was 2.9 ppm, that is, within the 1.38–3.54 ppm range typical of intrusive metaultramafic rocks of the Olondo belt (Popov et al., 1990). However, it is apparent from the distribution pattern of REE normalized to chondrite that the diamondiferous olivine–talc–serpentinite rocks differ from intrusive peridotites occurring in Krasnaya Gorka blocks (Fig. 4). The former are characterized by a differentiated distribution of REE with (La/Yb)N = 4.16, whereas in the latter (La/Yb)N varies from 0.63 to 1.3. F.P. Lesnov (2007) has analyzed this unusual enrichment of dunites and harzburgites in LREE and shown that serpentinization is accompanied by subtraction of medium REE from ultramafic restites. With regard to the degree of rock metamorphism and associated mineral transformations (serpentinization, talc development, and appearance of hydromagnesite), this conjecture seems reasonable. Thus, diamondiferous olivine–talc–serpentine rocks result from metamorphic alteration of intrusive ultramafic rocks. The Sm-Nd isochronous age of Krasnaya Gorka peridotites from bulk samples is 3003 ± 117 Ma (εNd(T) = 1.12 ± 0.1) (Pukhtel and Zhuravlev, 1993). With allowance for errors, this result is close to values obtained by various methods for mafic and felsic volcanic rocks building the Olondo greenstone belt. They fall within the range 2950–3005 Ma (Bibikova et al., 1984; Nutman et al., 1992; Pukhtel and Zhuravlev, 1993). In other words, the penetration of metaultramafic rocks of intrusive habit and their diamondiferous varieties occurred about 3.0 Ga BP.
Table 1. Chemical composition of minerals in the diamondiferous olivine–talc–serpentine rock, wt.% Oxide
Olivine
Serpentine
Talc
Carbonate
Chromium magnetite
SiO2
40.9
41.1
44.1
43.0
62.9
63.0
–
–
–
–
Al2O3
–
–
2.28
2.44
–
–
–
–
–
–
Cr2O3
–
–
0.92
1.32
–
–
–
–
7.74
8.81
FeO
8.45
8.30
2.62
2.78
1.19
1.10
8.09
6.40
90.0
89.2
MnO
–
–
–
–
–
–
–
–
0.30
0.30
MgO
49.8
49.6
37.6
36.9
30.1
30.2
36.3
38.3
–
–
NiO
0.76
0.27
–
–
0.37
0.37
–
–
1.57
1.61
Total
99.9
99.3
87.5
86.4
94.6
94.7
44.4
44.7
99.6
99.9
XMg
0.91
0.91
0.96
0.96
0.97
0.98
0.89
0.91
–
–
Note. Analyses were performed with a JEOL JSM–6480LV scanning electron microscope equipped with the INCA Energy microanalysis system (Oxford Instruments). Analyst N.V. Leskova, Diamond and Precious Metal Geology Institute, Yakutsk.
Table 2. Chemical composition (wt.%) and contents of rare earth elements (ppm) in the diamondiferous olivine-talc-serpentine rock SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O P2O5 LOI Total 37.9
0.01
0.24
5.9
0.06
36.6
0.07
0.44
0.15
La
Ce
Pr
Nd
Sm Eu
Gd
18.6 100 (81.4) 0.57 1.19 0.14 0.53 0.12 0.03 0.1
Tb
Dy
Ho
Er
Tm Yb
Lu
0.01 0.07 0.02 0.05 0.01 0.05 0.01
Note. Analyses were performed by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS). Analyst I.V. Nikolaeva, Institute of Geology and Mineralogy, Novosibirsk.
A.P. Smelov et al. / Russian Geology and Geophysics 53 (2012) 1012–1022
Fig. 4. Chondrite-normalized REE distributions (Boyton, 1984) in the Olondo greenstone belt: 1, diamondiferous olivine–talc–serpentine rock; 2, metaultramafic rocks of intrusive habit.
Diamonds from metaultramafic rocks All diamonds isolated from metaultramafic rock sample 18/4 were fragments 0.3–0.5 mm in size (Gadiyatov et al., 2003a,b). Most grains were characterized by triangular fragments of flat faces of octahedra, triangular or foliated–stepped. In addition, characteristic triangular or conchoidal etch patterns were observed on grain surfaces. Morphological study of the diamond collection received by us has revealed crystals of laminar octahedra with transition shapes to variety I dodecahedroids according to Yu.L. Orlov’s (1984) classification, crystals with polycentric faces, and spinel twins (Fig. 5). The grains were transparent and colorless, sometimes with yellowish tint and intense adamantine luster. The photoluminescence of the crystals was diverse: blue, green, yellow, red, or albescent. Its color range was similar to that of diamond crystals from komatiites of Paleoproterozoic greenstone belts (Smith et al., 2010) and lamprophyres of Archean ones (Kopylova et al., 2010a). The appearance of green, yellow, and red luminescence in diamonds is associated with changes of nitrogen aggregation caused by temperature and pressure during metamorphism (Kopylova et al., 2010b).
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Twenty-three diamond crystal fragments were examined by Fourier transform IR spectroscopy. The results indicated that nitrogen impurity in the diamonds studied was present mainly in the aggregated form. The diamonds contained paramagnetic nitrogen in the form of single carbon-substituting atoms in the diamond lattice (C-center) at concentrations not detectable by IR if any. Characteristic absorption bands revealed A and B1 nitrogen defects and tabular structures, B2 defects (Table 3). These data allow the crystals to be assigned to type IaA/B, widespread in the nature. Tabular structures in plane {100} (platelets, B2 defects) were recognized in most crystals from absorption band ~1370 cm–1. This band, with significant amplitude, was accompanied by a smaller one at 1430 cm–1. The location of the main band maximum varied among crystals from 1361 to 1379 cm–1. This variation is likely to be associated with linear sizes of the plates (Sobolev et al., 1968). The integrated intensity of the ~1370 cm–1 band also shows a linear correlation with the total platelet area (Sumida and Lang, 1982). The overall nitrogen content (calculated concentrations in the form of A and B1 defects) in the diamonds studied varied from < 100 to 3800 ppm. The percentage of nitrogen in the form of B1 defects varied from 0 to 94%. Moreover, some IR spectra had broad (peak widths at half heights 200 cm–1 or more) absorption bands with maximums at ~3440 and ~1650 cm–1, attributed to dispersed water impurities (Chrenko et al., 1967). Most of the crystals showed additional absorption bands: 3107, 1405, 3237, and 2785 cm–1. It is known from comprehensive studies of a large number of natural diamonds that bands 3107 and 1405 cm–1 for the type Ia crystals are practically universal in diamonds. They have been proven to belong to C–H bond vibrations (Davies et al., 1984; Woods and Collins, 1983). According to various models, the 3107 and 1405 cm–1 bands are associated with vibrations in vinylidene or ethylene groups (Sobolev et al., 1972; Woods and Collins, 1983) adsorbed on the surface of microscopic inclusions in diamonds. The strongest absorbance recorded at 3107 cm–1 in crystals studied was ~22 cm–1 (sample 18/4-2). It is conjectured that bands 3231 and 2785 cm–1 stem from vibrations in acetylene groups (–C≡CH) (Iakoubovskii and Adriaenssens, 2002; Woods and Collins, 1983). Absorption bands 3107, 1405, 3237, and 2785 cm–1 are widespread in natural diamonds from kimberlites and lamproites, being
Fig. 5. Morphology of diamond crystals from metaultramafic rock: A, fragment of a laminar octahedron; B, twin aggregate of crystals; C, spinel twin of laminar octahedra transitive to dodecahedroids, D, laminar octahedron with polycentric faces.
