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Abstract

The recognition of abundant microdiamonds included along with coesite in primary rock-forming minerals and zircons of metamorphic rocks from the Kokchetav massif, Northern Kazakhstan (Sobolev & Shatsky, 1990; Shatsky et al., 1991; Sobolev et al., 1991) indicates that crustal segments reached pressures of the order of at least 40 kbar (4 GPa), implying their subduction to depths greater than 100 km. The Kokchetav massif became internationally recognised as the type locality of diamondiferous metamorphic rocks; the petrological study of diamondiferous ultrahigh pressure (UHP) rocks provides a unique insight into the formation of diamond in crustal rocks at very high pressures. Along with the finding of coesite (Chopin, 1984) the discovery of diamond in supracrustal rocks has drastically changed the current ideas concerning the limits of UHP metamorphism of supracrustal rocks. The specific features and significance of such unique ultrahigh pressure metamorphism have been extensively discussed in different works (e.g. Coleman & Wang, 1995) and in numerous subsequent publications (Ernst & Liou, 2000). However, some workers preferred a hypothesis of a metastable diamond growth at low P–T parameters (Ekimova et al., 1994). The most recent collection of papers devoted specifically to petrotectonic characteristics of the Kokchetav massif is published as a special issue of island Arc (Liou & Banno, 2000). The significance of metamorphic processes at the origin of a new type of diamond has been extensively discussed (e.g. Haggerty, 1999). it is important to note that the Kumdy-Kol microdiamond deposit covers only a small portion of a more than 200 square km large area in which diamondiferous rocks are distributed in the Kokchetav massif (Shatsky et al., 1991; Dobretsov et al., 1999a). The proven microdiamond reserves of this deposit exceed 3 billion carats (e.g. Haggerty, 1999), making it an absolutely unique phenomenon worldwide. Apart from the Kokchetav massif, occurrences of microdiamonds in other UHP metamorphic terranes elsewhere are less well documented because of the need of bulk extraction and the lack of an unambiguous confirmation of microdiamond in situ (Xu et al., 1992; Dobrzhinetskaya et al., 1995). Another microdiamond locality in gneisses, confirmed by direct observations of thin sections, is from Erzgebirge, Germany, which has been suggested to be similar to the type locality of the Kokchetav massif (Massonne, 1999; Stöckhert et al., 2001).

Introduction

The recognition of abundant microdiamonds included along with coesite in primary rock-forming minerals and zircons of metamorphic rocks from the Kokchetav massif, Northern Kazakhstan (Sobolev & Shatsky, 1990; Shatsky et al., 1991; Sobolev et al., 1991) indicates that crustal segments reached pressures of the order of at least 40 kbar (4 GPa), implying their subduction to depths greater than 100 km. The Kokchetav massif became internationally recognised as the type locality of diamondiferous metamorphic rocks; the petrological study of diamondiferous ultrahigh pressure (UHP) rocks provides a unique insight into the formation of diamond in crustal rocks at very high pressures. Along with the finding of coesite (Chopin, 1984) the discovery of diamond in supracrustal rocks has drastically changed the current ideas concerning the limits of UHP metamorphism of supracrustal rocks. The specific features and significance of such unique ultrahigh pressure metamorphism have been extensively discussed in different works (e.g. Coleman & Wang, 1995) and in numerous subsequent publications (Ernst & Liou, 2000). However, some workers preferred a hypothesis of a metastable diamond growth at low P–T parameters (Ekimova et al., 1994). The most recent collection of papers devoted specifically to petrotectonic characteristics of the Kokchetav massif is published as a special issue of island Arc (Liou & Banno, 2000). The significance of metamorphic processes at the origin of a new type of diamond has been extensively discussed (e.g. Haggerty, 1999). it is important to note that the Kumdy-Kol microdiamond deposit covers only a small portion of a more than 200 square km large area in which diamondiferous rocks are distributed in the Kokchetav massif (Shatsky et al., 1991; Dobretsov et al., 1999a). The proven microdiamond reserves of this deposit exceed 3 billion carats (e.g. Haggerty, 1999), making it an absolutely unique phenomenon worldwide. Apart from the Kokchetav massif, occurrences of microdiamonds in other UHP metamorphic terranes elsewhere are less well documented because of the need of bulk extraction and the lack of an unambiguous confirmation of microdiamond in situ (Xu et al., 1992; Dobrzhinetskaya et al., 1995). Another microdiamond locality in gneisses, confirmed by direct observations of thin sections, is from Erzgebirge, Germany, which has been suggested to be similar to the type locality of the Kokchetav massif (Massonne, 1999; Stöckhert et al., 2001).

Geological outline

The Kokchetav massif is situated within the Caledonides of the Central Asian fold belt (Fig. 1). This was formed as a single structure at the end of Paleozoic due to collision of the Siberian continent with the North China, Tarim, Tadzhik, Karakorum and Kazakhstan-North Tien Shan massifs (Zonenshain et al., 1990). The Kokchetav massif is commonly considered as a fragment of Kazakhstan-North Tien Shan massif which has been broken up in the Late Riphean-Vendian (Zonenshain et al., 1990).

Fig. 1.

Major tectonic features of Central Asia for the region surrounding and including the Kokchetav massif, with the location of Figure 2 shown (Dobretsov et al., 1999a).

Fig. 1.

Major tectonic features of Central Asia for the region surrounding and including the Kokchetav massif, with the location of Figure 2 shown (Dobretsov et al., 1999a).

According to Dobretsov et al. (1995), the Kokchetav Massif is a contrasting megamelange composed of slices and blocks of ultrahigh (UHP) and high pressure (HP) (Units I and II), medium pressure (MP) (Barrow-type metamorphism, unit III), and low pressure (LP) rocks. In turn, the UHP and HP rocks are subdivided into two domains (Dobretsov et al., 1998): western (Kumdy-Kol) and eastern (Kulet) (Fig. 2), which are characterised by different P–T conditions of metamorphism and deformation types, suggesting different mechanisms of exhumation in the western and eastern blocks (Theunissen et al., 2000a,b). The ultrahigh pressure Unit I is composed of diamond-bearing metasedimentary rocks and gneisses with eclogites in the core of a large antiform covered or substituted by Unit II, an UHP eclogite-bearing mélange with predominant micaschists in the western and/or orthogneisses in the central parts of megamelange.

Fig. 2.

The Kokchetav Megamelange with tectonic units 1-3 and adjacent domains I-V (Dobretsov et at., 1998). 1: ultrahigh pressure/HP unit with high temperature eclogites and diamond-bearing rocks (Kumdy-Kol) and with whiteschist, coesite-bearing micaschist and relatively low temperature eclogites (Kulet); 2: medium pressure unit with Al-rich metasediments and coronite (Enbek-Berlyk); 3 : low pressure unit (Daulet). Domain I: Neoproterozoic sequences mainly belonging to the Vendian-Early Cambrian island arc; Domain II: “Kokchetav Microcontinent Domain” is composed of several blocks. In various amounts each of this blocks includes: a) a gneissic basement, b) its sedimentary cover (black shale and dolomite in the lower part and metasandstone in the upper part, c) tectonic slices of Vendian-Early Cambrian (?) ophiolite and volcanics, d) Ordovician rocks and e) Devonian granites; Domain III: the “Megamelange Domain”; Domain IV: the “White Lake Domain” - amphibolites, amphibole schist and quartzites, assumed to represent fragments of an oceanic crust; Domain V: the “Granite Dome Domain” is polyphase and composed of the Imantau and Zerenda granite domes with Cambrian gabbro, Late Ordovician-Silurian diorite, granodiorite and granite, and Devonian leucocratic granite and granosyenite.

Fig. 2.

The Kokchetav Megamelange with tectonic units 1-3 and adjacent domains I-V (Dobretsov et at., 1998). 1: ultrahigh pressure/HP unit with high temperature eclogites and diamond-bearing rocks (Kumdy-Kol) and with whiteschist, coesite-bearing micaschist and relatively low temperature eclogites (Kulet); 2: medium pressure unit with Al-rich metasediments and coronite (Enbek-Berlyk); 3 : low pressure unit (Daulet). Domain I: Neoproterozoic sequences mainly belonging to the Vendian-Early Cambrian island arc; Domain II: “Kokchetav Microcontinent Domain” is composed of several blocks. In various amounts each of this blocks includes: a) a gneissic basement, b) its sedimentary cover (black shale and dolomite in the lower part and metasandstone in the upper part, c) tectonic slices of Vendian-Early Cambrian (?) ophiolite and volcanics, d) Ordovician rocks and e) Devonian granites; Domain III: the “Megamelange Domain”; Domain IV: the “White Lake Domain” - amphibolites, amphibole schist and quartzites, assumed to represent fragments of an oceanic crust; Domain V: the “Granite Dome Domain” is polyphase and composed of the Imantau and Zerenda granite domes with Cambrian gabbro, Late Ordovician-Silurian diorite, granodiorite and granite, and Devonian leucocratic granite and granosyenite.

Relatively poor outcrop conditions, limited geochronological information and few metamorphic data that is largely restricted to UHP remnant sequences did not allow to present more precise information on early exhumation processes and associated structural settings.

UHP rocks of the Kumdy-Kol and Barchi locations (Unit I)

According to drilling data, the diamondiferous rocks in the western block occur in the form of slices in granite-gneisses, whose thickness is within a few hundreds of meters (Fig. 3). Our studies have established considerable variations in equilibrium P–T parameters of eclogites in units I and II (Shatsky et al., 1989a). It was noticed that the highest temperature eclogites occur at the Kumdy-Kol area (T = 920−1000 °C). Further studies have revealed high temperature eclogites and diamondiferous metamorphic rocks near Pakhar village (17 km northwest of the Kumdy-Kol) (Shatsky et al., 1991) and Lake Barchi (15 km west from Kumdy-Kol) (Korsakov et al., 1998). All these types belong to unit I of the Kumdy-Kol romb-horst domain (Theunissen et al., 2000b). Microdiamonds have been found only in metasediments of unit I: garnet-biotite gneisses and schists, garnet-muscovite-kyanite schists, garnet-pyroxene rocks and dolomitic marbles. It should be emphasised that eclogites do not contain diamonds (Shatsky & Sobolev, 1993).

