Precambrian geology of Finland: key to the evolution of the Fennoscandian shield

Developments in Precambrian Geology, 14

PRECAMBRIAN GEOLOGY OF FINLAND KEY TO THE EVOLUTION OF THE FENNOSCANDIAN SHIELD

DEVELOPMENTS IN PRECAMBRIAN GEOLOGY Advisory Editor Kent Condie Further titles in this series 1. 2. 3. 4. 5. 6. 7.

8. 9. 10. 11. 12.

13.

B.F. WINDLEY and S.M. NAQVI (Editors) Archaean Geochemistry D.R. HUNTER (Editor) Precambrian of the Southern Hemisphere K.C. CONDIE Archean Greenstone Belts A. KRÖNER (Editor) Precambrian Plate Tectonics Y.P. MEL’NIK Precambrian Banded Iron-formations. Physicochemical Conditions of Formation A.F. TRENDALL and R.C. MORRIS (Editors) Iron-Formation: Facts and Problems B. NAGY, R. WEBER, J.C. GUERRERO and M.SCHIDLOWSKI (Editors) Developments and Interactions of the Precambrian Atmosphere, Lithosphere and Biosphere S.M. NAQVI (Editor) Precambrian Continental Crust and Its Economic Resources D.V. RUNDQVIST and F.P. MITROFANOV (Editors) Precambrian Geology of the USSR K.C. CONDIE (Editor) Proterozoic Crustal Evolution K.C. CONDIE (Editor) Archean Crustal Evolution P.G. ERIKSSON, W. ALTERMANN, D.R. NELSON, W.U. MUELLER and O. CATUNEANU (Editors) The Precambrian Earth:Tempos and Events T.M. KUSKY (Editor) Precambrian Ophiolites and Related Rocks

Developments in Precambrian Geology, 14

PRECAMBRIAN GEOLOGY OF FINLAND KEY TO THE EVOLUTION OF THE FENNOSCANDIAN SHIELD

Editors:

M. LEHTINEN University of Helsinki, Finland

P.A. NURMI Geological Survey of Finland Espoo, Finland ..

..

O.T. RAMO University of Helsinki, Finland

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TABLE OF CONTENTS Table of contents ...............................................v Preface ........................................................ xiii 1. Overview ....................................................1 (M. Vaasjoki, K. Korsman, T. Koistinen) 1. Location, subdivision, timing, and general charateristics................................................4 2. Regional geographic nomenclature .............7 3. The Archean bedrock ................................13 4. Faulting of Archean crust and emplacement of Paleoproterozoic cover rocks .......13 5. The Svecofennian bedrock ........................13 6. Rapakivi magmatism and the Jotnian period ........................................................15 7. The Vendian period and the Paleozoic era 15 8. Late events affecting the bedrock ..............16 2. Archean rocks .........................................19 (P. Sorjonen-Ward, E.J. Luukkonen) 1. Introduction to the Archean of Finland .....22 1.1. The extent of the Archean in Finland .............................................22 1.2. Classifying and subdividing the Archean bedrock of Finland .............26 2. The Karelian domain in eastern Finland ......28 2.1. Ilomantsi terrain ..............................28 Hattu supracrustal belt .....................29 Kovero supracrustal belt...................36 Nunnanlahti and Ipatti supracrustal belts ......................................36 Lieksa complex – granitoids and high-grade gneisses ...................37 Granitoids intruding the Hattu and Kovero supracrustal rocks .........38 2.2. Kianta terrain ...................................40 Suomussalmi greenstone belt ...........43 Kuhmo greenstone belt ....................44 Tipasjärvi greenstone belt ................47 Granitoids, gneisses, and crustal evolution in the Kianta terrain .........48 Nurmes gneiss complex ...................52 2.3. Iisalmi terrain ...................................53 Proterozoic reworking and the boundaries of the Iisalmi terrain ......53 Origin of the present metamorphic

3.

4. 5.

6.

zonation pattern ...............................56 Varpaisjärvi granulite complex ........57 Rautavaara complex .........................58 2.4. Ranua terrain ....................................59 Oijärvi greenstone belt .....................60 Siurua granulite complex .................60 The Karelian domain in northern Finland.......................................................61 3.1. Koillismaa terrain .............................62 3.2. Napapiiri terrain ...............................62 Suomu terrain ...................................63 3.3. Tuntsa terrain....................................64 Granitoid complexes ........................65 Tuntsa and Tulppio supracrustal belts .................................................68 3.4. Pomokaira terrain .............................68 3.5. Muonio terrain .................................69 3.6. Ropi terrain ......................................69 The Kola domain in Finland......................70 4.1. Inari terrain.......................................71 4.2. Sørvaranger terrain...........................71 Insights into the deeper Archean crust in Finland ..................................................73 5.1. Exhumed deep crustal sections in Finland? ............................................73 5.2. Distribution and composition of buried Archean crust .......................75 5.3. Xenoliths and deep seismic studies...............................................76 Discussion and synthesis ...........................78 6.1. Archean thermal regimes and tectonic consequences ......................78 6.2. Regional scenarios and correlations ...............................................81 6.3. Comparisons and contrasts between Archean and Svecofennian crustal processes .......................83

3. Layered mafic intrusions of the Tornio–Näränkävaara belt ...................101 (M. Iljina, E. Hanski) 1. Introduction .............................................104 2. Geologic setting of the Tornio–Näränkävaara belt .............................................104 3. Cumulus sequences .................................106 3.1. General characteristics ...................106

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3.2. 3.3. 3.4. 3.5.

Kemi intrusion................................106 Penikat intrusion ............................108 Portimo layered igneous complex ..111 Koillismaa layered igneous complex ..........................................114 4. Parental magmas and isotope studies ......118 4.1. Parental magmas ............................118 4.2. Isotope studies ................................120 5. Mineral deposits ......................................120 5.1. Ore types ........................................120 5.2. Kemi chromite deposit ...................122 5.3. Mustavaara Fe-Ti-V oxide deposit, Koillismaa complex .......................123 5.4. PGE reefs of the Penikat intrusion .123 5.5. Marginal series Cu-Ni-PGE and reef-type mineralization of the Koillismaa complex .......................124 5.6. Diverse Cu-Ni-PGE mineralizations in the Portimo complex ......125 5.7. PGE geochemistry .........................130 6. Summary and discussion .........................131

4.1. ~2440 Ma intrusions in Lapland ....167 Akanvaara intrusion .......................167 Koitelainen intrusion ......................169 Parental magma ..............................170 Isotope geology ..............................171 4.2. ~2220 Ma differentiated sills .........171 4.3. ~2050 Ma intrusions ......................172 Keivitsa intrusion ...........................172 5. Lapland granulite belt .............................174 5.1. Metamorphic conditions ................175 5.2. Radiogenic isotopes .......................175 6. Summary and discussion .........................176 6.1. Mantle plume(s) and cracking of the craton ........................................176 6.2. Cratonic sedimentation and volcanism .......................................177 6.3. Primitive volcanism and deepening basins ............................177 6.4. Breakup of a supercontinent? .........178 6.5. Ocean floor volcanism ...................179 6.6. Acid magmatism related to obduction? ..........................................181 6.7. Foreland basin ................................181 Terrestrial sedimentation and volcanism .......................................181 Correlation with Svecofennian sedimentation and volcanism .........182 Relationship to the exhumation of granulites ........................................182 7. Conclusions .............................................183

4. Central Lapland greenstone belt ........139 (E. Hanski, H. Huhma) 1. Introduction .............................................142 2. Main geologic units of northern Finland.142 3. Central Lapland greenstone belt .............144 3.1. General features .............................144 3.2. Lithostratigraphy ............................144 3.3. Salla Group ....................................146 Geochemistry and Nd isotopes ......149 Geochronology ...............................149 3.4. Onkamo Group ...............................150 Geochemistry and Nd isotopes ......150 Geochronology ...............................154 3.5. Sodankylä Group............................154 Geochemistry and geochronology. .155 3.6. Savukoski Group ............................156 Geochemistry .................................156 Geochronology ...............................157 3.7. Kittilä Group ..................................158 Stable isotopes................................158 Field characteristics and geochemistry of mafic metavolcanic rocks ..159 Geochronology of mafic rocks .......160 Nuttio serpentinites and related dikes ...............................................161 Felsic rocks.....................................162 3.8. Lainio and Kumpu Groups .............164 Metasediments ...............................164 Metavolcanic rocks ........................165 Isotope studies of conglomerate clasts and detrital minerals ............166 4. Mafic plutonism ......................................167

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5. Paleoproterozoic mafic dikes in NE Finland .......................................195 (J. Vuollo, H. Huhma) 1. Introduction .............................................198 2. Geological background ...........................201 3. Mafic dike swarms ..................................203 3.1. ~2.45 Ga dike swarms ....................204 Boninite–norite dikes .....................205 Gabbronorite dikes .........................205 Low-Ti tholeiitic dikes ...................205 Fe-tholeiitic dikes ...........................207 Orthopyroxene-plagioclase-phyric dikes ...............................................207 3.2. Geochemical and isotopic characteristics ..........................................207 3.3. ~2.32 Ga dike swarm and intrusions ........................................211 3.4. ~2.2 Ga layered sills and dikes.......212 3.5. ~2.1 Ga dike swarms ......................215 3.6. ~1.98 Ga dike swarm .....................223 4. Tectonic significance of the dike swarms.....................................................226 4.1. Paleoproterozoic rifting events in OF

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the Archean Kuhmo block ..............226 4.2. Uplifted Archean high-grade terranes ...........................................228 6. Ophiolites ..............................................237 (P. Peltonen) 1. Introduction .............................................240 2. Significance of ancient ophiolites ..........241 3. Age constrains for Finnish ophiolites .....243 4. The Jormua ophiolite ..............................244 4.1. The crustal unit ..............................246 Petrology of the basalts ..................246 Gabbros and plagiogranites ............251 4.2. The mantle section .........................252 Serpentinites...................................253 Clinopyroxenitic and hornblenditic mantle dikes of the western block ...............................................254 5. Outokumpu-type ultramafic massifs .......255 5.1. Ultramafic rocks .............................257 5.2. Basaltic rocks .................................258 5.3. Cu-Co-Zn-Ni±Au sulfide deposits .261 6. The Nuttio serpentinite belt ....................262 7. Comparative geochemistry of the Finnish ophiolites ....................................264 7.1. Metaperidotites...............................264 7.2. Lavas and dikes ..............................266 8. Environments of ophiolite formation ......268 9. Concluding remarks ................................273 7. Karelian supracrustal rocks ................279 (K. Laajoki) 1. Introduction ............................................282 2. Geological setting and basin classification ....................................................282 2.1. Regional distribution of the supracrustal belts ...................................282 2.2. Metamorphism ...............................285 2.3. Tectonic features ............................287 2.4. Basin classification .........................287 3. Sumi tectofacies ......................................290 3.1. Supracrustal rocks ..........................290 3.2. 2440 Ma layered intrusions ............292 4. Sub-Sariola unconformity .......................292 5. Sariola tectofacies ...................................295 5.1. North Karelia .................................295 Glaciogenic rocks of the Urkkavaara Formation..............................295 5.2. Eastern part of the Kainuu belt ......297 Glaciogenic rocks of the Honkajärvi Group .....................................297 Kurkikylä Group ............................297 5.3. Western part of the Kainuu belt .....299 Puolankajärvi Formation ................299

6. 7.

