(EARLY) PRECAMBRIAN CONVECTION CELL IN THE ...komati.mbnet.fi/ConvectionCell.pdfSignificance in Northern Finland, the Fennoscandian (Baltic) Shield. Saverikko (1990). The Lopian and

Originally unpublished appendix to my unaccepted academic dissertation at 1991. 1 (26)

A supplementary review to Saverikko, Matti, 1992: Komatiitic Explosive Volcanism, Volcanoes, and Its Tectonic

Significance in Northern Finland, the Fennoscandian (Baltic) Shield.

(Early) Precambrian Convection Cell In The Fennoscandian Shield?

by Matti Saverikko

SAVERIKKO, MATTI, 1992: (Early) Precambrian convection cell in the Fennoscandian Shield?

A supplementary review to Saverikko, Matti, 1992: Komatiitic Explosive Volcanism, Volcanoes, and

Its Tectonic Significance in Northern Finland, the Fennoscandian (Baltic) Shield.

Http://koti.mbnet.fi/komati/convectioncell.pdf.

The Fennoscandian (Baltic) Shield comprises an Archean granite–greenstone terrain and an

early Proterozoic crustal segment of discrepant geosynclinal nature and obscure early history.

The Saamian sialic crust, which cratonized at 3.1-3.0 Ga and split into a mosaic of

megablocks, was involved as a coherent plate in domal uplift and broke up along a mantle

diapir. Mantle-activated rifting at 2.7-2.6 Ga was accelerated with widespread explosive

volcanism of komatiites. The divergence remained at the incipient stage and the cratonic area

restabilized at 2.6 Ga.

The early Proterozoic segment, which differentiated from mantle to crust at 1.9-1.8 Ga, now

exists mainly as a continental granitoid province containing about 10% completely associated

Archean crustal material. A long intraplate basin with volcanic borders of rocks of mixed and

contaminated composition is here called the Birkala mobile belt. This belt was very active

1.90-1.89 Ga ago but shows signs of prolonged sagduction by older ( 2.75 Ga) crustal

material.

The late Archean linear mantle diapir and the Birkala mobile belt form a pair of tectonic

belts of shield dimensions, implying a mantle-convection cell (< 1000 km on the ground

level). The hypothesis requires that convection-cell mechanism was feasible only under

broken continental nuclei when the free high-thermal energy escaped from surrounding

"oceanic" provinces. The mantle-convection cell was exceedingly long in duration (2.8-2.3

Ga) and resulted by mantle currents which were active at 3.0-1.9 Ga in periodic systems

linked to global-scale mantle activities.

Key words. Mantle convection. Endogenic processes. Exogenic processes. Plate tectonics. Archean.

Proterozoic. Finland. Fennoscandia. Baltic Shield.




1 Introduction

The Precambrian bedrock in Fennoscandia is

composed of an Archean basement complex in

the northeastern part of the Shield and an early

Proterozoic crust in the southwestern part.

Contradictory interpretations of the strati-

graphy, particularly of the greenstone belts in

the Archean domain, have been presented;

although the Finnish bedrock (Fig. 1) provides

the most complete geologic profile across the

Fennoscandian (Baltic) Shield, yet Finnish

geologists are not unanimous about its

chronostratigraphy. This essay examines the

geochronological events, in particular those

tentatively attributed to an isotopic resetting

associated with previously overlooked mantle

activities, and proposes an unused plate-

tectonic pattern to the Fennoscandian Shield: I

need to explain the role of a linear mantle rise

(Saverikko 1990) in the Precambrian evolution

of the Shield!

Fig. 1. Lithostratigraphic features of the

Finnish bedrock, compiled after Simonen

(1980), Gaál (1986), Luukkonen and

Lukkarinen (1986), and Saverikko (1987).

Greenstone belts: 1. Lapland, 2. Kuhmo –

Suomussalmi, 3. Ilomantsi. The southwestern

border of the Raahe(R)–Ladoga(L) tectonic

belt is marked with a dashed line.




2 Finnish Precambrian in brief

2.1 Archean domain

The Archean granite-gneiss complex

includes unknown quantities of the Saamian

granitoids dated at 3.1-2.8 Ga (Simonen 1980,

Kröner et al. 1981, Luukkonen and Lukkarinen

1986, Paavola 1986, Huhma 1987) and,

indirectly, at 3.2 Ga from a detritus of Archean

metasediments (Huhma 1987). Nd-isotopic

data on the rare protolith indicate derivation

from the mantle at least 3.5 Ga ago (Jahn et al.

1984). The Saamian sialic crust, in Russia if

not elsewhere, was cratonized at 3.1-3.0 Ga

(Musatov et al. 1984), as suggested also by

cratonic sedimentation after 3.0 Ga in northern

Finland (Saverikko 1987). The high scatter in

2.8-2.6 Ga dates (Simonen 1980) results from

geochronological overprints in the above-

mentioned rocks (e.g. Kröner et al. 1981,

Paavola 1986). Granitoids generated at that

time occur in and around the Archean

greenstone belts (Gaál 1986, Luukkonen and

Lukkarinen 1986), some of them being gene-

tically linked to the adjacent volcanics (Martin

1987). Hence, the Archean granitoid complex

contains the Saamian body and a younger

granitic contribution.

The Cwenan greenstone-belt genesis, 3.0-

2.5 Ga in maximum age span (Saverikko

1987), has been dated at 2.9-2.5 Ga in the

Kuhmo–Suo-mussalmi area (Luukkonen and

Lukkarinen 1986, Vaasjoki 1988). The

greenstone-belt associations in Lapland and

Kuhmo–Suomussalmi (Ilomantsi included) are,

broadly speaking, similar (Saverikko 1990): the

tripartite supracrustal sequences are composed

of bimodal volcanics in the lower part and of

multimodal volcanics in the upper part, with

shallow-water clastic sediments, Fe-rich

tholeiitic basalts, and graphitic debris in

between (e.g. Taipale et al. 1983, Barbey and

Martin 1987, Saverikko 1987, Tuukki et al.