Fig. 3.

Detailed geological map of diamond-bearing metasedimentary rocks mined, drilled and trenched (Scheshkel et al.,pers. comun.). 1: granite-gneisses; 2: biotite gneisses; 3: granite-gneisses and gneisses alternation; 4: fine grained chlorite-tremolite-quartz rocks; 5: migmatites; 6: garnet-muscovite, kyanite-muscovite schists; 7: pyroxene-carbonate rocks; 8: garnet pyroxenites; 9: eclogites and amphibolites; 10: dykes of diorite porphyrites.

Fig. 3.

Detailed geological map of diamond-bearing metasedimentary rocks mined, drilled and trenched (Scheshkel et al.,pers. comun.). 1: granite-gneisses; 2: biotite gneisses; 3: granite-gneisses and gneisses alternation; 4: fine grained chlorite-tremolite-quartz rocks; 5: migmatites; 6: garnet-muscovite, kyanite-muscovite schists; 7: pyroxene-carbonate rocks; 8: garnet pyroxenites; 9: eclogites and amphibolites; 10: dykes of diorite porphyrites.

Eclogites

Eclogites occur within kyanite-mica schists, plagiogneisses and diamondiferous metasediments. The eclogites are medium-grained, pinkish to green rocks of granoblastic texture consisting of garnet (40-50%), pyroxene (30-40%), quartz (5-10%). The secondary minerals are represented by amphibole (10-20%), biotite (3-5%), plagioclase (4-2%). Rutile occurs as an accessory mineral. The replacement of omphacite by plagioclase-pyroxene symplectite is one of the peculiarities of eclogite. In some samples, omphacite is completely transformed into pyroxene-plagioclase symplectite. In the host rocks the following mineral associations are present: qzt + (chl) + gr, qtz + gr + bi + ph, qtz + amph + zo + ep + (chl) + gr, qtz + amph + ort + cpx + tit, qtz + fsp + bi + gr + cpx. These associations formally indicate a mixture of amphibolite, granulite, epidote-amphibolite and greenschist facies in the host rocks. But relics of UHP minerals show that the polystage history of these rocks is similar to eclogite history.

Biotite gneisses and garnet-pyroxene-quartz rocks

Biotite gneisses are the most common diamondiferous rocks among the metasediments (Fig. 3). Lenses and bands of calc-silicate metacarbonate rocks, garnet-pyroxene rocks, relics of cataclastic massive garnet-pyroxene-quartz and garnet-quartz rocks, and finegrained chlorite-actinolite-quartz rocks occur within biotite gneisses (Shatsky et al., 1995; Shatsky & Sobolev, 1993). Cataclastic garnet-biotite gneisses contain relics of ultrahigh pressure garnets (Shatsky et al., 1995) as well as biotite, plagioclase, potassic feldspar, muscovite and amphibole. All varieties of transitional rocks from massive garnet-pyroxene-quartz rocks through banded garnet-biotite gneisses and chlorite schists occur in both sites. Banding at both large (0.5-1 m) and small scales (0.5-1 cm) is observed in diamondiferous rocks, particularly in garnet-biotite gneisses. This banding is caused by alternating layers composed mainly of varying amounts of garnet, biotite, pyroxene, quartz and plagioclase.

Clinopyroxene found in some types of biotite gneisses and garnet-pyroxene-quartz rocks could be referred to the earlier high pressure paragenesis. However, the jadeitic component in these types of rocks is lower in comparison with pyroxene from inclusions in garnet and zircon. We found that omphacite included in zircons from cataclastic garnet-biotite gneiss contains up to 50% jadeite component. It should be mentioned that clinopyroxene with 74% jadeite was identified as inclusion in zircon from pelitic gneisses by Katayama et al. (1998). This can indicate that the jadeitic component in pyroxenes of the matrix is reduced as a result of the retrograde metamorphism under the conditions of the amphibolite facies (Shatsky & Sobolev, 1985).

Garnets from these types of rocks vary greatly in their composition (Figs. 4 and 5). The fields of the non-diamondiferous and diamondiferous rocks partly overlap. Garnets from eclogites plot in the field of diamondiferous rocks. The contents of Ca and Fe sometimes increase from the core to the rim when zoning is present, but in most cases garnets from diamondiferous rocks do not show zoning.

Fig. 4.

Compositional variations of garnet from the metamorphic rocks of Kumdy-Kol. 1: diamondiferous gneiss, garnet-pyroxene-quartz rocks and schist; 2: garnet-pyroxene rocks; 3: dolomitic marbles; 4: diamond-free gneiss; 5: eclogite; 6: inclusions in zircons from dolomitic marble (K9-16); 7: garnet from matrix (K9-16); 8: diamondiferous garnet-kyanite-muscovite schist.

Fig. 4.

Compositional variations of garnet from the metamorphic rocks of Kumdy-Kol. 1: diamondiferous gneiss, garnet-pyroxene-quartz rocks and schist; 2: garnet-pyroxene rocks; 3: dolomitic marbles; 4: diamond-free gneiss; 5: eclogite; 6: inclusions in zircons from dolomitic marble (K9-16); 7: garnet from matrix (K9-16); 8: diamondiferous garnet-kyanite-muscovite schist.

Fig. 5.

Compositional variation of garnet from zircons. 1: dolomitic marble; 2: biotite plagiogneiss; 3: garnet-pyroxene rocks; 4: pyroxene-biotite gneiss; 5: migmatites; 6: garnet from matrix of all rock types.

Fig. 5.

Compositional variation of garnet from zircons. 1: dolomitic marble; 2: biotite plagiogneiss; 3: garnet-pyroxene rocks; 4: pyroxene-biotite gneiss; 5: migmatites; 6: garnet from matrix of all rock types.

Biotite intergrowth with diamonds in garnets from garnet-biotite gneisses and garnet-pyroxene-quartz cataclastic rocks has a lower Fe/(Fe + Mg) compared with biotite in the host rocks (Sobolev & Shatsky, 1990; Vavilov et al., 1991). Si contents in phengites included in zircon and garnets vary from 3.22 to 3.56 p.f.u. (Shatsky et al., 1995; Korsakov et al., 2002). Phengite from matrix, as usual, contain less celadonite component than that included in zircon (3.2-3.32 Si p.f.u.).

Coesite inclusions have been found in zircon grains in more than thirty samples of diamondiferous metapelites, including biotite-kyanite-zoisite gneisses and schists (Sobolev et al., 1998; Korsakov et al., 2002). These finds were confirmed by subsequent studies by Katayama et al. (2000). One bimineralic inclusion representing intergrowth of coesite and diamond was documented (Sobolev et al., 1994). Coesite was also found as inclusions in zircon grains from a high temperature eclogite from the drilling hole near Lake Barchi (Korsakov et al., 1998). In addition to zircon inclusions, single crystals of coesite occur also in garnets from three samples of zoisite gneisses of the Lake Barchi site (Korsakov et al., 2002). In all these samples, the presence of coesite was confirmed by Raman spectroscopy. The presence of coesite inclusions in garnets from zoisite gneisses is an unusual feature compared with typical relicts of coesite described from elsewhere.

Metasomatic rocks

Very fine-grained rocks with granoblastic structure are treated as a separate group. Great variations in the ratios of the constituent minerals are observed with dominant quartz, amphibole and chlorite. These minerals are accompanied by biotite, muscovite, zoisite and tourmaline. Some garnet grains are replaced by amphibole, chlorite and mica. Amphibole is replaced by chlorite and biotite. Zoisite forms large grains with chlorite and amphibole inclusions. Some areas in these rocks are composed of large grains of K-feldspar.

Garnet-pyroxene rocks

Diamondiferous and diamond-free garnet-pyroxene rocks consist mainly of garnet and clinopyroxene with a small amount of carbonate and K-feldspar. Amphibole and chlorite occur as retrograde minerals. Garnet-pyroxene rocks can be subdivided into two groups according to the composition of garnet and clinopyroxene. The first one is made up of Mg-rich clinopyroxene (12-17% MgO) and grossular-pyrope-almandine garnet (Figs. 4 and 5). High K2O and a lack of or low Na2O distinguishes the pyroxenes in this type of rock. Orthoclase lamellae (1-20 μm) and silica needles are also found in the pyroxene grains in the matrix and sometimes in clinopyroxene included in garnet (Shatsky et al., 1985, 1995; Katayama et al., 2000). When exsolution structures are absent, pyroxenes in the groundmass and those included in garnets have the same composition. Lamellae of highly aluminous titanite were also found in clinopyroxene from these rocks. Garnet-pyroxene rocks of this type generally contain diamonds but samples without diamonds are also available. Microdiamonds are widely scattered as inclusions in garnet, clinopyroxene and zircon. Taking into account that diamonds are unevenly distributed within samples and even within thin sections, in specific cases it is difficult make conclusions about diamond occurrences within certain samples. In contrast, garnet-pyroxene rocks of the second type never contain diamonds. These rocks are made up of grossular-almandine garnet and clinopyroxene enriched in iron. The most characteristic features are the absence of K impurities in this clinopyroxene.

Dolomitic marbles

Marbles containing a variable proportion of dolomite, Mg-calcite, diopsidic pyroxene and garnet are typical of the Kumdy-Kol microdiamond deposit. As previously, two types of marbles can be recognised based on a composition of clinopyroxene. Clinopyroxene from the first type of marbles similar to clinopyroxene from the first type of pyroxene-garnet rocks, containing lamellae of K-feldspar and quartz needles. K-rich clinopyroxene is also found in the matrix. These types of rocks contain diamonds.

Garnets from dolomites differ significantly in their composition from the garnets of other rock types (Figs. 4 and 5). They are notable for their high calcium content and high Mg/(Mg + Fe) ratio. Compositional profiles along the grains of garnet in most cases show that the central part is not zoned, but a change in the chemical composition occurs in the 100-125 μm wide rim (Shatsky et al., 1995). It should be noted that when zoning is absent in garnets, their composition can vary within a thin section. Strong compositional variation was found for garnets included in zircons from single samples of marbles. Calcium contents vary from 27% to 49%, and the Fe/(Fe + Mg) ratio ranges from 0.437 to 0.336. At the same time, in the garnet from the matrix the compositional variations are insignificant (Shatsky et al., 1995).