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A.P. Smelov et al. / Russian Geology and Geophysics 53 (2012) 1012–1022
Table 3. Major impurity centers and additional absorption bands in diamonds according to IR spectroscopy results Sample
Crystal morphology
NA, ppm
NB1, ppm %B
Ntot, ppm
αB2
vB2
α3107
Additional bands, v, cm–1
cm–1 18/4–1
Fragment
22
111
83
133
2
1377
–
2852, 2923, 2954, 1735, 1456, 1375
18/4–2
Fragment
271
3495
93
3766
50
1372
22
1405, 3237, 2785, 1430
18/4–3
Fragment
150
0
0
150
–
–
–
2852, 2923, 2954, 1456, 1735
18/4–4
Fragment
131
0
0
131
–
–
2.5
~3700
18/4–5
Fragment
–
–
–
–
–
–
–
2852, 2923, 1742, 1558, 960, 1066, 798
18/4–6
Fragment
No spectrum, weak transmission
18/4–7
Fragment
30
211
88
241
8.5
1363
14.5
1405, 3660, 1640, 829
18/4–8
Fragment
94
194
67
288
1.5
1362
1.5
~3440, 1640, 1735
18/4–9
Fragment
719
193
21
912
–
–
1.5
2850, 2920
18/4–10
Fragment
76
350
82
426
no data
no data
4
1650, ~3540, 2382, 1735, 2852, 2923, 2954
18/4–11
Fragment
95
271
64
266
3.5
1361
2
2850, 2921, 1735, 880
18/4–13
Fragment
341
506
60
847
13
1363
16
1405, 3237, 2785, 2852, 2923
18/4–14
Fragment
116
319
73
435
8
1363
12
1405, 3237, 2785, 788, 2852, 2923
indicative of C–H (and, probably, N–H) bonds. Some crystals studied may contain carbonates. Results and conclusions The FTIR data on the Archean diamonds studied indicate that their IR spectra are close to those of diamonds from kimberlites and lamproites. Like diamonds from the Abitibi-Wawa greenstone belt in the Superior craton (Stachel et al., 2006), Olondo diamonds broadly vary in the degree of nitrogen aggregation. Their specific feature is that their total nitrogen content can reach 3800 ppm. In contrast to Paleoproterozoic diamonds from lamprophyre dikes in the Nunavut Territory, Canada, Olondo diamonds have a high degree of nitrogen aggregation (Cartigny et al., 2004). Unlike other diamonds from lamprophyres, the diamonds considered here have a lighter carbon isotope ratio: δ13C = –26.0‰. Crystals with such properties have been found only in placers in the northeastern Siberian craton (Shatsky et al., 2011), for which the ore bodies have not been found or proven (Grakhanov et al., 2010). It is reasonable to infer from the nitrogen impurity and isotopic composition of the diamonds that their carbon source was subducted rocks of the Earth’s crust. However, in contrast to metamorphogenic diamonds of the Kokchetav block (Sitnikova and Shatsky, 2009) or diamonds from lamprophyre dikes in Nunavut, Canada (Cartigny et al., 2004), the Olondo diamonds stayed for long in the mantle at a high temperature. The presence of diamonds in rocks of the Olondo greenstone belt points to the existence of a diamondiferous lithospheric mantle beneath the southeastern portion of the North-Asian craton in the Mesoarchean, which is also proven by calculation of PT parameters of the formation of barophilic
minerals in the Middle Paleozoic Manchara kimberlite pipe. It was discovered by geologists of Yakutskgeologiya enterprise in 2007 (Smelov et al., 2009, 2010; Zaitsev et al., 2009). We are grateful to V.G. Gadiyatov, Yu.G. Tyllar†, and V.I. Pavlov for their long-term search for diamond ore bodies in the Olondo greenstone belt and for the collection they provided to us. We also express our gratitude to N.V. Popov and N.N. Dobretsov for detailed geological and petrographical mapping of the belt. This study was supported by the Ministry of Education and Science, State Contract 02.740.11.0328, and RAS Program No. 24.1.
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A.P. Smelov et al. / Russian Geology and Geophysics 53 (2012) 1012–1022 Dobretsov, N.L. (Ed.), 1986. Early Precambrian in Southern Yakutia [in Russian]. Nauka, Moscow. Dobretsov, N.L., Pokhilenko, N.P., 2010. Mineral resources and development in the Russian Arctic. Russian Geology and Geophysics (Geologiya i Geofizika) 51 (1), 98–111 (126–141). Dobretsov, N.L., Dobretsov, N.N., Popov, N.V., Dobretsova, L.V., Smelov, A.P., 1986. Mineralogy and geochemistry of the komatiite series in the Olondo structure of the Aldan shield, in: Geochemistry of Volcanic Rocks of Various Geodynamic Settings [in Russian]. Nauka, Novosibirsk, pp. 34–49. Dobretsov, N.N., Popov, N.V., Smelov, A.P., Bogomolova, L.M., Moskovchenko, N.I., Barton, J.M., 1997. Aldan-Stanovik Shield, in: De Wit, M.J., Ashwal, L.D. (Eds.) Greenstone Belts. Clarendon Press, Oxford, pp. 710–725. Drugova, G.M., Kharitonov, A.L., Bushmin, S.A., Chukhonin, A.P., Shuleshko, I.K., 1983. Structural–metamorphic development of the Olondo Group (East Siberia), in: Neelov, A.N. (Ed.), Precambrian Metamorphism in the Baikal–Amur Mainline Region [in Russian]. Nauka, Leningrad, pp. 49–65. Drugova, G.M., Pukhtel, I.S., Shustova, L.E., Berezhnaya, N.G., 1988. The Olondo Greenstone Belt (Aldan Shield). Izv. AN SSSR. Ser. Geol., No. 8, 40–56. Evans, T., 1992. Aggregation of nitrogen in diamond, in: Field, J.E. (Ed.), The Properties of Natural and Synthetic Diamond. London, pp. 259–290. Gadiyatov, V.G., 2005. Gemstone associations in Northeastern Asia [in Russian]. Izd. Voronezhskogo Universiteta, Voronezh. Gadiyatov, V.G., Marshintsev, V.K., Tyllar, Yu.G., Pavlov, V.I., 2003a. Diamonds in effusive komatiites of the Olondo greenstone belt. Rudy i Metally, No. 2, 47–51. Gadiyatov, V.G., Marshintsev, V.K., Tyllar, Yu.G., Pavlov, V.I., Chaban, N.N., Asoskov, V.M., Tombasov, I.A., 2003b. Diamond potential of komatiites in the Olondo greenstone belt. Otechestvennaya Geologiya, No. 2, 30–31. Grakhanov, S.A., Smelov, A.P., Egorov, K.N., Golubev, Yu.K., 2010. The terrigenous–volcanogenic nature of the bottom of the Carnian Stage, the source of diamonds of the northeastern Siberian craton. Otechestvennaya Geologiya, No. 5, 3–12. Henning, A., Kiviets, G., Kurszlaukis, S., Barton, E., Mayaga-Mikolo, F., 2003. Early Proterozoic metamorphosed kimberlites from Gabon, in: Extended Abstracts of 8th Int. Kimberlite Conf. Victoria, BC, Canada, June 22–27th (CD). Iakoubovskii, K., Adriaenssens, G.J., 2002. Optical characterization of natural Argyle diamonds. Diamond and Related Materials 11, 125–131. Khaidarov, R.A., Chechetkin, V.S., 1990. Prospects of finding new types of high-abrasive materials in the Vitim–Aldan shield, in: Geology and Evaluation of Deposits of Non-Ore Materials in the Baikal–Amur Mainline Region [in Russian]. Nedra, Moscow, pp. 90–94. Kopylova, M., Bruce, L., Ryder, J., 2010a. Diamonds in an Archean greenstone belt: Diamond suites in unconventional rocks of Wawa, Northern Ontario (Canada). Vienna, Austria, 02–07 May. Geophysical Research