Directory UMM :Data Elmu:jurnal:P:Precambrian Research:Vol104.Issue3-4.2000:

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Precambrian Research 104 (2000) 147 – 174

Archaean – Proterozoic transition: geochemistry, provenance

and tectonic setting of metasedimentary rocks in central

Fennoscandian Shield, Finland

Raimo Lahtinen *

Geological Sur6ey of Finland,P.O.Box96,FIN-02151Espoo,Finland

Received 8 July 1999; accepted 5 May 2000

Abstract

The central part of the Fennoscandian Shield in Finland is composed of the Palaeoproterozoic Svecofennian domain and the Archaean Karelian craton with a Palaeoproterozoic allochthonous and autochthonous cover. A cryptic suture separating these areas and another tentative suture dividing the Svecofennian into central and southern parts have been proposed. The chemical composition of sedimentary rocks (N=300) within the study area, including the effects of palaeoweathering, hydraulic sorting, depositional environment and post-depositional processes, have been studied in order to delineate sediment source components. The main proposed source components for the Archaean sedimentary rocks are weathered 3.0 – 3.2 Ga greenstone+granite9TTG and local 2.7 Ga sources. Autochthonous 2.2 – 1.9 Ga cover rocks were mainly derived from a mixture of chemically weathered palaeosol (2.2 – 2.35 Ga), sedimentary rocks derived from the palaeosol, and mafic dykes and plateau volcanics (mainly 2.2 – 2.1 Ga) although in places locally derived non-weathered Archaean sources dominated. Archaean crust and 2.0 – 1.92 low-K bimodal rocks from a primitive island arc are the proposed source for the allochthonous Western Kaleva cover rocks. These formed in a subsiding foredeep during initial collision from orogenic detritus in the same oblique collision zone. The central Svecofennian sedimentary rocks can be divided into local arc-derived rocks (51.89 Ga) and older (]1.91 Ga) rocks from a mixture of Western Kaleva sources and a 1.91 – 2.0 Ga mature island arc/active continental margin source. Rifting followed by increased subsidence during initial collision in the NE and subsequent arc reversal caused rapid erosion from the mountain belt, exposing diverse source compositions as seen in the large variation of Th/Sc (2 – 0.5), and deposition into an oblique hinterland basin further developing into a subduction related foredeep. Mature greywackes from the southern Svecofennian in the study area resemble passive margin sediments with a source dominated by inferred alkaline-affinity complexes and Archaean rocks. Less mature rocks also occur and had sources dominated either by island arc/active continental margin rocks or local picritic rocks. In the sedimentary record the Archaean – Proterozoic transition up to 2.1 Ga was dominated by input of mainly mafic plateau-type volcanic contribution to the Archaean detritus. Palaeoproterozoic sediments having a crustal component (52.1 Ga) show higher Th/Sc, Th/Cr, and lower Sm/Nd and Eu/Eu* relative to the Archaean rocks but locally low Th/Cr ratios complicate the situation. Ba depletion relative to K, Rb and Th is a characteristic feature of the

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* Fax: +358-2055012.

E-mail address:[email protected] (R. Lahtinen).

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R.Lahtinen/Precambrian Research104 (2000) 147 – 174 148

sedimentary rocks of the central Fennoscandian Shield indicating high amounts of Ba lost from the clastic record during 2.3 – 1.9 Ga and further recycled back to the mantle forming a subduction component and an enriched mantle component. Ba depletion seems to have been especially characteristic of chemical weathering during 2.35 – 2.2 Ga under CO2-rich and low-O2atmosphere. Whether this strong Ba depletion is characteristic of the Archaean –

Keywords:Archaean; Palaeoproterozoic; Sedimentary rocks; Geochemistry; Provenance; Finland

1. Introduction

The geochemistry of clastic sedimentary rocks can be used as an indicator of crustal evolution (e.g. Taylor and McLennan, 1985) or to identify ancient tectonic settings in metamorphic terranes. Sedimentary rocks can be divided into those showing local sources and those having experi-enced effective mixing in large river marine sys-tems before deposition. The latter types sample large areas providing data for crustal-scale pro-cesses. The possibility of different crust-forming mechanisms during Archaean and Proterozoic times emphasizes the importance of the Ar-chaean – Proterozoic boundary where there might be a corresponding compositional change in the sedimentary record (e.g. Taylor and McLennan, 1985; McLennan and Taylor, 1991). Selective preservation of sedimentary rocks in the ancient record can on the other hand hamper their use in crustal evolution studies. Along with this limita-tion, other factors discussed below, should also be taken into account when using ancient sedi-ments to give information on the general prove-nance of the studied sedimentary unit.

The lithology of the provenance area essen-tially controls the chemical composition of the clastic sediments but other factors such as degree of palaeoweathering, hydraulic sorting (grain-size effects), organic and sulphide input, diagenesis and metamorphism (especially migmatization) may greatly modify or ultimately erase prove-nance memory. Sediment recycling is a common feature (e.g. Veizer and Jansen, 1985) and pro-duces a buffering effect where a small amount of new input can go unnoticed. Nevertheless, even though the interpretation of their compositions is more controversial than with igneous rocks, the

long ‘memory’ of sedimentary rocks can be quite powerful when modelling the tectonic settings and evolutionary histories of metamorphic ter-ranes.

The central part of the Fennoscandian Shield in Finland is composed of the Palaeoproterozoic Svecofennian domain and the Archaean Karelian craton with a Palaeoproterozoic allochthonous and autochthonous cover (Fig. 1). The occur-rence of a cryptic ‘suture’ (Fig. 1; Koistinen, 1981; Huhma, 1986) between the Karelian and Svecofennian domains is favoured by the obser-vation that no Archaean component is found in the 1.93 – 1.91 Ga gneissic tonalites and related felsic volcanics adjacent to the Archaean craton (Lahtinen and Huhma, 1997). Lahtinen (1994) proposed also the occurrence of a tentative ‘su-ture’ (Fig. 1) separating the central part of the Svecofennian domain from the southern Sve-cofennian. Studies on the geochemistry of sedi-mentary rocks in the study area are few and include a geochemical and isotopic study from the Archaean Hattu schist belt (O’Brien et al., 1993), a major element study from the northern part of the Ho¨ytia¨inen area (Kohonen, 1995), a regional correlation diagram study from the Savo province (Kontinen and Sorjonen-Ward, 1991) and a research concentrating on black schists (Loukola-Ruskeeniemi and Heino, 1996 and ref-erences therein).

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evolution of Fennoscandian Shield. The main source components and implications for the tec-tonic evolution of the central Fennoscandian shield are given with emphasis on proposed su-tures. Notes on the crustal evolution and Ar-chaean – Proterozoic transition in general, and on Ba depletion are also given. As all the studied sedimentary rocks are metamorphosed, the prefix ’meta’ has been dropped. The data set is available on request from the author.

2. Sampling and analytical methods

Sampling was done with a mini-drill with dia-mond bit. Each sample comprised four to six subsamples (altogether 1 – 1.5 kg) from the same lithological unit, if detection of unit boundaries was possible (sometimes this was impossible, e.g. in some migmatites). In the case of turbidites, the whole Bouma A, AB or ABC was sampled in most cases. A composite sample was taken from

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Fig. 2. Simplified geological map of the study area (Fig. 1) modified from Korsman et al. (1997). Sample locations are also indicated.

veined migmatites and pelitic rocks where layers were B5 cm thick and a more homogeneous unit was not available.

The analytical work was done in the laborato-ries of the Geological Survey of Finland. Samples were jaw crushed and splits were pulverized in a tungsten carbide bowl for X-ray fluorescence (XRF) analysis, and in a carbon steel bowl for inductively coupled plasma mass spectrometry (ICP-MS). Major elements and Cl, V, Cr, Ni, Zn, Rb, Sr, Y, Zr, Nb and Ba were determined by XRF, CGraf.by Leco CR-12 carbon analyzer, F by ion selective electrode, aqua regia leachable S and Cu by ICP-AES, and aqua regia leachable Au, Pd, Te, As, Ag, Bi, Sb and Se by GAAS (Sand-stro¨m, 1996). REE, Co, Nb, Hf, Rb, Sc, Ta, Th and U were determined by ICP-MS after dissolu-tion of the sample (0.2 g) with hydrofluoric acid-perchloricacid treatment completed by a lithium metaborate/sodium perborate fusion (Rautiainen et al., 1996). The estimated uncertainty is 1 – 5% for major elements and 3 – 10% for trace elements.

3. General geology

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(Pekkarinen, 1979; Pekkarinen and Lukkarinen, 1991; Kohonen and Marmo, 1992; Karhu, 1993). Otherwise lithostratigraphy and chronostratigra-phy of the Ho¨ytia¨inen area are not resolved (Ko-honen, 1995) but depositional ages from 2.1 to about 1.9 Ga are inferred.

The Suvasvesi area is characterized by the ‘Up-per Kaleva’ (Kontinen and Sorjonen-Ward, 1991) or Western Kaleva (Kohonen, 1995 a term adopted in this study) greywackes that occur as allochthonous units in thrust complexes charac-terized by associated ophiolites and related rocks (Koistinen, 1981 and references therein) though evidence for local deposition upon Archaean basement has also been noted (Ward, 1987). The increase in metamorphic grade from east to west (Fig. 2) is seen as an increase in quartz veins and the onset of segregational banding (quartz+ feldspar) leading finally to migmatites.