New types of lamellae have been observed in pyroxene from dolomitic marbles (Shatsky & Sobolev, 2001). One of the samples (94-275) of such a rock contains pyroxene grains with garnet lamellae along with grains containing K-feldspar lamellae and no lamellae at all (Table 1). Compositions of garnet lamellae are very similar to garnet rims in matrix crystals (Grs 31.8-33.2; Prp 38.3-45.7 and Alm 19.5-15.9). Prograde zonation is typical of garnets from the matrix. Grossular and pyrope components increase from core (Grs 32, Prp 29) to rim (Grs 36.5, Prp 37.9) with decreasing almandine content (from 35.6 to 18.5). All Cpx both with and without lamellae are diopside with low FeO (2 wt%), Al2O3 (1.32-1.89 wt%) and Na2O (0.11-0.16 wt%). No detectable K2O was found.

Table 1.

Representative analyses of minerals from pyroxene-carbonate rock 94–275

NameCpx1Lam1Lam2Gt1Gt2IncInc2Cpx2
corerim
SiO253.9040.3040.3040.3040.7041.0053.8038.6053.20
TiO20.220.280.220.360.280.330.230.730.22
Al2O31.8621.5021.6021.1021.4021.401.8617.301.76
Cr2O30.020.010.040.020.020.020.000.000.02
FeO2.0610.308.5510.2010.4010.902.414.322.03
MnO0.050.670.540.590.580.550.030.010.05
MgO16.4010.1012.109.9810.109.6116.6022.6016.00
CaO24.9015.4013.7015.6016.0015.9024.800.0425.00
Na2O0.120.000.000.000.000.010.170.060.12
K2O0.000.000.000.000.010.010.009.350.00
Total99.4898.5097.0098.1299.4399.6499.8993.0698.37
Si1.9673.0153.0223.0253.0233.0371.9592.7891.967
Ti0.0060.0160.0120.0200.0150.0180.0060.0400.006
Al0.0801.8951.9131.8691.8721.8660.0801.4750.077
Cr0.0000.0010.0020.0010.0010.0010.0000.0000.000
Fe2+0.0630.5660.4700.5380.5400.5690.0680.063
Fe3+0.0000.0820.0670.1010.1030.1060.0060.000
Mn0.0010.0430.0340.0380.0360.0350.0010.0010.002
Mg0.8901.1291.3511.1171.1121.0620.8992.4290.882
Ca0.9731.2341.1021.2581.2701.2640.9680.0030.988
Na0.0080.0000.0000.0000.0000.0010.0120.0090.009
K0.0000.0000.0000.0000.0000.0000.0000.8610.000
Fe tot0.0630.6500.5380.6420.6460.6780.0730.2610.063
Total3.9907.9807.9747.9687.9737.9593.9977.8673.993
f6.6036.5528.4936.5036.7338.987.559.706.64
NameCpx1Lam1Lam2Gt1Gt2IncInc2Cpx2
corerim
SiO253.9040.3040.3040.3040.7041.0053.8038.6053.20
TiO20.220.280.220.360.280.330.230.730.22
Al2O31.8621.5021.6021.1021.4021.401.8617.301.76
Cr2O30.020.010.040.020.020.020.000.000.02
FeO2.0610.308.5510.2010.4010.902.414.322.03
MnO0.050.670.540.590.580.550.030.010.05
MgO16.4010.1012.109.9810.109.6116.6022.6016.00
CaO24.9015.4013.7015.6016.0015.9024.800.0425.00
Na2O0.120.000.000.000.000.010.170.060.12
K2O0.000.000.000.000.010.010.009.350.00
Total99.4898.5097.0098.1299.4399.6499.8993.0698.37
Si1.9673.0153.0223.0253.0233.0371.9592.7891.967
Ti0.0060.0160.0120.0200.0150.0180.0060.0400.006
Al0.0801.8951.9131.8691.8721.8660.0801.4750.077
Cr0.0000.0010.0020.0010.0010.0010.0000.0000.000
Fe2+0.0630.5660.4700.5380.5400.5690.0680.063
Fe3+0.0000.0820.0670.1010.1030.1060.0060.000
Mn0.0010.0430.0340.0380.0360.0350.0010.0010.002
Mg0.8901.1291.3511.1171.1121.0620.8992.4290.882
Ca0.9731.2341.1021.2581.2701.2640.9680.0030.988
Na0.0080.0000.0000.0000.0000.0010.0120.0090.009
K0.0000.0000.0000.0000.0000.0000.0000.8610.000
Fe tot0.0630.6500.5380.6420.6460.6780.0730.2610.063
Total3.9907.9807.9747.9687.9737.9593.9977.8673.993
f6.6036.5528.4936.5036.7338.987.559.706.64

Another type of lamellae was found in clinopyroxene from sample 98-6 (Table 2). Cpx from this sample contain lamellae of phengite (3.42 Si p.f.u.) and K-feldspar. K contents in Cpx varies from 0.05 to 0.54 wt%. On the basis of textural grounds it is concluded that topotaxial growth of mica lamellae in Cpx are best explained by exsolution from a former K and OH-bearing Cpx stabilised at UHP conditions.

Table 2.

Representative analyses of minerals from pyroxene-carbonate rock K 98-6

NameGrtIncCpx1Lam1Lam2Cpx2Mica
corerimcorerim
SiO239.3039.6054.1053.8054.0051.1063.5054.7039.20
TiO20.400.380.000.000.001.010.000.001.18
Al2O320.8021.302.111.620.8428.2916.801.8415.00
Cr2O30.020.010.000.000.000.000.000.000.00
FeO6.557.002.903.013.121.080.112.0010.40
MnO0.991.080.140.120.100.020.000.070.30
MgO3.373.8215.9016.2016.403.910.0416.4918.50
CaO27.6026.3024.3025.0024.500.240.2024.870.35
Na2O0.050.040.360.340.170.040.070.410.06
K2O0.010.020.360.050.545.9214.80.088.69
Total99.0499.62100.27100.0499.7091.5995.42100.4693.68
Si2.9922.9961.9691.9601.9823.4203.0511.8942.888
Ti0.0230.0220.0000.0000.0000.0510.0000.0240.065
Al1.8661.8990.0900.0700.0362.2290.9491.1141.307
Cr0.0010.0000.0000.0000.0000.0000.0000.0000.000
Fe2+0.3160.3700.0880.0660.0870.034
Fe3+0.1010.0730.0000.0250.0090.000
Mn0.0640.0690.0040.0040.0030.0010.0000.0000.019
Mg0.3820.4300.8640.8800.8960.3900.0030.2602.030
Ca2.2512.1340.9480.9750.9630.0170.0100.0680.028
Na0.0080.0060.0250.0240.0120.0050.0060.0030.009
K0.0000.0010.0170.0020.0250.5050.9070.2580.817
Fe tot0.4190.4440.0880.0920.0960.0600.0040.0340.642
Total8.0047.9994.0064.0054.0146.6744.9313.6567.806
f52.2750.779.279.479.6513.4260.8511.6224.04
NameGrtIncCpx1Lam1Lam2Cpx2Mica
corerimcorerim
SiO239.3039.6054.1053.8054.0051.1063.5054.7039.20
TiO20.400.380.000.000.001.010.000.001.18
Al2O320.8021.302.111.620.8428.2916.801.8415.00
Cr2O30.020.010.000.000.000.000.000.000.00
FeO6.557.002.903.013.121.080.112.0010.40
MnO0.991.080.140.120.100.020.000.070.30
MgO3.373.8215.9016.2016.403.910.0416.4918.50
CaO27.6026.3024.3025.0024.500.240.2024.870.35
Na2O0.050.040.360.340.170.040.070.410.06
K2O0.010.020.360.050.545.9214.80.088.69
Total99.0499.62100.27100.0499.7091.5995.42100.4693.68
Si2.9922.9961.9691.9601.9823.4203.0511.8942.888
Ti0.0230.0220.0000.0000.0000.0510.0000.0240.065
Al1.8661.8990.0900.0700.0362.2290.9491.1141.307
Cr0.0010.0000.0000.0000.0000.0000.0000.0000.000
Fe2+0.3160.3700.0880.0660.0870.034
Fe3+0.1010.0730.0000.0250.0090.000
Mn0.0640.0690.0040.0040.0030.0010.0000.0000.019
Mg0.3820.4300.8640.8800.8960.3900.0030.2602.030
Ca2.2512.1340.9480.9750.9630.0170.0100.0680.028
Na0.0080.0060.0250.0240.0120.0050.0060.0030.009
K0.0000.0010.0170.0020.0250.5050.9070.2580.817
Fe tot0.4190.4440.0880.0920.0960.0600.0040.0340.642
Total8.0047.9994.0064.0054.0146.6744.9313.6567.806
f52.2750.779.279.479.6513.4260.8511.6224.04

Note: Cpx1 - clinopyroxene contains lamellae of phengite

Dolomite, magnesite and magnesian calcite were found among the inclusions in zircon. MgCO3 content in magnesium calcite is about 23.5% (Shatsky et al., 1995). Inclusions of calcite and dolomite are found in garnet. Calcite inclusions predominate. The Mg admixture is minor in matrix calcite and calcite inclusions in garnets. Radial cracks are often observed around the calcite inclusions in garnet.

Microdiamonds

Microdiamonds are widely scattered as inclusions in garnet, zircon, and rare in clinopyroxene, kyanite and zoisite (Sobolev & Shatsky, 1987, 1990; Korsakov et al., 2002; Ogasawara et al., 2000) They are also very typical in rounded pseudomorphs after garnet, consisting of mica + chlorite + amphibole aggregates and tourmaline (Sobolev & Shatsky, 1990; Shatsky et al., 1995). Diamonds were not detected in intergranular space. As usual, diamonds coexist with graphite.