1987). The paleoresidue of the Saamian sialic

crust is preserved only in Lapland, where it is

found together with basal arkoses and the cra-

tonic quartzite–carbonate–schist suite; also

unique to Lapland are the pyroclastic koma-

tiites prevailing in the upper volcanic complex

(Saverikko 1987).

The Lapland greenstone belt extends into

northern Sweden and Norway (e.g. Saverikko

1987, 1990), where it is considered early Prote-

rozoic in age (e.g. Witschard 1984, Krill et al.

1985, Skiöld 1987).

High-grade metamorphic rocks in Lapland

form an arc-shaped granulite belt of the lower-

middle Lapponian (see Saverikko 1987). The

bimodal-volcanic–arkose–slate association

(Barbey et al. 1984) is 2.8-2.5 Ga old and may

have been uplifted at 2.5 Ga (Meriläinen 1976).

The identification by Bernard-Griffiths et al.

(1984) of a 1.9 Ga old volcanic rock at the

major thrust plane of the arc (Barbey et al.

1984) has encouraged Barbey et al. (1984) and

Barbey and Martin (1987) to argue that the

granulites are of early Proterozoic age.

Depositional evolution in Lapland advanced

from cratonic sedimentation through cratonic

rifting to mantle-activated rifting mainly in ter-

restrial settings (Saverikko 1987). The strata of

the granulite belt were the result of turbidity-

current deposition in an intraplate trench after

the cratonic rifting (Barbey et al. 1980, 1984).

The Kuhmo–Suomussalmi belt evolved in con-

tinental to oceanic trench environments of

intraplate rifting (Martin et al. 1984). The

intraplate trenches are also found in Russia

(Musatov et al. 1984, Rybakov 1988), where

Salop (1983, ps. 99, 136) proposed a

correlation between the Lopian and Sumian

sequences and the Kuhmoan and Lapponian

sequences, although not as suggested by




Saverikko (1990). The Lopian and the

Kuhmoan/Lapponian sequences are considered

by Muradymov et al. (1988) to be directly com-

parable in stratigraphy.

Paleomagnetic data imply that the

continental areas have formed a coherent plate

since at least 2.7 Ga (Pesonen and Neuvonen

1981). Nonetheless, the data are insufficient for

the drawing of separate paths for crustal blocks

(Pesonen et al. 1989), even though Mertanen et

al. (1989) suggest that no large-scale

movements have taken place between the

basement segments of Lapland and southern

Karelia. Preferably, the radial swarm of

greenstone-belt trenches (Fig. 2) should be

interpreted as an aulacogen net, a viewpoint

that is consistent with domal uplift of the

compact plate at the start of the Cwenan

diastrophism, i.e. at 3.0 Ga! Continental rifting

at 3.0 Ga has been recognized elsewhere, too

(Burke et al. 1985), and the Archean conti-

nental crust in general has been interpreted as a

mosaic of sialic megablocks (Kröner 1981).

.

Fig. 2. Continental breakup and mantle diapir in association with domal uplift in the Archean conti-

nent (Saverikko 1990). The Skellefte(–Raahe–Ladoga) tectonic belt is regarded as the marginal rift,




although the Kiruna(–Jällivaara) fault system (Witschard 1984) is also in agreement with Öhlander et

al. (1987).

The basic factor in the Cwenan tectonic re-

gime in Lapland was linear mantle diapirism of

shield dimensions (Fig. 2), apparent as the

large Solovetski mantle plume (Bylinski et al.

1977), a swarm of small ultramafic bodies

(Papunen and Idman 1982), the upper

Lapponian pyroclastic komatiite zone (Fig. 3)

and an auriferous province (Saverikko 1990).

The center of domal uplift correlates well with

the linear mantle diapir.

Fig. 3. The chain of komatiitic volcanoes ex-

posed in a zone of pyroclastic complexes

(Saverikko et al. 1985).

Bylinski et al. (1977) and Efimov et al.

(1977) described counterclockwise rotation of

the Kola megablock in association with the

Solovetski mantle plume after the Late

Archean; the divergence, which is revealed in

the Kantalahti rift, led to southwesterly overth-

rusting of the granulite belt in the Early

Proterozoic. The overthrust has also been

regarded as a continent-collisional structure at

2.0-1.9 Ga (e.g. Barbey et al. 1984, Marker

1985, Krill 1985), despite the lack of any deep-

crustal proof of a descending plate or tectonic

crustal thickening (von Knorring and Lund,

1989). Further, the inferred suture (Barbey et

al. 1984) includes a wedging-subsidence fault

system of the Kantalahti rift which ruptures

circular megastructures in the granitoid

basement (Bylinski et al. 1977), demonstrating

preferably the crustal split in this region.

Tectonic restabilization 2.6 Ga ago (Silven-

noinen 1985) appears to have continued until

early Karelian times, 2.5-2.3 Ga, when a

Sariolan conglomerate–arkosite–argillite–

greenstone association was laid down onto the

weathered granite–greenstone terrain as local

sediments of half-graben to platformal type

deposits (Pekkarinen 1979, Meriläinen 1980,

Marmo et al. 1988); the sandy-argillic suite

was partly glaciogenic (Marmo and Ojakangas

1984, Marmo et al. 1988). The widespread but

rare, thin exposures (see: Luukkonen and

Lukkarinen 1986, Saverikko 1987) may

indicate extensive crustal fissuring with minor

fault-block subsidences and subordinate

volcanism. The Sariolan lithogenesis was

similar to the lithosphere-activated rifting of

Condie (1982, pp.175-177).