8. 9.

10. 11.

12. 13.

14.

15. 16.

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5.4. Saari–Kiekki belt............................299 5.5. Sariola cover of the layered intrusions within the basement complexes..............................................300 5.6. Kuusamo belt .................................300 5.7. Peräpohja belt .................................300 5.8. Other Sariola occurrences ..............301 Sub-Kainuu unconformity.......................301 Kainuu tectofacies ...................................303 7.1. Kainuu belt .....................................303 Korvuanjoki Group in Kainuu .......303 Middle and Upper part of the Central Puolanka Group .................305 7.2. North Karelia .................................305 7.3. Kuusamo and Kuusijärvi ................306 7.4. Peräpohja ........................................306 7.5. Other occurrences ..........................306 Sub-Jatuli unconformity ..........................307 Jatuli tectofacies ......................................310 9.1. Koli and Kiihtelysvaara areas in North Karelia .................................310 9.2. East Puolanka Group and corresponding groups in Kainuu ............311 9.3. Kuusamo ........................................311 9.4. Peräpohja ........................................313 9.5. Other occurrences ..........................313 Sub-Lower Kaleva unconformity ............313 Lower Kaleva tectofacies ........................314 11.1. Kainuu belt .....................................314 11.2. Höytiäinen basin, North Karelia ....315 11.3. Kuopio area ....................................315 11.4. Salahmi belt....................................317 11.5. Kiiminki belt ..................................317 11.6. Peräpohja ........................................318 Sub-Upper Kaleva unconformity ............318 Upper Kaleva tectofacies ........................319 13.1. Upper Kaleva in Kainuu.................319 13.2. Upper Kaleva within the Outokumpu nappe complex and the Kuopio–Pielavesi area ....................319 Problematic younger Karelian formations ....................................................320 14.1. Vihajärvi Group and Haapalanmäki and Jokijyrkkä conglomerates ...........................................320 14.2. Pyssykulju Formation.....................321 14.3. Northern margin of the Peräpohja belt ..............................................323 14.4. Himmerkinlahti Member and Kolmiloukkonen Formation in Posio ...323 Karelian metadiabases.............................324 Previously proposed basin models ..........324 16.1. Continental and pericontinental Karelia (sensu stricto) basins .........324 GEOLOGY

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16.2. Kaleva basins .................................325 17. Paleogeographic reconstructions .............325 17.1. Continental–marginal Karelian sequences .......................................326 17.2. Kaleva sequences ...........................326 18. Synopsis ..................................................326 18.1. Karelia (sensu stricto) basin development .......................................326 18.2. Lower Kaleva development ............331 18.3. Upper Kaleva development ............331 18.4. Closing comments ..........................331

10. Discussion ..............................................388 10.1. Correlation of the Pohjanmaa belt to northern Sweden ........................388 10.2. Correlation of the Uusimaa belt to the Bergslagen field....................390 10.3. Correlation of the Häme and Uusimaa belts .................................391 10.4. Tiirismaa-type quartz arenites ........392 10.5. Angular unconformities? ...............393 11. Summary .................................................393 9. Svecofennian mafic–ultramafic intrusions ..............................................407 (P. Peltonen) 1. Introduction .............................................410 2. Classification of the intrusions ................410 3. Intrusions close to the craton margin (Group Ia) ................................................412 3.1. Laukunkangas ................................414 3.2. Kotalahti .........................................415 3.3. Lapinlahti gabbro–anorthosite .......416 4. Intrusions of the Tampere and Pirkanmaa belts (Group Ib) ...............................417 4.1. Ultramafic intrusions of the Vammala Ni province ............................419 4.2. Porrasniemi layered gabbro ............422 4.3. Kaipola layered intrusion ...............423 5. Synvolcanic intrusions of the Arc complex of southern Finland (Group II) ........426 5.1. Forssa gabbro .................................426 5.2. Hyvinkää layered intrusion ............426 6. Ti-Fe-P gabbros of the Central Finland granitoid complex (Group III).................428 6.1. Kauhajärvi gabbro province ...........428 Kauhajärvi gabbro ..........................429 Perämaa gabbro ..............................430 6.2. Koivusaarenneva layered intrusion .........................................430 7. Chemical and isotope composition of the mafic–ultramafic intrusions...........432 8. Economic aspects and petrogenesis of the ores ........................................................435 9. Concluding remarks ................................437

8. Svecofennian supracrustal rocks ........343 (Y. Kähkönen) 1. Introduction .............................................346 2. Geologic setting ......................................346 2.1. General aspects ..............................346 2.2. Proterozoic cover deposits of the Archean craton ..............................349 2.3. Division of the Svecofennian domain ............................................350 2.4. U-Pb zircon ages and Nd isotopes .351 3. Geochemical and tectonomagmatic characterization of the volcanic rocks ....354 4. Savo belt ..................................................355 4.1. General ...........................................355 4.2. Pielavesi–Pyhäsalmi region............356 4.3. Rautalampi region ..........................358 4.4. Volcanic rocks of the Virtasalmi region .............................................358 5. Pohjanmaa belt ........................................361 5.1. General ..........................................361 5.2. Evijärvi field...................................362 5.3. Ylivieska field ................................362 6. Tampere and Pirkanmaa belts .................365 6.1. General ...........................................365 6.2. Central Tampere belt ......................365 6.3. Western and eastern Tampere belt ..369 6.4. Pirkanmaa belt................................371 7. Supracrustal belts within the Central Finland granitoid complex .....................374 8. Häme belt and Saimaa area .....................375 8.1. General ...........................................375 8.2. Volcanic rocks of the Häme belt ....375 8.3. Volcanic rocks of the Saimaa area .377 8.4. Sedimentary rocks of the Saimaa area .................................................378 9. Uusimaa belt ...........................................380 9.1. General aspects ..............................380 9.2. Kemiö–Järvenpää field ...................380 9.3. Nauvo–Korppoo field .....................385 9.4. Pellinki field ...................................385 9.5. Sedimentary carbonates of the Uusimaa belt ..................................388

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10. Proterozoic orogenic granitoid rocks ...443 (M. Nironen) 1. Classification of plutonic rocks ...............446 2. Preorogenic rocks ....................................449 2.1. Preorogenic rocks of central Finland (1.93–1.91 Ga) ..................449 2.2. Preorogenic rocks of northern Finland (1.95–1.91 Ga) ..................449 3. Synorogenic rocks ...................................450 3.1 Synkinematic rocks of southern OF

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4.

5.

6.

7. 8.

and central Finland (1.89– 1.87 Ga) ..........................................451 3.2. Postkinematic rocks of central Finland (1.88–1.86 Ga) ..................452 3.3. Synorogenic rocks of northern Finland (1.89–1.86 Ga) ..................455 Lateorogenic granites ..............................456 4.1. Lateorogenic granites of southern Finland (1.84–1.81 Ga) ..................456 4.2. Lateorogenic granites of northern Finland (1.84–1.80 Ga) ..................458 Postorogenic rocks ..................................459 5.1. Postorogenic rocks of southern Finland (1.81–1.77 Ga) ..................459 5.2. Postorogenic granites of northern Finland (1.80–1.77 Ga) ..................462 Geochemical comparison and petrogenetic implications ................................462 6.1. Preorogenic rocks ...........................462 6.2. Synorogenic rocks ..........................468 Synkinematic rocks of southern and central Finland .........................468 Postkinematic rocks of central Finland ...........................................468 Synorogenic rocks of northern Finland ...........................................469 6.3. Lateorogenic granites .....................469 6.4. Postorogenic rocks .........................470 Discussion ..............................................471 Summary .................................................474

11. Paleoproterozoic tectonic evolution .....481 (R. Lahtinen, A. Korja, M. Nironen) 1. Introduction .............................................484 2. Geologic outline .....................................489 3. Pre-1.92 Ga crustal components and crustal-scale boundaries .........................493 3.1. Lapland–Kola area .........................494 3.2. Karelian craton ...............................496 3.3. Norrbotten Archean nucleus and attached island arcs .......................497 3.4. Keitele microcontinent and attached island arc ..........................498 3.5. Bothnia microcontinent and attached island arc ..........................498 3.6. Bergslagen microcontinent and Tavastia island arc ..........................499 4. Terminology related to the Paleoproterozoic tectonic evolution .........................499 5. Tectonic model ........................................500 5.1. Breakup of the Archean craton (or cratons) at 2.06 Ga ........................500 5.2. Lapland–Kola orogen .....................501 5.3. Lapland–Savo orogen .....................504

5.4. Subduction reversal and switchover: prelude to the Fennian orogeny at 1.90 Ga .........................505 5.5. Fennian orogeny: a north–south accretion stage at 1.89–1.87 Ga .....507 5.6. Attempted orogenic collapse and related magmatism at 1.89–1.87 Ga ................................508 5.7. The end of the Fennian orogeny at 1.87–1.85 Ga: orogenic collapse ....509 5.8. Svecobaltic orogeny: Andean-type active margin and continent–continent collision at 1.84–1.80 Ga .....511 5.9. The Nordic orogeny: continent– continent collision at 1.82–1.79 Ga ...................................................513 5.10. End of the Nordic orogeny and orogenic collapse at 1.79–1.77 Ga..514 6. Gothian evolution at 1.73–1.55 Ga .........515 7. Discussion ...............................................516 7.1. Comparison with modern analogues .............................................516 7.2. Comparison with earlier studies and models .....................................517 8. Concluding remarks ................................520 12. Rapakivi granites .................................533 (O.T. Rämö, I. Haapala) 1. Introduction .............................................536 2. What is rapakivi granite? ........................536 3. Distribution, mode of occurrence, and age ....................................................537 4. Lithologic association .............................539 4.1. Felsic plutonic rocks ......................540 4.2. Mafic plutonic rocks.......................545 4.3. Intermediate plutonic rocks............546 4.4. Dikes and volcanic rocks ...............546 5. Chemical composition.............................549 6. Origin of the rapakivi texture ..................550 7. Origin of the rapakivi magma .................552 8. Tectonic scenarios ...................................556 9. Future challenges ....................................557 13. Sedimentary rocks, diabases, and late cratonic evolution ...........................563 (J. Kohonen, O.T. Rämö) 1. Introduction .............................................566 2. Mesoproterozoic sedimentary sequences ................................................567 2.1. Regional setting..............................567 2.2. The Satakunta Formation and its submarine extensions .....................569 2.3. The Muhos Formation and its submarine extensions ....................573

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2.4. Minor occurrences .........................574 3. Mesoproterozoic igneous rocks ..............574 3.1. Introduction ...................................574 3.2. The ~1265 Ma magmatism ............574 Regional setting..............................574 Petrography and geochemistry .......575 Source characteristics and magmatic evolution ....................................576 3.3. The 1100–1000 Ma magmatism ....577 Regional setting..............................577 Geochemistry and source characteristics ...........................................579 4. Neoproterozoic and early Paleozoic sedimentary sequences ............................579 4.1 Regional setting..............................579 4.2. The Hailuoto Formation and its submarine extensions .....................580 4.3. The Lauhanvuori Formation...........580 4.4. The bottom of the Bothnian and Åland seas ......................................581 4.5. The Dividal Group of northwestern Lapland .............................582 4.6. Minor occurrences .........................583 5. Allochthonous rocks of the Finnish Caledonides .............................................584 5.1. Introduction and regional setting ...584 5.2. The Lower Allochthon (Jerta Nappe) ............................................585 5.3. The Middle Allochthon (Nalganas and Nabar Nappes) .........................585 5.4. The Upper Allochthon (Vaddas Nappe) ............................................585 6. Paleosols and Cenozoic sedimentary remnants ..................................................586 7. Tectonic evolution from the Mesoproterozoic to the Cenozoic ....................587 7.1. Introduction ....................................587 7.2. The intracratonic rift basin stage (~1600–1300 Ma) ..........................588 7.3. Crustal extension episodes and the Sveconorwegian orogeny (~1300– 900 Ma) ..........................................589 7.4. The Neoproterozoic exhumation stage (~900–600 Ma) .....................589 7.5. The stage of platform sedimentation (~600–420 Ma) ....................591 7.6. The Caledonian foreland stage (~420–350 Ma) and the final exhumation of the shield ................593 7.7. Concluding remarks .......................593