Each analysed diamond-bearing rock type from the Kokchetav massif, (garnet-clinopyroxene rocks, marbles, biotite gneisses and zoisite gneisses) contains microdiamonds with a distinctive range of morphologies (Shatsky et al., 1989b, 1998a, 1999b). Diamonds occurring in garnet-clinopyroxene calc-silicate rocks and marbles are predominantly of cuboid morphology. Biotite gneisses are characterised by a cubo-octahedral diamond population, although cuboids and growth forms transitional to cuboids also exist. Octahedral crystals are only observed in zoisite gneisses. Some zoisite gneisses carry octahedral diamonds, whereas others typically contain cuboids. Detailed petrological investigation of zoisite gneiss samples demonstrates that octahedral crystals appear only in rocks which do not contain symplectitic zoisite (Korsakov et al., 2002).

The greatest variety in diamond morphology within a single rock type is observed in biotite gneisses. in this rock type, the predominant crystal habit is the cubo-octahedron. intergrowths and aggregates are quite abundant. Shatsky et al. (1995) observed an octahedron and a cuboid in a single intergrowth. in addition to cuboidal and octahedral faces, some crystals have {110} faces. They consist of alternating small {111} steps and are believed to reflect a growth form. Coated diamonds also occur.

The predominant diamond type in garnet-pyroxene rocks and dolomitic marbles is the cuboid. Crystal surfaces display a large number of pits of variable configuration. Octahedral faces may be found at the edges of the cuboids. Diamonds from marbles, more likely than those of garnet-clinopyroxene rocks, often show octahedral microfaces on their cuboidal surfaces, indicating a change in the environmental conditions during diamond growth.

Zoisite gneisses devoid of symplectitic zoisite occur in the Barchi area. They include dominantly octahedral diamonds (more than 80%). Diamond size varies from 10 to 50 μm. In one rock sample, sharp-edged octahedra with octahedra displaying antiskeletal growth features are present. Twins are fairly frequent. In addition to octahedra, cuboids also occur. In general, in a single sample of zoisite gneiss devoid of symplectitic zoisite, all transition forms between cuboids and octahedra can be found. The occurrence of octahedral microfaces on cuboidal surfaces is also noted, but here this process ends with the formation of octahedra.

Cuboids from garnet-pyroxene rocks, as well as one octahedron from zoisite gneiss, have been investigated for their internal morphology using X-ray topography (Shatsky et al., 1998a). Cuboids from garnet-pyroxene rocks show a radiated structure resulting from space-filling columnar growth, also known as “fibrous growth”. An interesting feature of the inner structure of some crystals is the presence of a core, expressed as a slight blackening on the topograms. The octahedral diamond from zoisite gneiss has predominantly cubic growth sectors and a fibrous structure. Despite the small size of this crystal (30 μm), its central zone is well marked on the X-ray topography picture. This may suggest that either the central zone has no defects, or that it is disoriented relative to the main part of crystal. X-ray topography data for the octahedron support the conclusion, based on external morphology, that octahedral microdiamonds are in fact reshaped cuboids.

Diamonds from metamorphic rocks display all morphological types found for diamonds from kimberlites, lamproites, and alluvial deposits. The basic difference is however the relative abundance of the different diamond morphologies. In particular, diamonds from metamorphic rocks are dominated by cuboids, whereas in kimberlites and lamproites octahedra dominate. Moreover, the octahedra from metamorphic rocks initially grew as cuboids under conditions of high supersaturation and only at their final growth stages, when supersaturation decreased, did they acquire an octahedral shape. This is in contrast to octahedra from kimberlites and lamproites that in general grew by spiral or layer-by-layer growth mechanism at low degrees of carbon supersaturation (Sunagawa, 1984).

The mean δ13C and δ15N values of diamonds from garnet-clinopyroxene rock are −10.5 and +5.9%0, respectively. Diamonds from dolomitic marble have a mean δ13C value of −10.2 and a δ15N of +8.5%0 (Cartigny et al., 2001a). The isotopic values of nitrogen and carbon rather reflect a crustal origin. For mantle diamonds there is a strict relationship between the δ13C value and total nitrogen content (Cartigny et al., 2001b). The higher the nitrogen content, the higher the δ13C value. When the in situ diamonds with δ13C −10.5 and −10.2%0 are formed of mantle carbon, their maximum nitrogen content should be less than 1750 ppm. This is not confirmed by the nitrogen data obtained by infrared spectroscopy or combustion. Diamonds from garnet-clinopyroxene rocks are rich in nitrogen. Up to 2765 ppm of nitrogen is incorporated in their lattice (de Corte et al., 1998) and, moreover, up to 7000 ppm is present additionally as fluid inclusions of molecular nitrogen (Cartigny et al., 2001a). Diamonds from dolomitic marbles have a total nitrogen content of about 2000 ppm, of which maximum 1000 ppm is in the diamond structure and the remainder in fluid inclusions (Cartigny et al., 2001a). Therefore, the carbon source for the Kokchetav diamonds is interpreted as a mixture of crustal carbonates (δ13C ~ 0%0) and organic matter from carbonaceous metasediments (δ13C ≈ −25%0). The similar peculiarities were earlier observed for some kimberlitic and alluvial diamonds as a possible proof of a crustal origin of their carbon (Sobolev & Sobolev, 1980). The nitrogen isotopic composition of Kokchetav diamonds reflects the moderately to highly positive δ15N values of metamorphosed sedimentary rocks.

Diamonds from garnet-clinopyroxene rocks, marbles and zoisite gneisses contain water and carbonate inclusions (de Corte et al., 1998, 1999). They result from entrapment of fluid during the rapid crystallisation of the diamond fibres and hence provide evidence of diamond growth from a C-O-H fluid.

A great variety in the morphology of diamond crystals even in polycrystalline aggregates (Sobolev & Shatsky, 1990) may be due to variation in the degree of carbon supersaturation in the fluid phase as a consequence of contemporaneous melt generation. The presence of melt would vary the fluid phase composition. With free carbon in the rocks, and at constant oxygen fugacity, the fluid phase composition would be a function of the reaction of carbon and water: 

formula
When melt is present, the reactions given above would be shifted toward the right side resulting in precipitation of C from fluid phase. The reason is that the solubility of water in the melt is much higher compared with CO2 and CH4. Thus the appearance of a melt phase could lead to a mass crystallisation of diamonds or graphite.

The UHP and HP rocks of Unit II

In Unit II lower-temperature eclogites occur among phengite-kyanite-quartz and muscovite-garnet-quartz-plagioclase schists (Kulet, Chaglinka, Sulu-Tyube) and among mylonitic granitogneisses (Enbek-Berlyk, Lake Uyaly) (Shatsky et al., 1989a). Coesite inclusions have been found only in garnet from phengite schist from the Kulet area.

Sulu-Tyube area

The eclogites occur as several large bodies outcropping within the mica schists and gneisses of the second unit (Fig. 6). The largest (1.5×0.5×0.5 km) body of the Sulu-Tyube Hill represents a lens deformed into a synform. Strongly zoned garnet porphyroblasts, whose Fe/(Fe + Mg) ratio varies from 0.85 to 0.57, contain quartz, rutile, titanite, zoisite and epidote inclusions. Zoisite and amphibole are in textural equilibrium with omphacite.

Fig. 6.

Geological map of the Sulu-Tyube area based on geophysical, drilling and structural data (compiled by Dobretsov & Zayachovsky; Dobretsov et al, 1999b). 1 : traced boundaries of eclogites; 2 : blocks of fresh eclogites corresponding to highest density geophysical bodies; 3 : the LP Daulet unit; 4: small eclogite bodies; 5 : structural lines; 6: faults visible (a) and supposed under quaternary sediment (b).

Fig. 6.

Geological map of the Sulu-Tyube area based on geophysical, drilling and structural data (compiled by Dobretsov & Zayachovsky; Dobretsov et al, 1999b). 1 : traced boundaries of eclogites; 2 : blocks of fresh eclogites corresponding to highest density geophysical bodies; 3 : the LP Daulet unit; 4: small eclogite bodies; 5 : structural lines; 6: faults visible (a) and supposed under quaternary sediment (b).

During the detailed investigation the Sulu-Tyube eclogites were found to have a heterogeneous structure. Varieties containing porphyroblasts of garnet (up to 7 mm) are interbeds of medium-grained eclogite with garnet grains up to 1 mm. The eclogite modal composition is: garnet (25-30%), omphacite (20-25%), quartz (5-10%), amphibole (20-25%), zoisite (5-10%). Among the secondary minerals there are plagioclase, chlorite, muscovite, epidote and calcite. These rocks are massive and porphyroblastic. Drilling of the eclogite body revealed that the composition of eclogites does not change up to a depth of 550 m (A. Zayachkovsky, pers. comm.).

Besides the large body of the Sulu-Tyube region, there are numerous outcrops of fine and medium-grained eclogites, forming separate boudins tracing the outline of original layers. The host rocks include both the white mica (“silver”) schists and medium-grained gneisses of dark brown colour, composed of biotite, quartz, plagioclase, muscovite, potassium feldspar and sometimes kyanite.

Enbek-Berlyk area

Amphibolised eclogite and gabbroid bodies occur in this area. According to some researchers (Dobretsov & Sobolev, 1970; Udovkina, 1985), eclogitisation reactions can be observed in gabbro. Because of these observations a detailed investigation of metabasites at this site was made. Garnet-biotite-sillimanite-kyanite schists and metagabbro-garnet amphibolite boudins are characteristic for the upper unit (Fig. 7). The lower unit contains retrograde eclogite and coarse-grained gneiss inclusions among garnet mica schists with transition to sheared orthogneisses with boudinaged metadolerite dykes.

Fig. 7.

Structural geological sketch map of the Enbek-Berlik area. 1 : amphibolised dolerite; 2: amphibolite; 3: eclogite and amphibolised eclogite; 4: pyroxenite and harzburgite; 5: coronite; 6: diopside-plagioclase rock; 7: metapelites and the strike of their schistosity; 8: granite; 9: boundary of recent deposits; 10: fault; 11: dip of bedding (Reverdatto, 1999).

Fig. 7.