A late Karelian (2.3-2.0 Ga) transgressive

sedimentation followed the erosional period,

and Jatulian quartzites and a Marine-Jatulian

slate–dolomite–black-slate association

deposited as an extensive sheet with distinct

depositional subprovinces (e.g. Ojakangas

1965, Pekkarinen 1979, Meriläinen 1980,

Perttunen 1985, Kontinen 1986, Luukkonen

and Lukkarinen 1986). This anorogenic period

was accompanied by mafic volcanism (Meriläi-

nen 1980, Simonen 1980, Aro and Laitakari




1987).

2.2 Raahe–Ladoga tectonic belt

The Raahe–Ladoga tectonic belt (Fig. 4) is

composed of major NW-trending dextral

wrench faults, the Main Sulfide Ore Belt, a

negative gravity anomaly, and a deep-seated

fracture system; its Svecokarelian (2.5-1.75

Ga) origin is suggested by the Svecofennian

(2.0-1.75 Ga) intrusions, volcanics, and fault-

block movements, and by the Jatulian (2.3-2.0

Ga) eruption fissures, and Sariolan (2.5-2.3 Ga)

graben faults (see references in: Gaál 1986,

Huhma 1986). It is clear that, as a tectonically

active continental margin (Gaál 1982), the belt

separated the Archean and early Proterozoic

crustal segments (Huhma 1986). However, the

proof of Archean tectono-magmatic activities

has been disregarded: the Archean greenstones

of subalkaline to calc-alkaline nature

(Kähkönen et al. 1986, Luukkonen and

Lukkarinen 1986, Tuukki et al. 1987) and the

block-faulting processes linked to an Archean

carbonatite (Talvitie 1971, Puustinen and

Kauppinen 1989) and to vertical movements

(Paavola 1984, 1986) suggest tectonic activity

at the margin as early as during the Archean.

The marginal rift was part of the Archean rift

swarm with NW trend (Saverikko 1990).

Fig. 4. The Raahe–Ladoga tectonic belt bound-

ing the Svecofennian crustal segment with vol-

canic arcs, in the northeast.

The Karelian (2.5-2.0 Ga) tectonic evolution

of the belt, too, has remained obscure largely

on account of the Kalevan graywacke–slate

cover deposited at 2.0-1.9 Ga. The Kalevan

metasediments contain materials derived from

both Karelian epicontinental supracrustals and

the Archean basement (Huhma 1987); arkosic

conglomerates and other phenoclastics at the

base indicate marginal rifting and base-of-slope

environment (Honkamo 1985, Ward 1987,

references in Luukkonen and Lukkarinen

1986). The sedimentation, which was partly

controlled by the N-trending fractures of the

Archean basement (Bowes et al. 1984), was

characterized by high relief and rapid erosion

(Huhma 1987). The Kalevan province with

distinct depositional basins (Ward 1988) and

local Jatulian-Kalevan unconformities

(Simonen 1980) may indicate continental shelf

environments of basin-and-range nature

(Saverikko 1990).




The Raahe–Ladoga marginal rift extends to

Sweden in the form of the Skellefte fault

system and metallogenic province (Adamek

and Wilson 1979, Lund 1987). The

contribution of a Svecofennian suture to its

development has been advocated by some

authors, especially by Gaál (1986) and by Gaál

and Gorbatschev (1987), largely in view of the

limited subhorizontal thrusting of supracrustal

cover, the tectonic crustal thickening, and the

volcanism. Park (1985) suggested an accretion

model similar to that of the Phanerozoic.

According to Vaasjoki and Sakko (1988), the

Raahe–Ladoga belt underwent different plate-

tectonic processes, when a Proterozoic oceanic

plate collided with the Archean continent

between 1.93 and 1.85 Ga ago. However,

Welin (1987) has abandoned the plate-tectonic

paradigm, because no seismic evidence for a

subducted plate has been found. [Obs! A

descendent infracrustal fault surface is found

by sounding south of Skellefte (P. Heikkinen,

oral commun., 18.4.1991)]. Instead he

proposes an intracontinental rift development.

As is stated later in this paper, the Sveco-

fennian sialic segment evolved mainly at 1.90-

1.87 Ga. Mafic-ultramafic plutons of that age

in the Raahe–Ladoga belt display chemo-

petrologic affinity of continental intrusions

emplaced at moderate pressure (Mäkinen

1987). After this crustal event, probably during

the uplift of some fault blocks (Korsman et al.

1988) 1.88-1.84 Ga ago (Neuvonen et al.

1981), the tectonic belt was active between the

sialic/continental segments. The collision

proper with the Archean segment may have

been a later event, coinciding with reactivation

of NW-trending strike-slip faults that

controlled the emplacement of some granitoids

1.84-1.83 Ga ago (Nironen 1989), and with the

tectonic crustal thickening at 1.82-1.80 Ga

(Hölttä and Korsman 1986, Vaasjoki and

Sakko 1988); the metamorphic effect at 1.82-

1.80 Ga became weaker into the Archean

segment (Haudenschild 1990). The

subhorizontal thrust, which is evident in

distinct allochthons (Koistinen 1981), appears

to have been related to the fault-block

movements, because only the block adjacent to

the allochthons is rotated relative to its

surroundings in the Raahe–Ladoga belt

(Neuvonen et al. 1981).

2.3 Svecofennian (2.0-1.75 Ga) domain

According to Welin (1987), this crustal seg-

ment formed a bowl-shaped basin for the

Svecofennian sea (Fig. 5). Its Archean-Karelian

(> 2.0 Ga) geologic history is unknown.

Thickening of the strata (Lundqvist 1987,

Lundström 1987) and paleocurrent patterns

(Ojakangas 1986) are directed on a depocenter

which may have been located north of the

Åland Islands. The geosynclinal sequence has a

maximum thickness of 8-10 km within the vol-

canic belt in southern Finland and on the west

coast of the Gulf of Bothnia (Simonen 1980,

Lundqvist 1987).

