Archaean–Proterozoic transition: geochemistry, provenance ...directory.umm.ac.id/Data Elmu/jurnal/P/Precambrian...Precambrian Research 104 (2000) 147–174 Archaean–Proterozoic

Precambrian Research 104 (2000) 147–174

Archaean–Proterozoic transition: geochemistry, provenanceand tectonic setting of metasedimentary rocks in central

Fennoscandian Shield, Finland

Raimo Lahtinen *Geological Sur6ey of Finland, P.O. Box 96, FIN-02151 Espoo, Finland

Received 8 July 1999; accepted 5 May 2000

Abstract

The central part of the Fennoscandian Shield in Finland is composed of the Palaeoproterozoic Svecofenniandomain and the Archaean Karelian craton with a Palaeoproterozoic allochthonous and autochthonous cover. Acryptic suture separating these areas and another tentative suture dividing the Svecofennian into central and southernparts have been proposed. The chemical composition of sedimentary rocks (N=300) within the study area, includingthe effects of palaeoweathering, hydraulic sorting, depositional environment and post-depositional processes, havebeen studied in order to delineate sediment source components. The main proposed source components for theArchaean 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 2.2–2.1Ga) although in places locally derived non-weathered Archaean sources dominated. Archaean crust and 2.0–1.92low-K bimodal rocks from a primitive island arc are the proposed source for the allochthonous Western Kaleva coverrocks. These formed in a subsiding foredeep during initial collision from orogenic detritus in the same obliquecollision 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/activecontinental margin source. Rifting followed by increased subsidence during initial collision in the NE and subsequentarc reversal caused rapid erosion from the mountain belt, exposing diverse source compositions as seen in the largevariation of Th/Sc (2–0.5), and deposition into an oblique hinterland basin further developing into a subductionrelated foredeep. Mature greywackes from the southern Svecofennian in the study area resemble passive marginsediments with a source dominated by inferred alkaline-affinity complexes and Archaean rocks. Less mature rocksalso occur and had sources dominated either by island arc/active continental margin rocks or local picritic rocks. Inthe sedimentary record the Archaean–Proterozoic transition up to 2.1 Ga was dominated by input of mainly maficplateau-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 lowTh/Cr ratios complicate the situation. Ba depletion relative to K, Rb and Th is a characteristic feature of the

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R. Lahtinen / Precambrian Research 104 (2000) 147–174148

sedimentary rocks of the central Fennoscandian Shield indicating high amounts of Ba lost from the clastic recordduring 2.3–1.9 Ga and further recycled back to the mantle forming a subduction component and an enriched mantlecomponent. Ba depletion seems to have been especially characteristic of chemical weathering during 2.35–2.2 Gaunder CO2-rich and low-O2 atmosphere. 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. Allrights reserved.

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

1. Introduction

The geochemistry of clastic sedimentary rockscan be used as an indicator of crustal evolution(e.g. Taylor and McLennan, 1985) or to identifyancient tectonic settings in metamorphic terranes.Sedimentary rocks can be divided into thoseshowing local sources and those having experi-enced effective mixing in large river marine sys-tems before deposition. The latter types samplelarge areas providing data for crustal-scale pro-cesses. The possibility of different crust-formingmechanisms during Archaean and Proterozoictimes emphasizes the importance of the Ar-chaean–Proterozoic boundary where there mightbe a corresponding compositional change in thesedimentary record (e.g. Taylor and McLennan,1985; McLennan and Taylor, 1991). Selectivepreservation of sedimentary rocks in the ancientrecord can on the other hand hamper their use incrustal evolution studies. Along with this limita-tion, other factors discussed below, should alsobe 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 theclastic sediments but other factors such as degreeof palaeoweathering, hydraulic sorting (grain-sizeeffects), organic and sulphide input, diagenesisand metamorphism (especially migmatization)may greatly modify or ultimately erase prove-nance memory. Sediment recycling is a commonfeature (e.g. Veizer and Jansen, 1985) and pro-duces a buffering effect where a small amount ofnew input can go unnoticed. Nevertheless, eventhough the interpretation of their compositions ismore controversial than with igneous rocks, the

long ‘memory’ of sedimentary rocks can be quitepowerful when modelling the tectonic settingsand evolutionary histories of metamorphic ter-ranes.

The central part of the Fennoscandian Shieldin Finland is composed of the PalaeoproterozoicSvecofennian domain and the Archaean Kareliancraton with a Palaeoproterozoic allochthonousand autochthonous cover (Fig. 1). The occur-rence of a cryptic ‘suture’ (Fig. 1; Koistinen,1981; Huhma, 1986) between the Karelian andSvecofennian domains is favoured by the obser-vation that no Archaean component is found inthe 1.93–1.91 Ga gneissic tonalites and relatedfelsic 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 theSvecofennian domain from the southern Sve-cofennian. Studies on the geochemistry of sedi-mentary rocks in the study area are few andinclude a geochemical and isotopic study fromthe Archaean Hattu schist belt (O’Brien et al.,1993), a major element study from the northernpart of the Höytiäinen area (Kohonen, 1995), aregional correlation diagram study from the Savoprovince (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 courseof a regional bedrock geochemical survey under-taken by the Geological Survey of Finland in-cluding the 300 metasedimentary samplesdiscussed here. The samples range from Archaeanto Palaeoproterozoic, were formed in a variety oftectonic settings, and are thus suitable for study-ing the Archaean–Proterozoic transition and the

R. Lahtinen / Precambrian Research 104 (2000) 147–174 149

evolution of Fennoscandian Shield. The mainsource components and implications for the tec-tonic evolution of the central Fennoscandianshield are given with emphasis on proposed su-tures. Notes on the crustal evolution and Ar-chaean–Proterozoic transition in general, and onBa depletion are also given. As all the studiedsedimentary rocks are metamorphosed, the prefix’meta’ has been dropped. The data set is availableon 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 sixsubsamples (altogether 1–1.5 kg) from the samelithological unit, if detection of unit boundarieswas possible (sometimes this was impossible, e.g.in some migmatites). In the case of turbidites, thewhole Bouma A, AB or ABC was sampled inmost 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).


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 layerswere B5 cm thick and a more homogeneous unitwas not available.

The analytical work was done in the laborato-ries of the Geological Survey of Finland. Sampleswere jaw crushed and splits were pulverized in atungsten carbide bowl for X-ray fluorescence(XRF) analysis, and in a carbon steel bowl forinductively coupled plasma mass spectrometry(ICP-MS). Major elements and Cl, V, Cr, Ni, Zn,Rb, Sr, Y, Zr, Nb and Ba were determined byXRF, CGraf. by Leco CR-12 carbon analyzer, F byion selective electrode, aqua regia leachable S andCu by ICP-AES, and aqua regia leachable Au,Pd, Te, As, Ag, Bi, Sb and Se by GAAS (Sand-ström, 1996). REE, Co, Nb, Hf, Rb, Sc, Ta, Thand U were determined by ICP-MS after dissolu-tion of the sample (0.2 g) with hydrofluoric acid-perchloricacid treatment completed by a lithiummetaborate/sodium perborate fusion (Rautiainenet 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 2.76–2.73Ga; Vaasjoki et al., 1993) and Palaeoproterozoiccratonic 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 toformation of ophiolitic sequences (1.95 Ga; Pel-tonen et al., 1996). The cratonic cover in theHöytiäinen and Suvasvesi areas (Fig. 2) are domi-nated by autochthonous and allochthonous rocks,respectively. The Höytiäinen area or rift basin(Ward, 1987) includes the Tohmajärvi volcaniccomplex (2105915 Ma; Huhma, 1986) and asso-ciated coarse clastic deposits but is dominated bymica schists representing metamorphosed thinlylaminated pelites to massive turbidites (Ward,1987; Kohonen, 1995). The formal lithostrati-graphic procedure has been applied only to theautochthonous Sariola, Jatuli and Ludian groupsin the eastern margin of the Höytiäinen area


(Pekkarinen, 1979; Pekkarinen and Lukkarinen,1991; Kohonen and Marmo, 1992; Karhu, 1993).Otherwise lithostratigraphy and chronostratigra-phy of the Höytiäinen area are not resolved (Ko-honen, 1995) but depositional ages from 2.1 toabout 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 termadopted in this study) greywackes that occur asallochthonous units in thrust complexes charac-terized by associated ophiolites and related rocks(Koistinen, 1981 and references therein) thoughevidence for local deposition upon Archaeanbasement has also been noted (Ward, 1987). Theincrease in metamorphic grade from east to west(Fig. 2) is seen as an increase in quartz veins andthe onset of segregational banding (quartz+feldspar) leading finally to migmatites.