Structural geological sketch map of the Enbek-Berlik area. 1 : amphibolised dolerite; 2: amphibolite; 3: eclogite and amphibolised eclogite; 4: pyroxenite and harzburgite; 5: coronite; 6: diopside-plagioclase rock; 7: metapelites and the strike of their schistosity; 8: granite; 9: boundary of recent deposits; 10: fault; 11: dip of bedding (Reverdatto, 1999).

The gabbro-norites (hyperstene granulites) are medium-grained, dark grey rocks, consisting of plagioclase (25–30%), clinopyroxene (30–40%), orthopyroxene (10–15%) and ilmenite (2–3%). Secondary amphibole and biotite are found. The texture is ophitic, poikiloblastic in some places. The structure is massive. At the margins of gabbro-norite boudins, garnet amphibolites are formed: garnet (25–30%), newly formed pyroxene (10–20%), amphibolite (20–30%), quartz (5–10%) and plagioclase (3–5%). Rutile, ilmenite and titanite are presented as accessory minerals, the secondary minerals are zoisite and biotite. The texture is granoblastic, the structure is massive. Between these two rock types, transitional varieties are observed. it has been mentioned by earlier investigators that the corona gabbro in the marginal parts of the bodies is interpreted as the result of eclogitisation of gabbroid rocks (Dobretsov & Sobolev, 1970). Our investigations do not confirm this conclusion, since the newly formed pyroxene has low omphacite content and the metamorphic assemblage is represented by sodic augite, amphibole and plagioclase.

Transition to garnet amphibolite involves the following stages. Corona stage: rims of garnet at the contact between plagioclase and primary pyroxene. Second stage: growth of garnet grains is observed within plagioclase. Simultaneously primary plagioclase is replaced by a very fine-grained quartz-plagioclase aggregate, the new plagioclase is more sodic (An15–20). Newly formed pyroxene replacing augite has low jadeite content (5.5% Ja).

The eclogites of the Enbek-Berlyk area, as a rule, are intensely altered. Garnets (25–30%), omphacite (15–17%) and quartz (5–10%) are the primary minerals. Plagioclase (15–17%), amphibole (10–20%), zoisite, biotite and epidote also are present as secondary minerals. Accessory minerals are rutile and ilmenite. Pyroxenes from eclogite are almost completely replaced by pyroxene-plagioclase symplectite or by amphibole. The analyses of the fresh grains show that pyroxene contains 21% Jd. in symplectite pyroxene the jadeitic component decreases down to 5%.

The value of Fe/(Fe + Mg) ratios of garnets from amphibolites and corona gabbro ranges from 0.80 to 0.94, while that of the eclogite garnets ranges from 0.60 to 0.80. The garnets in all metabasites are zoned. The decrease of Mg content and the increase of Fe and Mn from the core to the rim are observed in eclogites. The garnets from amphibolite have the same zoning patterns. This zoning is regressive. The garnets from corona gabbro are either non-zoned, or their Mg content increases and Ca content decreases from the core to the rim. it should be noted that the composition of garnets from the rims, surrounding pyroxene and the garnets growing inside plagioclase grains remains the same.

The metasedimentary schists are characterised by garnet porphyroblasts in a finegrained matrix of biotite (20-25%), quartz (15-20%), kyanite (10-15%), and muscovite (5%). Kyanite is replaced by sillimanite.

Kulet area

Bodies of fine and medium-grained eclogites occur among garnet-muscovite-kyanite-quartz and muscovite-garnet-quartz-plagioclase schists (Fig. 8). This structure possibly includes several tectonic sheets similar to the Enbek-Berlyk area.

Fig. 8.

Detailed geological map of the central part of the South Zheltau (Kulet) area (Dobretsov et al., 1999b).

Fig. 8.

Detailed geological map of the central part of the South Zheltau (Kulet) area (Dobretsov et al., 1999b).

Two eclogite varieties are interlayered on the scale of 10 to 160 metres. They are dark green and light pink fine-grained eclogites, with predominant omphacite and garnet, respectively. The main minerals are pyroxene, garnet, quartz, porphyroblastic zoisite, pale brown amphibole and rare phengite. Some areas of thin sections consist of larger foliated omphacite grains. The increase in omphacite grain size is observed near zoisite grains and quartz veins, where idiomorphic pyroxene forms. The country rocks are represented mainly by garnet-mica schists, with interbeds of talc-kyanite-garnet schists (Udovkina, 1985). Garnet and pyroxene are replaced by dark green amphibole at the eclogite boudin margins. In some cases eclogite transforms into amphibolite. Talc-garnet-kyanite rocks are intercalated with amphibole-garnet-kyanite assemblages. The other rocks present are: amphibole-garnet-zoisite, garnet-biotite-kyanite-amphibole and garnet-zoisite-kyanite rocks.

A recent discovery of coesite relics in garnet from phengite schist near Lake Kulet is of particular significance (Shatsky et al., 1998b). The rocks are composed of garnet, phengite and quartz. Moreover, there are relics of kyanite after which white mica develops. Phengite in turn is replaced by biotite in the rims. Rutile and tourmaline occur as accessory minerals. Two generations of garnet are present there. The first generation garnet forms large subhedral grains often containing abundant inclusions of quartz and rutile. The second generation garnet occurs as 50-150 μm long grains. Large garnet grains display zoning, which is expressed as an increase in Fe and Mg and a decrease in Ca and Mn contents from core to rim. The Fe/Mg value decreases from 9.5 to 7.9. The second generation garnets exhibit weak zoning. In composition they are similar to the rims of large garnet porphyroblasts, though their Fe/Mg value is still lower (to 5.9). Phengite has a high content of celadonite component (3.43 p.f.u. Si). The Si content in muscovite in a previously studied mica schist from the Kulet area, however, was no more than 3.2 p.f.u., whereas phengite was found only in eclogites (3.46 p.f.u. Si). Quartz occurs in garnet both as polycrystalline inclusions and single crystals. Inclusions are often surrounded by radial cracks. Coesite inside a polycrystalline quartz aggregate was found in only one inclusion. Coesite was identified both by petrography and by the characteristic Raman spectral band at 520 cm−1. Whiteschists rich in talc and Mg-rich garnet are usual in the Kulet area. Phengite can occur instead of talc.

P–T path

The results of thermobarometry are summarised in Table 3. One must keep in mind that diamondiferous rocks have minerals formed at least during three metamorphic stages (UHP, amphibolite, and greenschist facies). We are using only the minerals included within garnet porhyroblasts and zircon for the estimation of P–T parameters of the UHP stage related to diamond formation.

Table 3.

P–T conditions of western and eastern domains of the Kokchetav Massif

We have shown previously that eclogites of the Kumdy-Kol area have equilibrium temperature of 800-950 °C at P minimum 18 kbar (Shatsky & Sobolev, 1985). The study of diamondiferous rocks gives us evidence that the pressure of metamorphism has been significantly higher.

The greatest difficulties arise in the estimation of pressure. Earlier we stressed that we can estimate only the lower limit of pressure, using the graphite-diamond equilibrium line (Sobolev & Shatsky, 1990; Shatsky et al., 1995). To substantiate higher pressures (70 kbar), the Japanese researchers (Kaneko et al., 2000) use experimental results by Luth (1997) on K impurity in pyroxene. But as shown by Harlow (1999), the partition coefficient of K between pyroxene and melt depends not only on pressure but also on melt composition. Therefore, at present there is no direct possibility to estimate pressures from K content in pyroxene.

At the same time, inclusions of magnesite, along with calcite and dolomite, found in zircons from the carbonate rock, can be evidence of pressures of more than 70 kbar (Sato & Katsura, 2001). The relationships of these phases remain, however, to be understood. Provided that zircons grow at both progressive and regressive stages of metamorphism (Hermann et al., 2001; Katayama et al., 2001), another reaction can be responsible for the appearance of magnesite, for example: En + Dol = Di + Mgs.

Fe-Mg partitioning between garnet and clinopyroxene, garnet and biotite (Hodges & Spear, 1981) or garnet and phengite (Green & Hellman, 1982) was used to calculate the temperatures at pressure 40 kbar for diamondiferous rocks and eclogites of Unit I. The calibration of Powell (1985) and Ellis & Green (1979) yields temperatures within the interval 760-1050 °C. For the majority of samples the estimation exceeds 950 °C. The lowest temperatures are obtained for gneisses (760-850 °C). As mentioned above, we suppose that pyroxene from the matrix recrystallised during the retrograde path. The equilibrium temperature obtained by using inclusions of omphacite in zircon is much higher (970 °C). It should be mentioned that a great variety of the garnet composition in zircons from dolomitic marble can indicate a major variation in the temperature during their crystallisation. The values obtained are between 670 and 945 °C. It gives us additional evidence that zircon may be formed at the prograde path of the metamorphic evolution within a wide interval of temperature.

The study of pyroxene-plagioclase symplectites, which developed after omphacite in eclogites of the Kumdy-Kol deposit, shows that these rocks experienced conditions of granulite facies metamorphism (T = 800−820 °C, P = 10 kbar; Shatsky & Sobolev, 1985). This is also confirmed by the investigation of inclusions in zircons from diamondiferous rocks (Hermann et al., 2001). The granulite facies conditions are also inferred from pyroxene-spinel symplectites developed around grossular-pyrope garnets in dolomite marbles as well as from finding of sapphirine and corundum in them (Sobolev et al., 2001). Newly formed pyroxene is characterised by a high content of the Ca-Tschermak component (up to 23%).

In the Kulet domain, the high-pressure rocks contain no diamonds, but coesite inclusions are established in them (Shatsky et al., 1998b; Parkinson, 2000). The equilibrium temperature of the eclogites from Unit II indicates that metamorphic rocks from this unit were metamorphosed at temperatures not higher than 650-800 °C (Table 3). The recent discovery of coesite as inclusions in garnet from garnet-mica schist of the Kulet area indicates that some blocks of Unit II were metamorphosed at pressures higher than 26 kbar (Shatsky et al., 1998b). The estimation of the equilibrium temperature of eclogite, garnet-amphibole-zoisite and garnet-kyanite-biotite- amphibole rocks show that the temperature of the UHP metamorphism did not exceed 800 °C. The presence of primary amphibole in eclogites indicate that pressure did not exceeded 27 kbar (Poli & Schmidt, 1995).