Abstracts, Vol. 12, EGU2010-6835 (CD). Kopylova, M., Bruce, L., Longo, M., Ryder, J., Dobrzhinetskaya, L., 2010b. Cathodoluminescence of diamond as an indicator of its metamorphic history. Vienna, Austria, 02–07 May. Geophysical Research Abstracts, Vol. 12, EGU2010-6822-1 (CD). Kovach, V.P., Kotov, A.B., Sal’nikova, E.B., Smelov, A.P., 1995. Systematics of highly metamorphic supracrustal assemblages of the Aldan Shield according to Sm–Nd isotopic composition, in: Main Milestones in the Precambrian Evolution of the Earth and Their Isotopic-Geochronological Substantiation. Abstracts [in Russian]. St. Petersburg, p. 31. Lesnov, F.P., 2007. Rare Earth Elements in Ultramafic and Mafic Rocks and Their Minerals. Vol. 1: Major Rock Types. Rock-Forming Minerals [in Russian]. Akademicheskoe Izd. Geo, Novosibirsk. Levchenkov, O.A., Morozova, I.M., Drugova, G.M., Duk, V.L., Levskii, L.K., 1987. Uranium–lead dating of ancient bodies of the Aldan shield, in: Levskii, L.K. (Ed.), Isotope Dating of Metamorphism and Metasomatism Processes [in Russian]. Nauka, Moscow, pp. 116–138. Logvinova, A.M., Virt, R., Tomilenko, A.A., Afanas’ev, V.P., Sobolev, N.V., 2011. The phase composition of crystal-fluid nanoinclusions in alluvial
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diamonds in the northeastern Siberian Platform. Russian Geology and Geophysics (Geologiya i Geofizika) 52 (11), 1286–1297 (1634–1648). Mironyuk, E.P., Lyubimov, V.K., Magnushevskii, E.L., 1971. Geology of the Western Part of the Aldan Shield [in Russian]. Nedra, Moscow. Neymark, L.A., Kovach, V.P., Nemchin, A.A., Morozova, I.M., Kotov, A.B., Vinogradov, D.P., Gorokhovsky, B.M., Ovchinnikova, G.V., Bogomolova, L.M., Smelov, A.P., 1993. Late Archaean intrusive complexes in Olekma granite-greenstone terrain (Eastern Siberia): geochemical and isotopic study. Precambrian Res. 62, 453–472. Nutman, A.P., Chernyshev, I.V., Baadsgaard, H., Smelov, A.P., 1992. The Aldan Shield of Siberia USSR: the age of its Archean components and evidence for widespread reworking in the mid-Proterozoic. Precambrian Res. 54, 195–209. Orlov, Yu.L., 1984. Diamond Mineralogy [in Russian]. Nauka, Moscow. Popov, N.V., Smelov, A.P., Dobretsov, N.N., Bogomolova, L.M., Kartavchenko, V.G., 1990. The Olondo Greenstone Belt [in Russian]. Izd. Yakutskogo Nauchnogo Tsentra SO AN SSSR, Yakutsk. Popov, N.V., Dobretsov, N.N., Smelov, A.P., Bogomolova, L.M., 1995. Tectonics, metamorphism, and the problems of evolution of the Olondo greenstone belt. Petrology 3, 73–86. Pukhtel, I.S., Zhuravlev, D.Z., 1993. Petrology of mafic and ultramafic rocks and rocks associated with them in the Olondo greenstone belt, Aldan shield. Petrologiya 1 (3), 306–344. Reutskii, V.N., Logvinova, A.M., Sobolev, N.V., 1999. Carbon isotope ratio in polycrystalline diamond aggregates containing chromite inclusions in the Mir kimberlite pipe, Yakutia. Geokhimiya, No. 11, 1191–1196. Shatsky, V.S., Zedgenizov, D.A., Ragozin, A.L., Kalinina, V.V., Reutskii, V.N., 2011. Local variations in carbon isotopes and nitrogen contents in diamonds from placers of the northeastern portion of the Siberian Platform. Dokl. Earth Sci. 440, 1282–1286. Sitnikova, E.S., Shatsky, V.S., 2009. New FTIR spectroscopy data on the composition of the medium of diamond crystallization in metamorphic rocks of the Kokchetav Massif. Russian Geology and Geophysics (Geologiya i Geofizika) 50 (10), 842–849 (1095–1103). Skuzovatov, S.Yu., Zedgenizov, D.A., Shatsky, V.S., Ragozin, A.L., Kuper, K.E., 2011. Composition of cloudy microinclusions in octahedral diamonds from the Internatsional’naya kimberlite pipe (Yakutia). Russian Geology and Geophysics (Geologiya i Geofizika) 52 (1), 85–96 (107–121). Smelov, A.P., 1989. Metamorphic Evolution of the Olekma Granite–Greenstone Region [in Russian]. Nauka, Novosibirsk. Smelov, A.P., 1996. Archean and Proterozoic Metamorphism in the Aldan– Stanovoy Shield. Doctoral (Geol. & Mineral.) Dissertation. Institute of Geology and Geophysics, Novosibirsk. Smelov, A.P., Timofeev, V.F., 2007. The age of the North Asian Cratonic basement: an overview. Gondwana Res. 12, 279–288. Smelov, A.P., Zedgenizov, A.N., Timofeev, V.F., 2001. Aldan–Stanovoy Shield, in: Tectonics, Geodynamics, and Metallogeny in the Sakha Republic (Yakutia) [in Russian]. MAIK Nauka/Interperiodica, Moscow, pp. 81–104. Smelov, A.P., Ashchepkov, I.V., Oleinikova, O.B., Surnin, A.A., Babushkina, S.A., Polufuntikova, L.I., Koroleva, O.V., 2009. Chemical composition and P–T conditions of the formation of barophilic minerals from the Manchara kimberlite pipe (Central Yakutia). Otechestvennaya Geologiya, No. 5, 27–30. Smelov, A.P., Andreev, A.P., Altukhova, Z.A., Babushkina, S.A., Bekrenev, K.A., Zaitsev, A.I., Izbekov, E.D., Koroleva, O.V., Mishnina, V.M., Okrugin, A.V., Oleinikov, O.B., Surnin, A.A., 2010. Kimberlites of the Manchary pipe: a new kimberlite field in Central Yakutia. Russian Geology and Geophysics (Geologiya i Geofizika) 51 (1), 121–126 (153–159). Smith, C., Bulanova, G., Ushkov, V., Gobbo, L., 2010. A tale of two diamond hosts—Kimozero and Dachine. Vienna, Austria, 02–07 May. Geophysical Research Abstracts, Vol. 12, EGU2010-6486-1 (CD). Sobolev, E.V., Lisoivan, V.I., 1972. On the nature of properties of intermediate-type diamonds. Dokl. Akad. Nauk 204, 88–90. Sobolev, E.V., Lisoivan, V.I., Lenskaya, S.V., 1968. Tabular structures in natural diamonds. Zh. Strukturnoi Khimii 9 (6), 1029–1033. Stachel, T., Banas, A., Muehlenbachs, K., Kurszlaukis, S., Walker, E.C., 2006. Archean diamonds from Wawa (Canada): samples from deep cratonic
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Editorial responsibility: N.P. Pokhilenko
Russian Geology and Geophysics 53 (2012) 1012–1022 www.elsevier.com/locate/rgg
Diamondiferous Archean rocks of the Olondo greenstone belt (western Aldan–Stanovoy shield) A.P. Smelov a,*, V.S. Shatsky b, A.L. Ragozin b, V.N. Reutskii b, A.E. Molotkov a a
Diamond and Precious Metal Geology Institute, Siberian Branch of the Russian Academy of Sciences, ul. Lenina 39, Yakutsk, 677980, Russia b V.S. Sobolev Institute of Geology and Mineralogy, Siberian Branch of the Russian Academy of Sciences, pr. Akademika Koptyuga 3, Novosibirsk, 630090, Russia Received 31 October 2011; accepted 23 March 2012