The boundary zone (BZ) includes migmatitic sedimentary rocks (Korsman et al., 1984) and a 1.93 – 1.91 Ga volcano-plutonic formation (Lahti-nen, 1994 and references therein). The Svecofen-nian is divided into the central SvecofenSvecofen-nian including the Central Finland Granitoid Complex (CFGC) and Bothnian Belt (BB), and the south-ern Svecofennian including the Rantasalmi – Haukivuori area (RH). The tentative sedimentation ages for the central Svecofennian, based on data available from the Tampere Schist Belt (Lahtinen, 1996 and references therein), are

]1.91 and 1.89 – 1.87 Ga for rocks correlated to basement- and arc-related groups in the Tampere Schist Belt, respectively. The southern Svecofen-nian, including the Rantasalmi – Haukivuori area, is characterized by granite migmatites, which is a clear difference to the central Svecofennian, boundary zone and Suvasvesi area, which are characterized by tonalite migmatites (Korsman et al., 1999 and references therein).

4. Results

Because lithostratigraphic division of tary rocks is rarely available, division of sedimen-tary rocks into different groups within domains is based mainly on lithotype and geochemical

char-acteristics. All elements analyzed have been used but the main weight has been put on the REE, Th, Sc, Cr and major elements where the REE, Th and Sc are considered as most reliable ele-ments in monitoring the average source composi-tion (Taylor and McLennan, 1985; McLennan et al., 1990). The arc-related (upper) central Sve-cofennian rocks of this study (Fig. 2), not dis-cussed in detail, show CaO, MnO, P2O5, Sr, Ba and Sb enrichment, which is characteristic of sed-imentary rocks derived from high-K calc-alkaline to shoshonitic volcanics (Lahtinen, 1996). Strongly altered or mineralized samples are ex-cluded from discussion as are minor groups of sedimentary rocks either having undefined origins or a large non-clastic component (e.g. iron forma-tions and carbonate rocks).

The group characteristics were also studied by using normalized diagrams (Fig. 3). Archaean sedimentary groups are normalized to Archaean crust (AC1), autochthonous and allochthonous groups to average Karelian craton (KC1) and boundary zone and Svecofennian groups to West-ern Kaleva WK1 (Table 1). The AC1 is a first approximation of the average composition of Ar-chaean crust in Finland at its present erosion level based solely on the data from the study area. The granitoid-dominated nature of the exposed Ar-chaean part of the study area is seen in higher LILE and LREE and lower MgO, Cr and Ni compared to the Late Archaean (3.5 – 2.5 Ga) restoration model for average juvenile upper con-tinental crust (Table 4 in Condie, 1993). The Karelian craton includes a large contribution from Palaeoproterozoic mafic dykes and volcanics (2.2 – 1.97 Ga; Vuollo, 1994) relative to the Ar-chaean crust average (Fig. 3).

4.1. Archaean sedimentary rocks

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Average chemical composition of estimated Archaean crust (AC1) and Karelian craton (KC1), and selected sedimentary rock groups (non-migmatized, except groups BZ1–BZ2)a

BZ1

Ar1 H1 H2 H3 WK1 WK1frag WK2 BZ2

KC1 AC1

(N=4) (N=8)

(N=156)

(N=129) (N=11) (N=5) (N=9) (N=47) (N=17) (N=6) (N=5)

67.23 69.85 69.58 63.23 65.15

68.60 56.42

SiO2(%) 65.18 63.64 65.15 60.16

0.51 0.65 0.76 0.80 0.62 0.68 0.69 0.83 0.72 1.08

TiO2 (%) 0.72

12.87 14.86 14.74 13.11 13.27 15.42 15.16 17.68

14.68

Al2O3(%) 15.19 15.15

5.20 4.95 4.93 6.64 6.05

7.90 9.24

6.60 6.27

FeO (%) 4.71 5.73

0.06

0.08 0.10 0.08 0.08 0.07 0.07 0.07 0.08 0.08 0.11

MnO (%)

5.19

2.34 2.81 3.55 2.91 2.52 2.26 2.33 3.23 2.84 4.29

MgO (%)

1.46 2.22 2.42 2.36 2.34

1.68 2.59

CaO (%) 3.39 4.06 1.46 0.87

4.24 3.89 2.37 1.24 1.98 2.98 2.76 2.84 2.92 2.93

Na2O (%) 2.18

3.44 2.37 2.41 3.36 3.34

3.87 3.44

2.76

K2O (%) 2.46 2.35 2.71

0.15

0.18 0.18 0.12 0.13 0.11 0.16 0.15 0.16 0.14 0.11

P2O5(%)

0.34

(0.05) (0.05) (0.10) 0.09 0.13 (0.22) (0.29) (0.07) (0.05) (0.05)

Cgraf.(%)

0.21 0.067 0.082 0.061 0.21

1.24 0.23

S (%) 0.054 0.061 0.41 0.12

0.070 0.054 0.053 0.085 0.078 0.094

F (%) 0.055 0.051 0.045 0.062 0.094

62.64 54.4 54.7 55.8 55.6 57.8

62.9 62.5

49.3 61.9

CIA 50.0

36.2 31.8 15.2 31.1 32.0 31.6 30.6 33.2 36.7 44.3

La (ppm) 23.4

71.2 63.4 32.7 62.2 63.2 62.9 60.9 65.4 73.2 86.5

Ce (ppm) 47.9

7.27 7.43 7.29 8.02 8.60

7.44 10.1

Pr (ppm) 8.23 7.42 4.12 5.67

21.5 27.9 26.7 27.3 26.6 28.9 31.4 37.1

27.3

Nd (ppm) 29.7 15.3

5.17 5.13 4.98 5.55 5.72

5.49 6.44

3.10 4.28

Sm (ppm) 4.90 4.76

1.14

1.02 1.07 0.94 0.91 0.96 1.06 1.03 1.15 1.13 1.44

Eu (ppm)

4.91

3.86 4.00 2.96 3.88 4.27 4.63 4.47 5.04 5.26 6.13

Gd (ppm)

0.66 0.68 0.66 0.75 0.73

0.73 0.90

Tb (ppm) 0.50 0.55 0.48 0.61

3.36 3.68 3.42 4.12 3.75

Dy (ppm) 2.40 2.76 2.95 3.36 4.21 5.01

0.66 0.73 0.68 0.79 0.71

0.79 1.00

0.58 0.67

Ho (ppm) 0.45 0.53

2.31

1.26 1.50 1.76 1.94 1.88 2.12 2.04 2.31 2.03 3.06

Er (ppm)

0.33

0.18 0.21 0.26 0.28 0.27 0.31 0.30 0.32 0.29 0.47

Tm (ppm)

1.79 2.16 1.95 2.19 1.94

2.23 3.13

Yb (ppm) 1.17 1.38 1.73 1.84

0.27 0.32 0.30 0.32 0.27 0.46

Lu (ppm) 0.18 0.21 0.25 0.28 0.35

570 489 508 613 704

348 712

Ba (ppm) 858 742 371 392

116 157 127 58.1 48.8 79.4 100 139 139 172

Cl (ppm) 52.2

14.1 18.8 32.0 30.0 16.8 14.1 14.4 21.3 18.9 30.2

Co (ppm) 21.7

110 106 104 137 120

238 172

Cr (ppm) 77.7 80.6 294 180

3.50 5.02 5.01 4.46 4.92

Hf (ppm) 3.78 3.63 3.53 4.63 3.91 5.10

10.2 9.20 9.13 11.2 12.2

9.75 14.6

5.74b _8.70

Nb (ppm) 5.54 5.70

149

35.6 41.3 145 111 52.4 44.9 45.4 65.3 53.6 90.9

Ni (ppm)

138

84.0 74.0 84.5 104 138 82.5 89.1 117 122 135

Rb (ppm)

15.4 15.3 14.9 20.5 17.9

22.0 29.2

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Table 1 (Continued)

BZ1 H1

KC1 H2 H3 BZ2

AC1 Ar1 WK1 WK1frag WK2

(N=11) (N=5) (N=9) (N=47)

(N=4) (N=17) (N=6)

(N=156) (N=8) (N=5)

(N=129)

147 247 250 223 275 326

Sr (ppm) 495 437 180 111 108

0.80 0.68 0.66 0.76 0.82

0.68 0.74

Ta (ppm) _0.40 _0.41 0.42b 0.64

8.72 7.59 4.60 8.51 10.8 8.93 8.54 9.27 10.9 12.5

Th (ppm) 7.59

2.56 1.82 1.98 2.00 1.88

2.76 1.64

1.22 1.91

U (ppm) 1.49 1.32

196

94.9 127 160 142 120 128 128 164 143 222

V (ppm)

26.6

15.3 17.2 20.9 24.4 23.2 23.7 22.4 26.5 23.1 30.2

Y (ppm)

105 83.7 83.5 115 109

154 153

Zn (ppm)b _81.6 _88.1 ₁₂₈ ₁₀₈

144

Zr (ppm) 162 155 161 190 150 217 208 203 193 202

0.082 0.067 0.044 0.061 0.055

0.053 0.059

Ag (ppm)b _0.047b _0.052 _0.068 _0.16

0.86 0.80 1.28 12.9 4.52 0.42 0.52 0.63 1.10 1.01

As (ppm)b _6.53

0.52 0.34 0.31 0.40 0.79

Au ppbb _0.78 _1.05 _0.47 _0.73 _0.42 _1.00

0.31 0.10 0.034 0.12 0.15

0.21 0.080

0.079

Bi (ppm) 0.072 0.22 0.20

84.0

23.8 42.7 61.3 42.7 41.9 25.6 25.1 31.7 37.3 88.8

Cu (ppm)b

1.71 3.88 (0.79) (0.26) (0.31) (0.39) (0.27) 1.0

Pd ppb (0.2) (0.2) 1.80

0.031 0.028 0.021 0.021 0.046

0.095 0.041

0.029 0.035

Sb (ppm) 0.037 0.036

0.56

0.053 0.075 0.31 0.15 0.22 0.13 0.13 0.15 0.20 0.45

Se (ppm)

28.2 42.2 25.0 12.7 13.5 16.7 22.6 49.6

9.46

Te ppb 9.56 47.4

a_{WK1frag is the average of mica gneiss fragments in migmatites. Values in parentheses include many determinations below the detection limit (C}

graf0.05% and Pd 0.2 ppm) and show either the detection limit value or averages excluding values below detection limits.