In general agreement with Simonen (1953),

Hietanen (1975) stated that the belt was a vol-

canic island arc, that continued over to

Sweden. However, the volcanic province in

central-southern Sweden is discontinuous

eastwards, being in some way related to a

continental crust (Lundström 1987). Oen

(1987) has described a graben-basinal

development, with the spreading-subsiding

trough directed to the inferred depocenter. The

mainly felsic volcanism at 1.90-1.86 Ga (Oen

1987, Welin 1987) was associated with this

intraplate destruction zone with no evidence of

an oceanic crust or its subduction (Welin

1987). Hence, the volcanic arc is limited to

southern Finland, though its strike may turn to

the north along a nickeliferous magmatic belt

(Kahma 1973) that continues in the Ni-Cu

province south to Skellefte in Sweden (Nilsson

1985) (see Fig. 6).




The arc framework is indirectly confirmed

by a combined zone of younger rapakivi

granites and associated mafic rocks (see Figs. 4

and 6), as these rock assemblages generally

constituted elongated arrangements in post-

orogenic rifting propagated by linear mantle

processes (Emslie 1978). The mafic rocks

appear as dikes and small intrusions within and

beside the rapakivi granites (Aro and Laitakari

1987) and in a 4 km thick sheet beneath the

largest massif (Tuomi 1988). The basic

magmatism was tholeiitic: the dikes are

chemically similar to continental flood basalts

and the gabbroid–anorthosite suite is typical of

the Proterozoic massive anorthosites (Rämö

1991). The rapakivi granites, which derived

from reworked 1.9 Ga Svecofennian crust

(Huhma 1986, Rämö 1991), indicate mantle–

crust interaction along the line of crustal

sinking. Previous, but weaker, action of the

mantle is evidenced by (hot)spot-like centers

of low-pressure granulitization (see Westra

1988) and by small komatiite exposures in

southernmost Finland (Ehlers et al. 1986,

Schreurs et al. 1986, Laine 1988). Komatiites

are found also south of Skellefte, in Sweden

(Nilsson 1985).

Fig. 5. Svecofennian supracrustal provinces in

Finland and Sweden, completed after Welin

(1987).

The supracrustal sequence is thickest (Simo-

nen 1980, Lundqvist 1987) within the volcanic

belt, which appears as synclines with

subvertical axial planes and gently plunging

fold axes (see references in Kähkönen 1989). A

marine trench is suggested by a mainly

subaqueous paleoenvironment (Kähkönen

1989), basaltic volcanics of oceanic affinity

(see below), lithologic provinces with different




metallogenic patterns (Kinnunen and Saltikoff

1989), and mafic-ultramafic rocks with chemo-

petrologic signs of low-pressure marine

intrusions (Mäkinen 1987).

The volcanic belt is characterized by calc-

alkaline acidic to intermediate volcanics with

predominant pyroclastics; tholeiitic rocks with

pillow structures are subordinate (Gaál 1986).

The rocks in chemo-stratigraphic units show a

wide dispersion in chemical composition and

vary from basalts to rhyolites, from subalkaline

types to alkaline types, and from low-K rocks

to very high-K rocks (Kähkönen 1989). The

volcanic-sedimentary evolution occurred at

1.90-1.89 Ga (Patchett and Kouvo 1986,

Kähkönen et al. 1989). The major phases of

deformation, metamorphism, and migmatiza-

tion also occurred at those times (Hopgood et

al. 1983).

At first sight, the island-arc hypothesis

seems to be confirmed by the presence of

andesitic volcanics with the mixed and conta-

minated composition typical of island-arc

affinity (Patchett and Kouvo 1986). These vol-

canics were coupled with a thick volcaniclastic

apron (Laitala 1973) similar to arc-basinal

sediments (cf. Okada 1980). The volcanics in

the belt host massive sulfide ores, just like

those of island arcs (Latvalahti 1979, K.

Mäkelä 1980, Hangala 1987) and related to

early rifting episodes (U. Mäkelä 1989).

Additionally, a melange was developed at the

southern margin of the volcanic belt (Edelman

and Jaanus-Järkkälä 1983), and the synkine-

matic rocks of the granitoid complex of central

Finland (Nurmi and Haapala 1986), on the

opposite side of the belt, display many of the

features of subduction-related material (Huhma

1986, Front and Nurmi 1987); they also imply

diapiric emplacement closely related to major

regional compression (Nironen 1985).

Hietanen (1975) maintained that the oceanic

crust descended to the northeast – an idea

shared by many others as well.

The evidence against the island-arc

hypothesis is stronger, however. The volcanic

belt exhibits zonality in that the volcanogenic

margins are separated by the Tampere schist

area. The "inner-arc" abounds with calc-alka-

line intermediate volcanics that resemble

present-day mature arcs or arcs near or at active

continental margins (Kähkönen 1987). The

lowermost pillowed lavas of mantle origin

(Vaasjoki and Huhma 1987) are mid-ocean

ridge basalts (K. Mäkelä 1980) or marginal-

basin basalts (Kähkönen 1989). The basaltic

rocks do not usually represent primary mantle

melts but have undergone fractionation of

olivine ± Cr-spinel ± pyroxene (Kähkönen

1989). The low ISr values of the metavolcanics

suggest that their protoliths separated from

upper mantle sources within a short time

interval (Kähkönen et al. 1989).

The "outer-arc" contains tholeiitic amphibo-

lites with close geochemical similarity to the

within-plate basalts and calc-alkaline basalts at

continental margins (Ehlers et al. 1986, Ehlers

and Lindroos 1986). The mafic-ultramafic vol-

canic association also displays MORB or OIB

affinity (Ehlers et al. 1986, Schreurs et al.

1986).

The LREE-depleted asthenospheric mantle

reservoir beneath the present shield area was

similar to the source of MORB today (Huhma

1986). Volcanics with geochemical characteris-

tics of mid-ocean ridge and/or within-plate ba-

salts occur also elsewhere, outside the volcanic

belt (e.g. Vaarma 1990).

