Petrogenetic constraints for the genesis of Archaean sanukitoid suites: geochemistry and isotopic evidence from Karelia, Baltic Shield

Lithos 79 (2005) 147 – 160 www.elsevier.com/locate/lithos

Petrogenetic constraints for the genesis of Archaean sanukitoid suites: geochemistry and isotopic evidence from Karelia, Baltic Shield A. Kovalenkoa,b, J.D. Clemensb,*, V. Savatenkova a Institute of Precambrian Geology and Geochronology, RAS, Makarova, 2, 199034, St-Petersburg, Russia School of Earth Sciences and Geography, CEESR, Kingston University, Penrhyn Rd, Kingston-upon-Thames, Surrey KT1 2EE, UK

b

Received 25 September 2003; accepted 2 September 2004 Available online 27 October 2004

Abstract New Sm–Nd isotope data are presented for late Archaean sanukitoid suites of the Karelian granite-greenstone terrane of the Baltic Shield. Within the terrane, there are regional variations in the Nd isotope compositions of the sanukitoids. Sanukitoids in the youngest Central Karelian domain have Nd isotope characteristics similar to depleted mantle, whereas those from the older West Karelian and Vodlozero domains have lower initial e Nd values. The isotopic heterogeneity of the sanukitoids is interpreted to result from variations in the time interval between mantle source enrichment and the partial melting that produced the sanukitoid magmas. We therefore suggest a two-stage model of the generation of sanukitoid magmas. During first stage, the mantle was metasomatised by fluids and/or melts generated during subduction or tectonic underplating. Later, between 2.74 and 2.70 Ga, a tectonothermal anomaly generated the sanukitoid magmas by melting of the previously metasomatised mantle. Most of the sanukitoid intrusions are cut by calc-alkaline lamprophyre dykes that have the similar geochemical signatures to their sanukitoid wall rocks. New Sm–Nd isotope data for these lamprophyres suggest a genetic kinship between the sanukitoids and the lamprophyres. Comparison between chemical signatures of sanukitoids and modern, subduction-related magmas shows that Archaean sanukitoids actually have no modern analogues. Sanukitoid melts can be modelled as partial melts of mantle that had been metasomatised a short time previously. The melts could have been in equilibrium with residues containing Ol, Cpx plus minor Phl and Grt, at a pressure of about 2.5 GPa and a temperature of 1150–1230 8C. The small percentage of residual phlogopite would retain Rb in the mantle and explain the peculiar signature of low-Rb and high Ba and Sr, characteristic of the sanukitoids. We envisage a temporal evolution of the composition of the mantle in this region. This pattern of evolution controlled production of the sanukitoid melts and explains why the sanukitoids have such a restricted range of ages. Sanukitoid magma production must have followed from a major, global mantle enrichment event. These magmas could not have been produced earlier because the Archaean mantle had not become sufficiently enriched before 2.7 Ga. We suggest that their scarcity after the * Corresponding author. Tel.: +44 20 8547 7023; fax: +44 20 8547 7497. E-mail address: [email protected] (J.D. Clemens). 0024-4937/$ - see front matter D 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.lithos.2004.05.006

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major period of sanukitoid magmatism may be due to the increased thickness of overlying crust, which can inhibit mantle enrichment. D 2004 Elsevier B.V. All rights reserved. Keywords: Archaean; Baltic shield; Sanukitoid; Sm–Nd; Rb–Sr

1. Introduction Archaean sanukitoid intrusions have been intensively investigated over the last 20 years and have been discovered in the Archaean shield of Canada (Superior Province; Stern and Hanson, 1991; Stevenson et al., 1999), the Baltic and Ukrainian shields (O’Brien et al., 1993; Chekulaev, 1999; LobachZhuchenko et al., 2000a; Kovalenko and Savatenkov, 2003; Lobach-Zhuchenko et al., 2003; Chekulaev et al., 2003; Artemenko et al., 2003; Halla, 2003), Australia (Pilbara Craton; Smithies and Champion, 2000), India (Dharwar Craton; Moyen et al., 2001) and South Africa (Zimbabwe Craton; Bagai et al., 2002). Sanukitoids are K-rich, high-Mg, intrusive rocks. When differentiated, the intrusions contain rocks that vary from ultramafic to felsic. They have some geochemical characteristics similar to both mantle- and crust-derived magmatic rocks. These features include high Cr and Ni (up to 400 ppm and up to 250 ppm, respectively) accompanied by LREE enrichment with steeply fractionated REE patterns, Ba and Sr N1000 ppm, P2O5 up to 1.5 wt.%, and Rb, Nb and Ti depletion. Sm–Nd and Rb–Sr isotope data show that there has been major juvenile input into sanukitoid magmas. Worldwide, these rocks form late Archaean (2.95–2.54 Ga), post-tectonic intrusions that are spatially and temporally associated with shear zones. The high Mg/(Mg+Fe) of sanukitoids (N0.6) and their enrichment in Cr and Ni suggest that the sanukitoid source region was peridotitic rather than basaltic. This inference is supported by Nd isotope data, which suggests a major juvenile input into sanukitoid magmas. On the other hand, the high LREE, Ba, Sr and P in sanukitoids cannot be caused by crustal contamination because the average Archaean crust has lower concentrations of these elements than the sanukitoids (Taylor and McLennan, 1985). Some researchers attempt to explain the LREE, Ba, Sr

and P enrichment by the addition of these elements through fluid- or melt-mediated metasomatism of the mantle source region, in a subduction environment (Stern and Hanson, 1991; Hattory et al., 1996; Shimoda et al., 1998). Another idea is that this enrichment is attributable to extensive metasomatism of the mantle source region, by addition of about 40% of a TTG-like melt, during tectonic underplating (Smithies and Champion, 2000). In this study, we focus on Sm–Nd isotope data for the sanukitoids of Karelia (Baltic Shield) and suggest petrogenetic models for their genesis.