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14. Kimberlites, carbonatites, and alkaline rocks ...............................................605 (H.E. O’Brien, P. Peltonen, H. Vartiainen) 1. Introduction .............................................608 2. Description of alkaline rock complexes of Finland ................................................608 2.1. The Archean Siilinjärvi carbonatite ...............................................608 2.2. Proterozoic Kortejärvi and Laivajoki intrusions.................................611 2.3. Proterozoic lamprophyre dikes ......615 2.4. Proterozoic Halpanen carbonatite ..617 2.5. Proterozoic Group II kimberlites – olivine lamproites (K2L) ................617 2.6. Neoproterozoic Group I kimberlites .................................................619 2.7. Devonian Sokli carbonatite complex ..........................................621 2.8. Devonian Sokli ultramafic lamprophyre dikes ..........................627 2.9. Devonian Iivaara alkaline complex .628 3. Geochemistry of kimberlites, carbonatites, and alkaline rocks .......................629 4. Isotope composition of kimberlites, carbonatites, and alkaline rocks ..............633 5. The kimberlite mantle sample .................636 5.1. Mantle xenoliths .............................636 5.2. Mantle xenocrysts ..........................638 5.3. Diamonds .......................................639 15. Drift history of the shield......................645 (S. Mertanen, L.J. Pesonen) 1. Introduction .............................................648 2. Remanent magnetization in the Fennoscandian Shield .......................................648 3. Fennoscandian drift history in the Precambrian ..................................................650 3.1. Neoarchean.....................................651 3.2. Continental rifting at 2.4 Ga ..........652 3.3. Jatulian rifting and magmatism at 2.2–2.0 Ga ..................................653 3.4. Onset of the Svecofennian orogeny at 2.0–1.9 Ga ..................................654 3.5. Svecofennian orogeny at 1.9–1.8 Ga ......................................654 3.6. Subjotnian magmatic interval at 1.65–1.5 Ga ....................................655 3.7. Postjotnian time at ~1.26 Ga ..........655 3.8. Dike magmatism at 1.1–1.0 Ga ......655 4. Position of the Fennoscandian Shield in the continental assemblies of the Precambrian ............................................656 4.1. Early Paleoproterozoic ...................656

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4.2. Middle Paleoproterozoic ................658 4.3. Late Paleoproterozoic ....................659 4.4. Middle Mesoproterozoic ................660 4.5. Late Mesoproterozoic ....................661 5. Conclusions .............................................661 16. Paleoproterozoic carbon isotope excursion ................................................669 (J.A. Karhu) 1. Introduction .............................................672 2. Early records ...........................................672 3. Fennoscandian δ13C data .........................673 4. Global δ13C data ......................................675 5. Discussion ..............................................676 6. Conclusions .............................................678

3. Research organizations............................686 3.1. From the Geological Commission to the Geological Survey ................686 3.2. Universities ....................................687 3.3. Mining enterprises .........................689 3.4. Other research organizations ..........690 4. Main fields of research ............................691 4.1. Petrology and physical geology......691 4.2. Geochemistry and isotope geology 695 4.3. Mineralogy .....................................696 4.4. Economic geology..........................698 5. Synopsis ..................................................699 Contributors .................................................703 Index of persons and institutions ................707

17. History of Finnish bedrock research ...681 (I. Haapala) 1. Introduction .............................................684 2. Finnish geology in the 19th century ........684

Locality index ...............................................709 Subject index ................................................715

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PREFACE The Fennoscandian (or Baltic) Shield represents the largest outcropping domain of Precambrian bedrock in Europe, covering more than a million km2 throughout Norway, Sweden, Finland, and northwestern Russia. This book focuses on Finland, which occupies the central part of the shield and which, since the advent of modern geology in the 19th century, has been instrumental in a number of fundamental insights and advances in understanding Earth processes. Wilhelm Ramsay, who was the Professor of Geology and Mineralogy at the University of Helsinki in 1899–1928 and who introduced the term Fennoscandia, made an outstanding contribution to the understanding of alkaline rocks through his studies of the Devonian Kola province in the northeasternmost part of the shield. Meanwhile, J.J. Sederholm, Director of the Geological Survey of Finland in 1893–1933, pioneered the application of actualistic principles to Precambrian terrains and the systematic study of Precambrian granites, introducing the concepts of migmatites and anatexis in 1907, and published acclaimed monographs on orbicular textures and the rapakivi granite association. Pentti Eskola, who succeeded Ramsay in the Chair of Geology and Mineralogy at Helsinki in 1929–1953, is particularly renowned for defining the metamorphic facies concept, based initially on the Orijärvi district near Helsinki, and which now underpins studies in metamorphic petrology worldwide. Further developments in analytical chemistry and elemental and isotope geochemistry, by Th.G. Sahama and Kalervo Rankama, paved the way for isotopic calibration of Precambrian rocks and events, which has been essential to attaining our present understanding of crustal evolution. Concurrent advances in geophysical techniques and instrumentation, while driven mainly by exploration applications, have played an equally significant role in mapping the country in recent decades, especially in poorly exposed areas, by providing detailed airborne survey as well as deep seismic sounding data. As a consequence, the Finnish part of the Fennoscandian Shield can rightfully be considered as one of the best-documented Precambrian terrains in the world. This compilation provides the first modern account of the geology of Finland. The seventeen chapters of the book have been written by geologists and geophysicists who have actively contributed to the research in their respective fields. In addition to a general overview chapter on the Precambrian of Finland and an account of the history of Finnish bedrock research, the book contains twelve chapters on specific lithologic and crustal entities (the Archean in the eastern part of the country; Paleoproterozoic supracrustal belts, mafic and PRECAMBRIAN

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ultramafic intrusions, mafic dike swarms, ophiolites, and granitoid rocks; the rapakivi granites in their type terrain, and subsequent supracrustal successions and mafic magmatism; Neoproterozoic/Phanerozoic kimberlites, carbonatites, and alkaline rocks), as well as chapters on Paleoproterozoic tectonic evolution, carbon isotope stratigraphy, and the paleomagnetically defined drift history of the shield. The aim of the book is thus to provide the international geological community with an up-to-date account of the geologic framework and conceptual interpretation of the bedrock of Finland and to serve as a basis for future research. The book will also be a valuable reference for exploration activities, which at present are focused on gold, platinum-group metals, nickel, and diamonds in particular. This book would not have been possible without the contribution from the Geological Society of Finland (the society published a precursor to this book in Finnish in 19981), the commitment of the authors, and help from devoted reviewers (Andrey Bekker, Walter Boyd, Carl Ehlers, Sten-Åke Elming, Roland Gorbatschev, Eero Hanski, Yrjö Kähkönen, Jarmo Kohonen, Asko Kontinen, Raimo Lahtinen, Laura Lauri, Matti I. Lehtonen, Arto Luttinen, Hannu Makkonen, Satu Mertanen, Heikki Niini, Hugh O’Brien, Richard W. Ojakangas, Juhani Ojala, Heikki Papunen, Riku Raitala, Peter Sorjonen-Ward, Matti Vaasjoki, Pär Weihed, Alan Woolley). We would also like to thank Kent Condie, the Series Editor, for accepting this volume to be included in Elsevier’s Developments in Precambrian Geology Series, and Patricia Massar and Friso Veenstra for excellent collaboration in technical and administrative matters. Our special thanks go to Sakari Haapaniemi, who patiently manufactured the final electronic manuscript of the book in the course of an overly long and tedious editorial process.

Martti Lehtinen

Pekka A. Nurmi

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Lehtinen, M., Nurmi, P., Rämö, T. (Eds.), 1998. Suomen kallioperä–3000 vuosimiljoonaa. Geological Society of Finland, Helsinki. xiv

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Chapter 1

OVERVIEW

M. Vaasjoki, K. Korsman, T. Koistinen 1

Cover page: Paleoproterozoic migmatic and gneissic granodiorite containing gabbro fragments crosscut by tiny granite pegmatite dikes (in the background). Porkkalanniemi, Kirkkonummi, ~30 km west of Helsinki. Photo: Jari Väätäinen.

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Vaasjoki, M., Korsman, K., Koistinen, T., 2005. Overview. In: Lehtinen, M., Nurmi, P.A., Rämö, O.T. (Eds.), The Precambrian Geology of Finland – Key to the Evolution of the Fennoscandian Shield. Elsevier B.V., Amsterdam, pp. 1–18. © 2005 Elsevier B.V. All rights reserved.

The bedrock of Finland belongs to the Precambrian East European craton of northern and eastern Europe and northwestern Russia. Precambrian crystalline rocks crop out only in the northern and southwestern parts of the craton, in the Fennoscandian and Ukrainian shields, respectively; elsewhere they are covered by platform sediments. In Sweden and Norway, the Fennoscandian Shield is delimited by the Caledonides. In Estonia in the south and Russia in the southeast, the Precambrian bedrock plunges at a shallow angle under Phanerozoic sedimentary rocks. The most important events during the evolution of the Finnish bedrock occurred at 2800–2700 Ma and 1900–1800 Ma. In those times, continental crust was segregated from the Earth’s mantle in two major (probably multiphase) orogenies. The resultant Archean and Paleoproterozoic crust of Finland is divided into 25 areas with characteristic lithologic traits. This chapter gives an overview of Finland’s bedrock and its evolution from the Mesoarchean to the present time.