The Tampere schist area, which is here re-

garded as the basinal fill of the central trough,

is a Bothnian deep-sea province (Kinnunen and

Saltikoff 1989) and consists of intervolcanic

turbidites of submarine fans with westerly pa-

leocurrents (Ojakangas 1986). Rifting events

(Ehlers et al. 1986, Colley and Westra 1987,

Kähkönen 1987), which may have occurred in

local or temporal extensional tectonic regimes,

were not restricted to a single episode in the




evolution of these convergent-like plate

margins (Kähkönen 1987). The extensional

basins exhibit rifting-spreading events not

necessarily related to subduction (Kähkönen

1989). The lithologic zones with characteristic

metallogenic patterns display paleotectonic

provinces which do not correspond to present-

day plate-tectonic belts (Kinnunen and

Saltikoff 1989).

A further argument against the island-arc

hypothesis is that the Svecofennian crustal seg-

ments on either(!) side of the volcanic belt have

nNd from -1 to +3, suggesting an Archean crus-

tal contribution to the new mantle-derived

material, and implying substantial mantle-to-

crust differentiation at 1.9-1.8 Ga (Huhma

1986, Patchett and Kouvo 1986, Patchett et al.

1987). Moreover, despite the low-pressure

granulitic areas (Hölttä 1986, Schreurs and

Westra 1986), the paired high-P/high-T

metamorphic pattern (and high-pressure

granulites in general) is lacking in the volcanic

belt (Gorbatschev and Gaál 1987). The mafic-

ultramafic intrusions were also emplaced under

low-pressure conditions (Mäkinen 1987), and

at least some granitoids were first emplaced

passively through faults and fractures (Nironen

1989). The melange at the southern margin

may be an intrusive breccia (Edelman et al.

1986).

Differing from present-day convergent plate

margins in the rapidity of crustal growth

(Kähkönen 1987), the Svecofennian granitic

crust developed between 1.90 and 1.86 Ga

(Front and Nurmi 1987, Patchett et al. 1987),

but mostly from 1.89-1.87 Ga (Patchett and

Kouvo 1986). During the major volcanism, the

new continental crust beside the volcanic arc

was more than 20 km thick (Kähkönen 1987).

Thus, the Bothnian volcanic belt present in

the form of shallow-to-deep sea provinces

(Kinnunen and Saltikoff 1989) evolved within

one or between two or more sialic/continental

segments of newly formed crust and underwent

"oceanic" crustal development before the

segments rewelded. A high-thermal gradient

over large areas implies a relatively thin

differentiated acidic-to-intermediate crust with

a substantial heat source over large areas

(Westra 1988). As was noted by Branigan

(1987), the belt displayed signs of an ensialic

mobile belt during the subsequent long period

of mantle–crust interaction and crustal

adjustment by shearing (Hubbart and Branigan

1987). These post-orogenic processes after

1.83 Ga led to granitic magmatism with

forceful injection and refilling of collapsed

magma chambers (Ehlers and Bergman 1984)

and, at 1.70-1.54 Ga, to polyphase intrusion of

rapakivi granites

(Vaasjoki 1977). The rapakivi granites and re-

lated mafic rocks derived from anorogenic

magmas of partial melting in the upper-mantle

and Svecofennian continental lower-crust

(Rämö 1991, Branigan 1989).

3 Convection cell

The apparent conformity in strike between

the late Archean linear mantle diapir and the

early Proterozoic mobile belt (Fig. 6) suggests

genetic linkage, despite the ostensible

difference in their temporal and spatial

appearance (Fig. 7). If a mantle upwelling of

shield dimensions occurred, the rising mantle

currents must have been compensated by

descending currents of regional

extent responding to the crustal downwarp. The

inferred sagduction zone is here called the

Birkala mobile belt, without any chronostrati-

graphic connection. (Birkala was a region near

to Tampere from which the Swedish-Fin-




nish/"Svecofennian" people of the Viking Age

taxed and robbed Lappish/"Lapponian" or

"Saamian" people to the north).

Fig. 6. The Birkala mobile belt, which manifests itself as geologic features 1.90-1.89 Ga in age and as

younger rapakivi granites, may also show older crustal sagduction to compensate for Archean mantle

upwelling to the north. Layered intrusions (see Papunen et al. 1985) indicate frontal spreading of the

thermal zones (dotted areas) associated with the rising mantle currents.




Fig. 7. Crustal evolutionary periods in Fenno-

scandia and their connection with mantle

activities attributed to a convection cell of one

or more cycles.

3.1 Rising mantle currents

The linear mantle diapir, which manifests it-

self as the Archean mid-continental chain of

komatiitic volcanoes and the Solovetski mantle

plume, must have extended its thermo-tectonic

influence over the surroundings. The

continental province in Russian Karelia

possesses attributes of the mantle-related

endogenic processes that shifted from east to

west at 3.0-2.65 Ga (Lobach-Zhuchenko et al.

1986) when the greenstone-belt genesis in the

Kuhmo–Suomussalmi area in Finland started

with small-scale mantle convection and

continued with a rise of the mantle–crust

boundary (Engel and Diez, 1989): the

volcanism repeatedly discharged felsic to

ultramafic material with increasing MgO con-

tents of komatiites owing to advancing crust

and mantle melting (Taipale 1988). At the

western continental margin the tectono-

metamorphic period started at 3.0 Ga and

culminated at 2.7 Ga (Paavola 1986, 1988).

The widespread dates of 2.8-2.6 Ga determined

as overprints in the Saamian granitic crust

define also strong heat-flow episodes.