The boundary zone (BZ) includes migmatiticsedimentary rocks (Korsman et al., 1984) and a1.93–1.91 Ga volcano-plutonic formation (Lahti-nen, 1994 and references therein). The Svecofen-nian is divided into the central Svecofennianincluding the Central Finland Granitoid Complex(CFGC) and Bothnian Belt (BB), and the south-ern Svecofennian including the Rantasalmi–Haukivuori area (RH). The tentativesedimentation ages for the central Svecofennian,based on data available from the Tampere SchistBelt (Lahtinen, 1996 and references therein), are]1.91 and 1.89–1.87 Ga for rocks correlated tobasement- and arc-related groups in the TampereSchist Belt, respectively. The southern Svecofen-nian, including the Rantasalmi–Haukivuori area,is characterized by granite migmatites, which is aclear difference to the central Svecofennian,boundary zone and Suvasvesi area, which arecharacterized by tonalite migmatites (Korsman etal., 1999 and references therein).

4. Results

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

acteristics. All elements analyzed have been usedbut 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 etal., 1990). The arc-related (upper) central Sve-cofennian rocks of this study (Fig. 2), not dis-cussed in detail, show CaO, MnO, P2O5, Sr, Baand Sb enrichment, which is characteristic of sed-imentary rocks derived from high-K calc-alkalineto shoshonitic volcanics (Lahtinen, 1996).Strongly altered or mineralized samples are ex-cluded from discussion as are minor groups ofsedimentary rocks either having undefined originsor a large non-clastic component (e.g. iron forma-tions and carbonate rocks).

The group characteristics were also studied byusing normalized diagrams (Fig. 3). Archaeansedimentary groups are normalized to Archaeancrust (AC1), autochthonous and allochthonousgroups to average Karelian craton (KC1) andboundary zone and Svecofennian groups to West-ern Kaleva WK1 (Table 1). The AC1 is a firstapproximation of the average composition of Ar-chaean crust in Finland at its present erosion levelbased solely on the data from the study area. Thegranitoid-dominated nature of the exposed Ar-chaean part of the study area is seen in higherLILE and LREE and lower MgO, Cr and Nicompared to the Late Archaean (3.5–2.5 Ga)restoration model for average juvenile upper con-tinental crust (Table 4 in Condie, 1993). TheKarelian craton includes a large contributionfrom 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 micaschists/gneisses have been divided into two maingroups (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 andYoung, 1982) shows the effects of weathering inthe source area and the REE, major and trace


Tab

le1

Ave

rage

chem

ical

com

posi

tion

ofes

tim

ated

Arc

haea

ncr

ust

(AC

1)an

dK

arel

ian

crat

on(K

C1)

,an

dse

lect

edse

dim

enta

ryro

ckgr

oups

(non

-mig

mat

ized

,ex

cept

grou

psB

Z1–

BZ

2)a

BZ

1A

r1W

K2

H1

WK

1fra

gH

2W

K1

H3

BZ

2K

C1

AC

1(N

=4)

(N=

8)(N

=15

6)(N

=12

9)(N

=5)

(N=

9)(N

=47

)(N

=5)

(N=

17)

(N=

11)

(N=

6)

67.2

369

.85

69.5

863

.23

65.1

568

.60

56.4

2Si

O2

(%)

60.1

665

.15

63.6

465

.18

0.51

0.80

0.62

0.68

0.69

0.83

0.72

1.08

0.65

0.76

TiO

2(%

)0.

7212

.87

14.8

614

.74

13.1

113

.27

15.4

215

.16

17.6

814

.68

Al 2

O3

(%)

15.1

515

.19

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

0.07

0.07

0.07

0.08

0.08

0.11

0.10

0.08

0.08

MnO

(%)

5.19

2.34

2.52

2.26

2.33

3.23

2.84

4.29

2.81

3.55

2.91

MgO

(%)

1.46

2.22

2.42

2.36

2.34

1.68

2.59

CaO

(%)

0.87

1.46

4.06

3.39

4.24

1.24

1.98

2.98

2.76

2.84

2.92

2.93

3.89

2.37

Na 2

O(%

)2.

183.

442.

372.

413.

363.

343.

873.

442.

76K

2O

(%)

2.71

2.35

2.46

0.15

0.18

0.11

0.16

0.15

0.16

0.14

0.11

0.18

0.12

0.13

P2O

5(%

)0.

34(0

.05)

0.13

(0.2

2)(0

.29)

(0.0

7)(0

.05)

(0.0

5)(0

.05)

(0.1

0)0.

09C

gra

f.(%

)0.

210.

067

0.08

20.

061

0.21

1.24

0.23

S(%

)0.

120.

410.

061

0.05

40.

070

0.05

40.

053

0.08

50.

078

0.09

4F

(%)

0.05

50.

051

0.04

50.

062

0.09

462

.64

54.4

54.7

55.8

55.6

57.8

62.9

62.5

49.3

61.9

CIA

50.0

36.2

31.1

32.0

31.6

30.6

33.2

36.7

44.3

31.8

15.2

La

(ppm

)23

.471

.262

.263

.262

.960

.965

.473

.286

.563

.432

.7C

e(p

pm)

47.9

7.27

7.43

7.29

8.02

8.60

7.44

10.1

Pr

(ppm

)5.

674.

127.

428.

2321

.527

.926

.727

.326

.628

.931

.437

.127

.3N

d(p

pm)

15.3

29.7

5.17

5.13

4.98

5.55

5.72

5.49

6.44

3.10

4.28

Sm(p

pm)

4.90

4.76

1.14

1.02

0.96

1.06

1.03

1.15

1.13

1.44

1.07

0.94

0.91

Eu

(ppm

)4.

913.

864.

274.

634.

475.

045.

266.

134.

002.

963.

88G

d(p

pm)

0.66

0.68

0.66

0.75

0.73

0.73

0.90

Tb

(ppm

)0.

610.

480.

550.

503.

363.

683.

424.

123.

75D

y(p

pm)

5.01

2.40

2.76

2.95

3.36

4.21

0.66

0.73

0.68

0.79

0.71

0.79

1.00

0.58

0.67

Ho

(ppm

)0.

450.

532.

311.

261.

882.

122.

042.

312.

033.

061.

501.

761.

94E

r(p

pm)

0.33

0.18

0.27

0.31

0.30

0.32

0.29

0.47

0.21

0.26

0.28

Tm

(ppm

)1.

792.

161.

952.

191.

942.

233.

13Y

b(p

pm)

1.84

1.73

1.38

1.17

0.27

0.32

0.30

0.32

0.27

0.46

Lu

(ppm

)0.

180.

210.

250.

280.

35

570

489

508

613

704

348

712

Ba

(ppm

)39

237

174

285

811

658

.148

.879

.410

013

913

917

215

712

7C

l(p

pm)

52.2

14.1

30.0

16.8

14.1

14.4

21.3

18.9

30.2

18.8

32.0

Co

(ppm

)21

.711

010

610

413

712

023

817

2C

r(p

pm)

180

294

80.6

77.7

3.50

5.02

5.01

4.46

4.92

Hf

(ppm

)5.

103.

783.

633.

534.

633.

9110

.29.

209.

1311

.212

.29.

7514

.65.

74b

8.70

Nb

(ppm

)5.

545.

7014

935

.652

.444

.945

.465

.353

.690

.941

.314

511

1N

i(p

pm)

138

84.0

138

82.5

89.1

117

122

135

74.0

84.5

104

Rb

(ppm

)15

.415

.314

.920

.517

.922

.029

.2Sc

(ppm

)16

.321

.915

.011

.4


Tab

le1

(Con

tinu

ed)

BZ

1H

1K

C1

H2

BZ

2H

3A

C1

WK

1W

K1f

rag

WK

2A

r1(N

=11

)(N

=5)

(N=

9)(N

=47

)(N

=4)

(N=

17)

(N=

6)(N

=15

6)(N

=5)

(N=

8)(N

=12

9)

147

247

250

223

275

326

Sr(p

pm)

495

437

180

111

108

0.80

0.68

0.66

0.76

0.82

0.68

0.74

Ta

(ppm

)0.

640.

42b

0.41

0.40

8.72

8.51

10.8

8.93

8.54

9.27

10.9

12.5

7.59

4.60

Th

(ppm

)7.

592.

561.

821.

982.

001.

882.

761.

641.

221.

91U

(ppm

)1.