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1. Location, subdivision, timing, and general characteristics Finland forms about one third of the Fennoscan dian Shield which crops out among younger sedimentary rocks and the Caledonian mountain chain. It can be divided into four areas clearly deviating from each other: the Archean, the Svecofennian, and the Sveconorwegian domains, and the Transscandinavian igneous belt lying between the latter two (Figure 1.1). The northern and eastern parts of Finland belong to the >2.5 Ga Archean domain, divided usually into the Kola and Karelia blocks, while the central and southern parts comprise the Svecofennian Paleoproterozoic rocks, 1.93–1.80 Ga in age. Only a small part of the Finnish bedrock is younger than 1.8 Ga; the most significant of the younger formations are the 1.65–1.54 Ga rapakivi granites. After the intrusion of the rapakivi batholiths no major magmatism has occurred in Finland, but considerable graben formation took place during the Mesoproterozoic and at least southern Finland was covered by Paleozoic–Mesozoic sediments. The first isotope datings from Finland were carried by Olavi Kouvo during his stay in the United States in the mid-1950’s, and his doctoral thesis (1958) caused a fundamental change in the understanding of the Finnish Precambrian. It had been generally accepted that there were two great Precambrian orogenies in Finland: the older Svecofennian and the younger Karelian, but Kouvo’s results showed that the lithologic units associated with these orogenies were in fact coeval and that the granite-gneiss domain northeast of Karelides was much older than the southwestern part of the country. The existence of an ancient plate boundary along the Raahe–Ladoga zone became an accepted fact, not a mere working hypothesis, during the 1960’s (Simonen, 1971). The laboratory for isotope geology at the Geological Survey of Finland was established in 1964, and since then the amount of age de4

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Mesoproterozoic, Neoproterozoic, and Phanerozoic rocks Permo-Carboniferous igneous rocks including the Oslo rift Vendian to Cambrian and Devonian alkaline igneous rocks Caledonian orogenic belt Lower Paleozoic intrusive rocks Caledonian supracrustal rocks Fennoscandian Shield Mesoproterozoic to Paleoproterozoic rocks Supracrustal rocks, predominantly metasedimentary Sveconorwegian igneous and metamorphic rocks Rapakivi granites and coeval igneous rocks Paleoproterozoic rocks (1.96–1.75 Ga) Migmatizing granites TIB 1 and Revsund granites Granitoids and metavolcanic rocks Supracrustal rocks Paleoproterozoic rocks in the Lapland–Kola orogen Granulite, amphibolite, anorthosite Paleoproterozoic rocks (2.50–1.96 Ga) Intrusive rocks, mainly mafic and ultramafic Supracrustal rocks Archean rocks TTG-complex Greenstone belts

terminations and other isotope measurements has steadily increased. Figure 1.2 depicts the current data base for igneous rocks on chronograms, where the age results are plotted simply in an ascending order. On this kind of presentation, plateaus represent clusters in ages,

Kola Block

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E B Svecofennian TIB C

Sveconorwegian F

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W TIB

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Fig. 1.1. Simplified geological map of the Fennoscandian Shield after Koistinen et al. (2001). TIB denotes the Transscandinavian igneous belt. The subdivisions of the Svecofennian are: (A) The Primitive arc complex of central Finland; (B) The Accretionary arc complex of central and western Finland; (C) The Accretionary arc complex of southern Finland; (D) The Skellefte district; (E) The Bothnian basin; and (F) The Bergslagen district.

while gaps indicate times with no significant igneous activity. The data are mainly based on U-Pb zircon analyses, but include also baddeleyite and columbite U-Pb data as well as some Sm-Nd results. Details of the data compilation can be obtained on request from

the Geological Survey of Finland. The border zone between the Archean and Paleoproterozoic rocks is sharp and has been accurately delineated by geological, isotope geological, and geophysical methods. Archean rocks are found in northern and CHAPTER

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Fig. 1.2. Chronograms showing published U-Pb zircon and baddeleyite ages from igneous rocks in Finland (data compiled at the Geological Survey of Finland; details available from the Survey upon request). The results of these analyses are interpreted as indicating the times of intrusion or extrusion of the rocks.

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eastern Finland, whereas the bedrock of central and southern Finland consists of rocks of the Svecofennian. The latter are divided on the current 1:1,000,000 bedrock map [Korsman et al., 1997; based on the 1:400,000 (whole country) and the 1:100,000 mapping (~2/3 of the country) as well as abundant special studies] into the primitive, central Finland and southern Finland arc complexes. Paleoproterozoic metasedimentary and metavolcanic rocks cover large areas of the Archean domain, which is also penetrated by 2.5–2.0 Ga, mainly mafic igneous rocks emplaced while the Archean crust was rifted and eroded. There is no sign of a major inherited Archean component within the igneous rocks of the Svecofennian domain, which has led to the conclusion that the Svecofennian bedrock represents new continental crust segregated from the mantle (Huhma, 1986). The Lapland granulite belt in northern Finland is a geologically significant formation, which has been thrusted from lower continental crust into its present environment. In the early 1980’s evidence on plate tectonic activity in early Precambrian times was insufficient. When the almost completely preserved 1950 Ma ophiolite at Jormua in eastcentral Finland was discovered in the 1980’s, it constituted strong evidence for the operation of plate tectonic processes already in Paleoproterozoic times (Kontinen, 1987). Within the Svecofennian island arc systems an unusually large amount of granites formed and the upper parts of the crust reached a high temperature. This caused an intense metamorphism of the volcanic and sedimentary rocks. In its course, the rocks partly melted and migmatites were formed. Thus migmatites and granites are the most widespread rocks in southern Finland. According to J.J. Sederholm, about 53% of the Finnish bedrock are granites and about 22% migmatites. Mafic igneous rocks, schists, quartzites, and limestones form a relatively small fraction. Metavolcanic rocks are more frequent in Lapland than in southern

Finland. The Precambrian mountain chains of the Fennoscandian Shield have been leveled a long time ago and only ~3% of the bedrock is directly visible. Therefore, it has been difficult to delineate the continuity of rock formations and to obtain a three dimensional picture of the bedrock by geological methods alone. The mapping and study of the bedrock is assisted by high quality geophysical data (Figures 1.3 and 1.4) and has required close collaboration between geophysicist and bedrock geologists.

of both ages, and at lest some of these are Proterozoic and were deposited upon Archean crust. Gabbros and granodiorites of 1.95–1.93 Ga age are found as conformable bodies in the Proterozoic gneisses.

2. Regional geographic nomenclature

3. Enontekiö area. The northwestern part, divided by a broken line, is covered by Caledonian assemblages. The Archean rocks in the northwest are granitoid gneisses with small greenstone belts and ultramafic bodies. The Proterozoic rocks in the southeast are mafic and felsic volcanic rocks as well as arkosic rocks and quartzites that are crosscut by ~1.88 Ga monzonites and granodiorites.

As probably in most other countries, Finnish geological literature is plagued by a multitude of regional names, often used for overlapping areas and sometimes with conflicting meanings. In this volume an attempt has been made towards consistency in this respect, and it has been chosen to apply the terminology proposed by an ad hoc working group (Nironen et al., 2002; Figure 1.5). It should be emphasized, that the names are lithological-geographical and do not have a genetic connotation, hence rocks of similar age and origin may be found in several areas. The names were given according to the oldest rocks, generally supracrustal ones, in each area. ‘Belt’ defines an area with linear shape and internal structures, and ‘complex’ means a fault-bounded part of bedrock, or an igneous complex. The areas cover the Archean and Paleoproterozoic bedrock; Mesoproterozoic and younger lithologic units are separated by broken lines in Figure 1.5. A short description of each area is given below. 1. Inari area. The area consists of paraand orthogneisses that are Archean (2.7–2.6 Ga) in the east and Proterozoic in the west. Greenstone belts are found among gneisses

2. Lapland granulite belt. The rocks of the belt are felsic, generally intensely deformed garnet and pyroxene gneisses that have been metamorphosed at granulite facies. The gneisses are migmatitic especially in the center of the belt. Mafic, pyroxene-bearing 1.93–1.91 Ga igneous rocks of are found as elongate bodies among the gneisses.

4. Central Lapland area. In the northeastern part of the belt there are felsic gneisses and amphibolites that are considered Archean. Moreover, Archean (3.1–2.7 Ga) gneisses are found as tectonic windows among the Proterozoic assemblages. In the eastern part, there are mafic–ultramafic layered intrusions with an age range of 2.44–2.05 Ga. Most of the Proterozoic supracrustal rocks were deposited upon Archean crust. Lowermost in the sequence are mafic volcanic rocks, overlain by arkosic rocks and mica schists. Two groups of mafic volcanic rocks, with an age range of 2.1–2.0 Ga, constitute the large greenstone belt in the western part of the belt: the first were erupted in a rift zone and the second upon oceanic crust. These rocks are crosscut by ~1.88 Ga monzonites and granodiorites. Quartz arenites and conglomerates were deposited after 1.88 Ga in the southern part of the belt. CHAPTER

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Fig. 1.3. Generalized aeromagnetic map of Finland after Ruotoistenmäki (1992).

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Fig. 1.4. Generalized gravity anomaly map of Finland after Elo (1992).

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Fig. 1.5. The geographic distribution of various geological regions of Finland according to Nironen et al. (2002). Note that the divisions have been arrived at on lithological and geographic grounds only and bear no genetic connotations.

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5. Eastern Lapland complex. The Archean complex mainly consists of 2.8–2.7 Ga tonalitic gneisses. In addition to these gneisses there is a belt of gneissic sedimentary rocks and several greenstone belts, consisting of ultramafic and mafic volcanic rocks as well as sedimentary rocks. Archean granitoid intrusions crosscut the gneisses.

amphibolitic migmatites metamorphosed at high grade in large areas. The complex also contains Archean paragneisses and an Archean carbonatite complex. Proterozoic granites and diabase dikes have intruded the gneisses, and Proterozoic deformation and alteration have locally strongly overprinted the gneisses.

6. Central Lapland granitoid complex. This poorly studied complex mainly consists of 1.8 Ga granites that migmatize and crosscut mica schists and arkosic gneisses. There are also Proterozoic mafic plutonic rocks and remnants of Archean gneisses within the complex.

11. Eastern Finland complex. This large complex mainly consists of 2.85–2.69 Ga granitoids and migmatites. In addition, there are paragneiss-dominated areas as well as several greenstone belts. Proterozoic granites and diabase dikes have intruded the gneisses, and Proterozoic deformation and alteration have locally caused strong overprinting especially in the western part of the complex.

7. Peräpohja belt. The rocks of this belt were deposited and extruded upon Archean crust. There is a swarm of 2.44 Ga mafic layered intrusions along the southern boundary. The rest of the belt consists of mica schists and quartzites with dolomites, metaconglomerates, black schists, and mafic volcanic rocks as interlayers. These rocks are crosscut by ~1.88 Ga monzonites.

12. Kuhmo belt. The greenstone belt consists mainly of volcanic rocks. The marginal parts consist of 2.97 Ga mafic and intermediate volcanic rocks, and 2.79 Ga mafic lavas with ultramafic parts and iron-formations as well as interlayers of mica schist are found in the central parts.