2. Geological setting The Karelian granite-greenstone terrane is a part of the Baltic Shield (Fig. 1), which is largely composed of greenstone belts and rocks of the tonalite–trondhjemite–granodiorite (TTG) series. Three groups of volcano-sedimentary belts have been recognised within Karelia. Two groups of older greenstone belts (3.02–2.93 and 2.91–2.85 Ga) comprise mainly ultramafic, mafic and intermediate volcanic rocks, with minor amounts of felsic volcanic and metasedimentary rocks. The youngest belt (b2.8 Ga) is mainly composed of felsic volcanic and sedimentary rocks, and has been termed the dmetasedimentary beltT (Chekulaev et al., 2004). Based on geological, geochemical and isotopic characteristics, five major stages of Archaean granitoid magmatism have been distinguished within the Karelian granite-greenstone terrane. These have the following ages: (1) 3.2–3.1, (2) 2.9–2.85, (3) 2.8–2.75, (4) 2.74–2.70 and (5) 2.7 Ga (Lobach-Zhuchenko et al., 2000a). During the first three magmatic events, predominantly TTG rocks were formed, while the fifth stage was characterised by potassic granitoid magmatism. Sanukitoid magmas were intruded during the relatively narrow time interval of the fourth magmatic stage (2.74–2.70 Ga).

A. Kovalenko et al. / Lithos 79 (2005) 147–160

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Fig. 1. Geological map of the Karelian granite-greenstone terrane, showing the locations of the sanukitoid intrusions and the ranges of their initial e Nd values. Average T NDDM values for the crustal rocks are from Chekulaev et al. (1997, 2002) and Lobach-Zhuchenko et al. (2000b).

Partly on the basis of Sm–Nd isotope data, the Karelian granite-greenstone terrane has been previously divided into three domains (Fig. 1), each with

its separate history of crustal growth (LobachZhuchenko et al., 2000b). The ancient Vodlozero domain is composed of rocks with model ages

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A. Kovalenko et al. / Lithos 79 (2005) 147–160 0.8

12

a Na2O+K2O, wt.%

0.7

mg#

0.6 0.5 0.4 0.3

Karelian sanukitoids

d

10 8 6 4

0.2 2

0.1

Typical Karelian TTG-series

0.0

0 40

50

60

70

80

40

50

SiO2, wt.%

60

70

80

SiO2, wt.%

2500

3500

b

e

3000

Ba, ppm

Sr, ppm

2000 1500 1000

2500 2000 1500 1000

500

500 0

0 40

50

60

70

40

80

50

SiO2, wt.% 2

70

450

c

1.8 1.6

80

f

400 350

1.4 1.2

Cr, ppm

P2O5, wt.%

60

SiO2, wt.%

1 0.8 0.6

300 250 200 150 100

0.4 0.2

50 0

0 40

50

60

70

40

80

50

SiO2, wt.% 1000.0

rock/chondrites

60

70

80

SiO2, wt.%

g

100.0

10.0

1.0

0.1 La Ce

Nd

SmEu

Tb

Yb Lu

Fig. 2. (a–f ) Harker variation diagrams and (g) chondrite-normalised REE plot comparing the chemistry of the Karelian sanukitoids to the Karelian TTG series. Data for TTG series from Lobach-Zhuchenko et al. (2004a). See text for discussion.

A. Kovalenko et al. / Lithos 79 (2005) 147–160

(TDMNd)N3 Ga. The younger, West and Central Karelian domains have TDMNd of 2.9–2.8 and b2.8 Ga, respectively. Sanukitoids occur mainly in youngest Central Karelian domain, but a few also intruded the older West Karelian domain and the western part of the Vodlozero domain. As for sanukitoids worldwide, the Karelian examples are also spatially and temporally shear zones, and mainly intrude metasedimentary belts. It should be noted that the metasedimentary belts only occur in the youngest Central Karelian domain, where most of the sanukitoids occur (Fig. 1). Geochemical features of the Karelian sanukitoids are shown in Fig. 2. The SiO2 contents range from ~40 to 72 wt.%. Over this great range of SiO2, Mg/ (Mg+Fe) remains nearly constant; specifically, it does not decrease with increasing of SiO2 (Fig. 2a). The high concentrations of Sr and Ba (up to 2000 or 3000 ppm), P2O5 and alkalis (Fig. 2b–f), and the highly fractionated REE patterns and LREE enrichment (Fig. 2g) show that these rocks are true sanukitoids. Some of the sanukitoid intrusions in this region are zoned and differentiated, with rock compositions ranging from ultramafic to quartz monzonitic; the felsic rocks form the cores of the intrusions (LobachZhuchenko et al., 2000a, 2003; Chekulaev et al., 2003). We suggest that this variation may be due to crystal fractionation of a monzogabbroic to monzodioritic parent magma (Lobach-Zhuchenko et al., this volume). Other intrusions contain more felsic rocks— granodiorites, quartz syenites or quartz diorites. This difference may be due to variations in the level of erosion; mafic rocks may well be present at depth in the felsic sanukitoid bodies. Most of the sanukitoid intrusions are cut by calc-alkaline lamprophyre dykes of uncertain age. Similar geochemical signatures and close spatial relationships with the sanukitoids suggest the possibility of a petrogenetic kinship between the sanukitoids and these lamprophyres (Lobach-Zhuchenko et al., 2000a).