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elements indicate a more mafic source compared to local Archaean bedrock at the present erosion level (Figs. 4 and 5, and Table 1). The Ar2 samples show variable REE and have higher CaO, Na2O and lower K2O, Cr and Rb compared to Ar1 (see Fig. 4 for K2O and Cr). The lower CIA indicates less weathering relative to Ar1 and low Th/Sc (0.09 – 0.17) favours a dominant mafic source.

4.2. Cratonic co6er

The Jatuli-type quartzites of this study show a strong increase in K2O with decreasing SiO2(Fig. 4), which is mainly due to variations in sericite/ muscovite content. One subarkose contains fresh K-feldspar also seen in a lower CIA value but otherwise high CIA is a characteristic feature. The

sedimentary rocks in the Hoÿtiaïnen basin are classified into high- and low-Cr groups H1 and H3, respectively (Fig. 4, Table 1). A distinct litho-logical unit (Huhma, 1975) of high-Cr rocks is classified as group H2 and a suspect group of low-Cr rocks, possibly related to the Western Kaleva (Kohonen, 1995), is classified as group H4. Samples outside the Hoÿtiaïnen area (Fig. 2), but that occur in autochthonous position to Ar-chaean dome rocks or are geochemically similar, are included in these groups. The H1 – H3 samples include quartz-rich greywackes and more typically pelites showing thin layering from 1 – 3 mm to 1 – 2 cm with thin psammitic interlayers occurring locally. The variation in element abundances in-side the H1 group is mainly explained by quartz dilution (Fig. 4). There is evidence of weathering in at least one component (CIA 54 – 70) and a

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Average chemical composition of selected sedimentary groups (non-migmatized, except CF3 average including also mica gneiss fragments in migmatites)a

CF3

RH2mig RH3/lCr RH4/hCr CF1 CF2 CF3mig

RH1 RH2

(N=4)

(N=4) (N=7) (N=6) (N=5) (N=6) (N=14) (N=12) (N=14)

72.37 69.70 63.71 63.71

70.75

SiO2(%) 76.50 64.95 61.99 67.94

0.53 0.73 0.79 0.69 0.52 0.60 0.73 0.74

TiO2 (%) 0.58

13.25 14.20 12.78 13.47 15.24 15.71

16.13

Al2O3(%) 11.75 17.84

3.78 4.56 5.97 6.33

5.25

7.18 4.51

FeO (%) 3.78 5.77

0.07

0.03 0.05 0.06 0.08 0.06 0.07 0.08 0.09

MnO (%)

2.78

1.38 2.30 3.17 2.05 1.56 2.21 3.04 2.92

MgO (%)

1.91 2.13 1.88 1.94

2.44

CaO (%) 0.52 1.23 0.89 2.04

2.91 2.54 2.97 2.94 2.59 2.57

2.18

Na2O (%) 1.67 1.67

2.58 2.59 3.41 3.28

2.55 2.71

K2O (%) 3.97 3.81 2.59

0.17

0.10 0.13 0.12 0.15 0.15 0.16 0.15 0.12

P2O5(%)

(0.05) (0.05) (0.08) 0.15

Cgraf.(%) (0.05) (0.25) (0.09) (0.05) (0.05)

0.023 0.033 0.11 0.082

0.23 0.10

S (%) 0.41 0.043 0.051

0.064

0.0488 0.085 0.12 0.052 0.052 0.064 0.075 0.076

F (%)

55.1 56.4 54.6 54.8 58.0 58.8

CIA 64.8 62.3 68.8

47.6 37.9 34.2 37.0

37.9

44.8 30.7

La (ppm) 31.4 44.1

94.3 74.8 69.1 74.3

Ce (ppm) 63.2 86.9 88.7 75.8 62.5

10.5 8.69 8.16 8.79

7.34

Pr (ppm) 7.42 10.1 10.4 8.64

37.9 32.0

Nd (ppm) 27.5 38.0 38.4 32.3 27.3 30.3 32.1

6.53 5.87 5.61 5.95

5.23

7.29 5.63

Sm (ppm) 4.98 6.85

1.12

0.94 1.33 1.19 1.06 1.21 1.17 1.10 1.10

Eu (ppm)

5.59 5.28 5.15 5.46

Gd (ppm) 4.32 6.08 6.46 5.03 4.44

0.80 0.75 0.75 0.80

0.65 0.63

Tb (ppm) 0.89 0.96 0.71

3.44

3.30 4.80 5.05 3.80 3.99 3.86 3.97 4.33

Dy (ppm)

0.67

0.64 0.94 0.95 0.72 0.81 0.77 0.78 0.89

Ho (ppm)

2.32 2.27 2.19 2.64

1.91

2.64 2.14

Er (ppm) 1.81 2.83

0.29

0.28 0.41 0.41 0.31 0.32 0.31 0.30 0.40

Tm (ppm)

1.85

1.78 2.84 2.52 2.03 2.12 2.09 2.22 2.60

Yb (ppm)

0.33 0.34 0.33 0.39

0.27 0.29

Lu (ppm) 0.40 0.39 0.30

408 771 639 534 640 618 630 595

Ba (ppm) 628

39.5 42.0 51.1 79.7

46.7 51.7

Cl (ppm) 91.8 59.5 75.0

16.3

8.93 14.1 19.2 13.2 9.86 14.2 17.9 19.6

Co (ppm)

158

107 116 149 97.3 81.2 92.9 119 126

Cr (ppm)

6.62 5.40 4.45 4.56

4.50

Hf (ppm) 5.10 5.01 4.41 5.46

9.59 9.48 11.4 12.0

Nb (ppm) 9.3 13.8 15.0 9.03 10.5

32.2 39.8 59.6 62.8

58.2

77.2 38.1

Ni (ppm) 42.8 60.2

108

115 155 208 101 104 107 144 145

Rb (ppm)

16.2

10.4 17.9 20.4 13.4 11.6 15.2 18.2 19.4

Sc (ppm)

301 294 242 238

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Table 2 (Continued)

CF3mig RH3/lCr

RH1 RH2 RH2mig RH4/hCr CF1 CF2 CF3

(N=6) (N=5) (N=6) (N=14)

(N=7) (N=12)

(N=4) (N=4) (N=14)

0.67

0.66 0.93 1.03 0.68 0.84 0.77 0.83 0.82

Ta (ppm)

10.4 8.12 15.2 11.2 10.3 11.1

Th (ppm) 9.6 12.9 13.8

2.39 3.18 3.32 1.98 2.74 2.56 2.56 2.21

U (ppm) 2.29

87.5 112 144 146

143

160 107

V (ppm) 88.9 153

22.2

21.3 30.1 33.1 23.2 24.0 23.1 23.9 27.8

Y (ppm)

100

94.3 157 166 69.9 64.0 78.1 101 116

Zn (ppm)

267 218 178 181

181 227

Zr (ppm) 203 175 225

0.067 0.061 0.039 0.044 0.063 0.071

Ag (ppm) 0.088 0.096 0.059

0.58 1.43 0.92 0.60

2.11 1.38

As (ppm)b _1.12 _0.56 _1.03

1.38

0.67 1.16 0.84 0.88 0.46 0.82 0.67 0.38

Au ppbb

0.17

0.14 0.24 0.23 0.12 0.056 0.12 0.18 0.045

Bi (ppm)

11.3 16.6 33.2 53.2

27.0

Cu (ppm) 30.8 32.4 24.0 19.0

(0.2) (0.29)

Pd ppbb _(0.25) _(0.82) _1.02 _(0.28) _(0.25) _0.48 _0.85

0.042 0.041 0.032 0.037

0.083

0.027 0.059

Sb (ppm) 0.045 0.089

0.10

0.18 0.56 0.12 0.053 0.053 0.10 0.13 0.18

Se (ppm)

17.4 10.6 6.6 14.8 23.5 28.2

Te ppbb _8.7 _36.0 _21.4

a_{The RH2mig and BB4mig are the averages of migmatites, respectively. Group RH3 have been divided into low-Cr (RH3/lCr) and high-Cr (RH/hCr) populations.}

Values in parentheses include many determinations below the detection limit (Cgraf0.05% and Pd 0.2 ppm) and show either the detection limit value or averages calculated excluding values below detection limits.