The mantle activity was apparent in one of

the most pronounced continental rotation at

2.7-2.6 Ga (Mertanen et al. 1989). The crustal

restabilization at 2.6 Ga (Silvennoinen 1985)

marks the end of the mantle-activated rifting

and of the associated strong heating that

probably brought about the limited granitic

diapirism accompanied by large-scale isotopic

homogenization (Halliday et al. 1988) at 2.75-

2.66 Ga (Luukkonen and Lukkarinen 1986),

and the low-pressure granulite metamorphism

at 2.65 Ga (Lobach-Zhuchenko et al. 1986).

The continued mantle activity is visible in

the layered intrusions, 2.45-2.44 Ga old, which

contain chromite, (vanadium)magnetite, and

PGE-bearing Cu-Ni ores (Piirainen et al. 1974,

Alapieti 1982, Söderholm and Inkinen 1982,

Lahtinen 1985, see Öhlander et al. 1987) and




which may derive from the upper mantle (Gaál

1986); the mantle activity is supported by

initial 87

Sr/86

Sr features of granites in the

surroundings (Halliday et al. 1988). The linear

distribution of the intrusions in the Archean

craton (Fig. 6) may indicate frontal shifting of

the rising mantle currents to the southwest. The

layered intrusions are also numerous in the

northeast, where they form an elongated

province with a WNW trend (Papunen et al.

1985). The mantle activity appears to have

slackened at the same time beneath the

paleotectonic mantle-diapir area in Lapland,

where only one layered gabbro complex

(Mutanen 1981) was emplaced. Diminishing in

strength, the mantle upwelling continued in

extended areas until 2.45-2.44 Ga. The

convective drift was nevertheless strong

enough to rotate the Kola megablock slowly

counterclockwise, causing the granulite belt to

be overthrust to the southwest (Bylinski et al.

1977) and uplifted at 2.5 Ga (Meriläinen 1976).

The thermal shifting advanced, as demon-

strated by isotopic homogenization (Halliday et

al. 1988) and K-Ar biotite blocking ages of 2.4-

2.2 Ga (Kallio et al. 1986) in the southwestern

periphery of the continental area in Finland.

This heat conduction culminated in the

remobilization of mantled gneiss domes in the

Raahe–Ladoga marginal rift at 2.5-2.3 Ga

(Brun 1980, Martynova 1980). The Birkala

mobile belt, in turn, showed sings of

sagduction in the form of the disappearance of

crust with an age of at least 2.3 Ga (see the

next section). This mantle-convection episode

was responsible for a clockwise rotation of the

shield at 2.4-2.2 Ga (see Mertanen et al. 1989).

3.2 Descending mantle currents

The unknown early history of the Sveco-

fennian (2.0-1.75 Ga) domain calls for specula-

tive introduction to the crustal processes before

1.90 Ga.

Although no Archean-Karelian (> 2.0 Ga)

crustal units have been found in the

Svecofennian domain, the newly formed crust

has a ubiquitous Archean component, which

accounts for an estimated 10% of the total

material (Patchett and Kouvo 1986, Patchett et

al. 1987). Owing to its uniform appearance, it

has been inferred that the main input of the

recycled Archean crustal material was in the

form of sediments (Patchett and Kouvo 1986,

Patchett et al. 1987). For this to be true, the

detritus would have to have been transported as

much as 500 km from the actual exposed

Archean crust. The distance is realistic for the

transportation of marine sediments (e.g.

Reineck and Singh 1980, p.474). However,

only a few synkinematic granitoids have

Al/(Ca/2+Na+K) ratios high enough to accord

with a pelitic component in their source (Front

and Nurmi 1987). The granitoids with the

highest content of Archean components, i.e.

about 20% (Patchett et al. 1987), are spatially

connected with the deposition areas of the

Bothnian (1.9-1.75 Ga) graywacke–slates,

which typically contain 25-50% Archean

material (Huhma 1985, Claesson 1987,

Patchett et al. 1987).

A few remains of micro organisms, mainly

stromatolites, in the early Proterozoic sea(s)

since the Marine-Jatulian stage (Tynni 1971,

Matisto 1974, Perttunen 1985, Tynni and Sara-

pää 1987) suggest that amounts of pelagic de-

bris, too, were inappreciable relative to the

other detritus.

A more likely agent for transporting the in-

ferred sediment cover of constant volume may

have been the Sariolan (2.5-2.3 Ga) continent-

wide glaciation proposed by Ojakangas (1985).

This requires the existence of a pre-Sariolan,

i.e. Archean, crust. The glaciomarine nearshore

deposits were composed of a sandy-argillic

association with a diamictite member (Marmo

and Ojakangas 1984, Marmo et al. 1988).

The inferred Archean crust and the absence




of clayey and pelagic detritus are features that

permit me to argue that the Archean

component, either partially or predominantly,

originated from a thin, solid crust that was

assimilated with the evolving early Proterozoic

crust; the Archean component was thoroughly

mixed, and the fair homogenization outside the

Bothnian sedimentary areas may have been due

to active, near-surface tectonic processes

(Patchett et al. 1987). Owing to the mixing, it

would be futile to look for a depleted mantle of

the conventional type, with åNd from +4 to +5,

beneath the Svecofennian domain (Patchett and

Kouvo 1986, Patchett et al. 1987).

In the Birkala mobile belt, Patchett and

Kouvo (1986) discerned features of mixing and

contamination with Archean material in several

volcanics 1.89 Ga old. Vaasjoki and Huhma

(1987) reported a common lead age of 1.99 Ga

for a volcanic suite of mantle-derived and

contaminated material that now lies at the

lowermost known supracrustal level

(Kähkönen 1989). Kähkönen (1986) described

a conglomerate with quartzitic and felsitic

pebbles near this level, and inferred the

existence of rocks older than 1.90-1.89 Ga.