3. Analytical methods Nd isotope analyses were obtained at the Institute of Precambrian Geology and Geochronology, Russian Academy of Sciences, in St. Petersburg. Before dissolution with HF and HNO3 mixtures, rock powders

151

were spiked with an isotopically enriched tracer of 150 Nd–149Sm. REE were separated, by cation exchange resins, from these solutions, and Sm and Nd were separated using HDEHP on Teflon power. Samples were analysed for Sm and Nd concentrations and 143 Nd/144Nd ratios on a Finnigan MAT-261 multicollector mass spectrometer equipped with eight cups. Total blanks were 0.05 ng for Sm and 0.1 ng for Nd. The results were normalised to 148Nd/144Nd=0.24157 and corrected to 143Nd/144Nd=0.511860 in the La Jolla isotope standard. The analytical results for the Karelian sanukitoids are presented in Table 1. The e Nd(t) calculations used present-day CHUR ratios of 143Nd/144Nd=0.512638 and 147Sm/144Nd=0.1967, and appropriate U–Pb zircon ages. Nd model ages (TDMNd) were calculated using the depleted mantle model of DePaolo (1981), which can be approximated by a second-order polynomial fit: e Nd(t)=0.25t2 3t+8.5 (where t is the age in Ma). It should be noted that the depleted mantle model of Na¨gler and Kramers (1998) (e Nd(t)= 0.164t 3 0.566t 2 2.79t+10.4) is similar to the DePaolo model, within the same time interval (~2.8–2.7 Ga). Thus, the calculated T DM Nd is independent of the depleted mantle model used. CSm/Nd is the enrichment factor of 147Sm to 144Nd in a sample relative to CHUR and was calculated as (147Sm/144Nd)/(147Sm/144NdCHUR) 1.

4. Results Fifty-five Sm–Nd isotope analyses of the 12 Karelian sanukitoid intrusions, and related lamprophyre dykes, are presented in Table 1 and Fig. 3. There is a regional pattern in the distribution of Nd isotope compositions. The initial e Nd values range from +0.7 to +2.1 in the Central Karelian domain (Fig. 3a), while the West Karelian and Vodlozero intrusions have initial e Nd in the range 1.3 to +0.7. The sanukitoid intrusions of Central Karelian domain have Nd isotope characteristics similar to the depleted mantle (DePaolo, 1981; Na¨gler and Kramers, 1998), whereas those of other parts of Karelia have lower initial e Nd values. The narrow Nd isotopic compositional range within individual intrusions indicates closed-system behaviour, for the Sm– Nd system at least (Table 1). It should be noted that

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Table 1 Sm–Nd isotopic data for the Karelian sanukitoid intrusions and spatially related lamprophyre dykes N

Rock

Central Karelian domain 222 granodiorite 256-1 Qz-monzonite 323 Qz-monzonite 205 Qz-monzonite 242 granodiorite 216-a granite 256 granodiorite 8-4 ultramafic 238 ultramafic 8-5 ultramafic 8-6 monzogabbro 208 monzodiorite 177a diorite 3 monzodiorite 207/b-2 monzodiorite 573-2 monzodiorite 174 monzodiorite 156 Qz-monzonite 138 Qz-monzonite 3-2 monzodiorite 221-a lamprophyre dyke 235-v lamprophyre dyke 21 lamprophyre dyke 100-2 Qz-monzonite 130 Qz-monzonite 100-9 syenite 100-12 syenite 100-8 lamprophyre dyke 505 ultramafic 505-a gabbro 507 monzodiorite 134a gabbroic diorite 184 gabbro 134 syenite 159 syenite 165 syenite West Karelian domain 39-1 diorite 14 granodiorite 37 monzonite 37-1 diorite 36-1 Qz-monzonite 36-2 monzonite 36-3 diorite 33 granodiorite 19-4

Qz-monzonite

19-3

lamprophyre dyke

Intrusion

Nuk Nuk Nuk Nuk Nuk Nuk Nuk Panozero Panozero Panozero Panozero Panozero Panozero Panozero Panozero Panozero Panozero Panozero Panozero Panozero Panozero Panozero Panozero Ust-Voloma Ust-Voloma Sjargozero Sjargozero Sjargozero Sharovalampi Sharovalampi Sharovalampi Huzhjarvi Huzhjarvi Huzhjarvi Huzhjarvi Huzhjarvi

Kurgelampi Kurgelampi Kurgelampi Kurgelampi Kurgelampi Kurgelampi Kurgelampi Kurgelampi area Kurgelampi area Kurgelampi area

Sm

Nd

147

143

144

144

Sm/ Nd

Nd/ Nd

F2r

e Nd(0)

e Nd(T)

f Sm/Nd

T ZR, Ga

T NdDM, Ga

3.70 6.22 4.64 4.28 4.42 3.71 5.39 15.47 21.99 16.77 10.38 11.22 6.71 5.78 9.65 11.6 7.99 4.78 4.56 7.19 5.79 4.71 9.3 6.43 3.8 11.5 29.25 5.13 71.3 85.54 90.66 18.21 15.72 13.30 10.42 7.04

23.95 53.99 27.00 27.60 27.30 25.49 35.08 77.28 118.84 85.11 56.97 62.80 38.90 35.29 55.39 65.80 45.48 27.80 26.16 38.62 30.76 21.70 50.00 41.08 23.27 74.50 183.77 27.90 12.42 15.57 15.35 113.27 92.57 80.06 63.59 43.62