Abstract Diamond from metaultramafic rocks of the Mesoarchean (2.96–3.0 Ga) Olondo greenstone belt, located in the western Aldan–Stanovoy shield, has been studied. Diamonds occur in lenses of olivine–serpentine–talc rocks within metaultramafic rocks of intrusive habit, whose composition corresponds to peridotite komatiites. All diamonds from the metaultramafic rocks are crystal fragments 0.3 to 0.5 mm in size. Morphological examination has revealed laminar octahedra, their transitional forms to dodecahedroids, crystals with polycentric faces, and spinel twins. The crystals vary in photoluminescence color: dark blue, green, yellow, red, or albescent. Characteristic absorption bands in crystals point to nitrogen impurities in the form of A and B1 defects and tabular B2 defects. The crystals studied belong to the IaA/B type, common among natural diamonds. The overall nitrogen content varies from <100 to 3800 ppm. The relative content of nitrogen in B1 centers varies from 0 to 94%, pointing to long stay in the mantle. The carbon isotope ratio in the diamonds, 13C = –26‰, is indicative of involvement of subducted crust matter in diamond formation in the Archean. © 2012, V.S. Sobolev IGM, Siberian Branch of the RAS. Published by Elsevier B.V. All rights reserved. Keywords: diamond; Archean; komatiites; greenstone belt; Aldan–Stanovoy shield
Introduction The mining of primary diamond deposits and placers in the Siberian craton yields 99% of the overall diamond winning in Russia (Dobretsov and Pokhilenko, 2010). Diamonds from these deposits have been studied by a wide range of up-to-date methods (Logvinova et al., 2011; Skuzovatov et al., 2011; Zedgenizov et al., 2011). However, it is expected that the exhaustion of the largest deposits, Mir and Udachnaya kimberlite pipes, will cause a decline in diamond output. Therefore, search for new diamond deposits is an urgent task (Dobretsov and Pokhilenko, 2010). As reported in (Khaidarov and Chechetkin, 1990), R.A. Khaidarov (Central Research Institute of the Geology of Non-Ore Minerals, Kazan, Russia) found two diamond crystals, 0.1 and 0.8 mm, in chip samples of metaultramafic rocks sampled in the Olondo Mesoarchean greenstone belt (Drugova et al., 1983, 1988; Popov et al., 1990, 1995; Dobretsov et al., 1997). In 1987–1989, a test survey crew of the Siberian
* Corresponding author. E-mail address: [email protected] (A.P. Smelov)
Research Institute of Geology, Geophysics and Mineral Resources, working under the auspices of the All-Russia Lamproite Program, found diamonds in alluvia of the Olondo and Taryn-Yuryakh Rivers, draining volcanic deposits of the belt: 3 grains 0.5–1.0 mm in size; 31 grains 0.2–0.5 mm; 1 crystal fragment 1.1 × 1.1 mm; and 1 rhombododecahedroid crystal 1.1 × 1.0 mm. At the same time, V.V. Levitskii took stream sediment samples from the Olondo River and its tributaries. These samples were bulked after washing and treated at the Irkutsk Institute of Rare and Precious Metals and Diamonds (Irkutsk, Russia). They yielded 18 diamond crystals. Later, the Olondo team of the Udokan mining enterprise conducted prospecting for diamonds in the catchment areas of the Olondo and Taryn-Yuryakh Rivers. Eight diamonds were found near the southwestern shore of Lake Tokko: two colorless transparent octahedroid crystals 0.7 × 0.8 mm in size and six diamond fragments (fraction –0.5 to +0.525 mm) with characteristic pale greenish-yellow color. Eleven diamonds of class 0.25 mm and two crystals 0.8 mm were found on the southern slope of the Krasnaya Gorka metaultramafic block in the upper reaches of Olondo. One of the larger diamonds was a flat-faced octahedron and the other, an irregularly
1068-7971/$ - see front matter D 201 2, V . S. S o bolev IGM, Siberian Branch of the RAS. Published by Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.rgg.2012.0 + 8.005
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Fig. 1. Map of terranes of the Aldan–Stanovoy shield according to (Smelov and Timofeev, 2007; Smelov et al., 2001), as subsequently amended: 1, deposits of the Siberian craton cover; 2–5, terranes: 2, granite–greenstone (WA—Western Aldan—AR, EBT—Batomga—Pt1), 3, tonalite–trondhjemite gneiss (TN—Tynda—ARPt1), 4, granulite–orthogneiss (ANM—Nimnyr—AR-Pt1, CG—Chogar—AR), 5, granulite–paragneiss (AST—Sutam—AR Archean, EUC—Uchur—Pt1); 6, tectonic mélange zones (am, Amga; kl, Kalar; tr, Tyrkanda); 7, faults (dj, Dzheltulak; ts, Taksakanda); 8, thrusts. The Batomga (EBT) and Uchur (EUC) terranes form the Eastern Aldan superterrane, and the Nimnyr (ANM) and Sutam (AST) ones form the Central Aldan superterrane.
shaped fragment. The smaller diamonds were transparent pale greenish-yellow irregularly shaped diamond fragments with rare preserved octahedron faces. A small sample taken from the talus of a small ultramafic block on the right bank of the Skvoznoi Creek (large right tributary of Taryn-Yuryakh) contained four pale greenish-yellow octahedral diamond crystals up to 0.25 mm in size (Yurgenson et al., 1999). Thus, about 80 diamond crystals were found within five years of prospecting works in the alluvial deposits. V.G. Gadiyatov inferred from analysis of available prospecting data and comparison of the geologic structure of the Olondo greenstone belt (Dobretsov et al., 1997; Drugova et al., 1983; Popov et al., 1990, 1995) with Dachine diamondiferous volcanic rocks in French Guiana (Capdevila et al., 1999) that komatiitic tuffs of the belt, dated to 2.96–3.01 Ga, were the diamond ore body. In 2002, he took samples from volcanic rocks of different compositions. Yu.G. Tyllar investigated rock samples from upper reaches of the Tokko River and found over 100 diamond grains in a single 150-g sample recognized in the field as komatiitic tuff (Gadiyatov, 2005; Gadiyatov et al., 2003a,b). Nearly all of them were irregular fragments 0.3–0.5 mm in size. Most grains were used up in diagnostics and external expertise. The physical parameters of the diamonds, as well as the mineralogical and petrochemical compositions of the host rocks remained unexplored. The very fact of the presence of ancient diamonds in rocks of the Aldan–Stanovoy shield is of great importance for understanding diamond formation and geodynamics in the early history of the Earth, because such findings are rare. The
oldest diamonds (3.0–2.8 Ga) are found in gold ore conglomerates of the Witwatersrand basin, South Africa. Relatively younger diamonds have been recorded in lamprophyres of the Abitibi–Wawa greenstone belt, Superior Craton (2.7 Ga; Stachel et al., 2006) and in metamorphosed kimberlites of Gabon, Central Africa (2.85 Ga; Henning et al., 2003). V.G. Gadiyatov kindly provided 25 diamond grains and crushed host rock (fraction 3–5 mm) from the Olondo greenstone belt for our study. Here we report the results of their investigation and analysis of the geologic location of the diamondiferous rocks.