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Fig. 4. Harker-type Cr, K2O, MgO and CIA (Nesbitt and Young, 1982) variation diagrams for Archaean, autochthonous and

allochthonous sedimentary rocks in the study area. Ar1 and Ar2-Archaean, Jqzt – Jatuli-type quartzites, H1 – H2-autochthonous high-Cr, H3-autochthonous low-Cr, H4- a low-Cr suspect group of Ho¨ytia¨inen area. WK1 – WK2 main field-allochthonous Western Kaleva. AC1 is the average of Archaean crust (Table 1).

large mafic component indicated by high contents of HREE, MgO and Pd. The H2 group has many compositional similarities with H1 but the H2 average shows higher levels of most elements (e.g. MgO) and lower SiO2(Fig. 4 and Table 1). Some H3 pelites show enrichment of felsic source com-ponents manifested as low MgO contents (Fig. 4). The K2O, Rb and Bi enrichment (not shown) favour a source dominated by a late-Archaean granite (Kutsu; see Fig. 3). The H4 is a heteroge-neous group that deviates to some extent from the WK1 main group in having higher K2O and lower Cr (Fig. 4).

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gneisses (WK2 migmatites). Both groups of migmatites only show the systematic depletion of Bi compared to non-migmatitic samples (Table 1).

4.3. Boundary zone and S6ecofennian sedimentary

rocks

The sedimentary rocks in the boundary zone (BZ; Fig. 2) have been divided into psammitic (BZ1) and pelitic (BZ2) groups. The BZ1 rocks are heterogeneous in chemical composition show-ing high variation, e.g. in HREE, CaO, K2O, Th and Nb and the average (Table 1) should be only considered as an areal average.

The southern Svecofennian sedimentary rocks in the Rantasalmi – Haukivuori area have been classified into three groups (RH1 – RH3). The non-migmatitic RH1 rocks are quartz-rich greywackes and the well-preserved RH2 rocks are more pelitic in character. Both RH1 and RH2 show rather similar patterns in Fig. 6 where the strong effect of weathering is seen in negative peaks of Ba, Sr, CaO, MnO and P2O5, and high CIA values (Table 2). The depletion of HREE, Sc, V, TiO2 and enrichment of K2O, Rb, Th and especially U is the main difference when com-pared to the Western Kaleva source. A relative

Fig. 5. Plots of La vs. Yb and Eu/Eu* vs. GdN/YbN for selected sedimentary rocks in this study. GdN and YbN are

chondrite-normalized values and Eu/Eu* has been calculated using Eu*=(SmN+GdN)/2. The Archaean average has been

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Fig. 6. Major- and trace-element distributions in averages of southern Svecofennian sedimentary rock groups RH1 – RH3 (Table 2) from the Rantasalmi – Haukivuori area normalized to the average of Western Kaleva psammites (WK1 in Table 1). RH3/lCr and RH3/hCr are averages of low- and high-Cr populations of RH3.

enrichment of Zn to Ni and Co is also a charac-teristic feature. The RH2 group shows the relative enrichment of CaO, Ba, Nb, V and Sc and low Cr/Sc ratio favouring a new additional mafic component in the RH2. The lower CIA values (Table 2), which are normally higher in more pelitic rocks, indicate that this additional compo-nent was less weathered. Compared to the RH1 and RH2 rocks the RH3 samples show lower CIA and higher CaO and Na2O with strong variation in the amount of mafic component (Fig. 6 and Table 2).

The RH1 – RH2 migmatites vary from gneisses with quartz veins and small melt patches cut by pegmatites to veined gneisses with abundant gran-ite leucosome. The main differences (Table 2) can be interpreted to show a more pelitic precursor

for migmatites but the slightly lower REE and especially deep negative Eu anomaly in some sam-ples ask for a loss of felsic component. The slight depletion in Ba, K2O and K/Rb can be related to a loss of a K-feldspar component and the enrich-ment of ferromagnesian components to the in-creased amount of restite. So it seems that these migmatites are mainly in situ migmatites that show a complex mixture of restite and a melt fraction in variable proportion in outcrop scale.

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5). The depletion of elements characteristic of mafic components and the relative enrichment of LREE, Sr, Th, U and Zr point to a larger felsic component relative to the WK psammites. The chemical composition of the CF2 group shows an enrichment of mafic components relative to CF1. CF3 is a heterogeneous group characterized by migmatites and thus the average (Table 2) in-cludes also mica gneiss fragments in migmatites. Mineralogically the CF3 rocks differ from the CF1 – CF2 in the ubiquitous occurrence of garnet. The more clay-rich nature of CF3 is seen in lower SiO2 and higher MgO and K2O (Table 2). The CF3 migmatites form an inhomogeneous group ranging from samples with HREE enrichment to samples with HREE depletion and Eu enrichment at low total REE abundances compared with less migmatitic CF3 samples. This is interpreted as different amounts of restite and leucosome in sampled outcrops.

5. Discussion

5.1. Palaeoweathering

Palaeoweathering in the source area is one of the most important processes affecting the com-position of sedimentary rocks. Sedimentary rocks sensu stricto are composed merely of weathering products and reflect the composition of weather-ing profiles, rather than bedrock (e.g. Nesbitt et al., 1996). Based on CIA values (Nesbitt and Young, 1982) the source rocks affected the most by weathering are those of Archaean group Ar1 (60 – 65), Jatulian quartzites (58 – 73), au-tochthonous groups H1 – H3 (54 – 70) and south-ern Svecofennian groups RH1 – RH2 (57 – 68) whereas the allochthonous WK1 – WK2 mostly show CIA values lower than 55 (Fig. 4). Most of the central Svecofennian psammitic rocks also have low CIA values (B55) with an increase up to (60 – 67) in CF3 pelitic rocks. This general increase in CIA with silica-poorer and more pelitic nature is a common feature and readily explained by the higher proportion of clays (weathering products) in pelites.

The CIA value is also affected by other pro-cesses than the clastic composition of the rock in question. Overestimation of Ca in carbonates can lead to too high CIA values if Mg-bearing car-bonates are present. Fortunately only a few sam-ples have over 0.5% CO2 and thus this is only problematic in limited cases but is especially cru-cial for quartz-rich samples. The other problem is related to the loss of CO2 and incorporation of liberated Ca in recrystallizing minerals (e.g. epi-dote and plagioclase) during metamorphism (cf. Lahtinen, 1996) a situation proposed for some samples in the Ho¨ytia¨inen area (Fig. 7).

The prevailing climatic conditions of the source areas during sediment formation are difficult to estimate especially if we consider the recycled nature of many sediments, possibly having older weathered components. The situation can be thus complex including mixing of a strongly weathered component (older sediments or deeply weathered palaeosol) with immature crust components be-fore deposition, forming a sedimentary rock showing moderate CIA values. Also the degree of weathering is related to the rate of erosion, which is high in tectonically active areas and thus in-hibiting extensive weathering even in high rainfall tropical conditions. The extent of weathering is determined primarily by the amount of rainfall (acids) on the weathering profile (Singer, 1980) where as the climatic effect on weathering trends is probably insignificant (Nesbitt and Young, 1989).

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by Hokkalampi Palaeosol. Sturt et al. (1994) con-cluded that widespread 2.35 Ga regolith (including the Ilvesvaara Formation) occurred on the Fennoscandian shield and was related to an arid or semi-arid palaeoenvironment. Although this might be the case for the Ilvesvaara Formation, the occurrence of the up to 80 m deep Hokkalampi Palaeosol (not mentioned by Sturt et al., 1994) with a minimum age of 2.2 Ga records intense chemical weathering under a tropical warm and humid climate (Marmo, 1992). The drift of Fennoscandian from 30°S at 2435 Ma to about 30°N at 2100 Ma (Pesonen et al., 2000) shows that Fennoscandian crossed the equator during this time favouring the interpretation of Marmo (1992). It has been suggested that the Hokkalampi Palaeosol and derived formations covered large areas of the stable Karelian craton (Kohonen and Marmo, 1992; Marmo, 1992) where they formed the bulk of detritus for the Palaeoproterozoic rift basins.

The chemical and mineralogical data of the Hokkalampi Palaeosol indicate a typical weather-ing sequence (cf. Nesbitt and Young, 1989; Condie et al., 1995) with an initial decrease in the amount of plagioclase followed by loss of K-feldspar and biotite seen as an increase in CIA values from about 60 – 70 (lowermost) to the highest values of 80 – 90 in the upper zone (Marmo, 1992). Potas-sium metasomatism of kaolinite to illite in palaeosol results in lowering of CIA values (Fedo et al., 1995). This possibility has been studied using an A – CN – K compositional space (Fig. 7) for the data of the Hokkalampi Palaeosol formed upon K-feldspar rich granitoid and sandstone. There is a slight amount of added potassium in lower palaeosol zones probably due to percolation of solutions from the leached uppermost potas-sium-depleted zone during weathering (Marmo, 1992). However, if the whole mass balance of the weathering profile is considered, no input of exter-nal potassium is needed.