0.0935 0.0696 0.1038 0.0937 0.0978 0.0880 0.0929 0.1210 0.1108 0.1191 0.1101 0.1080 0.1033 0.0989 0.1053 0.1069 0.1062 0.1030 0.1054 0.1125 0.1138 0.1311 0.1124 0.0945 0.0987 0.0923 0.0953 0.1102 0.1053 0.1110 0.1023 0.0963 0.1027 0.0995 0.0982 0.0967

0.510840 0.510442 0.511015 0.510891 0.510937 0.510786 0.510874 0.511362 0.511163 0.511321 0.511168 0.511098 0.511047 0.510948 0.511059 0.511089 0.511080 0.511011 0.511049 0.511204 0.511215 0.511543 0.511213 0.510872 0.510950 0.510843 0.510878 0.511177 0.511090 0.511152 0.510989 0.510915 0.511036 0.510984 0.510923 0.510921

15 8 10 11 9 9 8 7 7 8 11 13 10 7 7 11 7 19 10 7 11 7 7 11 8 12 9 9 7 10 9 9 6 12 9 8

35.1 42.8 31.7 34.1 33.2 36.1 34.4 24.9 28.8 25.7 28.7 30.0 31.0 33.0 30.8 30.2 30.4 31.7 31.0 28.0 27.8 21.4 27.8 34.4 32.9 35.0 34.3 28.5 30.2 29.0 32.2 33.6 31.3 32.3 33.5 33.5

0.9 1.5 0.7 1.9 1.3 1.8 1.8 1.8 1.5 1.7 1.9 1.3 1.9 1.5 1.4 1.5 1.5 1.3 1.2 1.7 1.4 1.7 1.8 1.5 1.5 1.8 1.4 1.9 2.1 1.2 1.1 1.8 1.9 2.0 1.3 1.8

0.52 0.65 0.47 0.52 0.50 0.55 0.53 0.38 0.44 0.39 0.44 0.45 0.47 0.50 0.46 0.46 0.46 0.48 0.46 0.43 0.42 0.33 0.43 0.52 0.50 0.53 0.52 0.44 0.46 0.44 0.48 0.51 0.48 0.49 0.50 0.51

2.711 2.711 2.711 2.711 2.711 2.711 2.711 2.742 2.742 2.742 2.742 2.742 2.742 2.742 2.742 2.742 2.742 2.742 2.742 2.742 2.73a 2.73a 2.73a 2.73a 2.73a 2.741 2.741 2.73a 2.74a 2.74a 2.74a 2.74a 2.74a 2.74a 2.74a 2.74a

2.79 2.75 2.81 2.73 2.77 2.73 2.73 2.77 2.78 2.77 2.76 2.80 2.75 2.78 2.79 2.79 2.78 2.80 2.80 2.77 2.79 2.77 2.75 2.77 2.77 2.76 2.79 2.75 2.74 2.80 2.80 2.76 2.76 2.75 2.79 2.76

4.88 4.20 6.60 5.50 4.46 6.53 6.03 3.73

30.31 26.70 40.87 33.90 27.23 40.00 35.64 24.25

0.0974 0.0942 0.0968 0.0974 0.0972 0.0987 0.1022 0.0929

0.510842 0.510826 0.510881 0.510873 0.510838 0.510904 0.510936 0.510793

12 26 10 4 7 9 7 9

35.0 35.3 34.3 34.4 35.1 33.8 33.2 36.0

0.4 0.4 0.6 0.2 0.4 0.4 0.2 0.2

0.50 0.52 0.51 0.50 0.51 0.50 0.48 0.53

2.711 2.711 2.711 2.711 2.711 2.711 2.711 2.71a

2.89 2.82 2.82 2.85 2.89 2.84 2.83 2.84

15.57

112.72

0.0828

0.510633

12

39.1

0.6

0.58

2.71a

2.80

14.72

69.67

0.1266

0.511384

14

24.5

0.1

0.36

2.70a

2.91

A. Kovalenko et al. / Lithos 79 (2005) 147–160

153

Table 1 (continued) N

Rock

West Karelian domain 20-4 lamprophyre dyke

Intrusion

1025 1060 Ja-9a

monzodiorite monzonite monzonite

Kurgelampi area Tulos Tulos Jalonvaara

Vodlozero 101 51 53 95/86 118 16 124

domain Qz-monzonite Qz-monzonite lamprophyre dyke diorite syenite Qz-monzonite monzodiorite

Bergaul Bergaul Bergaul Chalka Hautovaara Hautovaara Hautovaara

Sm

Nd

147

143

Nd/ Nd

F2r

144

144

Sm/ Nd

e Nd(0)

e Nd(T)

f Sm/Nd

T ZR, Ga

T NdDM, Ga

14.24

77.46

0.1111

0.511096

6

30.1

0.3

0.44

2.70a

2.90

5.10 3.84 9.70

29.98 21.22 56.87

0.1031 0.1094 0.1030

0.510870 0.511034 0.510880

15 18 12

26.9 26.9 26.9

1.9 0.8 1.6

0.48 0.44 0.48

2.71a 2.71a 2.71a

2.45 2.60 2.45

2.43 5.95 14.80 6.18 47.90 7.32 11.03

14.83 37.26 87.70 38.75 9.65 44.14 55.83

0.0990 0.0966 0.1011 0.0967 0.1218 0.0994 0.1183

0.510895 0.510847 0.510978 0.510761 0.511288 0.510910 0.511258

9 7 12 13 10 10 9

34.0 34.9 32.4 36.6 26.3 33.7 26.9

0.5 0.4 1.2 1.3 0.1 0.6 0.7

0.50 0.51 0.49 0.51 0.38 0.49 0.40

2.74a 2.74a 2.73a 2.743 2.741 2.741 2.741

2.85 2.86 2.79 2.98 2.92 2.84 2.85

Ages (TZR) from works: (1) Bibikova et al. (this volume), (2) Chekulaev et al. (2003), (3) Chekulaev et al. (1994). a Approximate age, based on field relationships.