491.

3219

694

.912

012

812

816

414

322

212

716

014

2V

(ppm

)26

.615

.323

.223

.722

.426

.523

.130

.217

.220

.924

.4Y

(ppm

)10

583

.783

.511

510

915

415

3Z

n(p

pm)b

108

128

88.1

81.6

144

Zr

(ppm

)21

716

220

820

319

320

215

516

119

015

0

0.08

20.

067

0.04

40.

061

0.05

50.

053

0.05

9A

g(p

pm)b

0.16

0.06

80.

052

0.04

7b

0.86

12.9

4.52

0.42

0.52

0.63

1.10

1.01

0.80

1.28

As

(ppm

)b6.

530.

520.

340.

310.

400.

79A

upp

bb1.

000.

781.

050.

470.

730.

420.

310.

100.

034

0.12

0.15

0.21

0.08

00.

079

Bi

(ppm

)0.

200.

220.

072

84.0

23.8

41.9

25.6

25.1

31.7

37.3

88.8

42.7

61.3

42.7

Cu

(ppm

)b

1.71

3.88

(0.7

9)(0

.26)

(0.3

1)(0

.39)

(0.2

7)1.

0P

dpp

b(0

.2)

(0.2

)1.

800.

031

0.02

80.

021

0.02

10.

046

0.09

50.

041

0.02

90.

035

Sb(p

pm)

0.03

70.

036

0.56

0.05

30.

220.

130.

130.

150.

200.

450.

075

0.31

0.15

Se(p

pm)

28.2

42.2

25.0

12.7

13.5

16.7

22.6

49.6

9.46

Te

ppb

9.56

47.4

aW

K1f

rag

isth

eav

erag

eof

mic

agn

eiss

frag

men

tsin

mig

mat

ites

.V

alue

sin

pare

nthe

ses

incl

ude

man

yde

term

inat

ions

belo

wth

ede

tect

ion

limit

(Cgra

f0.

05%

and

Pd

0.2

ppm

)an

dsh

owei

ther

the

dete

ctio

nlim

itva

lue

orav

erag

esex

clud

ing

valu

esbe

low

dete

ctio

nlim

its.

bO

neto

two

anom

alou

san

alys

esha

vebe

enex

clud

edfr

omso

me

grou

pav

erag

es.


elements indicate a more mafic source comparedto local Archaean bedrock at the present erosionlevel (Figs. 4 and 5, and Table 1). The Ar2samples show variable REE and have higherCaO, Na2O and lower K2O, Cr and Rb comparedto Ar1 (see Fig. 4 for K2O and Cr). The lowerCIA indicates less weathering relative to Ar1 andlow Th/Sc (0.09–0.17) favours a dominant maficsource.

4.2. Cratonic co6er

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

sedimentary rocks in the Höytiäinen basin areclassified into high- and low-Cr groups H1 andH3, respectively (Fig. 4, Table 1). A distinct litho-logical unit (Huhma, 1975) of high-Cr rocks isclassified as group H2 and a suspect group oflow-Cr rocks, possibly related to the WesternKaleva (Kohonen, 1995), is classified as groupH4. Samples outside the Höytiäinen 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 samplesinclude quartz-rich greywackes and more typicallypelites showing thin layering from 1–3 mm to1–2 cm with thin psammitic interlayers occurringlocally. The variation in element abundances in-side the H1 group is mainly explained by quartzdilution (Fig. 4). There is evidence of weatheringin 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 andKutsu-type granites normalized to Archaean Crust (AC1 in Table 1). The Karelian craton (KC1) and WK1 averages are from Table1 and the averages for Jatuli-type mafics (N=21) and Kutsu granites (N=8) are from Lahtinen (unpublished data).


Tab

le2

Ave

rage

chem

ical

com

posi

tion

ofse

lect

edse

dim

enta

rygr

oups

(non

-mig

mat

ized

,ex

cept

CF

3av

erag

ein

clud

ing

also

mic

agn

eiss

frag

men

tsin

mig

mat

ites

)a

CF

3R

H2m

igC

F2

RH

3/lC

rC

F1

RH

4/hC

rC

F3m

igR

H1

RH

2(N

=4)

(N=

4)(N

=14

)(N

=5)

(N=

6)(N

=6)

(N=

14)

(N=

7)(N

=12

)

72.3

769

.70

63.7

163

.71

70.7

5Si

O2

(%)

67.9

461

.99

64.9

576

.50

0.53

0.69

0.52

0.60

0.73

0.74

0.73

0.79

TiO

2(%

)0.

5813

.25

14.2

012

.78

13.4

715

.24

15.7

116

.13

Al 2

O3

(%)

17.8

411

.75

3.78

4.56

5.97

6.33

5.25

7.18

4.51

FeO

(%)

3.78

5.77

0.07

0.03

0.06

0.07

0.08

0.09

0.05

0.06

0.08

MnO

(%)

2.78

1.38

1.56

2.21

3.04

2.92

2.30

3.17

2.05

MgO

(%)

1.91

2.13

1.88

1.94

2.44

CaO

(%)

2.04

0.89

1.23

0.52

2.91

2.54

2.97

2.94

2.59

2.57

2.18

Na 2

O(%

)1.

671.

672.

582.

593.

413.

282.

552.

71K

2O

(%)

2.59

3.81

3.97

0.17

0.10

0.15

0.16

0.15

0.12

0.13

0.12

0.15

P2O

5(%

)(0

.05)

(0.0

5)(0

.08)

0.15

Cg

raf.

(%)

(0.0

5)(0

.25)

(0.0

9)(0

.05)

(0.0

5)0.

023

0.03

30.

110.

082

0.23

0.10

S(%

)0.

051

0.04

30.

410.

064

0.04

880.

052

0.06

40.

075

0.07

60.

085

0.12

0.05

2F

(%)

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

.444

.194

.374

.869

.174

.3C

e(p

pm)

63.2

86.9

88.7

75.8

62.5

10.5

8.69

8.16

8.79

7.34

Pr

(ppm

)8.

6410

.410

.17.

4237

.932

.0N

d(p

pm)

30.3

27.5

32.1

38.0

38.4

32.3

27.3

6.53

5.87

5.61

5.95

5.23

7.29

5.63

Sm(p

pm)

4.98

6.85

1.12

0.94

1.21

1.17

1.10

1.10

1.33

1.19

1.06

Eu

(ppm

)5.

595.

285.

155.

46G

d(p

pm)

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.

710.

960.

893.

443.

303.

993.

863.

974.

334.

805.

053.

80D

y(p

pm)

0.67

0.64

0.81

0.77

0.78

0.89

0.94

0.95

0.72

Ho

(ppm

)2.

322.

272.

192.

641.

912.

642.

14E

r(p

pm)

1.81

2.83

0.29

0.28

0.32

0.31

0.30

0.40

0.41

0.41

0.31

Tm

(ppm

)1.

851.

782.

122.

092.

222.

602.

842.

522.

03Y

b(p

pm)

0.33

0.34

0.33

0.39

0.27

0.29

Lu

(ppm

)0.

300.

390.

40

408

534

640

618

630

595

771

639

Ba

(ppm

)62

839

.542

.051

.179

.746

.751

.7C

l(p

pm)

75.0

59.5

91.8

16.3

8.93

9.86

14.2

17.9

19.6

14.1

19.2

13.2

Co

(ppm

)15

810

781

.292

.911

912

611

614

997

.3C

r(p

pm)

6.62

5.40

4.45

4.56

4.50

Hf

(ppm

)5.

464.

415.

015.

109.

599.

4811

.412

.0N

b(p

pm)

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

.860

.210

811

510

410

714

414

515

520

810

1R

b(p

pm)

16.2

10.4

11.6

15.2

18.2

19.4

17.9

20.4

13.4

Sc(p

pm)

301

294

242

238

282

Sr(p

pm)

240

141

181

96.0


Tab

le2

(Con

tinu

ed)

CF

3mig

RH

3/lC

rR

H1

RH

4/hC

rC

F1

CF

2R

H2m

igC

F3

RH

2(N

=6)

(N=

5)(N

=6)

(N=

14)

(N=

7)(N

=12

)(N

=4)

(N=

14)

(N=

4)

0.67

0.66

0.84

0.77

0.83

0.82

0.93

1.03

0.68

Ta

(ppm

)10

.48.