Geology of the Olondo greenstone belt The Aldan–Stanovoy shield is located on the southern margin of the Siberian craton. It is built mainly by deeply metamorphosed rocks of the granulite facies and, to a lesser extent, by rocks of the amphibolite and greenschist facies. As shown in Fig. 1, five terranes are recognized in the shield on the base of rock composition, metamorphism degree, and age: the Western Aldan granite–greenstone composite terrane, Central Aldan and Eastern Aldan superterranes in the north; Tynda tonalite–trondhjemite-gneiss composite terrane and Chogar granulite–orthogneiss terrane (Smelov et al., 2001) in the south. In a collision, which occurred 1.9 Ga BP, the composite terranes and superterranes formed the base of the North Asian craton (Smelov and Timofeev, 2007). The Western Aldan granite–greenstone terrane is the oldest com-
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Fig. 2. The center and east of the Western Aldan granite–greenstone terrane (Smelov, 1996). 1, Neoproterozoic–Cenozoic deposits of the platform cover and depressions of the Mesozoic and Cenozoic activity stages; 2, Paleoproterozoic metamorphosed subplatform deposits, the Udokan complex (graben–synclines: LKh, Lower Khani; Ol, Oldongso; Ug, Ugui); 3, outcrops of greenstone belt rocks (Subgan greenstone complex): I, Itchilyak fragment; Ya, Yaelakh fragment; Old, Olondo structure; Tm, Tasmiela structure; Sl, Syrylyr structure; Tn, Tungurcha structure; 4, tonalite–trondhjemite gneisses and migmatites of the Olekma complex; 5, granulite (Kurulta) complex (a, paragneisses; b, orthogneisses); 6, metagabbro–diorite–tonalite–trondhjemite complex (Amn, Amnunnakta block; Tu, Tungurchakan block); 7, tonalites and trondhjemites (UO, Ust’-Oldongso block); 8, granites and bastard granites, poorly defined; 9, geologic boundaries (a); faults with steep displacement surfaces: A, Amga; Kh, Khani; Ch, Chara (b); c, thrusts: S, South-Chulman thrust.
posite one, which was not reworked at granulite facies conditions in the course of the Paleoproterozoic collision. The Western Aldan composite terrane is 400 × 350 km in size. It borders the Baikal–Patom fold-and-thrust belt in the west. In the east and south, it borders the Amga and Kalar tectonic mélange zones, respectively. In the north, the terrane is overlain by Upper Riphean and Vendian deposits of the Siberian Craton sedimentary cover (Fig. 2). The terrane is built by Archean bodies of various types, which were metamorphosed in broad temperature and pressure ranges (Mironyuk et al., 1971; Smelov, 1989). Tonalite–trondhjemite orthogneisses are predominant there. They are bulked into the Olekma complex. They form several large linear blocks separated by four meridional belts 300 km in length and up to 30 km in width, where tectonic flakes of greenstone rocks of the Subgan complex are concentrated. The greenstone belts are typically confined by blastomylonites. The Kurulta granu-
lite complex forms several blocks separated by faults and tectonic flakes (Fig. 2). All these blocks and zones can be interpreted as individual terranes. Therefore, the Western Aldan terrane is determined as a composite one, consisting of several terranes. Their main feature is the presence of Archean greenstone bodies and tonalite–trondhjemite orthogneisses. The Olekma tonalite–trondhjemite complex includes compositionally uniform biotite, biotite–amphibole, and amphibole plagiogneisses and gneisses. They have Na : K ratios within 2.5–5.0 and belong mainly to the high-alumina type. They are enriched in LREE and Sr and depauperated in U (Dobretsov, 1986). The associated mafic schists and amphibolites constitute up to 10% of the complex (Cherkasov, 1979). The age of the tonalite–trondhjemite gneisses is 3.2–2.8 Ga (Baadsgaard et al., 1990; Neymark et al., 1993; Nutman et al., 1992). The progressive metamorphism of the rocks corresponds to conditions of the moderate pressure
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amphibolite facies (T = 650–700 °C, P = 5.5–7.0 kbar) or, less frequently, the epidote–amphibolite ones. Nearby the greenstone belts, orthogneisses have been reworked under the conditions of the epidote–amphibolite facies (T = 475– 535 °C, P = 5.0–4.0 kbar) to form heteroblastic, glomeroblastic, and blastomortar textures (Smelov, 1989). The Kurulta granulite complex, which builds up the Olomokit, Chara, and Kalar blocks, consists mainly of orthorocks: enderbite and charnokite gneisses. There are also paragneiss bodies, which include amphibole- and pyroxenecontaining gneisses and schists, garnet–biotitic gneisses, sometimes with cordierite and sillimanite. The blocks themselves are relatively thin tectonic flakes thrust over rocks of the trondhjemite–gneiss and greenstone complexes (Stognii et al., 1996). Little is known about the age of the rocks. It has been found that rocks with preliminary Nd age 3.5 Ga gave rise to protoliths of the garnet–biotite gneisses of the Olomokit block (Kovach et al., 1995), and protoliths of two-pyroxene crystalline schists formed 3.15 Ga BP (Levchenkov et al., 1987). The Subgan greenstone complex was recognized by E.P. Mironyuk and V.S. Fedorovskii in 1963–1968. Bodies of greenstone deposits vary in section patterns and metamorphism conditions (Smelov, 1989). The formation of the sedimentary and volcanogenic rocks of the greenstone belts is dated to two time spans: 3.2–3.0 and 3.0–2.7 Ga BP (Smelov et al., 2001). These time spans were separated by a stage of thrusting and metamorphism of young greenstone belts about 3.0 Ga BP (Nutman et al., 1992; Smelov, 1996). Stratigraphic relationships between the greenstone belts and the neighboring rocks of the tonalite–trondhjemite complex are unknown. The contacts are mainly tectonic or intrusive. The largest volumes of mafic and ultramafic volcanic rocks in combination with felsic and intermediate ones occur in the Tokko–Khani greenstone belt, which is located in the middle of a composite terrane (Fig. 2). The Chara–Tokko and Temulyakit–Tungurcha greenstone belts, located west and east of Tokko–Khani, are characterized mainly by tholeiitic volcanism, notable amounts of ferruginous quartzites, and appearance of carbonate–terrigenous rock assemblages (Temulyakit–Tungurcha and Saimagan) (Smelov, 1989). The southern portion of the Tokko–Khani belt has been best studied. It is referred to in the literature as the Olondo greenstone belt. The Olondo greenstone belt is a V-shaped body of submeridional strike in plan (Fig. 3). It is built by mafic varieties of volcanogenic and volcanosedimentary deposits, characteristic of Archean greenstone belts (Popov et al., 1990). Its section includes three metavolcanic rock associations differing in rock composition and spatial location. Comprehensive lithostratigraphical studies were insufficient for proper understanding the succession of primary superposition of rocks, which had experienced intense metamorphism under the conditions of the epidote–amphibolite facies of andalusite–sillimanite type (T = 550–570 °C and P = 3.5–4.5 kbar) and repeated deformations. Moreover, the belt contains numerous intrusion bodies of various ages and compositions, which also complicate the reconstruction of the stratigraphy of vol-