some samples from the West Karelian domain (NN 1025, 1060-Tulos intrusion and Ja-9a-Jalonvaara intrusion) have TDMNd ages younger than their probable crystallisation age. The Tulos intrusion occurs within a late-Archaean granulite-facies zone (Baykova et al., 1984) and the Jalonvaara intrusion is close to the Proterozoic Svecofennian Province (Fig. 1). Therefore, we suggest that the Sm–Nd systems of these samples demonstrate open-system behaviour and that the analytical results for these three samples should be treated with great caution. Table 1 also shows Sm–Nd isotope data for ultramafic to quartz monzonitic lamprophyre dykes that are spatially related to the sanukitoid intrusions of Panozero and Sjargozero in the Central Karelian domain, Kurgelampi in the West Karelian domain, and Bergaul in the Vodlozero domain. Except in the case of Bergaul, the e Nd and TDMNd values of the lamprophyres are similar to those of the related sanukitoid intrusions, suggesting petrogenetic kinship. The similarity between the compositions of the lamprophyres and the inferred parent sanukitoid magmas suggests that the lamprophyres may simply represent sanukitoid magmas quenched in the hypabyssal environment. There is a distinct difference in e Nd between the sanukitoids of the Bergaul intrusion (e Nd +0.4 to +0.5) and the spatially related lamprophyre dyke (e Nd +1.2). Thus, either the lamprophyre dykes in the Bergaul intrusion have no direct

petrogenetic kinship with the host sanukitoid, or the e Nd values of the sanukitoids have been lowered through assimilation of older Vodlozero crust. In any case, the isotope results for the Bergaul intrusion should be treated with caution because the intrusion occurs in the marginal zone, between two crustal domains of quite different age.

5. Discussion 5.1. Genesis of sanukitoid magmas in Karelia Most sanukitoids occur in the youngest Central Karelian domain, with fewer in the West Karelian and Vodlozero domains. Like sanukitoids elsewhere in the world, Karelian sanukitoids are late- to posttectonic intrusions and are generally neither temporally nor spatially related to TTG magmatism. In Karelia, only the West Karelian sanukitoids have temporal relationships with trondhjemites (LobachZhuchenko et al., 2004b). Sanukitoid intrusions occur mainly within broad shear zones and are spatially related to slightly older belts of metasedimentary and felsic volcanic rocks. They are not usually associated with greenstones. Sm–Nd isotope data provide evidence that the Karelian sanukitoids were produced by partial melting of a source region that had long-term LREE depletion

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Fig. 3. (a) An e Nd versus time diagram for the Karelian sanukitoid intrusions. Data for potassic granites and West Karelia and Vodlozero old crust from Chekulaev et al. (1997, 2002) and Lobach-Zhuchenko et al. (2000b). (b) Variation of T NdDM values for the Karelian sanukitoids and related lamprophyre dykes.

relative to chondrites. Strangely, sanukitoids also have high LREE concentrations and low 147Sm/144Nd ratios (CSm/Ndb 0.38). This feature is commonly interpreted as the result of enrichment of their mantle source region simultaneously with or shortly prior to

magma genesis. Isotope data for the sanukitoids of CK are consistent with this model because all the analyses plot close to the depleted mantle array. However, the West Karelian sanukitoids and those of the marginal part of Vodlozero domain have lower

A. Kovalenko et al. / Lithos 79 (2005) 147–160

e Nd(t) (Fig. 3a). This diversity could be due to contamination of the sanukitoid magmas of the West Karelian and Vodlozero domains by older crustal rocks. Alternatively, these features may be due to longer-term enrichment of the source region of West Karelia and the marginal part of the Vodlozero domain, leading to lower e Nd(t) values. Due to the high LREE abundance in sanukitoids (Nd between 20 and 180 ppm), the Sm–Nd isotope system in these rocks is insensitive to contamination by older TTG crustal rocks (with Ndb20 ppm). Thus, contamination by assimilation of older rocks is improbable as an explanation. Instead, we argue that the regional Nd isotope variations among the Karelian sanukitoids were primarily due to a protracted time interval between the enrichment of their mantle source region and its partial melting. Calculation of the difference between the crystallisation age (2.74–2.70 Ga) and TNdDM of the sanukitoids shows that the duration of enrichment was mainly b60 Myr, for the source region of the Central Karelian sanukitoids, and mainly b150 Myr for the source region of the sanukitoids in the older West Karelia and Vodlozero domains (Fig. 3b). We suggest a two-stage process for the genesis of the sanukitoid magmas. In the first stage, enriched mantle formed through crust–mantle interaction, possibly in a subduction environment, at around 2.9–2.85 Ga in the older West and Vodlozero domains and at a date younger than about 2.8 Ga in the Central Karelian domain. Thus, the timing of mantle metasomatism seems to be linked broadly to the timing of crust formation. In the second stage, a parental tectonothermal event at ~2.74–2.70 Ga generated the sanukitoid magmas, through partial melting of the enriched mantle. Given the narrow time interval and the wide geographical distribution, what kind of tectonothermal process might have been responsible? In the Karelian granite-greenstone terrane, the occurrence of late Archaean basaltic dykes is consistent with a plume origin (Arestova et al., 2002). Alternatively, delamination of the eclogitic portion of the lower crust may have followed thickening of the Karelian crust. This delamination would have allowed mantle upwelling and heating of the subcontinental lithosphere.