1215

.211

.210

.311

.1T

h(p

pm)

9.6

12.9

13.8

2.39

1.98

2.74

2.56

2.56

2.21

3.18

3.32

U(p

pm)

2.29

87.5

112

144

146

143

160

107

V(p

pm)

88.9

153

22.2

21.3

24.0

23.1

23.9

27.8

30.1

33.1

23.2

Y(p

pm)

100

94.3

64.0

78.1

101

116

157

166

69.9

Zn

(ppm

)26

721

817

818

118

122

7Z

r(p

pm)

225

175

203

0.06

70.

061

0.03

90.

044

0.06

30.

071

Ag

(ppm

)0.

088

0.09

60.

059

0.58

1.43

0.92

0.60

2.11

1.38

As

(ppm

)b1.

030.

561.

121.

380.

670.

460.

820.

670.

381.

160.

840.

88A

upp

bb

0.17

0.14

0.05

60.

120.

180.

045

0.24

0.23

0.12

Bi

(ppm

)11

.316

.633

.253

.227

.0C

u(p

pm)

19.0

24.0

32.4

30.8

(0.2

)(0

.29)

Pd

ppbb

0.48

(0.2

5)0.

85(0

.82)

1.02

(0.2

8)(0

.25)

0.04

20.

041

0.03

20.

037

0.08

30.

027

0.05

9Sb

(ppm

)0.

045

0.08

90.

100.

180.

053

0.10

0.13

0.18

0.56

0.12

0.05

3Se

(ppm

)17

.410

.66.

614

.823

.528

.2T

epp

bb36

.08.

721

.4

aT

heR

H2m

igan

dB

B4m

igar

eth

eav

erag

esof

mig

mat

ites

,re

spec

tive

ly.

Gro

upR

H3

have

been

divi

ded

into

low

-Cr

(RH

3/lC

r)an

dhi

gh-C

r(R

H/h

Cr)

popu

lati

ons.

Val

ues

inpa

rent

hese

sin

clud

em

any

dete

rmin

atio

nsbe

low

the

dete

ctio

nlim

it(C

gra

f0.

05%

and

Pd

0.2

ppm

)an

dsh

owei

ther

the

dete

ctio

nlim

itva

lue

orav

erag

esca

lcul

ated

excl

udin

gva

lues

belo

wde

tect

ion

limit

s.b

One

totw

oan

omal

ous

anal

yses

have

been

excl

uded

from

som

egr

oup

aver

ages

.


Fig. 4. Harker-type Cr, K2O, MgO and CIA (Nesbitt and Young, 1982) variation diagrams for Archaean, autochthonous andallochthonous sedimentary rocks in the study area. Ar1 and Ar2-Archaean, Jqzt–Jatuli-type quartzites, H1–H2-autochthonoushigh-Cr, H3-autochthonous low-Cr, H4- a low-Cr suspect group of Höytiäinen area. WK1–WK2 main field-allochthonous WesternKaleva. AC1 is the average of Archaean crust (Table 1).

large mafic component indicated by high contentsof HREE, MgO and Pd. The H2 group has manycompositional similarities with H1 but the H2average shows higher levels of most elements (e.g.MgO) and lower SiO2 (Fig. 4 and Table 1). SomeH3 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-Archaeangranite (Kutsu; see Fig. 3). The H4 is a heteroge-neous group that deviates to some extent from theWK1 main group in having higher K2O and lowerCr (Fig. 4).

The allochthonous Western Kaleva (WK) sedi-mentary rocks have been divided into WK psam-mites and SiO2-poor pelitic rocks (WK2). TheWK1 psammites (Table 1) form a geochemicallyhomogeneous group (Fig. 4) and most of thevariation can be explained by grain size variation.The more pelitic nature of WK2 is seen in enrich-ment of elements (e.g. Al2O3, MgO, FeO, K2O)that characterize clay minerals (Table 1) but theWK2 also seems to be enriched in a mafic sourceas seen in higher Sc and Cr relative to Th. TheWK1 migmatites are mainly psammitic fragmentsfloating in tonalitic (often trondjhemitic) veined


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

4.3. Boundary zone and S6ecofennian sedimentaryrocks

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

The southern Svecofennian sedimentary rocksin the Rantasalmi–Haukivuori area have beenclassified into three groups (RH1–RH3). Thenon-migmatitic RH1 rocks are quartz-richgreywackes and the well-preserved RH2 rocks aremore pelitic in character. Both RH1 and RH2show rather similar patterns in Fig. 6 where thestrong effect of weathering is seen in negativepeaks of Ba, Sr, CaO, MnO and P2O5, and highCIA values (Table 2). The depletion of HREE, Sc,V, TiO2 and enrichment of K2O, Rb, Th andespecially 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 arechondrite-normalized values and Eu/Eu* has been calculated using Eu*= (SmN+GdN)/2. The Archaean average has beencalculated 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-southern Svecofennian. CF1–CF3-central Svecofennian.


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 andRH3/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 relativeenrichment of CaO, Ba, Nb, V and Sc and lowCr/Sc ratio favouring a new additional maficcomponent in the RH2. The lower CIA values(Table 2), which are normally higher in morepelitic rocks, indicate that this additional compo-nent was less weathered. Compared to the RH1and RH2 rocks the RH3 samples show lower CIAand higher CaO and Na2O with strong variationin the amount of mafic component (Fig. 6 andTable 2).

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

for migmatites but the slightly lower REE andespecially deep negative Eu anomaly in some sam-ples ask for a loss of felsic component. The slightdepletion in Ba, K2O and K/Rb can be related toa loss of a K-feldspar component and the enrich-ment of ferromagnesian components to the in-creased amount of restite. So it seems that thesemigmatites are mainly in situ migmatites thatshow a complex mixture of restite and a meltfraction 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 highTh/Sc (]1) psammites, CF2 lower Th/Sc (51)psammites and CF3 silt-pelite rocks. The non-migmatitic CF1 samples show LREE enrichmentcompared to the Western Kaleva psammites (Fig.


5). The depletion of elements characteristic ofmafic components and the relative enrichment ofLREE, Sr, Th, U and Zr point to a larger felsiccomponent relative to the WK psammites. Thechemical composition of the CF2 group shows anenrichment of mafic components relative to CF1.CF3 is a heterogeneous group characterized bymigmatites and thus the average (Table 2) in-cludes also mica gneiss fragments in migmatites.Mineralogically the CF3 rocks differ from theCF1–CF2 in the ubiquitous occurrence of garnet.The more clay-rich nature of CF3 is seen in lowerSiO2 and higher MgO and K2O (Table 2). TheCF3 migmatites form an inhomogeneous groupranging from samples with HREE enrichment tosamples with HREE depletion and Eu enrichmentat low total REE abundances compared with lessmigmatitic CF3 samples. This is interpreted asdifferent amounts of restite and leucosome insampled outcrops.

5. Discussion

5.1. Palaeoweathering

Palaeoweathering in the source area is one ofthe most important processes affecting the com-position of sedimentary rocks. Sedimentary rockssensu stricto are composed merely of weatheringproducts and reflect the composition of weather-ing profiles, rather than bedrock (e.g. Nesbitt etal., 1996). Based on CIA values (Nesbitt andYoung, 1982) the source rocks affected the mostby 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 mostlyshow CIA values lower than 55 (Fig. 4). Most ofthe central Svecofennian psammitic rocks alsohave low CIA values (B55) with an increase upto (60–67) in CF3 pelitic rocks. This generalincrease in CIA with silica-poorer and morepelitic nature is a common feature and readilyexplained 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 inquestion. Overestimation of Ca in carbonates canlead 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 onlyproblematic in limited cases but is especially cru-cial for quartz-rich samples. The other problem isrelated to the loss of CO2 and incorporation ofliberated Ca in recrystallizing minerals (e.g. epi-dote and plagioclase) during metamorphism (cf.Lahtinen, 1996) a situation proposed for somesamples in the Höytiäinen area (Fig. 7).

The prevailing climatic conditions of the sourceareas during sediment formation are difficult toestimate especially if we consider the recyclednature of many sediments, possibly having olderweathered components. The situation can be thuscomplex including mixing of a strongly weatheredcomponent (older sediments or deeply weatheredpalaeosol) with immature crust components be-fore deposition, forming a sedimentary rockshowing moderate CIA values. Also the degree ofweathering is related to the rate of erosion, whichis high in tectonically active areas and thus in-hibiting extensive weathering even in high rainfalltropical conditions. The extent of weathering isdetermined primarily by the amount of rainfall(acids) on the weathering profile (Singer, 1980)where as the climatic effect on weathering trendsis probably insignificant (Nesbitt and Young,1989).