Similarly, Huhma (1987) reported a mean

crustal residence age of 2.22 Ga for the

Tampere metasediments, which, according to

Kouvo and Tilton (1966), contain 2.30-Ga-old

detrital zircons sufficiently coarse grained to be

evidence of a local crust at least this old. The

incorporation and digestion of slightly younger

material is corroborated by small inherited

zircon components 1.96 Ga and 1.91 Ga in age

within the 1.90 Ga old trondhjemites (Patchett

and Kouvo 1986). The only direct evidence for

the hypothetical mixing of Archean material is

the 2.75 Ga old relict zircon cores found within

a rapakivi granite 1.59 Ga old (Vaasjoki 1977).

The Birkala mobile belt was thus active be-

fore the Bothnian (1.9-1.75 Ga) volcanism, at

least since the Karelian times, when the 2.3 Ga

old crust disappeared. (If this crust disappeared

from sight, why could not the Archean crust

have done so, too, or dragged into the Birkala

mobile belt during sagduction). The surviving

evidence is poor as is that for the Archean

material at least 2.75 Ga old. It is worth noting

that the main stage of mantle diapirism in

Lapland, which is manifest in the mantle-

activated rifting, followed the cratonic rifting at

2.79-2.70 Ga (Saverikko 1987, 1988). The

mantle upwelling, beginning as early as 3.0 Ga

ago, appears to have reached such a high-

energy level after 2.8 Ga that the descending

mantle currents were probably in operation.

3.3 Duration

Providing the above concept is valid, the

Precambrian mantle-convection was a slow

process of very long duration, starting to evolve

3.0 Ga ago, accelerating at 2.8-2.6 Ga, and

continuing until at least the beginning of late

Karelian times (2.3 Ga) or of Svecofennian

times (1.96-1.90 Ga). But the slow rate and

protracted duration of Archean(-Proterozoic)

tectonic events are demonstrated by the mantle-

related tectono-metamorphic megapulse at 3.0-

2.7/2.65 Ga in the continental area (Lobach-

Zhuchenko et al. 1986, Paavola 1986, 1988)

and by the rotation of the Kola megablock after

the Late Archean and during Proterozoic times

(Bylinski et al. 1977, Efimov et al. 1977) when

the crustal opening stayed at an embryonic

stage of the Wilson cycle (see Saverikko 1990).

3.4 Mantle reactivation

The diminution in mantle activity permitted

the early Karelian (2.5-2.3 Ga) lithosphere-acti-

vated rifting, but the reactivity characterized

the late Karelian (2.3-2.0 Ga) period and

affected the entire Archean continent. The

process was thermally mild (Mertanen et al.

1989) despite its demonstration by the




numerous small mantle-derived intrusions of

2.2-2.1 Ga old gabbro–wehrlite association

(Hanski 1986) and the 2.0-1.9 Ga old mafic

granulites of upper-mantle origin (Bernard-

Griffiths et al. 1984) emplaced in the Archean

paleotectonic belts, and by the 2.2-1.95 Ga old

diabases originating from the mantle (Huhma

1986) and occurring in recently opened joints

throughout the continental area forming

parallel dense swarms mostly in a NW

direction (Aro and Laitakari 1987).

The Jatulian (2.3-2.0 Ga) sedimentation,

first alluvial and then marine, marked a

transgression of regional extent and the end of

mantle uplifting; the significant presence of

coarse-clastic arkosic detritus implies sharp

uplifts linked to rapid crustal collapse

(Pettijohn 1975, p.167). It was associated with

a high-kinematic activity in the continental

drift (Mertanen et al. 1989) and the crustal

extension was propably caused by lateral

stress-strain forces.

The continental margin underwent (half-

)riftal processes after 2.1 Ga (Kontinen 1986,

Koistinen 1986, Ward 1987), accompanied by

emplacement of the Jormua ophiolite at 1.96

Ga (Kontinen 1987) and of the Outokumpu op-

hiolitic rock suite with strata-bound Cu-Co-Zn

ores at 1.97 Ga (Koistinen 1981). The area as a

whole displays evolutionary records in agree-

ment with the non-plate-tectonic hotspot model

of Lambert (1981), as the medium-size mantle

plumes evolved during the main stage of devel-

opment in the continental setting (see

Saverikko 1990). Epicontinental deposition of

the Kalevan thick turbidity graywacke–slate

successions, 2.0-1.9 Ga ago, required high

differences in the vertical movements of crustal

blocks, which may have responded to limited

mantle upwelling.

Another hotspot-like product, in my

opinion, is the host rock of the Pechenga Ni-Cu

ores in Russia, i.e. the serpentinite–pyroxenite–

gabbro complex of mantle parentage

(Gorbunov et al. 1985, Hanski 1986) that lies

in a triple-junction setting (see Gorbunov et al.

1985, Saverikko 1990). Its comagmatic

relatives are 1.99(-1.97?) Ga old ultramafic

volcanics (Hanski et al. 1991) which display

chemical characters of within-plate basalts

(Hanski and Smolkin 1989). The hotspot

activity appeared thus only at the continental

periphery, owing either to its low intensity or to

the great thickness of the proximal crust.

The Svecofennian orogeny (2.0-1.75 Ga)

had a penetrative effect far into the Archean

terrane (Mertanen et al. 1989) and invoked

recrystallization and isotopic resetting of the

Saamian granitoid basement and the greensto-

ne-belt associations (see Simonen 1980,

Kröner et al. 1981, Siedlecka et al. 1985,

Paavola 1986, Mertanen et al. 1989), finely

demonstrated by convulsive granulite

metamorphism (Bernard-Griffiths et al. 1984)

and by basement reactivation (Witschard

1984).

The basement reactivation in a continental

continuation of the Birkala mobile belt in

northern Sweden is worth noting (see Fig. 6).

The distinct Archean granitoid areas preserved

there provide signs of geologic events at 2.83

Ga and 2.67 Ga; the contribution of Archean

material to the granitic remobilization at 1.89-

1.84 Ga decreases toward the Archean

continental margin (Öhlander et al. 1987), in

front of which only a minimal Archean

component occurs (Skiöld and Öhlander 1989)

and the mantle efflux now occurs as komatiites

of the 1.88 Ga old Ni-Cu ore province south of

Skellefte (Nilsson 1985, see Lundqvist 1987).

