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

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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E-mail address:raimo.lahtinen@gsf.fi (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 –

Protero-zoic transition and quiet supercontinent stages in general remains to be determined. © 2000 Elsevier Science B.V. All rights reserved.

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).

The study area has been sampled in the course of a regional bedrock geochemical survey under-taken by the Geological Survey of Finland in-cluding the 300 metasedimentary samples discussed here. The samples range from Archaean to Palaeoproterozoic, were formed in a variety of tectonic settings, and are thus suitable for study-ing the Archaean – Proterozoic transition and the

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 149

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

Fig. 1. Simplified geological map of Finland and surrounding areas modified from Sorjonen-Ward (1993), Korsman et al. (1997). The study area is outlined (see Fig. 2).

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 150

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

The cratonic part of the study area (Fig. 1) includes rocks from Archaean (mainly 2.76 – 2.73 Ga; Vaasjoki et al., 1993) and Palaeoproterozoic cratonic stage (2.5 – 2.1 Ga) with coeval and sub-sequent multiple rifting (e.g. Vuollo, 1994; Ko-honen, 1995) in which the latest phase led to formation of ophiolitic sequences (1.95 Ga; Pel-tonen et al., 1996). The cratonic cover in the Ho¨ytia¨inen and Suvasvesi areas (Fig. 2) are domi-nated by autochthonous and allochthonous rocks, respectively. The Ho¨ytia¨inen area or rift basin (Ward, 1987) includes the Tohmaja¨rvi volcanic complex (2105915 Ma; Huhma, 1986) and asso-ciated coarse clastic deposits but is dominated by mica schists representing metamorphosed thinly laminated pelites to massive turbidites (Ward, 1987; Kohonen, 1995). The formal lithostrati-graphic procedure has been applied only to the autochthonous Sariola, Jatuli and Ludian groups in the eastern margin of the Ho¨ytia¨inen area

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 151

(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

The Archaean metagreywackes and mica schists/gneisses have been divided into two main groups (Ar1 – Ar2). The Ar1 rocks have a homo-geneous composition indicating a thorough mix-ing of source components. The elevated CIA (Chemical Index of Alteration; Nesbitt and Young, 1982) shows the effects of weathering in the source area and the REE, major and trace

R . Lahtinen / Precambrian Research 104 (2000) 147 – 174 152 Table 1

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

R

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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.

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 154

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 Ho¨ytia¨inen 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 Ho¨ytia¨inen 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

Fig. 3. Major- and trace-element distributions in Karelian craton 1, Western Kaleva psammites (WK1), Jatuli-type mafics and Kutsu-type granites normalized to Archaean Crust (AC1 in Table 1). The Karelian craton (KC1) and WK1 averages are from Table 1 and the averages for Jatuli-type mafics (N=21) and Kutsu granites (N=8) are from Lahtinen (unpublished data).

R . Lahtinen / Precambrian Research 104 (2000) 147 – 174 155 Table 2

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

282

R

.

Lahtinen

/

Precambrian

Research

104

(2000)

147

–

174

156

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.

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 157

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).

The allochthonous Western Kaleva (WK) sedi-mentary rocks have been divided into WK psam-mites and SiO2-poor pelitic rocks (WK2). The WK1 psammites (Table 1) form a geochemically homogeneous group (Fig. 4) and most of the variation can be explained by grain size variation. The more pelitic nature of WK2 is seen in enrich-ment of eleenrich-ments (e.g. Al2O3, MgO, FeO, K2O) that characterize clay minerals (Table 1) but the WK2 also seems to be enriched in a mafic source as seen in higher Sc and Cr relative to Th. The WK1 migmatites are mainly psammitic fragments floating in tonalitic (often trondjhemitic) veined

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 158

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

calculated from the average in the Table 1 and Jatuli-type mafics from the average (N=21) in Lahtinen (unpublished data). Ar – Ar2-Archaean groups, Jqtz – Jatuli-type quartzites, H1 – H2-autochthonous high-Cr, H3-autochthonous low-Cr, RH1 – RH2-southern Svecofennian, RH3-RH2-southern Svecofennian. CF1 – CF3-central Svecofennian.

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 159

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.