The REE, Th and HFSE (especially Sc) areconsidered least susceptible to fractionation byexogene processes including weathering (Taylorand 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 whenwhole weathering profiles were considered. Deple-tion of Sc has been postulated during weatheringunder low-O2 atmosphere (Maynard et al., 1995).

The Palaeoproterozoic autochthonous unitsabove Archaean basement in the study area (Ko-honen and Marmo, 1992 and references therein)start with the Ilvesvaara Formation overlain bythe glaciogenic Urkkavaara Formation followed


by Hokkalampi Palaeosol. Sturt et al. (1994) con-cluded that widespread 2.35 Ga regolith (includingthe Ilvesvaara Formation) occurred on theFennoscandian shield and was related to an aridor semi-arid palaeoenvironment. Although thismight be the case for the Ilvesvaara Formation,the occurrence of the up to 80 m deep HokkalampiPalaeosol (not mentioned by Sturt et al., 1994)with a minimum age of 2.2 Ga records intensechemical weathering under a tropical warm andhumid climate (Marmo, 1992). The drift ofFennoscandian from 30°S at 2435 Ma to about30°N at 2100 Ma (Pesonen et al., 2000) shows thatFennoscandian crossed the equator during thistime favouring the interpretation of Marmo(1992). It has been suggested that the HokkalampiPalaeosol and derived formations covered largeareas of the stable Karelian craton (Kohonen andMarmo, 1992; Marmo, 1992) where they formedthe bulk of detritus for the Palaeoproterozoic riftbasins.

The chemical and mineralogical data of theHokkalampi Palaeosol indicate a typical weather-ing sequence (cf. Nesbitt and Young, 1989; Condieet al., 1995) with an initial decrease in the amountof plagioclase followed by loss of K-feldspar andbiotite seen as an increase in CIA values fromabout 60–70 (lowermost) to the highest values of80–90 in the upper zone (Marmo, 1992). Potas-sium metasomatism of kaolinite to illite inpalaeosol results in lowering of CIA values (Fedoet al., 1995). This possibility has been studiedusing an A–CN–K compositional space (Fig. 7)for the data of the Hokkalampi Palaeosol formedupon K-feldspar rich granitoid and sandstone.There is a slight amount of added potassium inlower palaeosol zones probably due to percolationof solutions from the leached uppermost potas-sium-depleted zone during weathering (Marmo,1992). However, if the whole mass balance of theweathering 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 andautochthonous groups of this study. (A) Data for Hokkalampi palaeosol formed upon a K-feldspar-rich granitoid (granitoid zones2–3) 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 inFig. 7A. Dashed line encloses possible source end members for autochthonous sedimentary rocks. (C) Data for Jatuli-type quartzitesand autochthonous groups H1–H3. Note the shift of some samples towards the sodium-rich (A–K)–N-line indicating thatalbitization has possibly affected these samples. Horizontal arrows for some samples indicate the amount of Ca input due to theinferred occurrence of carbonates followed by CO2 loss. Averages of palaeosol zones 1–3 in Fig. 7B and C are calculated usingmixtures of sandstone zones (50%) and granite zones (50%). J and K are calculated averages of Jatuli-type mafics and Kutsu-typegranites, respectively.


The autochthonous Höytiäinen H1–H3 groupsshow characteristic depletion of CaO, Na2O,MnO, P2O5, Sr and Ba, and low K/Rb, which aretentatively proposed to have an ultimate source inthe chemically weathered palaeosol. The southernSvecofennian RH1–RH2 groups also show deple-tion of elements normally lost during weathering(Fig. 6) but the CIA values of other groups aremoderately low (B60) and no clear weatheringtrends 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 inpelites relative to associated sands (e.g. Korsch etal., 1993). The situation of pure quartz dilution isthe ultimate case and most easily interpreted as adecrease in all other elements and an increase inSiO2. The situation is more complex when acces-sory minerals (zircon, monazite, apatite, spheneand allanite), ferromagnesian minerals, feldsparsand lithic fragments are also sorted. The Th/Scratio remains nearly constant in some cases butoften muds can have significantly lower Th/Scratios indicating a preferential incorporation ofmafic volcanic material in the finer fractions (e.g.McLennan et al., 1990). Considering a simpletwo-component mixture of mature weathered ma-terial (quartz+clays) and immature rock debris(separate minerals+ lithic fragments) the result ispsammites enriched in immature rocks debrisshowing complex sorting patterns and pelites en-riched in mature weathered material. This prefer-ential sorting can lead to REE fractionationmaking interpretation of Sm–Nd isotope system-atics difficult (Zhao et al., 1992) but this is mainlyeffective when considering sedimentary materialfrom unweathered coarse-grained granitoids with,e.g., allanite hosting LREE and Th.

The wide range of SiO2 (Fig. 4) the HöytiäinenH1–H3 groups exhibit is clearly an effect ofsorting (cf. Kohonen, 1995) dominated by quartzdilution seen as abundant quartz clasts. Sortingenhanced enrichment of mafic component wasnoticed, e.g. in Western Kaleva and southern

Svecofennian pelites over psammites. The varia-tion of Zr (normally 160–350 ppm) found inWestern Kaleva psammites indicate zircon sortingbut there is no correlation between Zr and HREEor U showing that the zircon control on theseelements is minor. The effect of hydraulic sortingis readily observed in the studied samples but inmany cases it also sorts different source compo-nents into different grain size classes. This is adisadvantage when using only shales (on averagemore 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 theinterpretation of the depositional environment ofancient sediments using black shales/schists.These include pyrite formation, S/C ratios, degreeof pyritization (Berner, 1984; Berner andRaiswell, 1984; Raiswell and Berner, 1986) andenrichment of U and V (e.g. Jones and Manning,1994; Breit and Wanty, 1991). The averagepresent S/C ratio of normal marine sediments is0.36 (0.23–0.77) but age dependent variation oc-curs and, for example, early Palaeozoic marinesediments show significantly higher S/C ratio ofabout 2 (Berner and Raiswell, 1984; Raiswell andBerner, 1986). In fresh or low-salinity brackishwater low sulfate level is the limiting factor forpyrite formation and sediments show low S/Cratios without any inter-element correlation(Berner and Raiswell, 1984). According toThompson and Naldrett (1984) mantle-derivedmagmatic S/Se ratios are generally lower (B10 000) than in sedimentary sulphides (\10 000),which can be used to discriminate hydrothermalinfluxes of sulphur.

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


Fig. 8. Plot of Cgraf. vs S for autochthonous (H1–H3) andallochthonous (WK1–WK2) sedimentary rocks in this studydivided 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 arethermal history and time, where rapid burial com-pacts sediments quickly (dewatering) and blanketsany thermal changes (Lee and Klein, 1986). Thuslong-lived basins, like the Höytiäinen basin (Ko-honen, 1995), should show more pronounced ef-fects of diagenesis compared to allochthonousWestern Kaleva-type rocks that were deposited asmassive units in an active tectonic setting. Thevery limited element variation in the WK rocksfavours this and although small-scale diageneticchanges within WK samples are possible, a large-scale redistribution of elements is not evident.Similar arguments hold for most of the centralSvecofennian rocks but, for example, the deposi-tional environment and the elapsed time beforedewatering and metamorphism of the Archaeanand southern Svecofennian mature rocks are un-known. Diagenetic reactions may include Na-, K-,Mg- and Fe-metasomatism (e.g. Nesbitt andYoung, 1989) while REE redistribution and frac-tionation have also been proposed (Awwiller andMack, 1991; Milodowski and Zalasiewicz, 1991;Ohr et al., 1991). There is not however consensusabout how common the redistribution of REEduring diagenesis is (cf. Hemming et al., 1995)and one critical question is that are the proposeddiagenetic reactions open or closed systems atsample scale.

Redistribution of alkalies during diagenesis hasbeen proposed for the Höytiäinen area rocks (Ko-honen, 1994) and to evaluate this possibility, thedata are plotted in the A–CN–K and (A–K)–C–N compositional spaces (Fig. 7; see also Fig. 4 forK2O). The data show scatter and there are severalfactors that may have been responsible for theobserved trends: (1) Sedimentary rocks have dif-ferent source components with different K2O/Na2O ratios (see differences in MgO contents andTh/Sc and Th/Cr ratios; Figs. 4 and 9). Theproblem lies also in the thinly layered nature ofpelites where chlorite-rich and biotite-rich layerswere noticed, possibly indicating that differentlayers were derived from different sources in somecases. (2) During grain-size sorting K-rich phases(illite and biotite-vermiculite) are enriched inpelites (K-feldspar is rare in these rocks) andplagioclase in sands forming a trend similar to

indicate anoxic conditions during deposition andif the S/C ratio of 3.5 is higher than found in thePalaeoproterozoic marine sediments during depo-sition, it could point out to euxinic environment.