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Summarising the above observations, we suggest that the genesis of the Karelian sanukitoids involved the following stages and processes: (1)

Between 2.7 and 2.8 Ga new crust formed between the West Karelian and Vodlozero continental fragments. The presence of arc-type magmatic rocks in this new crust suggests that it formed in an island arc environment (Chekulaev et al., 2004), and that the underlying mantle was metasomatised by slab-derived melts and/or fluids. The suggestion that sanukitoids were generated in a subduction setting is further supported by their spatial and temporal association with metasedimentary belts that are interpreted as having formed in an island arc environment (Chekulaev et al., 2004). Worldwide, there is a relationship between sanukitoid intrusions and metasedimentary basins (Hattory et al., 1996; Stevenson et al., 1999; LobachZhuchenko et al., 2000a; Smithies and Champion, 2000; Smithies, 2002). It has been suggested (Stevenson et al., 1999) that such metasedimentary belts are pull-apart basins, formed by strike-slip faulting during oblique plate movements. Thus, there is good reason to suggest a general link between sanukitoids, subduction and crust–mantle interactions. (2) In comparison with the Central Karelian domain, there are fewer sanukitoids in the older Western Karelian and Vodlozero domains. We therefore suggest that mantle metasomatism was a less significant process in these older domains, because the crust here was thicker and colder. (3) At the margins of the Vodlozero domain, the sanukitoid intrusions are spatially related to the Hautovaara greenstone belt, dated at ~ 2.95 Ga (Fig. 1). Here, the second-stage magmatism (~2.85 Ga; Lobach-Zhuchenko et al., 2004a) is represented by intermediate to felsic volcanic rocks that have island-arc affinities. From this, we suggest that subduction, at ~2.85 Ga, was also responsible for crust–mantle interactions that resulted in mantle enrichment. (4) Metasomatised mantle was partially melted under the influence of a thermal anomaly (e.g. plume-related uplift or delamination and decompression melting) at 2.74–2.70 Ga. This resulted

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in the generation of sanukitoid magmas, mostly in the Central Karelian domain. Later, the same thermal event produced the potassic granite magmas, through melting of older crustal rocks. This is the model most consistent with the geological relationships and the Nd isotope data for the sanukitoids. Smithies and Champion (2000) proposed a similar model for the genesis of Australian postkinematic sanukitoids. 5.2. Comparison of sanukitoids with modern subduction-related magmas Archaean rocks were first called bsanukitoidQ by Shirey and Hanson (1984), who compared the chemistry of these rocks to that of Miocene highMg andesites (known as sanukites) from the Setouchi Volcanic Belt in Japan. Based on the geochemical signatures of the Setouchi high-Mg andesites (Table 2) Shimoda et al. (1998) argued that the Archaean sanukitoids are analogues of the Setouchi high-Mg andesites. The high-Mg andesites, or adakites (Table 2), are intermediate to felsic volcanic rocks, found only in subduction-zone environments where either the downgoing slab is young and/or subduction is rapid (Drummond et al., 1996; Martin, 1999). Two types of high-Mg andesites have been distinguished —Piiptype and Adak-type (Yogodzinski et al., 1994, 1995). Piip-type high-Mg andesites are saturated in olivine and are thus interpreted as magmas derived from a source that was in near equilibrium with mantle peridotite. Their extreme Cr and Ni enrichment (N250 and N100 ppm, respectively) also suggests mantle affinities. The high-Mg andesites from the Setouchi belt of Japan, (Tatsumi and Ishizaka, 1982; Shimoda et al., 1998) are Piip-type high-Mg andesites (Yogodzinski, 1994). Adak-type high-Mg andesites have Mg/(Mg+Fe) similar to Piip-type (~0.70). However, they have lower Cr contents, are enriched in K2O, LREE, Ba and Sr, and have steeply fractionated REE patterns (Table 2). It has been suggested that Adak-type high-Mg andesites are formed by incomplete reaction between felsic slab melts and mantle peridotite (Yogodzinski et al., 1995). It is unlikely that Piip-type high-Mg andesites (or high-Mg andesites similar to those in the Setouchi

Table 2 Main geochemical features of subduction-related rock groups TTG

1 wt.% SiO2 Al2O3 K2 O Mg#

69.79 15.56 1.76 0.43

BADR Sanukitoids High-Mg andesites

2 63.87 17.02 2.37 0.41

3 53.94 14.75 3.26 0.54

ppm Rb 55 83 97 Sr 454 424 1547 Ba 690 501 1583 Cr 29 29 152 Ni 14 19 50 Nd 21 28 52 Yb 0.55 2.44 1.6 La/Yb 58 12 30

Adaktype

Piiptype

Setouchi

4

5

6

59.7 15.43 2.08 0.72

58.00 17.19 0.96 0.70

13 2366 320 161 126 40 0.62 49

9 384 87 262 127 b15 1.42 4.5

57.72 14.78 1.36 0.74

60a 235a 303a 513 181 b15 1.5 8.3

(1) Average of 355 TTGs from Martin (1994), (2) average continental arc from Drummond et al. (1996), (3) average mondiorite of Panozero sanukitoid intrusion, considered as primary melt (Lobach-Zhuchenko et al., this volume), (4) typical Adak-type high-Mg andesite, Cenozoic western Aleutian arc (Yogodzinski et al., 1995), (5) typical Piip-type high-Mg andesite from Piip Volcano, Cenozoic western Aleutian arc (Yogodzinski et al., 1995), (6) typical Miocene sanukite from the Setouchi volcanic belt of Japan (Tatsumi and Ishizaka, 1982). a Rb, Ba and Sr concentrations for Setouchi high-Mg andesites of similar SiO2, quoted by Davis et al. (1994).