Fig. 7. A – CN – K and (A – K) – C – K triangles (see Fedo et al., 1995, 1997) depicting trends in the Hokkalampi palaeosol and autochthonous groups of this study. (A) Data for Hokkalampi palaeosol formed upon a K-feldspar-rich granitoid (granitoid zones 2 – 3) and sandstone (sandstone zones 1 – 3), and an average of Archaean crust and Archaean sedimentary rocks (Ar1 – Ar2). Trajectories a and b represent weathering trends for sandstone and Archaean average crust predicted from kinetic leach rates (Nesbitt and Young, 1984). (B) Data for Jatuli-type quartzites and autochthonous groups H1 – H3. Trajectories a and b same as in Fig. 7A. Dashed line encloses possible source end members for autochthonous sedimentary rocks. (C) Data for Jatuli-type quartzites and autochthonous groups H1 – H3. Note the shift of some samples towards the sodium-rich (A – K) – N-line indicating that albitization has possibly affected these samples. Horizontal arrows for some samples indicate the amount of Ca input due to the inferred occurrence of carbonates followed by CO2 loss. Averages of palaeosol zones 1 – 3 in Fig. 7B and C are calculated using

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The autochthonous Ho¨ytia¨inen H1 – H3 groups show characteristic depletion of CaO, Na2O, MnO, P2O5, Sr and Ba, and low K/Rb, which are tentatively proposed to have an ultimate source in the chemically weathered palaeosol. The southern Svecofennian RH1 – RH2 groups also show deple-tion of elements normally lost during weathering (Fig. 6) but the CIA values of other groups are moderately low (B60) and no clear weathering trends are observable.

5.2. Hydraulic sorting

Clay minerals, enriched in most trace elements, and preferentially concentrated in the finer frac-tions during hydraulic sorting (grain size sorting) produce higher abundances of many elements in pelites relative to associated sands (e.g. Korsch et al., 1993). The situation of pure quartz dilution is the ultimate case and most easily interpreted as a decrease in all other elements and an increase in SiO2. The situation is more complex when acces-sory minerals (zircon, monazite, apatite, sphene and allanite), ferromagnesian minerals, feldspars and lithic fragments are also sorted. The Th/Sc ratio remains nearly constant in some cases but often muds can have significantly lower Th/Sc ratios indicating a preferential incorporation of mafic volcanic material in the finer fractions (e.g. McLennan et al., 1990). Considering a simple two-component mixture of mature weathered ma-terial (quartz+clays) and immature rock debris (separate minerals+lithic fragments) the result is psammites enriched in immature rocks debris showing complex sorting patterns and pelites en-riched in mature weathered material. This prefer-ential sorting can lead to REE fractionation making interpretation of Sm – Nd isotope system-atics difficult (Zhao et al., 1992) but this is mainly effective when considering sedimentary material from unweathered coarse-grained granitoids with, e.g., allanite hosting LREE and Th.

The wide range of SiO2 (Fig. 4) the Ho¨ytia¨inen H1 – H3 groups exhibit is clearly an effect of sorting (cf. Kohonen, 1995) dominated by quartz dilution seen as abundant quartz clasts. Sorting enhanced enrichment of mafic component was noticed, e.g. in Western Kaleva and southern

Svecofennian pelites over psammites. The varia-tion of Zr (normally 160 – 350 ppm) found in Western Kaleva psammites indicate zircon sorting but there is no correlation between Zr and HREE or U showing that the zircon control on these elements is minor. The effect of hydraulic sorting is readily observed in the studied samples but in many cases it also sorts different source compo-nents into different grain size classes. This is a disadvantage when using only shales (on average more mafic) or psammites (on average more fel-sic) in crustal evolution studies but is an advan-tage in characterizing source end members.

5.3. Effects of depositional en6ironment

Different methods have been applied to the interpretation of the depositional environment of ancient sediments using black shales/schists. These include pyrite formation, S/C ratios, degree of pyritization (Berner, 1984; Berner and Raiswell, 1984; Raiswell and Berner, 1986) and enrichment of U and V (e.g. Jones and Manning, 1994; Breit and Wanty, 1991). The average present S/C ratio of normal marine sediments is 0.36 (0.23 – 0.77) but age dependent variation oc-curs and, for example, early Palaeozoic marine sediments show significantly higher S/C ratio of about 2 (Berner and Raiswell, 1984; Raiswell and Berner, 1986). In fresh or low-salinity brackish water low sulfate level is the limiting factor for pyrite formation and sediments show low S/C ratios without any inter-element correlation (Berner and Raiswell, 1984). According to Thompson and Naldrett (1984) mantle-derived magmatic S/Se ratios are generally lower (B

10 000) than in sedimentary sulphides (\10 000), which can be used to discriminate hydrothermal influxes of sulphur.

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Fig. 8. Plot of Cgraf. vs S for autochthonous (H1 – H3) and

allochthonous (WK1 – WK2) sedimentary rocks in this study divided into low S/Se (B10 000) and high S/Se (B10 000) populations. The S/C ratio 0.36 is for normal marine sedi-ments after Berner and Raiswell (1984).

jor factors related to the degree of diagenesis are thermal history and time, where rapid burial com-pacts sediments quickly (dewatering) and blankets any thermal changes (Lee and Klein, 1986). Thus long-lived basins, like the Ho¨ytia¨inen basin (Ko-honen, 1995), should show more pronounced ef-fects of diagenesis compared to allochthonous Western Kaleva-type rocks that were deposited as massive units in an active tectonic setting. The very limited element variation in the WK rocks favours this and although small-scale diagenetic changes within WK samples are possible, a large-scale redistribution of elements is not evident. Similar arguments hold for most of the central Svecofennian rocks but, for example, the deposi-tional environment and the elapsed time before dewatering and metamorphism of the Archaean and southern Svecofennian mature rocks are un-known. Diagenetic reactions may include Na-, K-, Mg- and Fe-metasomatism (e.g. Nesbitt and Young, 1989) while REE redistribution and frac-tionation have also been proposed (Awwiller and Mack, 1991; Milodowski and Zalasiewicz, 1991; Ohr et al., 1991). There is not however consensus about how common the redistribution of REE during diagenesis is (cf. Hemming et al., 1995) and one critical question is that are the proposed diagenetic reactions open or closed systems at sample scale.

Redistribution of alkalies during diagenesis has been proposed for the Ho¨ytia¨inen area rocks (Ko-honen, 1994) and to evaluate this possibility, the data are plotted in the A – CN – K and (A – K) – C – N compositional spaces (Fig. 7; see also Fig. 4 for K2O). The data show scatter and there are several factors that may have been responsible for the observed trends: (1) Sedimentary rocks have dif-ferent source components with difdif-ferent K2O/ Na2O ratios (see differences in MgO contents and Th/Sc and Th/Cr ratios; Figs. 4 and 9). The problem lies also in the thinly layered nature of pelites where chlorite-rich and biotite-rich layers were noticed, possibly indicating that different layers were derived from different sources in some cases. (2) During grain-size sorting K-rich phases (illite and biotite-vermiculite) are enriched in pelites (K-feldspar is rare in these rocks) and plagioclase in sands forming a trend similar to indicate anoxic conditions during deposition and

if the S/C ratio of 3.5 is higher than found in the Palaeoproterozoic marine sediments during depo-sition, it could point out to euxinic environment. The Western Kaleva samples differ from the Ho¨ytia¨inen basin examples in that they do not show any clear correlation between S and C. The graphite-enriched (\0.5% C) psammites have low S/C ratios (B0.15) and S/Se ratios mostly

B10 000. Apart the graphite variation (0 – 1.6% C) there is no enrichment of studied elements. The occurrence of graphite-bearing thick psammites does not favour a direct hemipelagic origin and indicates mixing of carbonaceous matter into mass flows before deposition. The low S/C ratios could point to fresh water or brackish water environments, or to short intervals between depo-sition of mass flows preventing significant bacte-rial sulfate reduction. The lack of U and V enrichment indicates an oxygenated environment while a low S/C excludes an euxinic environment.

5.4. Diagenesis and metamorphism

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Ma-R.Lahtinen/Precambrian Research104 (2000) 147 – 174 164

that observed in the A – CN – K compositional space. (3) Albitization of K-feldspar in the sand-size fraction with immediate uptake of liberated K by kaolinite, chlorite, montmorillonite and/or smectite in the clay-rich fraction as proposed by Kohonen (1994). Based on Fig. 7C albite metaso-matism has occurred to some degree in some samples favouring Kohonen’s (Kohonen, 1994) interpretation. (4) Regional-scale potassic and sodic metasomatism affecting shales and silt-sand-size particles, respectively, has been proposed for the Palaeoproterozoic Serpent Formation (Fedo et al., 1997). The Serpent shales show ultimate potassium variation from 3.3 to 11.2% whereas the H1 – H3 pelites show only variation from 3 to 5% (Fig. 4) where the variation is mainly due to

the factors 1 – 3, as discussed above. Thus, the problem in depicting the amount of diagenetic redistribution in the Hoÿtiaïnen area rocks is that they show complicated mixing of source compo-nents associated with sorting and thus distinguish-ing purely diagenetic effects is difficult. Although not conclusive it seems that small-scale redistribu-tion of elements has occurred during diagenesis in the Hoÿtiaïnen area but no externally derived regional-scale metasomatism, at least for potas-sium, is observed.

Prograde metamorphic effects on REEs, except in areas of partial melting, are minor (Taylor et al., 1986) but the depletion of LILE elements (K, Rb, Ba) has been proposed for granulite terrains (e.g. Weaver and Tarney, 1983; Sheraton, 1984).