The Western Kaleva samples differ from theHöytiäinen basin examples in that they do notshow any clear correlation between S and C. Thegraphite-enriched (\0.5% C) psammites havelow S/C ratios (B0.15) and S/Se ratios mostlyB10 000. Apart the graphite variation (0–1.6%C) there is no enrichment of studied elements. Theoccurrence of graphite-bearing thick psammitesdoes not favour a direct hemipelagic origin andindicates mixing of carbonaceous matter intomass flows before deposition. The low S/C ratioscould point to fresh water or brackish waterenvironments, or to short intervals between depo-sition of mass flows preventing significant bacte-rial sulfate reduction. The lack of U and Venrichment indicates an oxygenated environmentwhile 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 thediagenetic history be based on geochemistry. Ma-


that observed in the A–CN–K compositionalspace. (3) Albitization of K-feldspar in the sand-size fraction with immediate uptake of liberated Kby kaolinite, chlorite, montmorillonite and/orsmectite in the clay-rich fraction as proposed byKohonen (1994). Based on Fig. 7C albite metaso-matism has occurred to some degree in somesamples favouring Kohonen’s (Kohonen, 1994)interpretation. (4) Regional-scale potassic andsodic metasomatism affecting shales and silt-sand-size particles, respectively, has been proposed forthe Palaeoproterozoic Serpent Formation (Fedoet al., 1997). The Serpent shales show ultimatepotassium variation from 3.3 to 11.2% whereasthe H1–H3 pelites show only variation from 3 to5% (Fig. 4) where the variation is mainly due to

the factors 1–3, as discussed above. Thus, theproblem in depicting the amount of diageneticredistribution in the Höytiäinen area rocks is thatthey show complicated mixing of source compo-nents associated with sorting and thus distinguish-ing purely diagenetic effects is difficult. Althoughnot conclusive it seems that small-scale redistribu-tion of elements has occurred during diagenesis inthe Höytiäinen area but no externally derivedregional-scale metasomatism, at least for potas-sium, is observed.

Prograde metamorphic effects on REEs, exceptin areas of partial melting, are minor (Taylor etal., 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 calculatedfrom the average AC1 in Table 1 and Jatuli-type mafics from the average (N=21) in Lahtinen (unpublished data). See Fig. 5.


The tonalite migmatites (veined gneisses, schollenmigmatites and diatexites) in the study area showvariable compositions due to differences in therelative amounts of restite and leucosome in sam-pled outcrops and those that represent totallymelted ‘in situ’ variants. A depletion of Bi is themain common feature and although migmatiteswith high proportions of restite component occurthere is no area showing large-scale depletion ofelements. In many cases the veined gneisses havemostly 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) withvariable amounts of restite and granite leucosomecomponents. This difference in the leucosomecomposition (tonalite–granite) has been at-tributed to the aluminium excess in the sourcerocks of migmatites having granite leucosomes(Korsman et al., 1999). This interpretation is fa-voured by the typical CIA values of 60–70 in theRH1–RH2 rocks compared to the typical CIAvalues below 60 in the source rocks of tonalitemigmatites. On the other hand water-rich condi-tions during tonalite migmatization favour theformation of plagioclase-enriched melts and wa-ter-rich conditions has been considered as themain cause for the formation of tonalitemigmatites (Lahtinen, 1996).

5.5. Main source components

The proposed main source components of sedi-mentary rocks of the Archaean craton and itscover, and Svecofennian domain are mainly basedon the geochemical differences but Sm–Nd resultsby Huhma (1986, 1987), O’Brien et al. (1993) arealso adopted. There are only a few detrital zirconage determinations from the FennoscandianShield (Huhma et al., 1991; Claesson et al., 1993)and thus the conclusions presented below are tosome extent tentative but serve as a workingmodel for future work.

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

shearing precludes further source componentinterpretation.

5.5.1. Archaean sedimentary rocksThe Archaean sedimentary rocks show very low

Th/Cr ratios, which discriminate them from otherrocks in this study (Fig. 9). O’Brien et al. (1993)concluded that greywackes in the eastern part ofthe study area (Ar2-type) with TDM ages from2.83 to 2.99 Ga normally show a local source. TheAr1 samples show a more homogenized sourceand higher degree of weathering of the sourcearea with higher MgO, Cr, K2O and SiO2. OneArchaean sediment has a TDM of 3.24 Ga(Huhma, 1987) favouring also the existence of anolder component (cf. Sorjonen-Ward, 1993). Twomain ages of source components with variableamounts 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 age. At least three different sourcerock types are indicated: komatiites (highMgO, Cr, Ni, Cr/Sc), tholeiite (high TiO2 andNb/Th) and felsic component (SiO2, K2O andRb). Intermediate to strong weathering in thesource area and thorough mixing has occurredbefore deposition. Possible sources aregreenstone+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 co6erLocal Archaean craton sources with contribu-

tions from Jatuli-type mafic volcanics and dykeshas been a common source model for the Höyti-äinen basin sedimentary rocks (Huhma, 1987;Ward, 1987; Kohonen, 1995). The results of thisstudy favour this general statement but the com-position of the H1–H2 groups is not explained bysimple mixing of the presently exposed erosionlevel of the Archaean crust and Jatuli-type mafics(Figs. 4, 5 and 9) because an additional Cr-richsource is needed. The simplest explanation ishigher amounts of Archaean sedimentary rocks(Cr-rich) in the average source area for the H1–H2 group samples. Some sedimentary rocks havehigh proportions of local Archaean cratonic


source dominated by felsic granitoids indicatedby high Th/Sc and Th/Cr ratios. The sedimen-tary rocks show TDM variation from 2.28 to 2.70Ga, which partly overlap with the WesternKaleva TDM variation of 2.29–2.40 Ga (Huhma,1986, 1987). The Sm–Nd data for the Höyti-äinen 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 uponArchaean crust and glaciogenic deposits. En-richment of Archaean sedimentary rocks (seethe 3.0–3.2 Ga component above).

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

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

The detrital zircon U–Pb isotopic data fortwo samples (Claesson et al., 1993) give age con-straints for granitoid components in the al-lochthonous Western Kaleva mica schists. Thesamples have 30–40% late Archaean zircons(2.5–2.8 Ga) and only a few crystals in the agerange between 2.6 and 2.1 Ga, which can also bemixture ages (Claesson et al., 1993). Both sam-ples have 50–60% zircons from a 2.0 to 1.92 Gaage group with a maximum deposition age ofabout 1.92–1.94 Ga. TDM ages of 2.3–2.4 Gabased on Sm–Nd data of Western Kaleva micaschists (Huhma, 1987) are in agreement with thedetrital zircon data. The Archaean componenthas been dominantly late-Archaean in age andwe can use the normalization to the AC1 of thisstudy to interpret the nature of the 1.92–2.0 Gacomponent (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 arctholeiitic origin for the mafic component. Thehigh Zr relative to K2O, Rb and REE favour alow-K felsic source also characterized by moder-ate to low La/Yb ratios. Two main componentsare proposed for the Western Kaleva sediments:

1. Archaean crust dominated by late-Archaeangranitoids mixed with a small contributionfrom Jatuli-type dykes. A small amount ofrecycled weathered component is possible.

2. 2.0–1.92 Ga bimodal source of low-K felsicrocks and tholeiitic volcanics derived fromprimitive island arc.