Volcanic Belt) represent analogues of the Archaean sanukitoids. The relatively low K2O, LREE, Ba and Sr abundances, low La/Yb ratios and very high Cr and Ni contents of the Piip-type high-Mg andesites suggest that their mantle source regions were not greatly metasomatised, and that the melts were not in equilibrium with residual garnet. Thus, we believe that it would be incorrect to call Archaean high-Mg diorites bsanukitoidsQ simply based on their elevated Mg/(Mg+Fe) ratios (cf. Shirey and Hanson, 1984). On the other hand, the Adak-type rocks have geochemical signatures generally similar to the sanukitoids (Table 2), suggesting possible kinship. Nevertheless, there are significant differences in the geochemistry of Adak-type and sanukitoid magmas. Specifically, the Ba and Rb contents are much lower and Mg/(Mg+Fe) is much higher in the adakitic rocks. We therefore conclude that adakites are not the direct analogues sanukitoids.

A. Kovalenko et al. / Lithos 79 (2005) 147–160

In addition to extreme Ba and Sr enrichment, sanukitoids are also characterised by Rb depletion. Rb is a highly incompatible element and is expected to behave in a fashion similar to other LILE. Thus, this depletion represents a problem with current models for sanukitoid magma genesis. The K enrichment in sanukitoids suggests that phlogopite-bearing peridotitic rocks formed the sources of the parent sanukitoid magmas. The metasomatism required to form such a rock may have involved mantle wedge enrichment by either slab derived fluids, or melts, or both. Prior to the onset of modern-style plate tectonics, the early Archaean oceanic crust may well have consisted of a patchwork of relatively small plates, separated by spreading ridges that were closely spaced, by present-day standards (Condie, 1997a,b). In this more chaotic subduction environment, hot, young, hydrothermally altered oceanic crust would have been rapidly recycled. The high temperatures in the slabs would certainly have caused efficient dehydration and potentially also melting reactions. The mantle above these slabs would thus have become progressively more enriched in elements such as Rb and K, as well as becoming hydrated. The formation of phlogopitic micas would be a natural consequence. With continued subduction, the composition of the enriched mantle would have evolved, progressively lowering its solidus temperature. At some point, the solidus would have shifted such that even a small thermal perturbation could trigger widespread partial melting, leading to production and extraction of magnesian, K-enriched magma. This is how we envisage the sanukitoids being generated. Such a temporal evolution of mantle composition, controlling melt production, would also explain why the sanukitoids have a restricted range of ages. Sanukitoid magma production must have represented a global mantle depletion event, but one restricted to the regions where the mantle had become enriched through intense metasomatism. These magmas could not have been produced earlier because the Archaean mantle had not yet reached a sufficient state of hydration and LILE enrichment. Their scarcity after the major period of sanukitoid magmatism may be due to the thicker overlying crust, which can inhibit mantle enrichment. Also, following the major sanukitoid magmatic event,

there must have been general depletion of the postsanukitoid upper mantle. In locations where mantle melting occurred periodically, or where subduction processes were not operating, more normal magmas could be developed. What could be the cause of the Rb depletion in sanukitoids? We believe that this question may have a rather simple answer. Wendlandt and Eggler (1980) studied the fluid-absent partial melting of a phlogopite peridotite containing about 67% olivine, 18% clinopyroxene, 5% orthopyroxene and 10% phlogopite. They showed that partial melting at 3.0 GPa produced a garnet-bearing residue but that garnet was not produced at 2.0 GPa. As a model for the source of sanukitoid magma, we could perhaps take such a peridotite at a pressure of 2.5 GPa. The phase diagram of Wendlandt and Eggler (1980, Fig. 1) shows that, at this pressure, phlogopite breakdown occurs over an interval of about 80 8C, beginning at about 1150 8C. Thus, if the mantle temperature had reached 1200 8C, for example, there would be some residual phlogopite left in the source (Fig. 4). The only constraint on the pressure of melting during sanukitoid genesis is from the REE evidence that garnet was probably stable in the source (suggestOl+Opx+Cpx+ Phl+Spl+Grt +melt

3 P ~ 2.5 GPa, T = 1150-1230 ºC

Pressure, GPa

5.3. Petrogenetic model

157

2

Ol+Opx+Cpx+ Phl+Spl+Grt

1

1000

Ol+Opx+Cpx+Spl+Grt+melt

1100

1200

1300

Temperature, ºC Fig. 4. Stability of phlogopite in spinel lherzolite+10 wt.% phlogopite, for fluid-absent conditions (after Wendlandt and Eggler, 1980). Standard mineral phase abbreviations after Kretz (1983). See text for discussion.