To summarise the available P–T data (Table 3), rocks exhumed from two levels can be recognised in the Kokchetav Massif. The first corresponds to diamondiferous rocks of the western block (P > 43 kbar, T ≈ 1000 °C). For the eclogites and country rocks of the eastern block the peak temperature is 650-800 °C at P = 26-28 kbar, and for amphibolites 650 °C at 10 kbar (Ota et al., 2000). There is a reason to think that amphibolites were formed during retrograde amphibolite facies (Shatsky & Sobolev, 1993; Hermann et al., 2001).

Geochemistry of metamorphic rocks and the age of the UHP metamorphism

Based on major element data we can conclude that the eclogite protoliths were rocks of basic composition related to the tholeiite series (Shatsky et al., 1993). Trace and REE elements data suggests that island arc or oceanic basalts can be considered as the protoliths of eclogites. As there are no rocks in the belt of UHP and HP rocks whose protoliths are andesites, island arc basalts can be excluded from consideration as possible protoliths of eclogites. When oceanic-type basalts are considered as protoliths of eclogites, according to the existing classification they are T-type basalts. This type of basalts occurs not only in mid-ocean ridges but also among products of basalt volcanism of passive margins.

The composition of diamondiferous metamorphic rocks varies widely. Shatsky et al. (1995) concluded that diamondiferous metamorphic rocks correspond to shales mixed with carbonate. Biotite-garnet kyanite schists and biotite schists from Unit ii and garnet-muscovite schists from Unit I have high Al2O3 contents and correspond to shale. Mica schists from Kulet (Unit II) have uncommonly high K2O and SiO2 contents; arkoses may be their protoliths. in most diamondiferous rocks the Th/U ratio is lower than in continental upper crust (Shatsky et al., 1999a). For the rocks from Unit II and garnet-muscovite schists and granito-gneisses from Kumdy-Kol this ratio is higher than in the continental upper crust.

The distribution of REE elements in gneisses and schists of Unit II and in granito-gnesses from Unit I are strongly fractionated and have a larger Eu anomaly than that of typical crustal rocks (Fig. 9). The diamondiferous metamorphic rocks are depleted in REE elements compared to upper crust compositions (Fig. 10). Based on the REE distribution patterns, the Kumdy-Kol rocks can be subdivided into 4 groups. The first group exhibits a depletion of the LREE ((La/Yb)N = 0.19) and a high Sm/Nd ratio (Fig. 10a), whereas the second group shows a variable abundance flat HREE pattern with a small Eu anomaly (Fig. 10b). Both groups are depleted (with respect to the average upper crust composition) in incompatible elements excluding Ti, K and Rb. The third rock group exhibits REE patterns similar to the second group but with a relatively large negative Eu anomaly (Eu*/Eu = 0.6, Fig. 10c). Compared to upper crust compositions these samples are depleted in REE, Ba, Ta and Th. The REE patterns of the fourth group are strongly fractionated and have a larger Eu anomaly (Eu*/Eu = 0.34) than that of typical crustal rocks (Fig. 10d). The lowest Sm/Nd ratios and highest Rb/Cs ratios are observed in these rocks. Rather numerous rock groups have characteristics intermediate between the groups III and IV, reflecting mixing. There is no correlation between bulk chemical composition and REE pattern. Only the group IV rocks have a rather narrow chemical compositional range including granite-gneisses and high-Mg garnet-muscovite-kyanite schists.

Fig. 9.

Chondrite-normalised REE patterns for granite-gneisses from Kulet and Sulu-Tyube. Thick line: upper crust.

Fig. 9.

Chondrite-normalised REE patterns for granite-gneisses from Kulet and Sulu-Tyube. Thick line: upper crust.

Fig. 10.

Chondrite-normalised REE patterns for selected groups of rocks. (a) REE abundance in group 1; (b) REE abundance in group 2; (c) REE abundance in group 3; (d) REE abundance in group 4; (e) REE pattern for trondhjemite.

Fig. 10.

Chondrite-normalised REE patterns for selected groups of rocks. (a) REE abundance in group 1; (b) REE abundance in group 2; (c) REE abundance in group 3; (d) REE abundance in group 4; (e) REE pattern for trondhjemite.

The gneisses and schists from Unit I and II have Sm-Nd model ages within the interval 2.08-2.65 Ga (Shatsky et al., 1999a). Zircon xenocrysts from the diamondiferous gneisses gave an age of 2 Ga (Claoue-Long et al., 1991). This age can be interpreted as the age of crust formation. The Sm–Nd model age of eclogites varies between wide limits (0.8–2.1 Ga) (Shatsky et al., 1993). Scattering in model ages of eclogites can be explained by the contamination of eclogites with Nd from country rocks during retrograde events. The Sm–Nd isochron diagrams for diamondiferous metamorphic rocks and eclogites show that whole rock, clinopyroxene and garnet do not form an isochron. This gives us additional evidence that Nd was exchanged in a later hydrothermal process. Mineral separates from the different types of rocks from Unit I and amphibole-garnet-zoisite rocks from Unit II (Kulet area) define a regression line with 524 ± 25 Ma. This is clear evidence that before UHP metamorphism all types of rocks in Units I and II, excluding eclogites, had a similar 143Nd/144Nd ratio. However deviations of the data from the isochron are far outside analytical error (MSWD-140). Mineral disequilibrium might cause a considerable error in Nd ages. At the same time this age is supported by a four-point mineral isochron for two high-temperature eclogites from Unit I (535 ± 3 Ma, Shatsky et al., 1999a) as well as by U–Pb dating of zircons (Claoue-Long et al., 1991; Hermann et al., 2001). The internal isochron for amphibole-garnet-zoisite rock (Kulet) gives 522 ± 27 Ma (MSWD-3.29). Whole rock analyses from Units I and II show a linear array which may represent a disturbed isochron. Diamond-bearing rocks show a considerable range of Sm/Nd ratio (0.205–0.962). The Sm/Nd ratios of metamorphic rocks from Unit II vary in a remarkably narrow interval (0.14–0.194) and do not differ significantly from the average value for continental crust (Shatsky et al., 1999a).

As mentioned above, matrix biotite and muscovite are found as late alteration products in diamondiferous gneisses of Kumdy-Kol. Ar–Ar age determination of micas from diamondiferous garnet-biotite gneisses yields an age of 517 Ma (Shatsky et al., 1999a).

High equilibrium temperatures of diamondiferous rocks suggest their partial melting (Shatsky et al., 1995; Hermann et al., 2001). Geochemical data confirm this supposition (Shatsky et al., 1995, 1999a). Diamondiferous rocks demonstrate a considerable scatter in Sm/Nd values (from 0.2 to 0.96; Shatsky et al., 1999a). The isotope data show that prior to high pressure metamorphism all varieties of rocks, with the exception of eclogites, had similar 143Nd/144Nd ratios (ϵNd value of −13.3). Confirmation of partial melting is received from distribution of rare earth elements. Diamondiferous rocks are depleted with respect to light REE ((La/Yb)N = 0,19–1). Trondhjemite, which cuts an eclogite body and differs from the other granite rocks by a high Na/K ratio, has a REE distribution consistent with the equilibrium of the melt with a garnet-enriched restite. The HREE depletion suggests that trondhjemite is a result of the partial melting of eclogite (Shatsky et al., 1999a).

As said above, diamondiferous rocks are interlayered with granite gneisses, and migmatised garnet-biotite gneisses are observed. It could be supposed that some part of granite gneisses and migmatite have been formed by the melting of diamondiferous rocks (Shatsky et al., 1999a).

Exhumation of high pressure rocks of the Kokchetav Massif

In discussing exhumation models, the plotting of a regressive P–T path play a crucial role (Fig. 11). But to plot exhumation paths is rather difficult, because in many cases only fragments of this trend are documented. Thus, the granulite facies stage of metamorphism is recorded only in some specimens of eclogites and Ca-silicate rocks. On the basis of high-Mg calcite present in pyroxenes of carbonate rocks, Ogasawara et al. (2000) distinguished a step with T > 800 °C and P ≈ 25 kbar at the regressive stage of metamorphism (stage 2, Fig. 11.). Thus, at the first stage of exhumation diamondiferous rocks should be cooled from 1000 °C to 800 °C with pressure decreased from > 42 to 25 kbar. The next stage corresponds to the isothermal uplift to pressures of about 10 kbar, with the subsequent nearly isobaric cooling to 600-650 °C (stage 3, Fig. 11). Final stages of exhumation correspond to greenschist facies metamorphism (stage 5, Fig. 11).

Fig. 11.

P–T retrograde path of Kumdy-Kol diamondiferous rocks. Calculated burial and exhumation P–T paths from Ernst & Peacock (1996); graphite-diamond transition after Bundy (1980); quartz-coesite transition after Mirwald & Massonne (1980); Bt-out, Opx-in curves after Vielzeuf & Montel (1994); Phl + Qtz = En + Sa + M after Vielzeuf & Clemens (1992). 1) biotite stability in KCMASH system (Hermann & Green, 2001); 2) biotite stability in metapelites (Vielzeuf & Holloway, 1988); the grey bend shows the field of phengite melting (Hermann & Green, 2001).

Fig. 11.

P–T retrograde path of Kumdy-Kol diamondiferous rocks. Calculated burial and exhumation P–T paths from Ernst & Peacock (1996); graphite-diamond transition after Bundy (1980); quartz-coesite transition after Mirwald & Massonne (1980); Bt-out, Opx-in curves after Vielzeuf & Montel (1994); Phl + Qtz = En + Sa + M after Vielzeuf & Clemens (1992). 1) biotite stability in KCMASH system (Hermann & Green, 2001); 2) biotite stability in metapelites (Vielzeuf & Holloway, 1988); the grey bend shows the field of phengite melting (Hermann & Green, 2001).