5.5.3. S6ecofennianThe central Svecofennian psammites show

large compositional variation (Figs. 5 and 9) in-dicating either different provenance areas orchanges 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 rathershort transport distances from a rapidly risingorogenic domain. The psammites of this study(CF1–CF2) and the basement-related sedimen-tary psammites (SG3–SG4) of Lahtinen (1996)from the Tampere–Hämeenlinna area have geo-chemical similarities as seen in Th/Sc ratios of2–0.7 and 1.5–0.5, respectively. The TDM of 2.2Ga (Huhma, 1987) from one sample is slightlyyounger 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 ispossible to use the WK1 as a normalizing valueto infer the nature of this additional component.The CF1 psammites are enriched in felsic com-ponent and thus the differences to the WK1should approximate the felsic composition. HighLREE, LaN/LuN and negative Eu anomalies withhigh Th and low K/Rb and Nb favour a matureintracrustal origin. The Th variation (13–19ppm) and Th/Ta ratios (typically ]15) in theCF1 are distinct from Th contents (B9 ppm,mostly B4 ppm) and Th/Ta ratios (59) foundin the 1.93–1.91 Ga primitive island arc felsicrocks (see Figure 26 in Lahtinen, 1994) excludingthem as a dominant felsic component in theCF1. The CF1 and CF2 groups are gradationalto each other and the low Cr/Sc ratio, low Nband only slight TiO2 enrichment relative to MgOfavour a mature island arc origin for the addedmafic–intermediate component. The proposedmain source components for the central Sve-cofennian sedimentary rocks are as follows:


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 WesternKaleva and central Svecofennian psammitespointing to different origins. High Zn/Co (about10) is a characteristic feature of RH1 and variablebut high Zn/Co also characterizes the RH2 sam-ples. The Zn/Co ratio is sensitive to changesduring weathering and sulphide precipitation butthere does not seem to be any relationship be-tween the existence of sulphides and Zn/Co indi-cating instead either source difference or aweathering effect. Similar Zn/Co enrichment wasnot noted in high CIA rocks from the Höytiäinenarea favouring a source origin for the high Zn/Co.Elevated Zn and low Co are characteristic fea-tures of alkaline-affinity intermediate–felsicwithin-plate-type granitoids (Lahtinen, unpub-lished data) and this type of magmatism in thesource area is one possible explanation for thehigh Zn/Co ratios. High Cr and Cr/Sc ratios inthe RH1 are interpreted to have their ultimatesources in an abundant komatiite or picriticcomponent.

The less mature greywackes (RH3) show mainlylow CIA values (B57) and thus resemble theWestern Kaleva psammites and psammites fromthe central Svecofennian. Although some sampleshave compositions close to those found in theWestern Kaleva psammites the RH3 rocks aretypically enriched in elements (LREE, Rb, Ba, Thand U) that characterize felsic source rocks. SomeRH3 rocks are enriched in elements that charac-terize mafic rocks especially seen in high Cr/Scratio (Fig. 6). This could indicate an Archaeankomatiite source but local Cr-rich lavas in theRantasalmi–Haukivuori area are more likely. Themain 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/Co2. Archaean crust with possibly high Cr/Sc (ko-

matiite component).

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

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

These tentative main source components char-acterize different groups differently; RH1 (19294), RH2 (1929394), RH3 (3+29491).The problem lies in depicting the origin of thehighly weathered component; Archaean versuspalaeoProterozoic.

5.6. Tectonic implications

Kohonen (1995) suggested that the syn-rift tur-bidites of the Höytiäinen rift basin (Ward, 1987)have a maximum depositional age of about 2.1Ga. The post-rift marine sediments probably in-cluded both passive margin and foredeep depositswhere the latter were deposited during foredeepmigration 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 volcanicsand dykes (mainly 2.2–2.06, and 1.96 Ga). Lahti-nen (1994) Palaeo-proterozoic Kohonen (1995)considered that the major rifting at 2.1–2.06 Gafinally lead to continental break-up (cf. Park etal., 1984; Gaál and Gorbatchev, 1987) and achange to a passive margin environment. A modelwith later continental break-up at 1.95 Ga hasalso been proposed (Peltonen et al., 1996). TheWestern Kaleva psammites have been describedas Svecofennian post-arc flysch (Park, 1985), amolasse from the Lapland granulite belt (Barbeyet al., 1984), pericontinental turbidites includingthe Kalevian as a whole (Laajoki, 1986), deep-wa-ter slope-rise greywackes related to uplift in Lap-land and the Kola Peninsula (Kontinen andSorjonen-Ward, 1991), a mixture of accretionprism sediments and derived foredeep sediments(Lahtinen, 1994) and axial foredeep deposits froma 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 aninferred 1.92–2.0 Ga low-K primitive island arccomponent and a non-weathered Archaean com-


ponent in the Western Kaleva psammites due toa rapidly rising orogene during oblique collisionstarting in the N (cf. Lahtinen, 1994). Aspresently understood these psammites have beendeposited both on Archaean basement andoceanic crust, and a foredeep origin associatedwith subsidence during initial collision is fa-voured and orogenic detritus either from thesame, oblique collision zone (mainly from theaccretionary prism) or a more distal orogenicdomain is proposed. One interesting feature is thepossible uptake of carbon-rich material, formedin 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 ofKohonen (1995) and deposition of organic matternear estuaries of large fresh water rivers.

The differences between the Western Kalevaand central Svecofennian sediments favour atleast partly different origins and different ages ofdeposition. The source for the central Svecofen-nian sediments included also mature island arcmaterial and the maximum deposition age wasabout 1.91 Ga for the main period of turbiditedeposition. Lahtinen (1994, 1996) has proposedthat ]1.91 Ga (possibly up to 1.95 Ga) riftingoccurred in the Tampere Schist Belt followed byincreasing subsidence during initial collision inthe NE and subsequent arc reversal. Abundanterosion from the mountain belt and depositioninto oblique hinterland basins that further devel-oped into a subduction related foredeep is theproposed model for the deposition of the mainsequences of turbidites in the central Svecofen-nian. The arc-related sediments are of localderivation and indicate deposition in small basinsbefore or during the 1.89 Ga collision (cf. Lahti-nen, 1994, 1996).

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

5.7. Crustal e6olution and Ba depletion in theArchaean–Proterozoic transition

Abrupt changes in the composition (REE, Th,Sc) of sedimentary rocks at the Archaean–Proterozoic transition has been proposed (Taylorand 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 to1.0–2.0, a decrease in the Sm/Nd ratio fromabout 0.21 to 0.19 and an increase in the Th/Scratio from about 0.5 to 1.0 (possibly only incontinental sediments). Secular changes in the Ar-chaean–Proterozoic transition, especially con-cerning the development of a Eu minimum, havebeen questioned and argued to be a consequenceof tectonic control resulting in biased sampling(e.g. Gibbs et al., 1986; Condie and Wronkiewicz,1990a; Gao and Wedepohl, 1995). Although theCr/Th ratio may not directly reflect the sourceratio, abrupt changes have been noticed in theArchaean–Proterozoic boundary reflecting thedecreasing amount of komatiites in the Protero-zoic (Taylor and McLennan, 1985; Condie andWronkiewicz, 1990b; Condie, 1993). No consen-sus exists about the importance of the Archaean–Proterozoic transition but komatiites andTTG-type rocks characterize the Archaean andtheir contribution to the sedimentary recordshould be distinguishable.

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

The Palaeoproterozoic sediments can be di-vided into cratonic sediments and sediments alsohaving Palaeoproterozoic crustal components.The cratonic sediments (H1–H3) have an inferredsource characterized by Archaean sources mixedwith variable amounts of Palaeoproterozoic mafic


material mainly from differentiated (e.g. low Cr/Sc) plateau volcanics and dykes. The otherPalaeoproterozoic sediments of this study alsoshow the contribution of intermediate to felsicigneous sources varying from low-K primitive is-land arc to mature active continental margintypes. An important feature is the almost totalabsence of 2.1–2.5 Ga mature crustal component(granitoids) in these sediments.

The data are somewhat scattered but thePalaeoproterozoic sediments having crustal com-ponents showing higher Th/Sc, Th/Cr, and lowerSm/Nd and Eu/Eu* relative to the Archaeanrocks (Figs. 5 and 9) as proposed in earlier studies(see references above) but the behaviour of GdN/YbN ratio is opposite to that proposed by McLen-nan and Taylor (1991). More data on Archaeansedimentary rocks from the Svecofennian shieldare evidently needed to see if the limited inputfrom TTG rocks (cf. Condie, 1993) is a commonfeature. The source variation seen in the centralSvecofennian sediments (e.g. Th/Sc 2–0.5 in theCF1–CF3 in Fig. 9) suggests exposure of differ-ent source components during erosion in thesource area. A slight grain-size induced preferen-tial separation of felsic source into the psammitesand mafic source into the pelites was also noted inthis study (cf. Lahtinen, 1996). Low Th/Cr ratioscharacterize the Archaean sedimentary rocks dueto the abundant komatiite component but lowerTh/Cr ratios can also be from a local picriticsource (e.g. some RH3 samples). These featuresreinforce the need for a large data set when usingsedimentary rocks in crustal evolution studies.

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

at 2.1–2.0 Ga based on detrital zircons from SäoFrancisco Shield (Machado et al., 1996) show thatthe age and nature of the Archaean–Proterozoictransition differ from shield to shield; a possibilityalso for the geochemical nature of associated sed-imentary rocks.