8. Kuusamo belt. The central part of the belt is occupied by 2.44 Ga intermediate and felsic volcanic rocks, followed by mafic and ultramafic volcanic rocks. The mafic rocks in the southern part were deposited upon Archean crust. They contain sericite and mica schist as well as carbonate rocks as interlayers, and on top of the strata there are quartzites as a thick pile. 9. Pudasjärvi complex. This poorly known complex consists of Archean gneisses and granitoids as well as amphibolites that are presumably remnants of Archean greenstone belts. Proterozoic granites and diabase dikes have intruded the gneisses. 10. Iisalmi complex. The complex consists of 3.2–2.6 Ga tonalitic gneisses and

13. Ilomantsi belt. The greenstone belt is part of a larger belt that extends to Russia. The predominant and oldest rocks are 2.75–2.70 Ga old and of sedimentary origin. Iron-formations are found higher in the sequence, and mafic lavas are the youngest rocks of the belt. 14. Kainuu belt. The eastern part of the belt mainly consists of autochthonous mafic volcanic rocks and conglomerates overlain by quartzites. The latter are unconformably overlain by mica schists with metaconglomerates, iron-formations, and black schists as interlayers. Highest in the strata are homogeneous mica schists. Part of the mica schists as well as the 1.95 Ga Jormua ophiolite complex are allochthonous. 15. Kiiminki belt. The metasediment-domCHAPTER

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inated belt contains conglomerates and arkosic rocks lowermost in the sequence. These are followed by a thick pile of turbiditic graywackes, and on top there are mafic volcanic rocks with quartzites, black schists, dolomite rocks, and iron-formations as interlayers. 16. Savo belt. The belt is characterized by numerous shear zones. The predominant rocks are mica gneisses, which contain volcanic rocks, graphite schists, black schists, and carbonate rocks as interlayers. The volcanic rocks in the center of the belt consist of two groups: a 1.92 Ga bimodal group, and a 1.89–1.88 Ga mafic–intermediate group. 1.92 Ga gneissic tonalites and 1.89–1.88 Ga granitoids are also found within this belt. 17. Höytiäinen belt. The northeastern part of the belt consists of autochthonous or parautochthonous conglomerates, arkosic rocks, and quartzites. The main part is dominated by turbiditic mica schists with some interlayers of conglomerates and mafic volcanic rocks. 18. Outokumpu area. The predominant rocks are homogeneous, turbiditic mica schists that contain interlayers of black schists. The rocks are migmatitic mica gneisses in the southwestern part of the area. The 1.97 Ga Outokumpu association, consisting of lensoid serpentinite bodies, carbonates, skarns, and sulfide mineralization, is in the center of the area. The whole-rock sequence is allochthonous. 19. Saimaa area. The predominant rocks in the area are turbiditic mica schists that grade into migmatitic mica gneisses and garnet-cordierite gneisses toward south. Mafic volcanic rocks are found mainly in the northern part of the area. Crosscutting 1.89–1.88 Ga granitoids are found throughout the area. Moreover, 1.84–1.81 Ga granites migmatize and crosscut the supracrustal rocks in the southern part.

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20. Central Finland granitoid complex. The complex consists of 1.89–1.88 Ga synkinematic tonalites, granodiorites, and granites, and 1.88–1.86 Ga postkinematic quartz monzonites and granites. In addition, there are minor areas of subvolcanic intermediate rocks, mafic igneous rocks, and remnants of supracrustal belts. 21. Pohjanmaa belt. The predominant rocks are turbiditic mica schists and gneisses, with mafic and intermediate volcanic rocks, black schists, metacherts, and carbonate rocks as interlayers. The conglomerates and arkosic rocks in the northern part represent the youngest sedimentation in the belt. Metamorphic grade increases in the center of the belt toward granulite facies. Granitoids of 1.88 Ga age crosscut the supracrustal rocks. 22. Tampere belt. The belt consists of 1.90–1.88 Ga intermediate and felsic volcanic rocks as well as turbiditic mica schists with conglomerate interlayers. Mafic volcanic rocks are found lowest and highest in the sequence. Granitoids of 1.88 Ga age crosscut the supracrustal rocks. 23. Pirkanmaa belt. The belt mainly consists of migmatitic, turbiditic mica gneisses with black schists and graphite-bearing schists as interlayers. Mafic and ultramafic plutonic rocks as well as 1.88 Ga granitoids crosscut the supracrustal rocks. 24. Häme belt. The belt is characterized by volcanic rocks which may be grouped into older, of intermediate and younger, of mafic–intermediate composition. The western part of the belt is dominated by metasedimentary rocks. 1.88 Ga granitoids of as well as 1.84–1.82 Ga granites crosscut and migmatize the supracrustal rocks. 25. Uusimaa belt. This sedimentarydominated belt contains mica schists and

gneisses with relatively common carbonate rock inter layers. Also felsic sedimentary rocks of volcanic provenance are typical of the belt. The volcanic rocks are generally mafic–intermediate in composition, but in the western part of the belt volcanism was bimodal. Granitoids of 1.88 Ga age as well as 1.84–1.82 Ga granites crosscut and migmatize the supracrustal rocks.

3. The Archean bedrock The oldest rocks in Finland lie within the Archean domain in the eastern and northern parts of the country, and several occurrences of rocks older than 3 Ga are known. However, they are all of local nature and lie widely dispersed from each other with emplacement ages ranging from 3.1 to 3.5 Ga (Figure 1.2). The oldest known rock is trondhjemite gneiss found at Siurua, where ionprobe results from zircons, supported by conventional zircon data and Sm-Nd whole-rock data, indicate an intrusion age of ~3.5 Ga (Mutanen and Huhma, 2003). There are, however, indirect Sm-Nd and common lead indications suggesting that the 3.5 Ga crust in Finland may have been more wide-spread. Greenstone belts formed by volcanic and sedimentary rocks are characteristic of all Archean terranes of the world. The mainly 2.8 Ga old greenstone belts especially in eastern Finland have been compressed into narrow sequences between Archean granitoid rocks, which are mainly ~2.7 Ga granodiorites and gneissose tonalites. This period of evolution is well evident in the isotope ages (Figure 1.2), although ion microprobe data suggest that some rocks both in the Suomussalmi and Ilomantsi areas contain also inherited zircons older than 3 Ga. A peculiarity of the Finnish Archean is the 2610 Ma carbonatite at Siilinjärvi, one of the oldest of its kind in the world.

4. Faulting of Archean crust and emplacement of Paleoproterozoic cover rocks When the Archean orogenic movements ceased, there commenced a period of peneplanation, which lasted for several hundred million years. However, crustal scale faulting with associated volcanic activity and formation of sedimentary basins occurred within the eroding and peneplaning Archean crust. A characteristic feature are numerous 2.44 Ga layered mafic intrusions in northern Finland and northwestern Russia. The faulting started to ease up about 2.4 Ga ago. At this time, weathering was well-advanced and the Archean bedrock was in many places covered by quartz sands, which later formed the so-called Jatulian quartzites. Volcanic activity occurred also during the Jatulian period, and is manifested as mafic lava flows and numerous diabase dikes that penetrated the Archean and its cover rocks 2.2–1.97 Ga ago. The cratonization of the Archean bedrock over a period of 500 Ma is especially diversely observable in Lapland. Fundamental atmospheric changes occurred at the same time as the rifting phase of the Archean continent ended. For the evolution of life most important was the increase of the oxygen contents of the atmosphere almost to its present level about 2.1 Ga ago. This information, relevant to the evolution of the entire Earth, has been obtained by careful stratigraphic and isotope geological studies of the Finnish Karelian formations (Karhu, 1993).

5. The Svecofennian bedrock The Jormua ophiolite demonstrates that oceanic mantle had formed and plate tectonics operated at least 1950 Ma ago, but, according to some interpretations, some kind of primitive Svecofennian continent may have CHAPTER

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formed already 2.1 Ga ago. However, so far no continental crust of that age has been found within the Fennoscandian Shield. The only indications are the zircon age distribution of younger metasedimentary rocks, Sm-Nd model ages, and some geochemical features suggesting that Svecofennian granites may have resulted from remelting of older crust, perhaps 2.1 Ga in age. The oldest Svecofennian volcanic rocks of primitive island arc type and associated gneisses are 1930–1920 Ma old and occur along the Archean–Proterozoic boundary in central Finland. Observations from the Lapland granulite complex indicate, however, that subduction was already occurring in that area, as the ocean in the (present) north had already closed and the granulites were being thrust from lower crustal levels into their present geological environment. This belt, called the Lapland–Kola orogen, formed more or less simultaneously with the Svecofennian orogeny, and extends from the granulite belt in Finland to the southern part of the Kola Peninsula. Evolved island arc volcanic rocks and associated metasediments in central and southern Finland are 1910–1890 Ma old. A particularly well-known volcano-sedimentary entity is the Tampere schist belt, where systematic studies have been carried out for over 100 years. Primary structures of the volcanic and sedimentary rocks have been preserved at many locations within the belt, facilitating conclusions on the origin of rock formations. The Svecofennian crust is exceptionally thick, up to 65 km in the Paleoproterozoic–Archean boundary zone. The crust was thickened first during the collision when the newly created crustal plates were thrust upon each other. There is little reliable information on the incipient part of the collision and its beginning can be timed only indirectly at about 1910–1900 Ma. It had concluded 1870 Ma ago, because at that time the Svecofennian bedrock was already attached to the Archean 14

• C H A P T E R 1 • OV E RV I E W

continent. During the collision and the ensuing tectonic thickening, molten rock material was injected into the collision zone from the underlying mantle. The mantle-derived magma caused melting of the lower crust, which lead to the intrusion of magmas close to the then existing erosional level. Thus the temperature even in the upper parts of the crust was raised, leading to recrystallization and partial melting of rocks. The metamorphism and the magmatism generated from the lower crust are coeval at ~1885 Ma in the collision zone between the Archean and Svecofennian domains. After this strong pulse of magmatism and recrystallization, cooling commenced within the collision zone. The collision of the Svecofennian island arc complex also affected the cratonized Archean continent. Easily observed evidence about the reactivation of the Archean continental crust during the Svecofennian orogeny are found up to 150 km from the collision zone: 1.9–1.86 Ga rocks with Archean Nd isotope signature, titanite and monazite U-Pb ages in the 1.9–1.8 Ga range, and reset biotite K-Ar ages in Archean granitoids. The migmatite-forming lateorogenic microcline granites in southern Finland form large, sheet-like bodies with usually diffuse contacts. They are about 1.83 Ga old, and their emplacement was associated either with the extensional collapse of Svecofennian orogen or transpressional faulting. In any case, the migmatization of the Svecofennian bedrock in southern Finland is best regarded as a quite separate event from the main phase of the Svecofennian orogeny. A special feature of the Svecofennian is also the survival of the 65 km thick crust, as the usual thickness of continental crust is about 40 km. Crust thickened during a collision of continents is in a disequilibrium. The light crust returns to equilibrium either by uplift or collapse, as is the case in the Phanerozoic mountain chains. There are signs of an incipient collapse within the Svecofen-

nian, but the process was left incomplete, as the light crust thickened by the collision was quickly stabilized by magmatism originated in the mantle. Due to this unusually quick isostatic equilibration the thick crust became permanent. It is still thick, although erosion has removed the top 15 km! The orogenic movements waned in southern Finland about 1.8 Ga ago. As the bedrock cooled, fissures opened and made way for deep-seated magmas, which crystallized in the upper crust as the so-called postorogenic (1.81–1.77 Ga) granites.

6. Rapakivi magmatism and the Jotnian period A period of 150 Ma of geological quiescence followed after the emplacement of the postorogenic granites. There are very few signs of strong bedrock movements from this time, which indicates that the crust was being peneplaned through erosion. The quiescence terminated when the rapakivi granites intruded into the rigid bedrock 1650–1540 Ma ago. More than ten rapakivi intrusions, often with associated gabbroic and anorthositic rocks, are known in southern Finland. The largest are the Wiborg, Åland, Laitila, and Vehmaa batholiths. Coeval with the rapakivi granites are tholeiitic (Subjotnian) diabase dikes. Rapakivi granites are not limited to the Finnish bedrock. They are found in all Precambrian shield areas, but the origin of the rapakivi magmas as remelted lower continental crust has been successfully explained in Finland (Rämö, 1991). According to the prevailing view, the formation of rapakivi granites was not a direct consequence of the Svecofennian orogeny. Some scientists have, however, considered the formation of rapakivi granites to reflect the last phase of the stabilization of the Svecofennian crust. Rapakivi granites intruded, at least par-

tially, into a bedrock on which the so-called Jotnian sediments had started to deposit in topographic shallows. The deepening of basins and sedimentation continued still long after the rapakivi magmatism. The Jotnian sandand claystones are preserved on the continent at Muhos and Satakunta, and the Satakunta sandstones continue into the Gulf of Bothnia covering large submerged areas. The Jotnian sedimentary rocks are cut by 1.26 Ga tholeiitic (Postjotnian) diabase dikes and sills. However, a recent result from the Valamo (Valaam) sill in the Ladoga basin, 1.46 Ga, suggests that this continental sedimentation at least in that area was well advanced much earlier on than belived so far. In Lapland, there are young dike rocks in local rifts: 1100 Ma at Salla and 1000 Ma at Laanila. These represent the youngest parts of the Finnish bedrock, because only rocks which were deposited or crystallized before the Vendian period (>650 Ma ago) are considered bedrock.