In addition to the P–T regressive trend, the following established facts must be taken into account in a model for exhumation. On the basis of SHRIMP dating of zircons (Claoue-Long et al., 1991), Sm-Nd dating of eclogites (535 ± 3 Ma), zircon domains containing diamond indusions from diamondiferous metapelites (530 ± 7 Ma) and dating of secondary biotite and muscovite from diamondiferous gneiss (517 ± 5 Ma), it was concluded that the time interval between high pressure metamorphism and the uplift of rocks to a crustal level could be no more than 10 Ma and the vertical rate of exhumation was not less than 1.2 cm/year (Dobretsov et al., 1995; Shatsky et al., 1999a). New ages for zircons from metamorphic diamondiferous rocks of the Kumdy-Kol and Barchi sites have been published recently (Hermann et al., 2001). Dating from separate zones of zircons containing mineral inclusions of both high pressure and retrograde stages of metamorphism shows that there is no systematic difference in ages between different domains in the zircons that belong to different stages of metamorphism. It means that the period of exhumation of high pressure rocks to depths corresponding to a pressure of 6-8 kbar is within the accuracy of the method, no more than 6 Ma, which connotes an exhumation rate of 1.8 cm/year. Additional evidence of rapid exhumation is a low degree of nitrogen aggregation in diamonds of metamorphic rocks. According to de Corte et al. (1999), these diamonds could not exist at 950 °C for more than 5 Ma and at 1000 °C for more than 0.1 Ma. Worthy of note is that there is some uncertainty in activation energy for nitrogen in diamond to aggregate from Ib-to-IaA. Using energies of activation of 5 eV, Finnie et al. (1994) obtained residence time of 0.2 Ma for 900 °C.

Thus, the estimates obtained in different ways indicate high rates of the first stage of exhumation of the Kokchetav high pressure rocks. The time interval between the climax of metamorphism and exhumation to the level of the Earth’s crust does not exceed 6 Ma. A decrease in temperature from 1000 to 800 °C occurred during the time interval of no more than 2-5 Ma, as inferred from the degree of nitrogen aggregation in diamond and zoning in garnets.

Structural studies showed that the deformation of diamondiferous rocks of the Kumdy-Kol deposit is insignificant (Theunissen et al., 2000a). The diamond inclusions in garnet are often intergrown with mica crystals carrying no traces of deformation. The low differential stress is supported by the fact that garnet aggregates in massive garnet-pyroxene-quartz rocks have foam structure indicating that their microstructure was controlled by grain boundary free energy. We believe all these facts could be explained by the partial melting of either metapelites or granitic rocks in the Kumdy-Kol domain.

The occurrence of the melt is responsible for an essential reduction of viscosity, and a density difference (Δρ) of crustal rocks and mantle material, and reduced friction between the upwelling sialic block, subducting and overriding plates.

Besides Δρ, the rate of exhumation seems to depend on the internal pressure in the subducting continental crustal block (Dobretsov & Kirdyashkin, 1992, 1998), which can be regarded as a viscous layer between subducting continental lithosphere and surrounding mantle.

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Figures & Tables

Fig. 1.

Major tectonic features of Central Asia for the region surrounding and including the Kokchetav massif, with the location of Figure 2 shown (Dobretsov et al., 1999a).

Fig. 1.

Major tectonic features of Central Asia for the region surrounding and including the Kokchetav massif, with the location of Figure 2 shown (Dobretsov et al., 1999a).

Fig. 2.

The Kokchetav Megamelange with tectonic units 1-3 and adjacent domains I-V (Dobretsov et at., 1998). 1: ultrahigh pressure/HP unit with high temperature eclogites and diamond-bearing rocks (Kumdy-Kol) and with whiteschist, coesite-bearing micaschist and relatively low temperature eclogites (Kulet); 2: medium pressure unit with Al-rich metasediments and coronite (Enbek-Berlyk); 3 : low pressure unit (Daulet). Domain I: Neoproterozoic sequences mainly belonging to the Vendian-Early Cambrian island arc; Domain II: “Kokchetav Microcontinent Domain” is composed of several blocks. In various amounts each of this blocks includes: a) a gneissic basement, b) its sedimentary cover (black shale and dolomite in the lower part and metasandstone in the upper part, c) tectonic slices of Vendian-Early Cambrian (?) ophiolite and volcanics, d) Ordovician rocks and e) Devonian granites; Domain III: the “Megamelange Domain”; Domain IV: the “White Lake Domain” - amphibolites, amphibole schist and quartzites, assumed to represent fragments of an oceanic crust; Domain V: the “Granite Dome Domain” is polyphase and composed of the Imantau and Zerenda granite domes with Cambrian gabbro, Late Ordovician-Silurian diorite, granodiorite and granite, and Devonian leucocratic granite and granosyenite.

Fig. 2.

The Kokchetav Megamelange with tectonic units 1-3 and adjacent domains I-V (Dobretsov et at., 1998). 1: ultrahigh pressure/HP unit with high temperature eclogites and diamond-bearing rocks (Kumdy-Kol) and with whiteschist, coesite-bearing micaschist and relatively low temperature eclogites (Kulet); 2: medium pressure unit with Al-rich metasediments and coronite (Enbek-Berlyk); 3 : low pressure unit (Daulet). Domain I: Neoproterozoic sequences mainly belonging to the Vendian-Early Cambrian island arc; Domain II: “Kokchetav Microcontinent Domain” is composed of several blocks. In various amounts each of this blocks includes: a) a gneissic basement, b) its sedimentary cover (black shale and dolomite in the lower part and metasandstone in the upper part, c) tectonic slices of Vendian-Early Cambrian (?) ophiolite and volcanics, d) Ordovician rocks and e) Devonian granites; Domain III: the “Megamelange Domain”; Domain IV: the “White Lake Domain” - amphibolites, amphibole schist and quartzites, assumed to represent fragments of an oceanic crust; Domain V: the “Granite Dome Domain” is polyphase and composed of the Imantau and Zerenda granite domes with Cambrian gabbro, Late Ordovician-Silurian diorite, granodiorite and granite, and Devonian leucocratic granite and granosyenite.

Fig. 3.

Detailed geological map of diamond-bearing metasedimentary rocks mined, drilled and trenched (Scheshkel et al.,pers. comun.). 1: granite-gneisses; 2: biotite gneisses; 3: granite-gneisses and gneisses alternation; 4: fine grained chlorite-tremolite-quartz rocks; 5: migmatites; 6: garnet-muscovite, kyanite-muscovite schists; 7: pyroxene-carbonate rocks; 8: garnet pyroxenites; 9: eclogites and amphibolites; 10: dykes of diorite porphyrites.

Fig. 3.

Detailed geological map of diamond-bearing metasedimentary rocks mined, drilled and trenched (Scheshkel et al.,pers. comun.). 1: granite-gneisses; 2: biotite gneisses; 3: granite-gneisses and gneisses alternation; 4: fine grained chlorite-tremolite-quartz rocks; 5: migmatites; 6: garnet-muscovite, kyanite-muscovite schists; 7: pyroxene-carbonate rocks; 8: garnet pyroxenites; 9: eclogites and amphibolites; 10: dykes of diorite porphyrites.

Fig. 4.

Compositional variations of garnet from the metamorphic rocks of Kumdy-Kol. 1: diamondiferous gneiss, garnet-pyroxene-quartz rocks and schist; 2: garnet-pyroxene rocks; 3: dolomitic marbles; 4: diamond-free gneiss; 5: eclogite; 6: inclusions in zircons from dolomitic marble (K9-16); 7: garnet from matrix (K9-16); 8: diamondiferous garnet-kyanite-muscovite schist.

Fig. 4.

Compositional variations of garnet from the metamorphic rocks of Kumdy-Kol. 1: diamondiferous gneiss, garnet-pyroxene-quartz rocks and schist; 2: garnet-pyroxene rocks; 3: dolomitic marbles; 4: diamond-free gneiss; 5: eclogite; 6: inclusions in zircons from dolomitic marble (K9-16); 7: garnet from matrix (K9-16); 8: diamondiferous garnet-kyanite-muscovite schist.

Fig. 5.

Compositional variation of garnet from zircons. 1: dolomitic marble; 2: biotite plagiogneiss; 3: garnet-pyroxene rocks; 4: pyroxene-biotite gneiss; 5: migmatites; 6: garnet from matrix of all rock types.

Fig. 5.

Compositional variation of garnet from zircons. 1: dolomitic marble; 2: biotite plagiogneiss; 3: garnet-pyroxene rocks; 4: pyroxene-biotite gneiss; 5: migmatites; 6: garnet from matrix of all rock types.

Fig. 6.

Geological map of the Sulu-Tyube area based on geophysical, drilling and structural data (compiled by Dobretsov & Zayachovsky; Dobretsov et al, 1999b). 1 : traced boundaries of eclogites; 2 : blocks of fresh eclogites corresponding to highest density geophysical bodies; 3 : the LP Daulet unit; 4: small eclogite bodies; 5 : structural lines; 6: faults visible (a) and supposed under quaternary sediment (b).

Fig. 6.

Geological map of the Sulu-Tyube area based on geophysical, drilling and structural data (compiled by Dobretsov & Zayachovsky; Dobretsov et al, 1999b). 1 : traced boundaries of eclogites; 2 : blocks of fresh eclogites corresponding to highest density geophysical bodies; 3 : the LP Daulet unit; 4: small eclogite bodies; 5 : structural lines; 6: faults visible (a) and supposed under quaternary sediment (b).

Fig. 7.

Structural geological sketch map of the Enbek-Berlik area. 1 : amphibolised dolerite; 2: amphibolite; 3: eclogite and amphibolised eclogite; 4: pyroxenite and harzburgite; 5: coronite; 6: diopside-plagioclase rock; 7: metapelites and the strike of their schistosity; 8: granite; 9: boundary of recent deposits; 10: fault; 11: dip of bedding (Reverdatto, 1999).

Fig. 7.

Structural geological sketch map of the Enbek-Berlik area. 1 : amphibolised dolerite; 2: amphibolite; 3: eclogite and amphibolised eclogite; 4: pyroxenite and harzburgite; 5: coronite; 6: diopside-plagioclase rock; 7: metapelites and the strike of their schistosity; 8: granite; 9: boundary of recent deposits; 10: fault; 11: dip of bedding (Reverdatto, 1999).