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canosedimentary strata and their facies transitions (Popov et al., 1995). The first rock association occurs in the margin of the V-shaped body (Fig. 3). It is characterized by alternating mafic and ultramafic metavolcanic rocks. Rocks of same lithological groups differ in geochemistry, depending on their spatial location. In the eastern part of the body, the planar structures and boundaries of rock bodies generally dip to the east. The following sequence is observed from the axis to margins of the body: intermediate and felsic metavolcanic rocks; metatholeiites; ultramafic metavolcanic rocks (chlorite–actinolite schists with interbeds of anthophyllite–carbonate rocks); and metatholeiites with interbeds of komatiites in the form of anthophyllite–carbonate rocks and intermediate and felsic volcanic rocks, as well as ferruginous quartzites and metapelites. The thickness of the eastern branch of the body is estimated to be about 400 m with regard to large folds notable in the map scale. The Sm–Nd isochronous age was established from bulk samples of ultramafic and mafic volcanic rocks to be in the range 2973 ± 48 to 2959 ± 17 Ma (εNd(T) = 2.2 ± 0.1) (Pukhtel and Zhuravlev, 1993), and the U–Pb age (SHRIMP) of zircons from interbeds of intermediate and felsic volcanic rocks, 2986 ± 6 to 3005 ± 5 Ma (Baadsgaard et al., 1990; Nutman et al., 1992). In the western branch of the body, the first association is dominated by amphibole–plagioclase schists (metatholeiites) with frequent thin interbeds of metavolcanic rocks of intermediate composition. Ultramafic metavolcanic rocks are absent. The thickness of this part of the body without regard to folding is 250–300 m. The second rock association builds the central part of the belt. It consists of amphibole and biotite–amphibole microgneisses of intermediate and felsic compositions. Their effusive origin is proven by well-preserved remains of primary structures and textures. Typically, the proportion of felsic rocks increases towards the eastern part of the body. The microgneisses host rare thin (5–10 m) interbeds of amphibole schists (metabasalts), sulfidized garnet amphibolites, and biotite–garnet (sometimes with staurolite) microgneisses. The association includes rocks formed from lavas, tuffs, and tuffaceous sandstones, as well as vein stones. The thickness of the body in the eastern part of the belt is estimated to be about 500 m. The U–Pb isochronous age of zircons from metadacites of this association is 2950 ± 50 Ma (Bibikova et al., 1984). The third rock association is confined to the axial part of the eastern branch of the belt. It consists of frequently alternating interbeds of metatholeiites, metatuffs, metasandstones, and metapelites. Less frequent are metaandesites, which presently include amphibole, amphibole–biotite, and biotite schists; microgneisses; and high-silica rocks. The metatuffaceous sandstones and metasandstones display a clear rhythmicity, emphasized by interbeds with carbonate cement. Rocks with well-preserved structures of tuff breccias containing volcanic rocks of various compositions and agglomerate
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Fig. 3. Geological outline of the Olondo greenstone belt according to (Popov et al., 1995). 1, Quaternary fluvioglacial deposits; 2, bifeldspathic granites; 3, terrigenous deposits of the Udokan complex; 4, amphibole, biotite–amphibole, and biotite microgneisses (metaandesites and metadacites, their tuffs, and metasandstones); 5, variegated deposits: intercalation of microgneisses and amphibole–plagioclase schists; 6, amphibole–plagioclase schists (metatholeiites); 7, actinolite–chlorite schists (komatiites and komatiite basalts); 8, carbonate–anthophyllite schists (komatiites); 9, gabbro–amphibolites and hornblendites; 10, olivinites, serpentinites, and talc rocks (blocks indicated with encircled numerals: 1, Krasnaya Gorka; 2, Tokko; 3, Taryn-Yuryakh); 11, differentiated bodies of gabbro–diorite–tonalite composition; 12, trondhjemites; 13, tonalite–trondhjemite gneisses and migmatites; 14, picrite dikes; 15, geologic boundaries; 16, (a) proven faults, (b) conjectured faults; 17, site of diamondiferous sample 18/4.
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tuffs occur in some exposures. The thickness of the body is tentatively estimated to be 200 m. Intrusions constitute no less than 30% of the volume of the Olondo greenstone belt. Ultramafic, mafic, and felsic rocks are recognized there. They belong to different age stages. The oldest intrusions include numerous bodies of ultramafic rocks of intrusive habit, sill-like bodies of gabbro–amphibolites, and differentiated blocks of gabbro–diorite–tonalites. The ultramafic rocks are metamorphosed dunites and peridotites. They have experienced profound serpentinization and, in places, secondary talc development and carbonatization. Secondary rock transformation destroyed the initial structures, textures, and minerals almost everywhere. The Sm–Nd isochronous age in bulk samples of peridotites of the Krasnaya Gorka block is 3003 ± 117 Ma (εNd(T) = 1.12 ± 0.1) (Pukhtel and Zhuravlev, 1993). Gabbro–amphibolites are predominant in the belt. The largest bodies are differentiated from actinolite–chlorite ultramafic rocks and monomineral amphibolites (hornblendites) at the bottom to mesocratic gabbro–amphibolites at the roof. These bodies are sill-shaped, 50 to 350 m in thickness. They are confined to boundaries between rock types. The rocks differ from metabasalts in being more massive in structure and containing remains of igneous textures. They are sheeted along contacts and within zones. A complex of differentiated gabbro–diorite–tonalite bodies was reliably recognized in the central part of the belt in the branching region (Fig. 3). Intrusive associations are observed between rocks of the complex contrasting in composition. The U-Pb SHRIMP age of zircons from diorites of the complex is 3018 ± 10 Ma (Baadsgaard et al., 1990; Nutman et al., 1992). The youngest intrusions are bodies of granitoids of tonalite– trondhjemite composition, widespread in the margin of the Olondo belt and forming large blocks in the center. The tonalites and trondhjemites are represented by amphibole and biotite varieties. They are sometimes migmatized. They contain inclusions of all types of volcanic and metaultramafic rocks and gabbro–amphibolites, varying in size. The U–Pb SHRIMP age of zircons from tonalites of the eastern contact of the belt is 2862 ± 14 Ma (Baadsgaard et al., 1990; Nutman et al., 1992). Dike-like bodies of ultramafic and mafic rocks, assigned to the picrite series (Dobretsov et al., 1986), vary in thickness from tens of centimeters to meters. They are confined to borders between rock associations and contacts with early intrusions. They are boudined and sheeted along contacts. The Sm–Nd isochronous age from bulk samples of magnesium-rich picrites is 2202 ± 41 Ma (εNd(T) = 2.2 ± 0.2) (Pukhtel and Zhuravlev, 1993).