An elevated level of both Th and Sc relative tomodern deep sea turbidites in the basement re-lated sediments in the Tampere–Hämeenlinnaarea was noted by Lahtinen (1996) and a similarsituation characterizes the central Svecofenniansediments of this study (not shown). A sourceenriched in bimodal volcanics and depleted insedimentary quartz was proposed (Lahtinen,1996). The elements released during weatheringare also lost from the clastic portion but can bepartly redeposited in separate units within thesedimentary sequence, as for example Ca inmarine carbonates and U with organic matter.Many elements are also recycled back to themantle during subduction and form a characteris-tic fingerprint for subduction-related magmas andenriched mantle components (e.g. Hawkesworthet al., 1991; Weaver, 1991). Lahtinen (1996) pro-posed that the Ba deficiency in the basement-re-lated sedimentary rocks and the Ba enrichment inthe Svecofennian enriched mantle component (seealso Lahtinen and Huhma, 1997) are related to Barelease during weathering (9diagenesis) and lateruptake in pelagic sediments (possibly as barite)that are further subducted into the mantle.

The source variation and mobile nature of Baproduces scatter in Fig. 10 but Ba depletion isnoticed in most samples. A striking feature is thestrong relative Ba depletion in most high-Cr Höy-tiäinen rocks (H1–H2) and Jatulian quartzitesimplying that the Ba depletion is related to thechemically weathered component (cf. Maynard etal., 1995). Major loss of alkali and alkaline earthmetals relative to more immobile elements (REE,Th, Sc) occur at CIA values of about 80 in theHokkalampi Palaeosol when the breakdown ofillite dominates (Marmo, 1997; personal commu-nication). Similar Ba depletion relative to Rb wasalso noticed (not shown). Potassium enrichmentin 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 averagein the Table 1 and Jatuli-type mafics from the average (N=21) in Lahtinen (unpublished data). The Archaean trend is approximatedfrom the data in this study and the Tampere Schist Belt (TSB) volcanics trend is from Lahtinen (1996). See Fig. 5.

depletion. The Hokkalampi Palaeosol showsslight potassium enrichment in the lower zone buta large-scale external potassium addition seemunlikely. The main part of the Ba depletion isassumed to derive from the chemically weatheredpalaeosol, especially from the highly weatheredpart (CIA \80).

Ba depletion is less pronounced in other groups(Ar1, RH1–RH2 and CF3) having also elevatedCIA values over 60. If the interpretation of Höyti-äinen sedimentary rocks is correct it indicatesmixing of deeply weathered Archaean source ma-terial (CIA 70–90) with less weathered Archaeancrustal and Palaeoproterozoic mafic sources (CIAB50) to produce H1–H2 rocks with CIA valuesin the range of 55–70. In this case the lack ofcomparable Ba depletion in other pelitic rockswith elevated CIA can be attributed to the lack ofextremely 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 characteristicfeature of the sedimentary rocks of centralFennoscandian Shield. This indicates a highamount of Ba lost from the clastic record during2.3–1.9 Ga and further incorporated, at leastpartly, into both a subduction component and theenriched mantle. The Fennoscandian shield seemsto have exemplified a cratonic stage during 2.6–2.1 Ga characterized by deep chemical weatheringabout 2.35–2.2 Ga ago, high burial rates of or-

ganic carbon and highly 13C-enriched sedimentarycarbonates (e.g. Karhu, 1993) about 2.2–2.1 Gaago, and multiply rifting from about 2.2 to 1.95Ga. One critical question is the possible effect ofCO2-rich and low-O2 atmosphere in the formationof weathering profiles before the significant rise inatmospheric oxygen levels at about 2.0 Ga (e.g.Karhu, 1993). If the Ba depletion has been espe-cially characteristic for the chemical weatheringduring 2.35–2.2 Ga it could imply that duringand after this time period high amounts of Bahave recycled back to the mantle forming a ‘peak’in the formation of enriched mantle component.

6. Conclusions

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

Archaean sedimentary rocks can be divided totwo main groups those that have a dominantcomponent from a weathered 3.0–3.2 Ga green-stone+granite9TTG and those a local 2.7 Gasource, respectively.

Autochthonous sedimentary rocks have in-ferred deposition ages 2.2–1.9 Ga and vary fromrift to passive margin and foredeep deposits (Ko-honen, 1995). Chemically weathered palaeosol


(2.2–2.35 Ga), and sedimentary rocks derivedfrom it, and mafic dykes and plateau volcanics(mainly 2.2–2.1 Ga) are the major sources butlocal non-weathered Archaean sources dominatein places. Anoxic (euxinic?) conditions pre-vailed during deposition of the most sulphide-richrocks.

Allochthonous Western Kaleva sedimentaryrocks were deposited both on Archaean basementand oceanic crust (1.95 Ga ophiolites). The mostcharacteristic feature of Western Kaleva sandygreywackes is an extreme compositional homo-geneity. Source components are only slightlyweathered and comprise Archaean crust and 2.0–1.92 low-K bimodal rocks from a primitive islandarc. A foredeep origin associated with subsidenceduring initial collision is favoured and orogenicdetritus either from the same, oblique collisionzone (mainly from an accretionary prism) or amore distal orogenic domain is proposed (cf.Lahtinen, 1994; Kohonen, 1995).

The central Svecofennian sedimentary rockscan be divided into local arc-derived rocks (51.89 Ga) and older (]1.91 Ga) rocks for which amixture of Western Kaleva sources and 1.91–2.0Ga mature island arc/active continental marginsource is proposed. Rifting (]1.91 Ga) followedby increasing subsidence during initial collision inthe NE and subsequent arc reversal causing abun-dant erosion from the mountain belt and exposingdifferent source compositions as seen in the varia-tion of Th/Sc (2–0.5), and deposition into obliquehinterland basin further developing into subduc-tion related foredeep is the proposed model forthe deposition of the main part of the olderturbidites in the central Svecofennian.

The southern Svecofennian (Rantasalmi–Haukivuori area) mature greywackes resemblepassive margin sediments and sources dominatedby inferred alkaline-affinity complexes is pro-posed. Less mature rocks occur also with sourcescharacterized either by island arc/active continen-tal margin domain or local picritic rocks. It isimportant to note the absence of the southernSvecofennian-type mature greywackes from thecentral Svecofennian, which favours the existenceof a suture between these areas as proposed byLahtinen (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 wasdominated by input of a mainly mafic plateau-type volcanic contribution into the sedimentaryrecord. Palaeoproterozoic sediments havingcrustal components (52.1 Ga) show higher Th/Sc, Th/Cr, and lower Sm/Nd and Eu/Eu* relativeto the Archaean rocks as proposed in earlierstudies (Taylor and McLennan, 1985; McLennanand Taylor, 1991; McLennan and Hemming,1992) but local low Th/Cr ratios complicate thesituation. The behaviour of GdN/YbN ratio is alsoopposite 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 thesedimentary rocks of the central FennoscandianShield indicating large amounts of Ba lost fromthe clastic record during 2.3–1.9 Ga. Ba depletionseems to have been especially characteristic forchemical weathering during 2.35–2.2 Ga underCO2-rich and low-O2 atmospheric conditions,which could imply that large amounts of Ba haverecycled back to the mantle forming a ‘peak’ inthe formation of enriched mantle component.Whether the strong Ba depletion is characteristicof the Archaean–Proterozoic transition globallyand quiet supercontinent stages in general is to bedetermined.

Acknowledgements

This work was carried out in the context ofregional rock geochemical study at the GeologicalSurvey of Finland and my fellow researches EskoKorkiakoski, Pekka Lestinen, Reijo Salminen andHeimo Savolainen are thanked for their effort toaccomplish the project and for their continuousinterest in the subject. I am also grateful to GaborGaál, Hannu Huhma, Jarmo Kohonen and HughO’Brien for critically reading an earlier version ofthe manuscript. Reviewers K.C. Condie and C.M.Fedo are thanked for their constructive commentson the manuscript. This work is published withthe permission of the Geological Survey ofFinland.


References

Awwiller, D.N., Mack, L.E., 1991. Diagenetic modification ofSm–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 theBelomorides (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. Geology12, 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 andmetamorphic zircons and monazites from a pre-Taltsonmagmatic zone basin at the western margin of RaeProvince. Can. J. Earth Sci. 31, 1353–1364.

Breit, G.N., Wanty, R.B., 1991. Vanadium accumulation incarbonaceous 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 thePrecambrian evolution of the Balt

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