The sedimentary rocks in the central Svecofen-nian have been divided to three groups (CF1 – CF3) where the CF1 includes high-SiO2 and high Th/Sc (]1) psammites, CF2 lower Th/Sc (51) psammites and CF3 silt-pelite rocks. The non-migmatitic CF1 samples show LREE enrichment compared to the Western Kaleva psammites (Fig.

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 160

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).

The REE, Th and HFSE (especially Sc) are considered least susceptible to fractionation by exogene processes including weathering (Taylor and McLennan, 1985; McLennan et al., 1990). REE mobility during weathering has been never-theless observed (Nesbitt, 1979; Duddy, 1980; Condie et al., 1995) although Nesbitt (1979), Duddy (1980) found no net losses or gains when whole weathering profiles were considered. Deple-tion of Sc has been postulated during weathering under low-O2 atmosphere (Maynard et al., 1995). The Palaeoproterozoic autochthonous units above Archaean basement in the study area (Ko-honen and Marmo, 1992 and references therein) start with the Ilvesvaara Formation overlain by the glaciogenic Urkkavaara Formation followed

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 161

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

mixtures of sandstone zones (50%) and granite zones (50%). J and K are calculated averages of Jatuli-type mafics and Kutsu-type granites, respectively.

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 162

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.

Only autochthonous (H1 – H3) and al-lochthonous (WK1 – WK2) groups have sufficient samples with carbonaceous matter (graphite) to be plotted in the S vs. C diagram (Fig. 8). The Ho¨ytia¨inen pelitic samples, especially H2 samples, show good correlation between S and C (S/C about 3.5). There is also a slight increase in U seen in the decrease of Th/U ratios from about 4 to about 2.5 in the H2 samples. These features

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 163

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

Monitoring the effects of diagenesis in meta-morphic rocks is a difficult task due recrystalliza-tion requiring that any evaluation of the diagenetic history be based on geochemistry.

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 Ho¨ytia¨inen 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 Ho¨ytia¨inen 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).

Fig. 9. Plots of Sm/Nd vs. Th/Sc and Th/Cr for selected sedimentary rocks in this study. The Archaean average has been calculated from the average AC1 in Table 1 and Jatuli-type mafics from the average (N=21) in Lahtinen (unpublished data). See Fig. 5.

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 165

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

Local Archaean craton sources with contribu-tions from Jatuli-type mafic volcanics and dykes has been a common source model for the Ho¨yti-a¨inen basin sedimentary rocks (Huhma, 1987; Ward, 1987; Kohonen, 1995). The results of this study favour this general statement but the com-position of the H1 – H2 groups is not explained by simple mixing of the presently exposed erosion level of the Archaean crust and Jatuli-type mafics (Figs. 4, 5 and 9) because an additional Cr-rich source is needed. The simplest explanation is higher amounts of Archaean sedimentary rocks (Cr-rich) in the average source area for the H1 – H2 group samples. Some sedimentary rocks have high proportions of local Archaean cratonic

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 166

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

The central Svecofennian psammites show large compositional variation (Figs. 5 and 9) in-dicating either different provenance areas or changes in the composition of source areas dur-ing erosion. The latter is favoured (cf. Lahtinen, 1996) and, if this is the case, it points to rather short transport distances from a rapidly rising orogenic domain. The psammites of this study (CF1 – CF2) and the basement-related sedimen-tary psammites (SG3 – SG4) of Lahtinen (1996) from the Tampere – Ha¨meenlinna area have geo-chemical similarities as seen in Th/Sc ratios of 2 – 0.7 and 1.5 – 0.5, respectively. The TDM of 2.2 Ga (Huhma, 1987) from one sample is slightly younger than that found in the WK psammites (2.3 – 2.4 Ga). Assuming that the central Sve-cofennian rocks are mixtures of Western Kaleva-type source and an additional source it is possible to use the WK1 as a normalizing value to infer the nature of this additional component. The CF1 psammites are enriched in felsic com-ponent and thus the differences to the WK1 should approximate the felsic composition. High LREE, LaN/LuNand negative Eu anomalies with high Th and low K/Rb and Nb favour a mature intracrustal origin. The Th variation (13 – 19 ppm) and Th/Ta ratios (typically ]15) in the CF1 are distinct from Th contents (B9 ppm, mostly B4 ppm) and Th/Ta ratios (59) found in the 1.93 – 1.91 Ga primitive island arc felsic rocks (see Figure 26 in Lahtinen, 1994) excluding them as a dominant felsic component in the CF1. The CF1 and CF2 groups are gradational to each other and the low Cr/Sc ratio, low Nb and only slight TiO2 enrichment relative to MgO favour a mature island arc origin for the added mafic – intermediate component. The proposed main source components for the central Sve-cofennian sedimentary rocks are as follows:

R.Lahtinen/Precambrian Research104 (2000) 147 – 174 167

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

Kohonen (1995) suggested that thesyn-rift tur-bidites of the Ho¨ytia¨inen rift basin (Ward, 1987) have a maximum depositional age of about 2.1 Ga. The post-rift marine sediments probably in-cluded both passive margin and foredeep deposits where the latter were deposited during foredeep migration from west to east during initial conti-nent-arc collision (Kohonen, 1995). The basic as-sumption is that autochthonous groups (H1 – H3) contain only cratonic detritus where the Palaeo-proterozoic component is from mafic volcanics and dykes (mainly 2.2 – 2.06, and 1.96 Ga). Lahti-nen (1994) Palaeo-proterozoic KohoLahti-nen (1995) considered that the major rifting at 2.1 – 2.06 Ga finally lead to continental break-up (cf. Park et al., 1984; Gaa´l and Gorbatchev, 1987) and a change to a passive margin environment. A model with later continental break-up at 1.95 Ga has also been proposed (Peltonen et al., 1996). The Western Kaleva psammites have been described as Svecofennian post-arc flysch (Park, 1985), a molasse from the Lapland granulite belt (Barbey et al., 1984), pericontinental turbidites including the Kalevian as a whole (Laajoki, 1986), deep-wa-ter slope-rise greywackes related to uplift in Lap-land and the Kola Peninsula (Kontinen and Sorjonen-Ward, 1991), a mixture of accretion prism sediments and derived foredeep sediments (Lahtinen, 1994) and axial foredeep deposits from a rising orogenic domain in the north during arc (Svecofennian) – continent (Karelian craton) colli-sion (Kohonen, 1995). The model of Kohonen (1995) could explain the occurrence of both an inferred 1.92 – 2.0 Ga low-K primitive island arc component and a non-weathered Archaean

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 Gd

N

/

Yb

N

ratio 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 (

9

felsic 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

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.

The source variation and mobile nature of Ba

produces scatter in Fig. 10 but Ba depletion is

noticed in most samples. A striking feature is the

strong relative Ba depletion in most high-Cr

Ho¨y-tia¨inen rocks (H1 – H2) and Jatulian quartzites

implying that the Ba depletion is related to the

chemically weathered component (cf. Maynard et

al., 1995). Major loss of alkali and alkaline earth

metals relative to more immobile elements (REE,

Th, Sc) occur at CIA values of about 80 in the

Hokkalampi Palaeosol when the breakdown of

illite dominates (Marmo, 1997; personal

commu-nication). Similar Ba depletion relative to Rb was

also noticed (not shown). Potassium enrichment

in palaeosols, attributed to diagenetic

overprint-ing, is not uncommon (e.g. Gall, 1992 and

refer-ences therein) and could cause relative Ba

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 highly

13

_{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

CO

2

-rich and low-O

2

atmosphere 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

+granite9

TTG and those a local 2.7 Ga

source, respectively.

Autochthonous sedimentary rocks have

in-ferred deposition ages 2.2 – 1.9 Ga and vary from

rift to passive margin and foredeep deposits

(Ko-honen, 1995). Chemically weathered palaeosol

(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 deposivaria-tion 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).

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 (

5

2.1 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 Gd

N

/Yb

N

ratio 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

CO

2

-rich and low-O

2

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

This work was carried out in the context of

regional rock geochemical study at the Geological

Survey of Finland and my fellow researches Esko

Korkiakoski, Pekka Lestinen, Reijo Salminen and

Heimo Savolainen are thanked for their effort to

accomplish the project and for their continuous

interest in the subject. I am also grateful to Gabor

Gaa´l, Hannu Huhma, Jarmo Kohonen and Hugh

O’Brien for critically reading an earlier version of

the manuscript. Reviewers K.C. Condie and C.M.