Negligible amounts of komatiitic magma also

discharged in the Finnish side of the Bothnian

volcanic belt (Ehlers et al. 1986, Schreurs et al.

1986, Laine 1988). The mantle-crustal

reactivation of the sialic basement took place in

association with the Bothnian "oceanic" crustal

opening of the Birkala mobile belt.

The Svecofennian crustal segment, which




developed mainly 1.90-1.87 Ga ago, was

uplifted coevally as shown by regression of the

Svecofennian sea at 1.90-1.85 Ga (Lundqvist

1987) and by a widespread randomly oriented

joint group of mafic dikes 1.90-1.80 Ga old

(Aro and Laitakari 1987). This crustal uplift is

consistent with

the Bothnian regional-scale magmatic

underplating caused by imbrication tectonics

and/or low-angle subduction as postulated by

Westra (1988) to explain the thermal history

and crustal thickness of that time in the Birkala

mobile belt: the crust thickened from 5-10 km

to 60-75 km during the Early Proterozoic.

4 Discussion

This essay attributes controversial age deter-

minations within the Archean domain to

thermal overprinting. The geotectonic

constraint is based on the findings of recent

investigations, which thus supersede most of

the previous plate-tectonic interpretations of

the Fennoscandian Shield. Although the studies

partly deal with the geochemical affinities of

Precambrian igneous rocks, it is beyond the

scope of this essay to deliberate on the

relevance of geochemical comparison between

the (early) Precambrian and Phanerozoic rocks.

Instead, the Archean linear mantle diapir, in

good agreement with domal uplift, and the co-

herency of the Archean continental plate

should be taken into account in future plate-

tectonic evaluations of the Fennoscandian

Shield. Furthermore, in view of the lack of

solid evidence for (early) Precambrian

subducted plates and other implications of

buoyancy-subduction dualism, the subduction

hypotheses may be in need of amendment. The

island-arc systems are not readily applicable to

the (Archean) craton-scale framework, or, more

precisely, to the aulacogenic nature of the

greenstone-belt trenches (cf. Windley 1984, ps.

87, 355).

On a global scale, the inferred (early)

Precambrian convection cell under the

Fennoscandian Shield is consistent with

Archean plate tectonics, along with Archean

sea-floor spreading (e.g. Helmstaed et al. 1986)

and the theoretical availability of the Archean

ultramafic oceanic crust for subduction (Arndt

1983). Admittedly, it is difficult to believe

without detailed specific studies that the mantle

convection was in continuous operation during

2.8-2.3 Ga (or even from 3.0 Ga to 1.9 Ga) and

that the continental plate stayed in position

with respect to the convection cell despite

strong mantle currents and the Precambrian

drift of the Fennoscandian Shield.

The Fennoscandian Shield underwent three

major driftal periods: a counterclockwise rota-

tion along with only moderate latitudinal shifts

at 2.7-2.6 Ga, a clockwise rotation at 2.4-2.2

Ga, and a counterclockwise rotation

accompanied by a considerable latitudinal shift

at 2.2-1.9 Ga, which all reflect extensional

tectonic regimes and may implicate shifting of

the Shield as an independent plate before

getting together with other shields (Mertanen et

al. 1989).

Because of the extremely thin or (partly) ab-

sent crust, might it not be possible that, in Ar-

chean "oceanic" provinces, high-thermal proc-

esses accompanied by the movements of

mantle-derived material resembled bubbling

porridge? And were the conditions for

descending mantle currents to form part of the

convection cell only suitable under continental

nuclei? The main continental ruptures in the

crustal screen should have liberated

asthenospheric thermal pressure (Fig. 8), with

the result that the mantle upwelling would have

remained stationary in relation to the continent

despite free migration of the continental

nucleus. Also, the lateral extent of a convection

cell, which in the present instance is inferred to




be much smaller than those of today, depended

on the size of the overlying continent.

Fig. 8. The exogenic features of the Fen-

noscandian Shield discussed in the text may be

attributed to the above mantle-convection

development of the (early) Precambrian: a

continental nucleus creates a thermal-pressure

shadow on the underside, making convection-

cell mechanism feasible as the crustal screen

breaks up and brings the asthenosphere into

disequilibrium. Descending mantle currents

lower the high geothermal gradient at the

continental edges and lead to crustal over-

growth. At the same time the convection cell is

enlarged along with expansion of the crustal

screen. The lithospheric periphery of crustal

weakness is liable to sagduction due to period-

ic acceleration of mantle convection by in-

creasing heat-producing processes.

Mantle convection operated periodically in

connection with global mantle activity. The ac-

celeration, which was due to continental rifting

(2.8-2.7 Ga) and mantle-activated rifting (2.7-

2.6 Ga), is consistent with a late thermal event

at 2.9-2.7 Ga in the North Atlantic Archean

craton (see Windley 1984, pp.21-22) and with

the major world-wide greenstone-belt

development at 2.7-2.6 Ga (see Condie 1981,

p.43). The later events – the emplacement of

numerous layered intrusions (2.45-2.44 Ga),

the termination(?) of convection cell activity

(2.3-2.2 Ga), and the Bothnian "oceanic"

crustal opening along with mantle-to-crust

differentiation at 1.90-1.87 Ga – are also

reflected in a rhythmic discharge of mafic dike

swarms and plateau basalts in the Canadian

Shield 2.5-1.9 Ga ago (see Baragar 1977).

Acknowledgement. I wish to thank Prof. J.

Kalliokoski of Michigan Technological

University for encouraging me to write this

paper. Special thanks are due to Drs. O.

Kouvo, M. Vaasjoki, and H. Huhma, of the

Geological Survey of Finland, for critically

reading the manuscript and adding spice to my

work by presenting contrary opinions.




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