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ing P N2 GPa, based on the phase diagram cited above). Thus, we do not suggest that melting occurred at 2.5 GPa, but only that this pressure is near the minimum required to satisfy the geochemical and phase equilibrium constraints. The temperature of 1200 8C seems plausible also. However, the main purpose of our modelling of the melting is to show what would happen if residual phlogopite remained in the source of the sanukitoid magmas. The partition coefficient for Rb distribution between phlogopite and coexisting mafic melt is about 2.5–3.0 (compilation of Rollinson, 1993; LaTourrette et al., 1995). Thus, any residual mica would highly concentrate the Rb, leaving the melt Rb-depleted, while allowing it to be enriched in K, Sr and Ba. The following example will serve to illustrate the partitioning of elements between residual mantle and possible sanukitoid melt, during batch partial melting. Let us take the Wendlandt and Eggler phlogopitic peridotite as a putative source material, and assume that the melting reaction destroys nearly all the clinopyroxene but leaves 2% residual phlogopite (out of the original 10%), corresponding to 26% partial melting. Phlogopitic micas contain close to 10 wt.% K2O, so the K2O content of the peridotite would be 1 wt.%. Since K is hosted only by phlogopite or melt, this model system would produce potassic mafic magma with a K2O content of about 3.1 wt.% (0.8/0.26), very similar to the composition of the inferred parental sanukitoid magma (Table 2; see Lobach-Zhuchenko, this volume). The calculated average parental sanukitoid magma has about 100 ppm Rb, 1500 ppm Sr and 1500 ppm Ba. Model batch partial melting equations, and traceelement partition coefficients from Rollinson (1993), can be used to calculate that our model parent peridotite would have contained about 33 ppm Rb, 410 ppm Sr and 433 ppm Ba. The crystalline residue, remaining after withdrawal of the partial melt, would contain 32 ppm Rb, 45 ppm Sr and 180 ppm Ba. Thus, the model predicts that most of the Sr and Ba in the source rock would be partitioned into the sanukitoid magma 92% and 74%, respectively. However, 75% of the Rb would remain behind in the melt-depleted source, concentrated in the residual mica. Thus, the presence of a small amount of residual, refractory phlogopite in their mantle sources provides the probable explanation for the rather strange LILE characteristics of sanukitoid magmas.

Table 2 shows that there are differences in Rb concentrations between the parental melts that produced the sanukitoids and those that produce the Adaktype high-Mg andesites (see above). These differences could be due to different degrees of mantle metasomatism prior to partial melting (different phlogopite contents), or to different degrees of partial melting of a similar source. In addition, the Ba contents of sanukitoids are higher than in Adak-type high Mg andesites. Since Ba partitions more strongly into biotite than phlogopite, this suggests that the Adak-type mantle source region may have contained more Ferich mica than the sanukitoid source rocks. Another significant distinction between sanukitoids and Adaktype high-Mg andesites is the relatively high Mg/ (Mg+Fe) values of the latter. We speculate that the lower Mg/(Mg+Fe) in sanukitoids may be due to lack of equilibrium with mantle olivine (i.e., possibly a more pyroxenitic source rock). It should be noted that late Archaean sanukitoids and modern adakites are volumetrically minor rock types in comparison with the Archaean TTG and modern BADR series. Thus, the Melzer and Wunder (2001) conclusion, that mantle metasomatism has only minor effects on the LILE signatures of the modern IAB, seems to be relevant in this context. In the late Archaean, as for the present day, mantle metasomatism may not have been particularly common. This effect may be produced mainly in cases where subduction of young, hot, oceanic crust promoted intense metasomatism of the mantle wedge, through addition of slabderived fluids or melts.

6. Summary and conclusions New isotope data for the sanukitoids of the Karelian granite-greenstone terrane led us to conclude that, over a narrow time interval in the late Archaean (2.74–2.70 Ga), sanukitoid magmas were emplaced in the various, geologically distinct parts of Karelia. This occurred through the partial melting of metasomatically enriched mantle, under the influence of some tectonothermal event. In addition, elemental analyses and Sm–Nd isotope data for lamprophyre dykes that cut most of the sanukitoid intrusions suggest petrogenetic kinship with the host sanukitoids. Geochemical features of the sanukitoids, and comparisons with modern arc-related

A. Kovalenko et al. / Lithos 79 (2005) 147–160

high-Mg andesites (adakites), show that the sanukitoids form a distinct geochemical group. Thus, the late Archaean sanukitoids appear to have no modern analogues. Using published experimental data on the fluid-absent partial melting phlogopite-bearing mantle rocks, we have modelled the unusual behaviour of Rb, Sr and Ba during the genesis of sanukitoid magmas. From this, we conclude that a small amount of residual phlogopite in the sanukitoid mantle source region can readily account for the relative Rb depletion and Ba and Sr enrichment of sanukitoid magmas. Combining our model for the generation of the Karelian sanukitoids, and comparing their geochemical and petrological features with those of modern highMg andesites, we can suggest a reason for the absence of modern analogues for the late Archaean sanukitoids. The key to understanding this issue is the differences in composition of the upper mantle and geothermal gradients between late Archaean and modern times. Firstly, the upper mantle was less depleted in the late Archaean. Thus, it required less heat input to initiate partial melting. Secondly, due to the more fertile composition of the Archaean upper mantle, less metasomatism was required to lower its solidus and permit partial melting. Thirdly, the much higher geothermal gradient in the early stages of Earth’s evolution would have facilitated the melting of the mantle. We have argued that some modern high-Mg andesite melts may be in equilibrium with residual biotite rather than phlogopite. The late Archaean, high geothermal gradient probably resulted in a higher degree of partial melting of the metasomatized mantle, and hence the generation of more mafic melts, in comparison with modern arc-related magmas. As a corollary we argue that the Sm–Nd isotope data provide evidence for a two-stage process in the generation of the Karelian sanukitoid magmas. In contrast, the geochemical data for some modern, arc-related, high-Mg andesites do not require the formulation of such a two-stage model.

Acknowledgements This work has benefited greatly through numerous discussions with Svetlana Lobach-Zhuchenko, Hugh Rollinson, Valery Chekulaev and Natalia Arestova, and through reviews by R.H. Smithies and J. Halla. We thank E. Bogomolov for his assistance in mass

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spectrometry. This research was supported by a NATO/Royal Society Fellowship (to AK), INTAS project N 01-0073, Russian Foundation of Fundamental Research project N 02-05-65052.

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