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350 microanalysis system (Oxford Instruments) at accelerating voltage 20 kV and current 1 nA. The accuracies of determination of elements (wt.%) were: Si Kα, 0.51; Al Kα, 0.13; Cr Kα, 0.18; Fe Kα, 0.28; Mn Kα, 0.14; Mg Kα, 0.42; Ni Kα, 0.37; O, 0.69. All analyses were done at the Diamond and Precious Metal Geology Institute, Yakutsk, by N.V. Leskova. The analyses of rock-forming, rare, and rare-earth elements in the rocks were done by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) by I.V. Nikolaeva, Institute of Geology and Mineralogy, Novosibirsk. Carbon isotope ratios in diamonds were analyzed at the Analytical Center, Institute of Geology and Mineralogy, Novosibirsk. As the diamonds were small, their carbon isotope ratios were determined in a 0.3-mg weight of five fragments. The weight was loaded into a platinum capsule and placed into a quartz reactor with purified copper oxide. The weight was combusted at 950 °C. The resulting carbon dioxide was purified and collected into a dismountable vacuum trap. Diamonds were prepared for isotopic studies as described in (Reutskii et al., 1999). Carbon isotope ratios were determined in carbon dioxide with a Finnigan MAT Delta mass spectrometer in the dual inlet mode. To ensure the accuracy of δ13C determination, it was done against the international USGS-24 standard (graphite, δ13C = –15.9). The reproducibility of the data with reference to the standard, including sample preparation, did not exceed 0.1 ‰ (2σ). As the amount of the material was limited, the result was obtained in a single measurement: δ13C = –26.0 ‰ PDB. Infrared absorbance was measured with an IFS 130v Fourier transform IR spectrometer (Bruker). Spectra were recorded within the wavelength range 2.5–12.5 µm (wavenumbers 4000–800 cm–1), beam aperture from 30 × 30 to 80 × 80 µm. The absorbances at the maximum of the two-phonon lattice absorption (2030 cm–1) were used as an internal standard (I2030 = 12.8 cm–1) to determine the net thickness of the sample. The interpretation of the IR spectra involved attribution of characteristic bands to impurity defects and quantitative analysis: determination of concentrations of nitrogen atoms belonging to various optically active centers. The following relationships exist between intensities of the main impurity-induced A and B1 bands and absorbances at 7.8 µm (1282 cm–1) and 8.5 µm (1175 cm–1): αA = 1.2α1282– 0.49α1175, αB1 = 1.2α1175–0.51α1282 (Bokii et al., 1986). Nitrogen concentration was calculated from spectral properties by using the relationships NA (ppm) = 17.5 × αA (Evans, 1992) and, NB1 (ppm) = 43 × αB1 (Sobolev and Lisoivan, 1972).
Mineralogy and geochemistry of diamondiferous metaultramafic rocks Materials and methods Relationships of minerals in diamondiferous rocks were determined by light microscopy in thin sections. Chemical compositions of minerals were determined with a jeol JSM6480LV electron microscope equipped with an INCA Energy
Sample 18/4 had been taken by V.G. Gadiyatov on the left bank of the Tokko River at the boundary between a thin body of metaultramafic rocks of intrusive habit and actinolite–chlorite schists (komatiites and komatiite basalts), as shown in Fig. 3. The mineralogy of the diamondiferous rocks was
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studied in thin sections from fragments 3–5 mm in size. The rock consisted of a fine-crystalline aggregate with olivine, serpentine, talc, and carbonate grains no larger than 200 µm and dust-like chromium magnetite. Typical compositions of these minerals are shown in Table 1. The olivine compositions had large values of XMg = 0.91 and NiO concentrations: 0.27–0.76 wt.%. The contents of magnesium and nickel and relationships with metamorphic minerals point to the primary igneous nature of the olivine. The serpentines had XMg = 0.96, and Cr2O3 contents within 0.92—1.32 wt.%. An admixture of 0.37 wt.% NiO was present in the talc. Judging from chemical composition, the carbonate mineral was ferruginous hydromagnesite (Table 1). Chromium magnetite particles in the rocks were no larger than 20 µm. They had high Cr2O3 contents: 7.74–8.81 wt.%. Rocks with similar mineral associations and chemical compositions (olivine with XMg = 0.88– 0.89 and NiO = 0.36–0.78 wt.%, serpentine with XMg = 0.94–0.95 and NiO = 0–0.25 wt.%, and chromium-magnetite with Cr2O3 = 4.7–8.4 wt.%) occur in the form of olivine– talc–serpentinite lenses in compact metaultramafic rocks of intrusive habit, which build the Taryn-Yuryakh, Krasnaya Gorka, and Tokko blocks (Fig. 3) (Dobretsov et al., 1986; Popov et al., 1990). The chemical composition of the diamondiferous olivine– talc–serpentine rock is shown in Table 2. As evident from its magnesium oxide content and proportions of other oxides (CaO/Al2O3 = 0.27, MgO/FeO = 6.2, and Al2O3/TiO2 = 21.8), it was close to peridotitic komatiites. According to the CPIW norm from dry residue, the olivine content was 63.9%, and hypersthene, 30.3%. The rest included normative albite,
magnetite, ilmenite, and apatite. The total REE content in the sample was 2.9 ppm, that is, within the 1.38–3.54 ppm range typical of intrusive metaultramafic rocks of the Olondo belt (Popov et al., 1990). However, it is apparent from the distribution pattern of REE normalized to chondrite that the diamondiferous olivine–talc–serpentinite rocks differ from intrusive peridotites occurring in Krasnaya Gorka blocks (Fig. 4). The former are characterized by a differentiated distribution of REE with (La/Yb)N = 4.16, whereas in the latter (La/Yb)N varies from 0.63 to 1.3. F.P. Lesnov (2007) has analyzed this unusual enrichment of dunites and harzburgites in LREE and shown that serpentinization is accompanied by subtraction of medium REE from ultramafic restites. With regard to the degree of rock metamorphism and associated mineral transformations (serpentinization, talc development, and appearance of hydromagnesite), this conjecture seems reasonable. Thus, diamondiferous olivine–talc–serpentine rocks result from metamorphic alteration of intrusive ultramafic rocks. The Sm-Nd isochronous age of Krasnaya Gorka peridotites from bulk samples is 3003 ± 117 Ma (εNd(T) = 1.12 ± 0.1) (Pukhtel and Zhuravlev, 1993). With allowance for errors, this result is close to values obtained by various methods for mafic and felsic volcanic rocks building the Olondo greenstone belt. They fall within the range 2950–3005 Ma (Bibikova et al., 1984; Nutman et al., 1992; Pukhtel and Zhuravlev, 1993). In other words, the penetration of metaultramafic rocks of intrusive habit and their diamondiferous varieties occurred about 3.0 Ga BP.
Table 1. Chemical composition of minerals in the diamondiferous olivine–talc–serpentine rock, wt.% Oxide
Olivine
Serpentine
Talc
Carbonate
Chromium magnetite
SiO2
40.9
41.1
44.1
43.0
62.9
63.0
–
–
–
–
Al2O3
–
–
2.28
2.44
–
–
–
–
–
–
Cr2O3
–
–
0.92
1.32
–
–
–
–
7.74
8.81
FeO
8.45
8.30
2.62
2.78
1.19
1.10
8.09
6.40
90.0
89.2
MnO
–
–
–
–
–
–
–
–
0.30
0.30
MgO
49.8
49.6
37.6
36.9
30.1
30.2
36.3
38.3
–
–
NiO
0.76
0.27
–
–
0.37
0.37
–
–
1.57
1.61
Total
99.9
99.3
87.5
86.4
94.6
94.7
44.4
44.7
99.6
99.9
XMg
0.91
0.91
0.96
0.96
0.97
0.98
0.89
0.91
–
–
Note. Analyses were performed with a JEOL JSM–6480LV scanning electron microscope equipped with the INCA Energy microanalysis system (Oxford Instruments). Analyst N.V. Leskova, Diamond and Precious Metal Geology Institute, Yakutsk.
Table 2. Chemical composition (wt.%) and contents of rare earth elements (ppm) in the diamondiferous olivine-talc-serpentine rock SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O P2O5 LOI Total 37.9
0.01
0.24
5.9
0.06
36.6
0.07
0.44
0.15
La
Ce
Pr
Nd
Sm Eu
Gd
18.6 100 (81.4) 0.57 1.19 0.14 0.53 0.12 0.03 0.1
Tb
Dy
Ho
Er
Tm Yb
Lu
0.01 0.07 0.02 0.05 0.01 0.05 0.01
Note. Analyses were performed by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS). Analyst I.V. Nikolaeva, Institute of Geology and Mineralogy, Novosibirsk.
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Fig. 4. Chondrite-normalized REE distributions (Boyton, 1984) in the Olondo greenstone belt: 1, diamondiferous olivine–talc–serpentine rock; 2, metaultramafic rocks of intrusive habit.