Fedo are thanked for their constructive comments

on the manuscript. This work is published with

the permission of the Geological Survey of

Finland.

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.

Kohonen, J., 1995. From continental rifting to collisional shortening — Paleoproterozoic Kaleva metasediments of the Ho¨ytia¨inen area in North Karelia, Finland. Geol. Surv. Finland Bull. 380, 79.

158.

Kontinen, A., Sorjonen-Ward, P., 1991. Geochemistry of metagraywackes and metapelites from the Paleoprotero-zoic Nuasja¨rvi Group, Kainuu Schist Belt and the Savo Province, north Karelia: Implications for provenance, lithostratigraphic correlation and depositional setting. In: Autio, S. (Ed.), Current Research 1989 – 1990. Geological Survey of Finland, Espoo, Finland, pp. 21 – 22 Special Paper 12.

Korsch, R.J., Roser, B.P., Kamprad, J.L., 1993. Geochemical, petrographic and grain-size variations within single turbid-ite beds. Sedimentary Geol. 83, 15 – 35.

Korsman, K., Ho¨ltta¨, P., Hautala, T., Wasenius, P., 1984. Metamorphism as an indicator of evolution and structure of the crust in eastern Finland. Geol. Surv. Finland Bull. 328, 40.

Korsman, K., Koistinen, T.J., Kohonen, J., Wennerstro¨m, M., Ekdahl, E., Honkamo, M., Idman, H., Pekkala, Y., 1997. Suomen kalliopera¨kartta — Berggrundskarta o¨ver Finland — Bedrock map of Finland 1: 1 000 000. Geological Sur-vey of Finland, Espoo, Finland.

GGT/SVEKA Working Group, Korsman, K., Korja, T., Pa-junen, M., Virransalo, P., 1999. The GGT/SVEKA Tran-sect — Structure and evolution of the continental crust in the Palaeoproterozoic Svecofennian orogen in Finland. Int. Geol. Rev. 41, 287 – 333.

Laajoki, K., 1986. The Precambrian supracrustal rocks of Finland and their tectono-exogenic evolution. Precambrian Res. 33, 67 – 85.

Lahtinen, R., Huhma, H., 1997. Isotopic and geochemical constraints on the evolution of the 1.93 – 1.79 Ga Svecofen-nian crust and mantle in Finland. Preccambrian Res. 82, 13 – 34.

Lahtinen, R., 1994. Crustal evolution of the Svecofennian and Karelian domains during 2.1 – 1.79 Ga, with special empha-sis on the geochemistry and origin of 1.93 – 1.91 Ga gneissic tonalites and associated supracrustal rocks in the Rauta-lampi area, central Finland. Geol. Surv. Finland Bull. 378, 128.

Lahtinen, R., 1996. Geochemistry of Palaeoproterozoic supracrustal and plutonic rocks in the Tampere – Ha¨meen-linna area, southern Finland. Geol. Surv. Finland Bull. 389, 113.

Lee, Y.I., Klein, G.deV., 1986. Diagenesis of sandstones in the back-arc basins of the western Pacific Ocean. Sedimentol-ogy 33, 651 – 675.

Loukola-Ruskeeniemi, K., Heino, T., 1996. Geochemistry and genesis of the black shale-hosted Ni – Cu – Zn deposit at Talvivaara, Finland. Econ. Geol. 91, 80 – 110.

Machado, N., Schrank, A., Noce, C.M., Gauthier, G., 1996. Ages of detrital zircon from Archean – Paleoproterozoic sequences: implications for Greenstone Belt setting and evolution of a Transamazonian foreland basin in Quadr-la´tero Ferrı´fero, southeast Brazil. Earth Planet. Sci. Lett. 141, 259 – 276.

Mineralogy and Energy Deposits. Springer, Berlin, pp. 41 – 66.

Maynard, J.B., Sutton, S.J., Robb, L.J., Ferraz, M.F., Meyer, F.M., 1995. A paleosol developed on hydrothermally al-tered granite from the hinterland of the Witwatersrand basin: characteristics of a source of basin fill. J. Geol. 103, 357 – 377.