7. The Vendian period and the Paleozoic era At the beginning of the Vendian period (~650 Ma ago) the Finnish bedrock had been eroded almost to its present level. Shallow-water sandstones were deposited on the continental peneplane. Cambrian sandstone is found in fissures in the southwest Finnish archipelago, at Lauhanvuori in northern Satakunta, and at Sulva (Söderfjärden) south of Vaasa. At Lumparn in the Åland Islands Ordovician limestones are known. At Muhos, the sedimentation, which had started in Jotnian times, lasted into the beginning of the Vendian period. Alkaline igneous rocks (e.g., kimberlites) were emplaced in eastern Finland at ~600 Ma. The Paleozoic sediments deposited west of Fennoscandia were folded against the craton 450–400 Ma ago. An overthrusted nappe of the Caledonides has been found in Finland only in the far northwestern part of the country. CHAPTER

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Other effects of the Caledonian orogeny on the Finnish bedrock are not well known. The 370–360 Ma alkaline intrusions at Iivaara and Sokli may have a causal relationship to the Caledonian orogeny, and faulting is likely to have occurred in the foreland of the Caledonides, i.e., in Finland.

8. Late events affecting the bedrock Although movements strongly affecting the bedrock waned decisively already ~1.8 Ga ago, many shear zones remained active for hundreds of millions of years after the Svecofennian orogeny. Some of them are weakly active even today, although the amount of movement is relatively small. The Svecofennian metasedimentary and metavolcanic rocks were deposited 1890 Ma ago, but subsided within a few million years to a depth of about 20 km within the crust, which demonstrates the rapidity of changes during ancient plate collisions. The present erosional level lay at a depth of 15 km even 1.8 Ga ago. The denudation which brought the Svecofennian metavolcanic and metasedimentary rocks back to surface lasted at least 200 Ma, as the intrusion of the rapakivi granites into the upper crust occurred at a depth of ~5 km. The present erosional level had been definitely reached at the onset of the Cambrian period about 600 Ma ago, as is demonstrated by the deposition of Cambrian sandstones and their preservation in bedrock cracks. The Pleistocene continental glaciation eroded the bedrock mainly by polishing the weathering surfaces and sharpening the shear zones. Preglacial weathering surfaces formed before the glaciation have survived in a few places only, most notably in Lapland. The shallow Finnish lakes are found mainly in shear zones dredged deeper by the continental ice sheet. The widening of the Atlantic Ocean and 16

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the postglacial isostatic uplift result in tensions within the bedrock which trigger earthquakes. The tremors are, however, so mild that they damage buildings or cause any alarm only in exceptional circumstances. Generally, recognizable traces of asteroids have survived only locally. There are at least ten positively identified impact craters in Finland, of which Lappajärvi (impact at 75 Ma), Söderfjärden (~530–510 Ma), Sääksjärvi (~515 Ma), Lumparn, Karikkoselkä, Suvasvesi, and Paasselkä are the most widely known (e.g., Lehtinen, 1976; Pesonen et al., 2000). The main features of the Finnish bedrock are ancient. As in many other Precambrian shield areas (e.g., Canada, Greenland, China) they were formed principally during late Archean and early Proterozoic times. Thus detailed results from the Fennoscandian Shield often have also a global bearing, which is one the reasons for the compilation of the present volume.

References Elo, S., 1992. Painovoima-anomaliakartat - Gravity anomaly maps. In: T. Koljonen (Ed.), Suomen geokemian atlas. Osa 2: Moreeni – The Geochemical Atlas of Finland. Part 2: Till. Geol. Surv. Finland, Espoo. 70–75. Huhma, H., 1986. Sm-Nd, U-Pb and Pb-Pb isotopic evidence for the origin of the early Proterozoic Svecokarelian crust in Finland. Geol. Surv. Finland, Bull. 337, 1–48. Karhu, J.A., 1993. Paleoproterozoic evolution of the carbon isotope ratios of sedimentary carbonates in the Fennoscandian Shield. Geol. Surv. Finland, Bull. 371, 1–87. Koistinen, T., Stephens, M.B., Bogatchev, V., Nordgulen, Ø., Wennerström, M., Korhonen, J. (Comps.), 2001. Geological map of the Fennoscandian Shield 1:2 000 000. Espoo : Trondheim : Uppsala : Moscow; Geol. Surv. Finland : Geol. Surv. Norway : Geol. Surv. Sweden : Min. Nat. Res. Russia. Kontinen, A., 1987. An early Proterozoic ophiolite – the Jormua mafic-ultramafic complex,

northern Finland. Precambrian Res. 35, 313–341. Korsman, K., Koistinen, T., Kohonen, J., Wennerström, M., Ekdahl, E., Honkamo, M., Idman, H., Pekkala, Y. (Eds.), 1997. Suomen kallioperäkartta - Berggrundskarta över Finland - Bedrock map of Finland 1:1 000 000. Geol. Surv. Finland, Espoo. Kouvo, O., 1958. Radioactive age of some Finnish Precambrian minerals. Bull. Comm. géol. Finlande 182, 1–70. Lehtinen, M., 1976. Lake Lappajärvi, a meteorite impact site in western Finland. Geol. Surv. Finland, Bull. 282, 1–92. Mutanen, T,. Huhma, H., 2003. The 3,5 Ga Siurua trondhjemite gneiss in the Archaean Pudasjärvi Granulite Belt, northern Finland. Bull. Geol. Soc. Finland 75, 51–68 Nironen, M., Lahtinen, R., Koistinen, T., 2002. Suomen geologiset aluenimet – yhtenäisempään nimikäytäntöön! Summary: Subdivision of Finnish bedrock – an attempt to harmonize terminology. Geologi 54 (1), 8–14. .Pesonen, L.J., Abels, A., Lehtinen, M., Plado, J.,

2000. Meteorite impact structures in Fennoscandia – a new look at the database. In: J. Plado, L.J. Pesonen (Eds.), Meteorite Impacts in Precambrian Shields. Programme and Abstracts, the 4th Workshop of the European Science Foundation Impact Programme, Lappajärvi - Karikkoselkä - Sääksjärvi, Finland, May 24-28, 2000. Geol. Surv. Finland and University of Helsinki. 20 p. Rämö, O.T., 1991. Petrogenesis of the Proterozoic rapakivi granites and related basic rocks of southeastern Fennoscandia: Nd and Pb isotopic and general geochemical constraints. Geol. Surv. Finland, Bull. 355, 1–161. Ruotoistenmäki, T., 1992. Magneettiset anomaliakartat - Magnetic anomaly maps. In: T. Koljonen (Ed.), Suomen geokemian atlas. Osa 2: Moreeni - The Geochemical Atlas of Finland. Part 2: Till. Geol. Surv. Finland, Espoo. 76–79. Simonen, A., 1971. Das finnische Grundgebirge. Geol. Rundschau 60 (4), 1406–1421.

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Chapter 2

ARCHEAN ROCKS

P. Sorjonen-Ward, E.J. Luukkonen

Cover page: Archean banded iron-formation. Ukkolanvaara, Ilomantsi. Photo: Peter Sorjonen-Ward.

Sorjonen-Ward, P., Luukkonen, E.J, 2005. Archean rocks.In: Lehtinen, M., Nurmi, P.A., Rämö, O.T. (Eds.), Precambrian Geology of Finland – Key to the Evolution of the Fennoscandian Shield. Elsevier B.V., Amsterdam, pp. 19–99. © 2005 Elsevier B.V. All rights reserved.

There have been few attempts in recent years to synthesize the nature and evolution of the Archean geological record in Finland. Therefore, the main purpose of this review is to describe the principal features of the Archean bedrock in Finland as currently known, primarily in terms of lithological units and structures. Through comparisons with the Proterozoic record of Finland, we then briefly consider whether the Archean bedrock of Finland reflects a distinctive style of crustal evolution, related to secular variations in thermal regime and rates of crustal growth and recycling. We are therefore also concerned with attempting to discriminate between processes relating to crustal formation and those that rework existing crust. For example, is the evolution of high-grade terrains in the deep crust level necessarily coeval with and complementary to lower grade supracrustal units, as for example in paired metamorphic belts in modern convergent accretionary settings? Alternatively, does the pattern of metamorphic grade represent a direct consequence of vertical crustal differentiation related to thermal and gravitational instability? Does crustal zonation with depth differ from that of younger continental crust and to what extent has the existence of Archean lithosphere predetermined subsequent crustal development? Although this review commences with brief descriptions of each of the various Archean rock units currently recognized, including a discussion of age relationships and possible correlations, we concentrate on those areas that are best known and which have begun to yield useful insights into Archean crustal processes. We conclude with a discussion of Archean thermal regimes and their tectonic consequences, the stabilization of the shield, and some regional scenarios and correlations, including a comparison between Archean and Paleoproterozoic crustal proceses in the Fennoscandian Shield.

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1. Introduction to the Archean of Finland 1.1. The extent of the Archean in Finland Although the distribution and nature of Archean rock types in Finland has been relatively well defined from regional reconnaissance scale mapping, a systematic framework for understanding Archean crustal evolution has yet to emerge. Indeed, in some cases there is still uncertainty over the age affinities of rock units. This applies particularly to extensive tracts of migmatitic gneisses intruded by Svecofennian potassic granite neosomes in the northern part of the country (Vaasjoki et al., 2001), as well as some metasedimentary complexes that contain exclusively Archean detrital zircons, but otherwise show evidence for reworking or partial melting during the Svecofennian orogeny (Huhma et al., 2000). Detailed studies addressing generic issues of crustal evolution are few and restricted largely to lower grade supracrustal greenstone belts which, by analogy with similar terrains elsewhere, are considered prospective for komatiite-hosted nickel and orogenic lode gold deposits. For example, a comprehensive commodity database for gold in Finland, prepared by Eilu (1999) includes attribute information for all known Archean occurrences and their geological context. In recent years attempts have also been made to understand the composition, thermal structure and evolution of the deeper crust and mantle lithosphere through seismic and other geophysical techniques and by studying exposed higher grade terrains (Hölttä, 1997; Hölttä and Paavola, 2000; Hölttä et al., 2000a,b) as well as xenolith suites sampled by Paleozoic kimberlites (Kukkonen and Peltonen, 1999; Hölttä et al., 2001). It is convenient, as first suggested by Gaál and Gorbatschev (1987), to consider the Archean and Paleoproterozoic history of the Fennoscandian Shield in terms of three 22

large crustal domains – the Kola, Karelian, and Svecofennian domains (Figure 2.1A). These three crustal units have shared a common history since amalgamation at about 1.8 Ga. The Karelian domain is the largest unit, forming a coherent late Archean (3.2–2.7 Ga) cratonic nucleus exceeding 200 000 km2 in area in eastern Finland and adjacent Russia (Figure 2.1B and 2.2). The Karelian domain is flanked to the northeast by the Kola domain, which represents a complex tectonic collage of Archean and early Proterozoic terranes, and to the southwest by the essentially Paleoproterozoic Svecofennian domain (Figures 2.1A and B). The Karelian domain is characterized by a number of narrow northerly trending low-pressure greenstone and metasedimentary belts (Figures 2.1B and 2.2), intruded by discrete plutons of dominantly granodioritic to monzogranitic compositions. Higher grade mediumpressure metasedimentary gneiss complexes are also present, some of which represent older relict enclaves with younger migmatites, while others appear to be coeval with the greenstone sequences. The Archean of the Kola domain includes granitoid gneisses, migmatites, char nockites, aluminous metasedimentary rocks, and iron-formations (Meriläinen, 1976; Gaál et al., 1989; Rundquist and Mitrofanov, 1993), and also a distinctive suite of alkaline intrusions and gabbro–anorthosite intrusions (Zozulya et al., 2001). The nature and age of the boundary zone between the Kola and Karelian domains in Russia has long been contentious, largely due to the presence of both Archean and Proterozoic isotope ages from medium- to high-pressure gneisses of the intervening Belomorian terrain (Figure 2.1A) (named from the Russian term for the White Sea). Intense deformation and medium-pressure metamorphism in unequivocally Proterozoic rocks, and widespread thermal resetting of U-Pb isotopes in titanites, demonstrate significant tectonic and thermal reworking of the Belomorian terrain between 1.9–1.8 Ga, which is attributed