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The tonalite migmatites (veined gneisses, schollen migmatites and diatexites) in the study area show variable compositions due to differences in the relative amounts of restite and leucosome in sam-pled outcrops and those that represent totally melted ‘in situ’ variants. A depletion of Bi is the main common feature and although migmatites with high proportions of restite component occur there is no area showing large-scale depletion of elements. In many cases the veined gneisses have mostly retained their original composition (cf. Lahtinen, 1996).

The southern part of the Rantasalmi – Haukivuori area (southern Svecofennian) is char-acterized by in situ migmatites (RH1 – RH2) with variable amounts of restite and granite leucosome components. This difference in the leucosome composition (tonalite – granite) has been at-tributed to the aluminium excess in the source rocks of migmatites having granite leucosomes (Korsman et al., 1999). This interpretation is fa-voured by the typical CIA values of 60 – 70 in the RH1 – RH2 rocks compared to the typical CIA values below 60 in the source rocks of tonalite migmatites. On the other hand water-rich condi-tions during tonalite migmatization favour the formation of plagioclase-enriched melts and wa-ter-rich conditions has been considered as the main cause for the formation of tonalite migmatites (Lahtinen, 1996).

5.5. Main source components

The proposed main source components of sedi-mentary rocks of the Archaean craton and its cover, and Svecofennian domain are mainly based on the geochemical differences but Sm – Nd results by Huhma (1986, 1987), O’Brien et al. (1993) are also adopted. There are only a few detrital zircon age determinations from the Fennoscandian Shield (Huhma et al., 1991; Claesson et al., 1993) and thus the conclusions presented below are to some extent tentative but serve as a working model for future work.

Boundary zone sedimentary rocks (BZ1 – BZ2) are probably related to the 1.92 Ga primitive island arc but the occurrence of numerous fault zones, extensive migmatization and complicated

shearing precludes further source component interpretation.

5.5.1. Archaean sedimentary rocks

The Archaean sedimentary rocks show very low Th/Cr ratios, which discriminate them from other rocks in this study (Fig. 9). O’Brien et al. (1993) concluded that greywackes in the eastern part of the study area (Ar2-type) with TDM ages from 2.83 to 2.99 Ga normally show a local source. The Ar1 samples show a more homogenized source and higher degree of weathering of the source area with higher MgO, Cr, K2O and SiO2. One Archaean sediment has a TDM of 3.24 Ga (Huhma, 1987) favouring also the existence of an older component (cf. Sorjonen-Ward, 1993). Two main ages of source components with variable amounts of intermixing are proposed for the Ar-chaean sediments in the study area:

1. Older main component with 3.0 – 3.2 Ga aver-age source aver-age. At least three different source rock types are indicated: komatiites (high MgO, Cr, Ni, Cr/Sc), tholeiite (high TiO2and Nb/Th) and felsic component (SiO2, K2O and Rb). Intermediate to strong weathering in the source area and thorough mixing has occurred before deposition. Possible sources are greenstone+granite9TTG.

2. Local source derived from the 2.76 – 2.73 Ga (Vaasjoki et al., 1993) magmatic event (cf. O’Brien et al., 1993).

5.5.2. Cratonic co6er

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source dominated by felsic granitoids indicated by high Th/Sc and Th/Cr ratios. The sedimen-tary rocks showTDM variation from 2.28 to 2.70 Ga, which partly overlap with the Western Kaleva TDM variation of 2.29 – 2.40 Ga (Huhma, 1986, 1987). The Sm – Nd data for the Ho¨yti-a¨inen basin is in general agreement with the geo-chemical data and suggest source components of: 1. Chemically weathered palaeosol, and sedi-mentary rocks derived from it, formed upon Archaean crust and glaciogenic deposits. En-richment of Archaean sedimentary rocks (see the 3.0 – 3.2 Ga component above).

2. Non-weathered Archaean crust. Local differ-ences, seen for example in the large amount of late-Archaean granite (Kutsu) component in some samples.

3. 2.2 – 1.96 Ga mafic magmatism, possibly volu-minous Jatuli-type plateau volcanism includ-ing presently exposed abundant dykes, to explain the high amount of mafic component in some rocks.

The detrital zircon U – Pb isotopic data for two samples (Claesson et al., 1993) give age con-straints for granitoid components in the al-lochthonous Western Kaleva mica schists. The samples have 30 – 40% late Archaean zircons (2.5 – 2.8 Ga) and only a few crystals in the age range between 2.6 and 2.1 Ga, which can also be mixture ages (Claesson et al., 1993). Both sam-ples have 50 – 60% zircons from a 2.0 to 1.92 Ga age group with a maximum deposition age of about 1.92 – 1.94 Ga. TDM ages of 2.3 – 2.4 Ga based on Sm – Nd data of Western Kaleva mica schists (Huhma, 1987) are in agreement with the detrital zircon data. The Archaean component has been dominantly late-Archaean in age and we can use the normalization to the AC1 of this study to interpret the nature of the 1.92 – 2.0 Ga component (Fig. 3). The relative TiO2, Nb (espe-cially Nb/Th ratio) and HREE enrichment with-out increase in the MgO level (slight depletion) and Cr/Sc ratio favour a primitive island arc tholeiitic origin for the mafic component. The high Zr relative to K2O, Rb and REE favour a low-K felsic source also characterized by moder-ate to low La/Yb ratios. Two main components are proposed for the Western Kaleva sediments:

1. Archaean crust dominated by late-Archaean granitoids mixed with a small contribution from Jatuli-type dykes. A small amount of recycled weathered component is possible. 2. 2.0 – 1.92 Ga bimodal source of low-K felsic

rocks and tholeiitic volcanics derived from primitive island arc.

5.5.3. S6ecofennian

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1. Western Kaleva-type source (see above). 2. Palaeoproterozoic (1.91 – 2.0 Ga) mature

is-land arc or active continental margin source. The southern Svecofennian mature metasedi-ments (RH1 – RH2) differ from the Western Kaleva and central Svecofennian psammites pointing to different origins. High Zn/Co (about 10) is a characteristic feature of RH1 and variable but high Zn/Co also characterizes the RH2 sam-ples. The Zn/Co ratio is sensitive to changes during weathering and sulphide precipitation but there does not seem to be any relationship be-tween the existence of sulphides and Zn/Co indi-cating instead either source difference or a weathering effect. Similar Zn/Co enrichment was not noted in high CIA rocks from the Ho¨ytia¨inen area favouring a source origin for the high Zn/Co. Elevated Zn and low Co are characteristic fea-tures of alkaline-affinity intermediate – felsic within-plate-type granitoids (Lahtinen, unpub-lished data) and this type of magmatism in the source area is one possible explanation for the high Zn/Co ratios. High Cr and Cr/Sc ratios in the RH1 are interpreted to have their ultimate sources in an abundant komatiite or picritic component.

The less mature greywackes (RH3) show mainly low CIA values (B57) and thus resemble the Western Kaleva psammites and psammites from the central Svecofennian. Although some samples have compositions close to those found in the Western Kaleva psammites the RH3 rocks are typically enriched in elements (LREE, Rb, Ba, Th and U) that characterize felsic source rocks. Some RH3 rocks are enriched in elements that charac-terize mafic rocks especially seen in high Cr/Sc ratio (Fig. 6). This could indicate an Archaean komatiite source but local Cr-rich lavas in the Rantasalmi – Haukivuori area are more likely. The main source components for the southern Sve-cofennian metasedimentary rocks in the Ran-tasalmi – Haukivuori area are as follows:

1. Alkaline-affinity complexes with high Zn and Zn/Co

2. Archaean crust with possibly high Cr/Sc (ko-matiite component).

3. Island arc/active continental margin type crust from an orogenic domain.

4. Local sources and, at least partly, picritic sources producing high Cr/Sc.

These tentative main source components char-acterize different groups differently; RH1 (19

294), RH2 (1929394), RH3 (3+29491). The problem lies in depicting the origin of the highly weathered component; Archaean versus palaeoProterozoic.

5.6. Tectonic implications

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com-R.Lahtinen/Precambrian Research104 (2000) 147 – 174 168

ponent in the Western Kaleva psammites due to a rapidly rising orogene during oblique collision starting in the N (cf. Lahtinen, 1994). As presently understood these psammites have been deposited both on Archaean basement and oceanic crust, and a foredeep origin associated with subsidence during initial collision is fa-voured and orogenic detritus either from the same, oblique collision zone (mainly from the accretionary prism) or a more distal orogenic domain is proposed. One interesting feature is the possible uptake of carbon-rich material, formed in an oxygenated and possibly brackish environ-ment, into the turbidite currents before deposi-tion of Western Kaleva sediment (this study). This could favour the axial foredeep model of Kohonen (1995) and deposition of organic matter near estuaries of large fresh water rivers.

The differences between the Western Kaleva and central Svecofennian sediments favour at least partly different origins and different ages of deposition. The source for the central Svecofen-nian sediments included also mature island arc material and the maximum deposition age was about 1.91 Ga for the main period of turbidite deposition. Lahtinen (1994, 1996) has proposed that ]1.91 Ga (possibly up to 1.95 Ga) rifting occurred in the Tampere Schist Belt followed by increasing subsidence during initial collision in the NE and subsequent arc reversal. Abundant erosion from the mountain belt and deposition into oblique hinterland basins that further devel-oped into a subduction related foredeep is the proposed model for the deposition of the main sequences of turbidites in the central Svecofen-nian. The arc-related sediments are of local derivation and indicate deposition in small basins before or during the 1.89 Ga collision (cf. Lahti-nen, 1994, 1996).