McLennan, S.M., Hemming, S., 1992. Samarium/neodymium elemental and isotopic systematics in sedimentary rocks. Geochim. Cosmochim. Acta 56, 887 – 898.

McLennan, S.M., Taylor, S.R., 1991. Sedimentary rocks and crustal evolution: tectonic setting and secular trends. J. Geol. 99, 1 – 21.

McLennan, S.M., Taylor, S.R., McCulloch, M.T., Maynard, J.B., 1990. Geochemical and Nd – Sr isotopic composition of deep-sea turbidites: crustal evolution and plate tectonic associations. Geochim. Cosmochim. Acta 54, 2015 – 2050. Milodowski, A.E., Zalasiewicz, J.A., 1991. Redistribution of

rare earth elements during diagenesis of turbidite/ hemipelagite mudrock sequences of Llandovery age from Central Wales. In: Morton, A.C., et al. (Eds.), Develop-ments in Sedimentary Provenance Studies. Geological So-ciety, pp. 101 – 124 Special Publication 57.

Nesbitt, H.W., Young, G.M., 1982. Early Proterozoic climates and plate motions inferred from major element chemistry of lutites. Nature 299, 715 – 717.

Nesbitt, H.W., Young, G.M., 1984. Prediction of some weath-ering trends of plutonic and volcanic rocks based on thermodynamic and kinetic consideration. Geochim. Cos-mochim. Acta 48, 1523 – 1534.

Nesbitt, H.W., Young, G.M., 1989. Formation and diagenesis of weathering profiles. J. Geol. 97, 129 – 147.

Nesbitt, H.W., Young, G.M., McLennan, S.M., Keays, R.R., 1996. Effects of chemical weathering and sorting on the petrogenesis of siliciclastic sediments, with implications for provenance studies. J. Geol. 104, 525 – 542.

Nesbitt, H.W., 1979. Mobility and fractionation of rare earth elements during weathering of a granodiorite. Nature 279, 206 – 210.

O’Brien, H.E., Huhma, H., Sorjonen-Ward, P., 1993. Petroge-nesis of the late Archean Hattu schist belt, Ilomantsi, eastern Finland: geochemistry and Sr, Nd isotopic compo-sition. In: Nurmi, P.A., Sorjonen-Ward, P. (Eds.), Geolog-ical development, gold mineralization and exploration methods in the late Archean Hattu schist belt, Ilomantsi, eastern Finland. Geological Survey of Finland, Espoo, Finland, pp. 147 – 184 Special Paper 17.

Ohr, M., Halliday, A.N., Peacor, D.R., 1991. Sr and Nd isotopic evidence for punctuated clay diagenesis, Texas Gulf Coast. Earth Planet. Sci. Lett. 105, 110 – 126. Park, A.F., Bowes, D.R., Halden, N.M., Koistinen, T.J., 1984.

Tectonic evolution at an early Proterozoic continental margin: the Svecokarelides of eastern Finland. J. Geody-nam. 1, 359 – 386.

Park, A.F., 1985. Accretion tectonism in the Proterozoic Sve-cokarelides of the Baltic Shield. Geology 13, 725 – 729. Pekkarinen, L.J., Lukkarinen, H., 1991. Paleoproterozoic

vol-canism in the Kiihtelysvaara – Tohmaja¨rvi district, eastern Finland. Geol. Surv. Finland Bull. 357, 30.

Pekkarinen, L.J., 1979. The Karelian formations and their depositional basement in the Kiihtelysvaara – Va¨rtsila¨ area, east Finland. Geol. Surv. Finland Bull. 301, 141. Peltonen, P., Kontinen, A., Huhma, H., 1996. Petrology and

geochemistry of metabasalts from the 1.95 Ga old Jormua Ophiolite, northeastern Finland. J. Petrol. 37, 1359 – 1383. Pesonen, L.J., Leino, M.A.H., Mertanen, S., 2000. Palaeomag-netism and the Finnish GGT-project, Geol. Soc. Finland Bull., in press.

Raiswell, R., Berner, R.A., 1986. Pyrite and organic matter in Phanerozoic normal marine shales. Geochim. Cosmochim. Acta 50, 1967 – 1976.