• CHAPTER 2 • ARCHEAN ROCKS

0

100 km

en

fot

200

Post-Archean rocks Phanerozoic sedimentary rocks Proterozoic and Caledonian orogenic domains

Bar ent

Lo

s Se a Murmansk

AY RW

O

N

Archean crustal domains Kola domain

RUSSIA SWEDEN

Karelian domain

Rovaniemi

Belomorian terrain Exposed Archean rocks in Sweden and Norway

Luleå

Boundary zone between Kola and Karelian domains Boundary between Karelian and Svecofennian domains Limit of isotopically defined Archean crust in Sweden

ia

n th

ulf

of

Bo

Kuopio

G

Petrozavodsk

FINLAND

Figure 2.12

B 32° E 66° N

Helsinki

A

Kemi

Figure 2.10

Oulu

Post-Archean rocks

Figure 2.9

Svecofennian orogenic domain Kajaani

Paleoproterozoic sequences within Karelian domain

Kuhmo Figure 2.5

Archean Karelian domain Granitoids, migmatites, and high-grade gneisses

Iisalmi

Lower grade supracrustal belts

Kuopio

Belomorian terrain Significant Svecofennian tectonic overprint

Joensuu 0

50 km

32° E 62° N

100

Fig. 2.1. Regional distribution of Archean rocks in the Fennoscandian Shield. (A) Principal crustal domains. (B) Distribution of greenstone belts and granitoid terrains within the Karelian domain in eastern Finland and adjacent Russia, showing locations of more detailed regional scale maps. CHAPTER

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Kuusamo

Oi

NORWAY

Koillismaa

Kemi

Sørvaranger Inari

Ranua Pudasjärvi

Inari Ivalo

Suo

RUSSIA

Ämmänsaari

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RUSSIA

Pomokaira Muonio

Kianta

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Kuopio 100 km

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Suomu Rovaniemi

Siilinjärvi

0

Kaavi

Kuusamo

Ranua Hat

Nun Kov Joensuu

Fig. 2.3. Archean terrains in northern Finland, as defined and described in this review.

Ilomantsi

Fig. 2. 2. Supracrustal greenstone belts and respective terrains within the Karelian domain in eastern Finland, as defined and described in this review. Oi–Oijärvi, Suo–Suomussalmi, Kuh–Kuhmo, Tip–Tipasjärvi, Nun–Nunnanlahti, Hat–Hattu, Kov–Kovero.

to collision between the Karelian and Kola domains (Bibikova et al., 1996, 2001; Daly et al., 2001). Nevertheless, there is considerable evidence accumulating to support the initial juxtaposition of the granitoid-greenstone terrains of the Karelian domain and high-pressure assemblages of the Belomorian terrain during the late Archean (Samsonov et al., 2001; Slabunov and Bibikova, 2001). The Belomorian terrain appears to be contiguous with the high-grade supracrustal gneisses of the Tuntsa terrain in northern Finland (Figures 2.1A and 2.3). Although this area has also been strongly affected by Paleoproterozoic deformation, it is evident that boundaries between Archean crustal units are discordant to Proterozoic trends, which lends further support to interpreting the Belomorian terrain as a higher grade unit 24

Savukoski

within the Karelian domain. On this basis, the granitoid gneisses exposed as basement windows beneath the Paleoproterozoic Lapland greenstone belt, assigned here to the Pomokaira terrain (Figure 2.3), also belong to the Karelian domain. The Kola domain is thus only represented in Finland by the Inari and Sørvaranger terrains, in northeast Lapland (Figures 2.1A, 2.3, and 2.16). These are separated from the Karelian domain by the Paleoproterozoic Lapland granulite belt (Figures 2.1A and 2.16), which has been thrust southwards along a gently dipping detachment surface that can be traced seismically to middle crustal depths (Gaál et al., 1989; Korja et al., 1989; Luosto et al., 1989). The Ropi terrain in the northwestern part of Finland (Figures 2.1A and 2.3) forms part of an extensive region in northern Sweden and Norway that is at least partly underlain by Archean rocks (Skiöld and Öhlander, 1989; Öhlander et al., 1993; Martinson et al., 1999). The Ropi terrain is separated from the western part of the Karelian domain by a highly strained zone of high-temperature–lowpressure metamorphism and abundant Svecofennian granitoids of both calc-alkaline and

• CHAPTER 2 • ARCHEAN ROCKS

potassic post-collisional affinity. This suggests that the current juxtaposition of these two Archean crustal units was a consequence of Svecofennian collisional tectonics. Tectonically and thermally reworked Archean rocks are also found farther south, in deep crustal windows and demonstrably allochthonous thrust sheets, along the western margin of the Karelian domain, recording deep crustal imbrication during the Svecofennian orogeny (Park and Bowes, 1983; Korsman et al., 1999). Although the Svecofennian event does not appear to have exposed deep Archean crustal sections, the remarkable possibility exists that the serpentinized harzburgites of the 1.95 Ga Jormua ophiolite complex represent Archean subcontinental lithospheric mantle exhumed during Paleoproterozoic extension and rifting along the southwest margin of the Karelian domain (see Kontinen, 1987; Peltonen et al., 1998; Chapter 6). The discovery of Archean zircons from dikes intruding these harzburgites (Peltonen et al., 2003) now provides an unparalleled opportunity for attempting to correlate late Archean deep lithospheric events with the magmatic record preserved in the Archean mafic and ultramafic greenstone sequences of eastern Finland. There is also considerable evidence for a Proterozoic thermal overprint across much of the Karelian domain itself (Kontinen et al., 1992; Kontinen and Paavola, 1996; Bibikova et al., 2001; Pajunen and Poutiainen, 1999). The boundary zone between the Karelian and Svecofennian domains in Finland is nevertheless considered to represent the true edge of the Archean crust of the Karelian craton – or at least coincides with one rifted margin of a formerly greater crustal block, as U-Pb, SmNd, and Hf-Lu isotope studies of crustally derived granites in the Svecofennian province tend to indicate early Proterozoic source ages (Patchett et al., 1981; Patchett and Kouvo, 1986; Huhma,1986). This Proterozoic tectonic and thermal reworking has complicated our understanding

of relationships between Archean structures and lithic units, caused confusion in dating events isotopically (Martin and Barbey, 1988; Vaasjoki, 1989; Tourpin et al., 1991; Gruau et al., 1992), and frustrated the characterization of hydrothermal fluids in Archean gold deposits (O’Brien et al., 1993a,b). On the other hand, the characteristics and distribution of Proterozoic magmatism and sedimentation, and deep xenoliths sampled by early Paleozoic kimberlites provide valuable insights into the stability of Archean lithosphere and the extent to which the deep crust might have been modified over time. For example, isotope and petrological investigations of lower crustal and mantle xenoliths from kimberlites penetrating both the Karelian and Kola provinces reveal a population of zircons that apparently record a thermal event around 1.8 Ga, immediately after the initial stages of Svecofennian collision (Hölttä et al., 2000a,b; Markwick and Downes, 2000). The Karelian domain was also widely and repeatedly affected by mafic magmatism throughout the early Proterozoic, as recorded by several sets of dike swarms, lava fields and layered intrusions. It is therefore likely that the Archean lower crustal and mantle lithosphere of Finland and indeed the Fennoscandian Shield in general has been at least partially modified by the addition of underplated magmas, cumulates or residual assemblages (Kempton et al., 2001). Hence, deep crustal and mantle characteristics inferred from deep seismic reflection and refraction and magnetotelluric surveys may not be representative of the state of the lithosphere during the Archean. Nevertheless, careful integration of these sources of data with surface observations will no doubt eventually lead to a four dimensional picture of the evolution of the cratonic lithosphere of Finland, even though at present this picture is fragmentary and far from clear.

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1.2. Classifying and subdividing the Archean bedrock of Finland Prior to the application of isotope dating techniques to the Precambrian of Finland (Kouvo and Tilton, 1966), uniformitarian principles had long been used in recognizing distinct and superimposed orogenic events (Sederholm, 1897). Following this example, Frosterus and Wilkman (1924) mapped a widespread unconformity separating a Proterozoic sedimentary cover sequence from a predominantly granitic and gneissose basement in eastern Finland. Within this basement terrain Frosterus and Wilkman (1920) further recognized that granites intruded enclaves of still older, variably metamorphosed supracrustal rock units, and so inferred another yet older orogenic cycle. Moreover, by identifying allochthonous and inverted basement-cover relationships and mapping intrusions that cut both the Proterozoic sediments and the older basement, they clearly demonstrated that both groups of rocks were affected by a younger Alpine style orogenic event – now known as the Svecofennian orogeny. It is important to realize however, that no depositionally unconformable relationships have been unequivocally demonstrated between any Archean rock units in Finland, even though evidence for derivation of sedimentary, volcanic, and granitoid rocks from older crustal sources is widepread. Instead, all exposed contacts between rock units are either highly strained or obviously intrusive. Neither have the tectonic elements and magmatic signatures of modern crustal accretionary and collisional processes been definitively recognized, although geochemical characteristics of volcanic and granitoid rocks have been used to infer paleotectonic settings and processes (Martin et al., 1983b, 1984; Jégouzo and Blais, 1995). The lack of a clearly defined foreland substrate or orogenic polarity has been an impediment to developing a coherent understanding of large scale 26