The southern Svecofennian mature greywackes resemble passive margin sediments but the more immature sediments contain arc-type material. The southern Svecofennian is characterized by abundant volcanics and, on the other hand, also by mature quartzites indicating both conti-nental margin-type and passive margin settings but more data are needed to explore these possi-bilities.

5.7. Crustal e6olution and Ba depletion in the

Archaean–Proterozoic transition

Abrupt changes in the composition (REE, Th, Sc) of sedimentary rocks at the Archaean – Proterozoic transition has been proposed (Taylor and McLennan, 1985; McLennan and Taylor, 1991; McLennan and Hemming, 1992). These in-clude an increase in negative Eu anomaly, a de-crease in the GdN/YbN ratio from \2.0 to 1.0 – 2.0, a decrease in the Sm/Nd ratio from about 0.21 to 0.19 and an increase in the Th/Sc ratio from about 0.5 to 1.0 (possibly only in continental sediments). Secular changes in the Ar-chaean – Proterozoic transition, especially con-cerning the development of a Eu minimum, have been questioned and argued to be a consequence of tectonic control resulting in biased sampling (e.g. Gibbs et al., 1986; Condie and Wronkiewicz, 1990a; Gao and Wedepohl, 1995). Although the Cr/Th ratio may not directly reflect the source ratio, abrupt changes have been noticed in the Archaean – Proterozoic boundary reflecting the decreasing amount of komatiites in the Protero-zoic (Taylor and McLennan, 1985; Condie and Wronkiewicz, 1990b; Condie, 1993). No consen-sus exists about the importance of the Archaean – Proterozoic transition but komatiites and TTG-type rocks characterize the Archaean and their contribution to the sedimentary record should be distinguishable.

The data of this study for Archaean sedimen-tary rocks are limited but some general remarks can be made. Archaean sedimentary rocks are characterized by low Th/Sc (B0.3) and Th/Cr (B0.018), and variable Eu/Eu* ratios that are normally higher than those in the Palaeoprotero-zoic sediments of this study (Fig. 5). The Sm/Nd ratios are also somewhat higher but the GdN/YbN ratios are lower than 2 and also lower than those found in many Palaeoproterozoic sediments of this study (Figs. 5 and 9).

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material mainly from differentiated (e.g. low Cr/ Sc) plateau volcanics and dykes. The other Palaeoproterozoic sediments of this study also show the contribution of intermediate to felsic igneous sources varying from low-K primitive is-land arc to mature active continental margin types. An important feature is the almost total absence of 2.1 – 2.5 Ga mature crustal component (granitoids) in these sediments.

The data are somewhat scattered but the Palaeoproterozoic sediments having crustal com-ponents showing higher Th/Sc, Th/Cr, and lower Sm/Nd and Eu/Eu* relative to the Archaean rocks (Figs. 5 and 9) as proposed in earlier studies (see references above) but the behaviour of GdN/ YbNratio is opposite to that proposed by McLen-nan and Taylor (1991). More data on Archaean sedimentary rocks from the Svecofennian shield are evidently needed to see if the limited input from TTG rocks (cf. Condie, 1993) is a common feature. The source variation seen in the central Svecofennian sediments (e.g. Th/Sc 2 – 0.5 in the CF1 – CF3 in Fig. 9) suggests exposure of differ-ent source compondiffer-ents during erosion in the source area. A slight grain-size induced preferen-tial separation of felsic source into the psammites and mafic source into the pelites was also noted in this study (cf. Lahtinen, 1996). Low Th/Cr ratios characterize the Archaean sedimentary rocks due to the abundant komatiite component but lower Th/Cr ratios can also be from a local picritic source (e.g. some RH3 samples). These features reinforce the need for a large data set when using sedimentary rocks in crustal evolution studies.

If the absence of 2.1 – 2.5 Ga mature subduc-tion-related material is true for most of the Fennoscandian shield (supercontinent stage) it im-plies that here the Archaean – Proterozoic transi-tion is characterized by the additransi-tion of only mafic magmatism (9felsic material in bimodal forma-tions) and the transition to Proterozoic crustal formation occurred about 2.1 Ga ago. A 2.4 – 2.3 Ga subduction event proposed for the western edge of Rae Province in Laurentia (Bostock and van Breemen, 1994), a roughly 2.2 Ga age for the onset of subduction-related Birimian magmatism (e.g. Davis et al., 1994 and references therein) and magmatic activity during 2.4 – 1.8 Ga with a mode

at 2.1 – 2.0 Ga based on detrital zircons from Sa¨o Francisco Shield (Machado et al., 1996) show that the age and nature of the Archaean – Proterozoic transition differ from shield to shield; a possibility also for the geochemical nature of associated sed-imentary rocks.

An elevated level of both Th and Sc relative to modern deep sea turbidites in the basement re-lated sediments in the Tampere – Ha¨meenlinna area was noted by Lahtinen (1996) and a similar situation characterizes the central Svecofennian sediments of this study (not shown). A source enriched in bimodal volcanics and depleted in sedimentary quartz was proposed (Lahtinen, 1996). The elements released during weathering are also lost from the clastic portion but can be partly redeposited in separate units within the sedimentary sequence, as for example Ca in marine carbonates and U with organic matter. Many elements are also recycled back to the mantle during subduction and form a characteris-tic fingerprint for subduction-related magmas and enriched mantle components (e.g. Hawkesworth et al., 1991; Weaver, 1991). Lahtinen (1996) pro-posed that the Ba deficiency in the basement-re-lated sedimentary rocks and the Ba enrichment in the Svecofennian enriched mantle component (see also Lahtinen and Huhma, 1997) are related to Ba release during weathering (9diagenesis) and later uptake in pelagic sediments (possibly as barite) that are further subducted into the mantle.

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Fig. 10. Plots of Ba vs. K2O for selected sedimentary rocks in this study. The Archaean average has been calculated from the average

in the Table 1 and Jatuli-type mafics from the average (N=21) in Lahtinen (unpublished data). The Archaean trend is approximated from the data in this study and the Tampere Schist Belt (TSB) volcanics trend is from Lahtinen (1996). See Fig. 5.

depletion. The Hokkalampi Palaeosol shows slight potassium enrichment in the lower zone but a large-scale external potassium addition seem unlikely. The main part of the Ba depletion is assumed to derive from the chemically weathered palaeosol, especially from the highly weathered part (CIA \80).

Ba depletion is less pronounced in other groups (Ar1, RH1 – RH2 and CF3) having also elevated CIA values over 60. If the interpretation of Ho¨yti-a¨inen sedimentary rocks is correct it indicates mixing of deeply weathered Archaean source ma-terial (CIA 70 – 90) with less weathered Archaean crustal and Palaeoproterozoic mafic sources (CIA

B50) to produce H1 – H2 rocks with CIA values in the range of 55 – 70. In this case the lack of comparable Ba depletion in other pelitic rocks with elevated CIA can be attributed to the lack of extremely strong chemical weathering (CIA \80) in the source area.

Different source areas have variable Ba/K ra-tios but the Ba depletion relative to K, Rb and Th (Lahtinen, 1996; this study) is a characteristic feature of the sedimentary rocks of central Fennoscandian Shield. This indicates a high amount of Ba lost from the clastic record during 2.3 – 1.9 Ga and further incorporated, at least partly, into both a subduction component and the enriched mantle. The Fennoscandian shield seems to have exemplified a cratonic stage during 2.6 – 2.1 Ga characterized by deep chemical weathering about 2.35 – 2.2 Ga ago, high burial rates of

or-ganic carbon and highly13_{C-enriched sedimentary} carbonates (e.g. Karhu, 1993) about 2.2 – 2.1 Ga ago, and multiply rifting from about 2.2 to 1.95 Ga. One critical question is the possible effect of CO2-rich and low-O2atmosphere in the formation of weathering profiles before the significant rise in atmospheric oxygen levels at about 2.0 Ga (e.g. Karhu, 1993). If the Ba depletion has been espe-cially characteristic for the chemical weathering during 2.35 – 2.2 Ga it could imply that during and after this time period high amounts of Ba have recycled back to the mantle forming a ‘peak’ in the formation of enriched mantle component.

6. Conclusions

The sedimentary rocks of the study area in central Finland can be divided to Archaean, au-tochthonous and allochthonous cover, and Sve-cofennian further divided into central and southern Svecofennian. The main conclusions are as follows:

Archaean sedimentary rocks can be divided to two main groups those that have a dominant component from a weathered 3.0 – 3.2 Ga green-stone+granite9TTG and those a local 2.7 Ga source, respectively.

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(2.2 – 2.35 Ga), and sedimentary rocks derived from it, and mafic dykes and plateau volcanics (mainly 2.2 – 2.1 Ga) are the major sources but local non-weathered Archaean sources dominate in places. Anoxic (euxinic?) conditions pre-vailed during deposition of the most sulphide-rich rocks.

Allochthonous Western Kaleva sedimentary rocks were deposited both on Archaean basement and oceanic crust (1.95 Ga ophiolites). The most characteristic feature of Western Kaleva sandy greywackes is an extreme compositional homo-geneity. Source components are only slightly weathered and comprise Archaean crust and 2.0 – 1.92 low-K bimodal rocks from a primitive island arc. A foredeep origin associated with subsidence during initial collision is favoured and orogenic detritus either from the same, oblique collision zone (mainly from an accretionary prism) or a more distal orogenic domain is proposed (cf. Lahtinen, 1994; Kohonen, 1995).