Rautiainen, I., Hagel-Brunnstro¨m, M.-L., Heiskanen, L., Kallio, E., 1996. Production oriented method for the deter-mination of rare earth and other trace elements in rocks by ICP-MS. Fresenius J. Anal. Chem. 355, 393 – 396. Sandstro¨m, H., 1996. The analytical methods and the precision

of the element determinations used in the regional bedrock geochemistry in the Tampere – Ha¨meenlinna area, southern Finland. Geol. Surv. Finland Bull. 393, 25.

Sheraton, J.W., 1984. Chemical changes associated with high-grade metamorphism of mafic rocks in the East Antarctic Field. Chem. Geol. 47, 135 – 157.

Singer, A., 1980. The paleoclimatic interpretation of clay minerals in soils and weathering profiles. Earth Sci. Rev. 15, 303 – 326.

Sorjonen-Ward, P., 1993. An overview of structural evolution and lithic units within and intruding the late Archean Hattu schist belt, Ilomantsi, eastern Finland. In: Nurmi, P.A., Sorjonen-Ward, P. (Eds.), Geological Development, Gold Mineralization and Exploration Methods in the Late Archean Hattu Schist Belt, Ilomantsi, Eastern Finland. Geological Survey of Finland, Espoo, Finland, pp. 9 – 102 Special Paper 17.

Sturt, B.A., Meleznik, V.A., Ramsay, D.M., 1994. Early Proterozoic regolith at Pasvik, NE Norway: palaeoenvi-ronmental implications for the Baltic Shield. Terra Nova 6, 618 – 633.

Taylor, S.R., McLennan, S.M., 1985. The Continental Crust: Its Composition and Evolution. Blackwell, Oxford, UK, p. 312.

Taylor, S.R., Rudnick, R.L., McLennan, S.M., Eriksson, K.A., 1986. Rare element patterns in Archean high-grade metasediments and their tectonic significance. Geochim. Cosmochim. Acta 50, 2267 – 2279.

Thompson, J.F.H., Naldrett, A.J., 1984. Sulphide – silicate re-actions as a guide to Ni – Cu – Co mineralization in central Maine, USA. In: Buchanan, D.L., Jones, M.L. (Eds.), Sulphide Deposits in Mafic and Ultramafic Rocks. Pro-ceedings IGCP projects 161 and 91, Third Nickel Sulphide Field Conference, Perth, Western Australia, 23 – 25 May, 1982. Institution of Mining and Metallurgy, London, pp. 103 – 113.

Vaasjoki, M., Sorjonen-Ward, P., Lavikainen, S., 1993. U – Pb age determinations and sulphide Pb – Pb characteristics from the late Archean Hattu schist belt, Ilomantsi, eastern Finland. In: Nurmi, P.A., Sorjonen-Ward, P. (Eds.), Geo-logical Development, Gold Mineralization and Explo-ration Methods in the Late Archean Hattu Schist Belt, Ilomantsi, Eastern Finland. Geological Survey of Finland, Espoo, Finland, pp. 103 – 131 Special Paper 17.

Veizer, J., Jansen, S.L., 1985. Basement and sedimentary recy-cling — 2: time dimensions to global tectonics. J. Geol. 93, 625 – 643.

Vuollo, J., 1994. Paleoproterozoic basic igneous events in eastern Fennoscandian Shield between 2.45 and 1.97 Ga. Ph.D. thesis. Acta Univ. Ouluensis Ser. A 250, 47. Ward, P., 1987. Early Proterozoic deposition and deformation

at the Karelian craton margin in southeastern Finland. Precambrian Res. 35, 71 – 93.

Weaver, B.L., Tarney, J., 1983. Elemental depletion in Ar-chaean granulite-facies rocks. In: Atherton, M.P., Gribble, C.D. (Eds.), Migmatites, Melting and Metamorphism. Shiva, pp. 250 – 263.

Weaver, B.L., 1991. The origin of ocean island basalt end-member compositions: trace element and isotopic con-straints. Earth Planet. Sci. Lett. 104, 381 – 397.

Zhao, J.X., McCulloch, M.T., Bennet, V.C., 1992. Sm – Nd and U – Pb zircon isotopic constraints on the provenance of sediments from the Amadeus Basin, central Australia: evi-dence for REE fractionation. Geochim. Cosmochim. Acta 56, 921 – 940.