Fig. 8.

Detailed geological map of the central part of the South Zheltau (Kulet) area (Dobretsov et al., 1999b).

Fig. 8.

Detailed geological map of the central part of the South Zheltau (Kulet) area (Dobretsov et al., 1999b).

Fig. 9.

Chondrite-normalised REE patterns for granite-gneisses from Kulet and Sulu-Tyube. Thick line: upper crust.

Fig. 9.

Chondrite-normalised REE patterns for granite-gneisses from Kulet and Sulu-Tyube. Thick line: upper crust.

Fig. 10.

Chondrite-normalised REE patterns for selected groups of rocks. (a) REE abundance in group 1; (b) REE abundance in group 2; (c) REE abundance in group 3; (d) REE abundance in group 4; (e) REE pattern for trondhjemite.

Fig. 10.

Chondrite-normalised REE patterns for selected groups of rocks. (a) REE abundance in group 1; (b) REE abundance in group 2; (c) REE abundance in group 3; (d) REE abundance in group 4; (e) REE pattern for trondhjemite.

Fig. 11.

P–T retrograde path of Kumdy-Kol diamondiferous rocks. Calculated burial and exhumation P–T paths from Ernst & Peacock (1996); graphite-diamond transition after Bundy (1980); quartz-coesite transition after Mirwald & Massonne (1980); Bt-out, Opx-in curves after Vielzeuf & Montel (1994); Phl + Qtz = En + Sa + M after Vielzeuf & Clemens (1992). 1) biotite stability in KCMASH system (Hermann & Green, 2001); 2) biotite stability in metapelites (Vielzeuf & Holloway, 1988); the grey bend shows the field of phengite melting (Hermann & Green, 2001).

Fig. 11.

P–T retrograde path of Kumdy-Kol diamondiferous rocks. Calculated burial and exhumation P–T paths from Ernst & Peacock (1996); graphite-diamond transition after Bundy (1980); quartz-coesite transition after Mirwald & Massonne (1980); Bt-out, Opx-in curves after Vielzeuf & Montel (1994); Phl + Qtz = En + Sa + M after Vielzeuf & Clemens (1992). 1) biotite stability in KCMASH system (Hermann & Green, 2001); 2) biotite stability in metapelites (Vielzeuf & Holloway, 1988); the grey bend shows the field of phengite melting (Hermann & Green, 2001).

Table 1.

Representative analyses of minerals from pyroxene-carbonate rock 94–275

NameCpx1Lam1Lam2Gt1Gt2IncInc2Cpx2
corerim
SiO253.9040.3040.3040.3040.7041.0053.8038.6053.20
TiO20.220.280.220.360.280.330.230.730.22
Al2O31.8621.5021.6021.1021.4021.401.8617.301.76
Cr2O30.020.010.040.020.020.020.000.000.02
FeO2.0610.308.5510.2010.4010.902.414.322.03
MnO0.050.670.540.590.580.550.030.010.05
MgO16.4010.1012.109.9810.109.6116.6022.6016.00
CaO24.9015.4013.7015.6016.0015.9024.800.0425.00
Na2O0.120.000.000.000.000.010.170.060.12
K2O0.000.000.000.000.010.010.009.350.00
Total99.4898.5097.0098.1299.4399.6499.8993.0698.37
Si1.9673.0153.0223.0253.0233.0371.9592.7891.967
Ti0.0060.0160.0120.0200.0150.0180.0060.0400.006
Al0.0801.8951.9131.8691.8721.8660.0801.4750.077
Cr0.0000.0010.0020.0010.0010.0010.0000.0000.000
Fe2+0.0630.5660.4700.5380.5400.5690.0680.063
Fe3+0.0000.0820.0670.1010.1030.1060.0060.000
Mn0.0010.0430.0340.0380.0360.0350.0010.0010.002
Mg0.8901.1291.3511.1171.1121.0620.8992.4290.882
Ca0.9731.2341.1021.2581.2701.2640.9680.0030.988
Na0.0080.0000.0000.0000.0000.0010.0120.0090.009
K0.0000.0000.0000.0000.0000.0000.0000.8610.000
Fe tot0.0630.6500.5380.6420.6460.6780.0730.2610.063
Total3.9907.9807.9747.9687.9737.9593.9977.8673.993
f6.6036.5528.4936.5036.7338.987.559.706.64
NameCpx1Lam1Lam2Gt1Gt2IncInc2Cpx2
corerim
SiO253.9040.3040.3040.3040.7041.0053.8038.6053.20
TiO20.220.280.220.360.280.330.230.730.22
Al2O31.8621.5021.6021.1021.4021.401.8617.301.76
Cr2O30.020.010.040.020.020.020.000.000.02
FeO2.0610.308.5510.2010.4010.902.414.322.03
MnO0.050.670.540.590.580.550.030.010.05
MgO16.4010.1012.109.9810.109.6116.6022.6016.00
CaO24.9015.4013.7015.6016.0015.9024.800.0425.00
Na2O0.120.000.000.000.000.010.170.060.12
K2O0.000.000.000.000.010.010.009.350.00
Total99.4898.5097.0098.1299.4399.6499.8993.0698.37
Si1.9673.0153.0223.0253.0233.0371.9592.7891.967
Ti0.0060.0160.0120.0200.0150.0180.0060.0400.006
Al0.0801.8951.9131.8691.8721.8660.0801.4750.077
Cr0.0000.0010.0020.0010.0010.0010.0000.0000.000
Fe2+0.0630.5660.4700.5380.5400.5690.0680.063
Fe3+0.0000.0820.0670.1010.1030.1060.0060.000
Mn0.0010.0430.0340.0380.0360.0350.0010.0010.002
Mg0.8901.1291.3511.1171.1121.0620.8992.4290.882
Ca0.9731.2341.1021.2581.2701.2640.9680.0030.988
Na0.0080.0000.0000.0000.0000.0010.0120.0090.009
K0.0000.0000.0000.0000.0000.0000.0000.8610.000
Fe tot0.0630.6500.5380.6420.6460.6780.0730.2610.063
Total3.9907.9807.9747.9687.9737.9593.9977.8673.993
f6.6036.5528.4936.5036.7338.987.559.706.64
Table 2.

Representative analyses of minerals from pyroxene-carbonate rock K 98-6

NameGrtIncCpx1Lam1Lam2Cpx2Mica
corerimcorerim
SiO239.3039.6054.1053.8054.0051.1063.5054.7039.20
TiO20.400.380.000.000.001.010.000.001.18
Al2O320.8021.302.111.620.8428.2916.801.8415.00
Cr2O30.020.010.000.000.000.000.000.000.00
FeO6.557.002.903.013.121.080.112.0010.40
MnO0.991.080.140.120.100.020.000.070.30
MgO3.373.8215.9016.2016.403.910.0416.4918.50
CaO27.6026.3024.3025.0024.500.240.2024.870.35
Na2O0.050.040.360.340.170.040.070.410.06
K2O0.010.020.360.050.545.9214.80.088.69
Total99.0499.62100.27100.0499.7091.5995.42100.4693.68
Si2.9922.9961.9691.9601.9823.4203.0511.8942.888
Ti0.0230.0220.0000.0000.0000.0510.0000.0240.065
Al1.8661.8990.0900.0700.0362.2290.9491.1141.307
Cr0.0010.0000.0000.0000.0000.0000.0000.0000.000
Fe2+0.3160.3700.0880.0660.0870.034
Fe3+0.1010.0730.0000.0250.0090.000
Mn0.0640.0690.0040.0040.0030.0010.0000.0000.019
Mg0.3820.4300.8640.8800.8960.3900.0030.2602.030
Ca2.2512.1340.9480.9750.9630.0170.0100.0680.028
Na0.0080.0060.0250.0240.0120.0050.0060.0030.009
K0.0000.0010.0170.0020.0250.5050.9070.2580.817
Fe tot0.4190.4440.0880.0920.0960.0600.0040.0340.642
Total8.0047.9994.0064.0054.0146.6744.9313.6567.806
f52.2750.779.279.479.6513.4260.8511.6224.04
NameGrtIncCpx1Lam1Lam2Cpx2Mica
corerimcorerim
SiO239.3039.6054.1053.8054.0051.1063.5054.7039.20
TiO20.400.380.000.000.001.010.000.001.18
Al2O320.8021.302.111.620.8428.2916.801.8415.00
Cr2O30.020.010.000.000.000.000.000.000.00
FeO6.557.002.903.013.121.080.112.0010.40
MnO0.991.080.140.120.100.020.000.070.30
MgO3.373.8215.9016.2016.403.910.0416.4918.50
CaO27.6026.3024.3025.0024.500.240.2024.870.35
Na2O0.050.040.360.340.170.040.070.410.06
K2O0.010.020.360.050.545.9214.80.088.69
Total99.0499.62100.27100.0499.7091.5995.42100.4693.68
Si2.9922.9961.9691.9601.9823.4203.0511.8942.888
Ti0.0230.0220.0000.0000.0000.0510.0000.0240.065
Al1.8661.8990.0900.0700.0362.2290.9491.1141.307
Cr0.0010.0000.0000.0000.0000.0000.0000.0000.000
Fe2+0.3160.3700.0880.0660.0870.034
Fe3+0.1010.0730.0000.0250.0090.000
Mn0.0640.0690.0040.0040.0030.0010.0000.0000.019
Mg0.3820.4300.8640.8800.8960.3900.0030.2602.030
Ca2.2512.1340.9480.9750.9630.0170.0100.0680.028
Na0.0080.0060.0250.0240.0120.0050.0060.0030.009
K0.0000.0010.0170.0020.0250.5050.9070.2580.817
Fe tot0.4190.4440.0880.0920.0960.0600.0040.0340.642
Total8.0047.9994.0064.0054.0146.6744.9313.6567.806
f52.2750.779.279.479.6513.4260.8511.6224.04

Note: Cpx1 - clinopyroxene contains lamellae of phengite

Table 3.

P–T conditions of western and eastern domains of the Kokchetav Massif

Contents

GeoRef

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