The central Svecofennian sedimentary rocks can be divided into local arc-derived rocks (5

1.89 Ga) and older (]1.91 Ga) rocks for which a mixture of Western Kaleva sources and 1.91 – 2.0 Ga mature island arc/active continental margin source is proposed. Rifting (]1.91 Ga) followed by increasing subsidence during initial collision in the NE and subsequent arc reversal causing abun-dant erosion from the mountain belt and exposing different source compositions as seen in the varia-tion of Th/Sc (2 – 0.5), and deposition into oblique hinterland basin further developing into subduc-tion related foredeep is the proposed model for the deposition of the main part of the older turbidites in the central Svecofennian.

The southern Svecofennian (Rantasalmi – Haukivuori area) mature greywackes resemble passive margin sediments and sources dominated by inferred alkaline-affinity complexes is pro-posed. Less mature rocks occur also with sources characterized either by island arc/active continen-tal margin domain or local picritic rocks. It is important to note the absence of the southern Svecofennian-type mature greywackes from the central Svecofennian, which favours the existence of a suture between these areas as proposed by Lahtinen (1996).

A supercontinent stage at 2.6 – 2.1 Ga is pro-posed for the Fennoscandian Shield and the Ar-chaean – Proterozoic transition up to 2.1 Ga was dominated by input of a mainly mafic plateau-type volcanic contribution into the sedimentary record. Palaeoproterozoic sediments having crustal components (52.1 Ga) show higher Th/ Sc, Th/Cr, and lower Sm/Nd and Eu/Eu* relative to the Archaean rocks as proposed in earlier studies (Taylor and McLennan, 1985; McLennan and Taylor, 1991; McLennan and Hemming, 1992) but local low Th/Cr ratios complicate the situation. The behaviour of GdN/YbNratio is also opposite to that proposed by McLennan and Tay-lor (1991).

Ba depletion relative to K, Rb and Th (cf. Lahtinen, 1996) is a characteristic feature of the sedimentary rocks of the central Fennoscandian Shield indicating large amounts of Ba lost from the clastic record during 2.3 – 1.9 Ga. Ba depletion seems to have been especially characteristic for chemical weathering during 2.35 – 2.2 Ga under CO2-rich and low-O2 atmospheric conditions, which could imply that large amounts of Ba have recycled back to the mantle forming a ‘peak’ in the formation of enriched mantle component. Whether the strong Ba depletion is characteristic of the Archaean – Proterozoic transition globally and quiet supercontinent stages in general is to be determined.

Acknowledgements

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References

Awwiller, D.N., Mack, L.E., 1991. Diagenetic modification of Sm – Nd model ages in tertiary sandstones and shales, Texas Gulf Coast. Geology 19, 311 – 314.

Barbey, P., Couvert, J., Moreau, B., Capdevila, R., 1984. Petrogenesis and evolution of an Early Proterozoic colli-sional orogenic belt: the granulite belt of Lapland and the Belomorides (Fennoscandia). Bull. Geol. Soc. Finland 56, 161 – 168.

Berner, R.A., Raiswell, R., 1984. C/S method for distinguish-ing freshwater from marine sedimentary rocks. Geology 12, 365 – 368.

Berner, R.A., 1984. Sedimentary pyrite formation: an update. Geochim. Cosmochim. Acta 48, 605 – 615.

Bostock, H.H., van Breemen, O., 1994. Ages of detrital and metamorphic zircons and monazites from a pre-Taltson magmatic zone basin at the western margin of Rae Province. Can. J. Earth Sci. 31, 1353 – 1364.

Breit, G.N., Wanty, R.B., 1991. Vanadium accumulation in carbonaceous rocks: a review of geochemical controls dur-ing deposition and diagenesis. Chem. Geol. 91, 83 – 97. Claesson, S., Huhma, H., Kinny, P.D., Williams, I.S., 1993.

Svecofennian detrital zircon ages — implications for the Precambrian evolution of the Baltic Shield. Precambrian Res. 64, 109 – 130.

Condie, K.C., Wronkiewicz, D.J., 1990a. A new look at the Archaean – Proterozoic boundary: Sediments and the tec-tonic setting constraint. In: Naqvi, S.M. (Ed.), The Pre-cambrian Continental Crust and its Economic Resources. Elsevier, Amsterdam, pp. 61 – 89.

Condie, K.C., Wronkiewicz, D.J., 1990b. The Cr/Th ratio in Precambrian pelites from the Kaapvaal Craton as an index of craton evolution. Earth Planet. Sci. Lett. 97, 256 – 267. Condie, K.C., Dengate, J., Cullers, R.L., 1995. Behavior of

rare earth elements in paleoweathering profile on granodi-orite in the Front Range, Colorado, USA. Geochim. Cos-mochim. Acta 59, 279 – 294.

Condie, K.C., 1993. Chemical composition and evolution of the upper continental crust: contrasting results from sur-face samples and shales. Chem. Geol. 104, 1 – 37. Davis, D.W., Hirdes, W., Schaltegger, U., Nunoo, E.A., 1994.

U – Pb age constraints on deposition and provenance of Birimian and gold-bearing Tarkwaian sediments in Ghana, West Africa. Precambrian Res. 67, 89 – 107.

Duddy, I.R., 1980. Redistribution and fractionation of rare earth and other elements in a weathering profile. Chem. Geol. 30, 363 – 381.

Fedo, C.M., Nesbitt, H.W., Young, G.M., 1995. Unraveling the effects of potassium metasomatism in sedimentary rocks and paleosols, with implications for paleoweathering conditions and provenance. Geology 23, 921 – 924. Fedo, C.M., Young, G.M., Nesbitt, H.W., Hanchar, J.M.,

1997. Potassic and sodic metasomatism in the Southern Province of the Canadian Shield: evidence from the Pale-oproterozoic Serpent Formation, Huronian Supergroup, Canada. Precambrian Res. 84, 17 – 36.

Gaa´l, G., Gorbatchev, R., 1987. An outline of the Precam-brian evolution of the Baltic Shield. PrecamPrecam-brian Res. 35, 15 – 52.

Gall, Q., 1992. Precambrian paleosols in Canada. Can. J. Earth Sci. 29, 2530 – 2536.

Gao, S., Wedepohl, K.H., 1995. The negative Eu anomaly in Archean sedimentary rocks: implications for decomposi-tion, age and importance of their granitic sources. Earth Planet. Sci. Lett. 133, 81 – 94.

Gibbs, A.K., Montgomery, C.W., O’Day, P.A., Erslev, E.A., 1986. The Archean – Proterozoic transition: evidence from the geochemistry of metasedimentary rocks of Guyana and Montana. Geochim. Cosmochim. Acta 50, 2125 – 2141. Hawkesworth, C.J., Hergt, J.M., McDermott, F., Ellam,

R.M., 1991. Destructive margin magmatism and the con-tributions from the mantle wedge and subducted crust. Aust. J. Earth Sci. 38, 577 – 594.

Hemming, S.R., McLennan, S.M., Hanson, G.N., 1995. Geo-chemical and Nd/Pb isotopic evidence for the provenance of the early Proterozoic Virginia Formation, Minnesota. Implications for the tectonic setting of the Animikie basin. J. Geol. 103, 147 – 168.

Huhma, H., Claesson, S., Kinny, P.D., Williams, I.S., 1991. The growth of Early Proterozoic crust: new evidence from Svecofennian zircons. Terra Nova 3, 175 – 179.

Huhma, A., 1975. Outokummun, Polvija¨rven ja Sivak-kavaaran kartta-alueiden kalliopera¨ (English summary). Geological map of Finland 1:100 000. Explanation to the map of rocks, sheet 4222, 4224, 4311. Geological Survey of Finland, Espoo, Finland, p. 151.

Huhma, H., 1986. Sm – Nd, U – Pb and Pb – Pb isotopic evi-dence for the origin of the Early Proterozoic Svecokarelian crust in Finland. Geol. Surv. Finland Bull. 337, 48. Huhma, H., 1987. Provenance of early Proterozoic and

Ar-chaean metasediments in Finland: a Sm – Nd isotopic study. Precambrian Res. 35, 127 – 143.

Jones, B., Manning, D.A.C., 1994. Comparison of geochemi-cal indices used for the interpretation of palaeoredox con-ditions in ancient mudstones. Chem. Geol. 111, 111 – 129. Karhu, J., 1993. Paleoproterozoic evolution of the carbon isotope ratios of sedimentary carbonates in the Fennoscan-dian Shield. Geol. Surv. Finland Bull. 371, 87.

Kohonen, J., Marmo, J., 1992. Proterozoic lithostratigraphy and sedimentation of Sariola and Jatuli-type rocks in the Nunnanlahti – Koli – Kaltimo area, eastern Finland; impli-cations for regional basin models. Geol. Surv. Finland Bull. 364, 67.

Kohonen, J., 1994. Post-depositional K-feldspar breakdown and its implications for metagraywacke provenance studies — an example from north Karelia, eastern Finland. In: Nironen, M., Ka¨hko¨nen, Y. (Eds.), Geochemistry of Proterozoic Supracrustal Rocks in Finland. Geological Survey of Finland, Espoo, Finland, pp. 161 – 171 Special Paper 19.