Diamonds from metaultramafic rocks All diamonds isolated from metaultramafic rock sample 18/4 were fragments 0.3–0.5 mm in size (Gadiyatov et al., 2003a,b). Most grains were characterized by triangular fragments of flat faces of octahedra, triangular or foliated–stepped. In addition, characteristic triangular or conchoidal etch patterns were observed on grain surfaces. Morphological study of the diamond collection received by us has revealed crystals of laminar octahedra with transition shapes to variety I dodecahedroids according to Yu.L. Orlov’s (1984) classification, crystals with polycentric faces, and spinel twins (Fig. 5). The grains were transparent and colorless, sometimes with yellowish tint and intense adamantine luster. The photoluminescence of the crystals was diverse: blue, green, yellow, red, or albescent. Its color range was similar to that of diamond crystals from komatiites of Paleoproterozoic greenstone belts (Smith et al., 2010) and lamprophyres of Archean ones (Kopylova et al., 2010a). The appearance of green, yellow, and red luminescence in diamonds is associated with changes of nitrogen aggregation caused by temperature and pressure during metamorphism (Kopylova et al., 2010b).
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Twenty-three diamond crystal fragments were examined by Fourier transform IR spectroscopy. The results indicated that nitrogen impurity in the diamonds studied was present mainly in the aggregated form. The diamonds contained paramagnetic nitrogen in the form of single carbon-substituting atoms in the diamond lattice (C-center) at concentrations not detectable by IR if any. Characteristic absorption bands revealed A and B1 nitrogen defects and tabular structures, B2 defects (Table 3). These data allow the crystals to be assigned to type IaA/B, widespread in the nature. Tabular structures in plane {100} (platelets, B2 defects) were recognized in most crystals from absorption band ~1370 cm–1. This band, with significant amplitude, was accompanied by a smaller one at 1430 cm–1. The location of the main band maximum varied among crystals from 1361 to 1379 cm–1. This variation is likely to be associated with linear sizes of the plates (Sobolev et al., 1968). The integrated intensity of the ~1370 cm–1 band also shows a linear correlation with the total platelet area (Sumida and Lang, 1982). The overall nitrogen content (calculated concentrations in the form of A and B1 defects) in the diamonds studied varied from < 100 to 3800 ppm. The percentage of nitrogen in the form of B1 defects varied from 0 to 94%. Moreover, some IR spectra had broad (peak widths at half heights 200 cm–1 or more) absorption bands with maximums at ~3440 and ~1650 cm–1, attributed to dispersed water impurities (Chrenko et al., 1967). Most of the crystals showed additional absorption bands: 3107, 1405, 3237, and 2785 cm–1. It is known from comprehensive studies of a large number of natural diamonds that bands 3107 and 1405 cm–1 for the type Ia crystals are practically universal in diamonds. They have been proven to belong to C–H bond vibrations (Davies et al., 1984; Woods and Collins, 1983). According to various models, the 3107 and 1405 cm–1 bands are associated with vibrations in vinylidene or ethylene groups (Sobolev et al., 1972; Woods and Collins, 1983) adsorbed on the surface of microscopic inclusions in diamonds. The strongest absorbance recorded at 3107 cm–1 in crystals studied was ~22 cm–1 (sample 18/4-2). It is conjectured that bands 3231 and 2785 cm–1 stem from vibrations in acetylene groups (–C≡CH) (Iakoubovskii and Adriaenssens, 2002; Woods and Collins, 1983). Absorption bands 3107, 1405, 3237, and 2785 cm–1 are widespread in natural diamonds from kimberlites and lamproites, being
Fig. 5. Morphology of diamond crystals from metaultramafic rock: A, fragment of a laminar octahedron; B, twin aggregate of crystals; C, spinel twin of laminar octahedra transitive to dodecahedroids, D, laminar octahedron with polycentric faces.
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Table 3. Major impurity centers and additional absorption bands in diamonds according to IR spectroscopy results Sample
Crystal morphology
NA, ppm
NB1, ppm %B
Ntot, ppm
αB2
vB2
α3107
Additional bands, v, cm–1
cm–1 18/4–1
Fragment
22
111
83
133
2
1377
–
2852, 2923, 2954, 1735, 1456, 1375
18/4–2
Fragment
271
3495
93
3766
50
1372
22
1405, 3237, 2785, 1430
18/4–3
Fragment
150
0
0
150
–
–
–
2852, 2923, 2954, 1456, 1735
18/4–4
Fragment
131
0
0
131
–
–
2.5
~3700
18/4–5
Fragment
–
–
–
–
–
–
–
2852, 2923, 1742, 1558, 960, 1066, 798
18/4–6
Fragment
No spectrum, weak transmission
18/4–7
Fragment
30
211
88
241
8.5
1363
14.5
1405, 3660, 1640, 829
18/4–8
Fragment
94
194
67
288
1.5
1362
1.5
~3440, 1640, 1735
18/4–9
Fragment
719
193
21
912
–
–
1.5
2850, 2920
18/4–10
Fragment
76
350
82
426
no data
no data
4
1650, ~3540, 2382, 1735, 2852, 2923, 2954
18/4–11
Fragment
95
271
64
266
3.5
1361
2
2850, 2921, 1735, 880
18/4–13
Fragment
341
506
60
847
13
1363
16
1405, 3237, 2785, 2852, 2923
18/4–14
Fragment
116
319
73
435
8
1363
12
1405, 3237, 2785, 788, 2852, 2923
indicative of C–H (and, probably, N–H) bonds. Some crystals studied may contain carbonates. Results and conclusions The FTIR data on the Archean diamonds studied indicate that their IR spectra are close to those of diamonds from kimberlites and lamproites. Like diamonds from the Abitibi-Wawa greenstone belt in the Superior craton (Stachel et al., 2006), Olondo diamonds broadly vary in the degree of nitrogen aggregation. Their specific feature is that their total nitrogen content can reach 3800 ppm. In contrast to Paleoproterozoic diamonds from lamprophyre dikes in the Nunavut Territory, Canada, Olondo diamonds have a high degree of nitrogen aggregation (Cartigny et al., 2004). Unlike other diamonds from lamprophyres, the diamonds considered here have a lighter carbon isotope ratio: δ13C = –26.0‰. Crystals with such properties have been found only in placers in the northeastern Siberian craton (Shatsky et al., 2011), for which the ore bodies have not been found or proven (Grakhanov et al., 2010). It is reasonable to infer from the nitrogen impurity and isotopic composition of the diamonds that their carbon source was subducted rocks of the Earth’s crust. However, in contrast to metamorphogenic diamonds of the Kokchetav block (Sitnikova and Shatsky, 2009) or diamonds from lamprophyre dikes in Nunavut, Canada (Cartigny et al., 2004), the Olondo diamonds stayed for long in the mantle at a high temperature. The presence of diamonds in rocks of the Olondo greenstone belt points to the existence of a diamondiferous lithospheric mantle beneath the southeastern portion of the North-Asian craton in the Mesoarchean, which is also proven by calculation of PT parameters of the formation of barophilic
minerals in the Middle Paleozoic Manchara kimberlite pipe. It was discovered by geologists of Yakutskgeologiya enterprise in 2007 (Smelov et al., 2009, 2010; Zaitsev et al., 2009). We are grateful to V.G. Gadiyatov, Yu.G. Tyllar†, and V.I. Pavlov for their long-term search for diamond ore bodies in the Olondo greenstone belt and for the collection they provided to us. We also express our gratitude to N.V. Popov and N.N. Dobretsov for detailed geological and petrographical mapping of the belt. This study was supported by the Ministry of Education and Science, State Contract 02.740.11.0328, and RAS Program No. 24.1.
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