Archean crustal processes and the evolution of the Archean crust in Finland. This is not to deny that such processes were involved in crustal formation and deformation; recent tectonic syntheses in adjoining Russian Karelia, invoke either interaction between plates and plumes (Puchtel et al., 1998, 1999) or collision and subduction of the the Karelian province beneath the Belomorian province at around 2720 Ma, with supporting evidence including a proposed intervening accretionary prism, the chemical characteristics of granitoid plutons and the fabric of deep crustal seismic reflectors (Berzin et al., 2001; Slabunov and Bibikova, 2001). This contrasts with earlier more traditional interpretations for the Russian part of the Karelian domain (Kratz and Mitrofanov, 1980) in which vertical crustal differentiation was seen as significant, and higher grade granulite terrains being generally considered older. This conceptual model led to the suggestion that there were two separate Archean orogenic events in Fennoscandia – the early Archean (3.1–2.9 Ga) Saamian cycle, represented by high grade metamorphic migmatite and granitoid terrains, and the late Archean (2.9–2.7 Ga) Lopian cycle, characterized by granitoid–greenstone complexes of lower metamorphic grade. In contrast, a simple statistical representation of available Archean U-Pb age determinations from Finland reveals that the majority of granitic and tonalitic intrusive rocks, as well as volcanics, formed within the interval 2.75–2.60 Ga, while a smaller population extends back as far as 3.2–3.3 Ga (Figure 2.4). However, because of the effective overlap in rocks of granitic and tonalitic and gabbroic composition, this approach does not provide a useful discriminant between primary crustal formation and intracrustal reworking processes. It also needs to be emphasized that at present, isotope ages within Finland do not appear to define any readily discernible regional spatial patterns. It is quite probable that even if the Archean

• CHAPTER 2 • ARCHEAN ROCKS

A

B

Fig. 2.4. Histograms displaying frequency distribution of Archean ages from Finland, based on compilation by Matti Vaasjoki of U-Pb analytical data, principally from multigrain zircon separates. (A) Histogram of total data, clearly indicating the importance of crustal formation and reworking between 2600 Ma and 2750 Ma. (B) Frequency distribution after discrimination of data according to rock type. There is considerable overlap between the various rock types, although there is a weak tendency for rocks of granitic composition to extend to younger ages, and conversely, tonalites to record older ages.

crust of Finland formed through plate tectonic accretionary crustal processes, original tectonic elements and terrain boundaries could have been significantly disrupted and obscured during subsequent crustal reworking. As noted

earlier, isotope studies do indicate that some detrital components in metasedimentary rocks and some inherited zircons in granitic rocks are derived from older crustal sources (Vaasjoki et al., 1993). Isotope studies have also

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•

27

revealed that structurally complex, highly strained migmatites and homogeneous weakly deformed intrusions cannot be discriminated on the basis of age alone (Vaasjoki et al., 1999). This underscores the need for careful and systematic documentation of structural evolution and intrusive sequences, rather than attempting to make correlations and infer tectonic settings simply on the basis of lithology, strain state or geochemical characteristics. The difficulties in defining boundaries between rock units in the field can be partly overcome by using magnetic signatures to delineate geophysical provinces. High-resolution magnetic data have also been important in detailed structural and stratigraphic mapping of the Archean (Sorjonen-Ward, 1993). At all scales however, it is necessary to take into account the effects of processes that generate or consume ferromagnetic minerals (Airo, 1999). The classification that follows reflects in a broad sense the thermal and strain history recorded by the Finnish bedrock, in that it is based on discriminating predominantly greenschist facies supracrustal sequences and higher grade terrains, supplemented by information from isotope studies of intrusive events and source rock ages. Another useful discriminant is the regional dip of the enveloping surface to major structures or lithic units, for this relates closely to the interaction between deformation and granitoid emplacement, which may also be a function of both crustal depth and the degree of thermal reworking. The hierarchy for classification used in this review proceeds from domain through terrain to complex or supracrustal belt; the use of the term terrain is descriptive only and is not intended to imply that adjacent crustal units were juxtaposed by accretionary plate tectonic processes. Indeed, in most cases the boundaries between terrains are either undefined in terms of kinematic history or obscured by younger rocks.

28

2. The Karelian domain in eastern Finland The Karelian domain is generally regarded as the cratonic nucleus to the Fennoscandian Shield. This is certainly valid from the perspective of Proterozoic deformation, since it acted as the foreland upon which Svecofennian nappes were emplaced, initially from the southwest around 1.9 Ga, as well as the Belomorian terrain and Lapland granulite belt, which were thrust from the opposite direction during the broadly coeval Kola–Lapland orogeny (Figures 2.1 and 2.16). The Karelian domain also formed a stable substrate for intracratonic volcanism and sedimentation throughout the Paleoproterozoic – indeed the earliest supracrustal units of the Lapland greenstone belt, deposited unconformably on the Karelian craton, strictly straddle the Archean–Proterozoic boundary (Manninen et al., 2001). In terms of Archean crustal growth, however, the Karelian domain records a complex pattern of ages dating back to at least 3.5 Ga, with a regionally coherent structural framework emerging only after 2.7 Ga.

2.1. Ilomantsi terrain The Ilomantsi terrain as defined here includes several well-preserved greenschist to amphibolite facies supracrustal sequences, namely the Hattu, Kovero, Nunnanlahti, and Ipatti supracrustal belts (Figures 2.2 and 2.5). The Hattu schist belt in particular has been studied in considerable detail because of its demonstrated potential for structurally controlled lode gold mineralization (Nurmi and Sorjonen-Ward, 1993). The distribution of preserved supracrustal sequences is principally controlled by variably sized and strained granodioritic and tonalitic plutons, with age ranges suggesting a close relationship between volcanism, deformation and pluton emplacement. The southwestern part of the Ilomantsi terrain is unconformably

• CHAPTER 2 • ARCHEAN ROCKS

overlain by Paleoproterozoic supracrustal sequences (Figure 2.5). Svecofennian granitoids with evolved Nd isotope characteristics (Figure 2.5) also indicate thermal reworking of deeply buried Archean crust farther to the southwest (Huhma, 1986). The western part of the Ilomantsi terrain, defined here as the Lieksa complex, includes abundant porphyritic granitoids, commonly containing pyroxene; granulite facies supracrustal enclaves indicate that an extensive high-grade terrain is present, as yet mapped only at reconnaissance scale. The ages of the granitoids, their source materials, and the metamorphism appear to be around 2.73–2.72 Ga (Halla, 1998, 2002) and therefore their development must be closely connected with the evolution of the adjacent lower grade supracrustal sequences. Migmatitic gneisses are also widespread, so that the transition between the Lieksa complex and the supracrustal gneisses of the Nurmes gneiss complex and the Kianta terrain is not precisely defined (Figure 2.5).

Hattu supracrustal belt The Hattu schist belt, located in the easternmost part of the Ilomantsi terrain, has been studied extensively in recent years, with emphasis on the structural architecture and its influence upon orogenic-style lode gold mineralization (Nurmi and Sorjonen-Ward, 1993). Isotope data indicate that deposition, deformation and granitoid intrusion were very closely related in time, the ages of the earliest supracrustal units effectively overlapping with those of syntectonic granitoids. All exposed contacts between the Hattu schist belt and these granitoids are intrusive (Figures 2.6 and 2.7), or else tectonically modified, and hence the granitoids cannot represent depositional basement to the greenstone belt. No other depositional basement to the Hattu schist belt has been identified, nor have any unconformities been recognized within the mapped sequence.

Zircon U-Pb zircon ages from the supracrustal sequence, ranging from 2754 ± 6 Ma for pyroclastic deposits low in the sequence to 2726 ± 15 Ma for porphyry clasts in conglomerate, overlap statistically with those for syntectonic granitoids (2746 ± 6 to 2725 ± 6 Ma). Although this indicates rapid crustal evolution, isotope studies of heterogeneous detrital and magmatic zircon suites indicate the presence of an older crustal component in granitoid source material as well as detrital sediments (Vaasjoki et al., 1993), with some xenocrysts from the Silvevaara granodiorite yielding ages up to 3.18 Ga (Sorjonen-Ward and Clauoé-Long, 1993). The involvement of older continental crust in the tectonic processes that deformed the Hattu schist belt is further attested to by the presence of some highly evolved granitoids, including tourmaline-muscovite leucogranites (Figure 2.8G), which appear to be analogous to those in Phanerozoic collisional belts. In spite of these features, no evidence for unconformities or any kind of depositional substrate to the Hattu schist belt has been found. The supracrustal sequence may commence with pillowed mafic volcanic rocks (Figure 2.8D) but consists predominantly of felsic pyroclastic and epiclastic deposits. Laterally persistent but volumetrically minor tholeiitic intercalations, and some komatiites occur in the upper part of the succession, typically associated with a variety of silicic and sulfidic banded iron-formations (Figure 2.8C). An abundance of depositional younging criteria indicate that the major folds in the greenstone belt are upward facing, thus tending to militate against interpretations invoking early recumbent folding (Figures 2.6 and 2.7). The only exceptions to this are local in nature and are currently ascribed to softsediment slumping. The well-defined stratigraphy and good correlation between magnetic properties and lithology, particularly at higher stratigraphic levels, has enabled the structural geometry to be further clarified (Figures 2.6 and 2.7A). Hence, in spite of locally intense

CHAPTER

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•

29

29°00’E

30°00’E

Kianta terrain

Tipasjärvi

Significant Svecofennian deformation Archean basement windows Boundaries between Archean rock units 0

Nurmes

50 km

Iisalmi terrain

Lieksa Ilomantsi terrain

Ipatti

Juuka Nunnanlahti

Hattu Figure 2.6 63°00’N

Kaavi Ags Pgm

εNd(2750)-2.1

Eno

εNd(1860)-6.0

Figure 2.7

Kontiolahti Outokumpu

Hattu

Ilomantsi

Sotkuma

Kovero Joensuu

Juojärvi

Kovero Suhmura

Kiihtelysvaara

Ilomantsi terrain Felsic volcanic and volcaniclastic rocks Turbiditic graywackes Mafic and ultramafic volcanics and sills Migmatites, leucocratic monzogranites, typically accompanying late orogenic transpressional deformation Biotite tonalitic and hornblende/pyroxene granodioritic plutons; granulite facies assemblages present in Lieksa complex

Oravisalo Pgk

εNd(1870)-3.6

εNd(1800)-6.9

Post-Archean rock units Proterozoic granites, typically recording isotope evidence for Archean crustal derivation Allochthonous Proterozoic supracrustal units, emplaced onto Karelian domain at 1.9 Ga Paleoproterozoic (2.4–2.0 Ga) sedimentary and volcanic units overlying the Karelian domain

Pgp

Kianta terrain Granitoids, gneisses, and migmatites Tipasjärvi greenstone belt Nurmes complex supracrustal gneisses

30°00’E

Iisalmi terrain Granitoids and supracrustal gneisses

Fig. 2.5. Regional synthesis of the Ilomantsi terrain in easternmost Finland. Extent of Proterozoic tectonic disruption and derivation of granitoids from buried Archean crust are also indicated. Ags–Silvevaara granodiorite, Pgk–Kermavesi granodiorite, Pgm–Maarianvaara granodiorite, and Pgp–Puruvesi monzogranite. Semitransparent gray shades relate to total magnetic intensity recorded by regional airborne surveys (reproduced from Geological Survey of Finland databases).

30

• CHAPTER 2 • ARCHEAN ROCKS

Pat High strain zones Undifferentiated metasediments

Kot 0

1 km

Vig Tat

Pampalo Formation Komatiitic pyroclastic flows and talc-chlorite-actinolite schists Medium- to coarse-grained massive metadolerite Intermediate to mafic volcanic rocks and volcaniclastic deposits Massive tholeiitic basalts Quartz-grunerite-magnetite banded iron-formation

Tiittalanvaara Formation

Pat

Thin-bedded metapelites Polymicitic conglomerates and feldspathic turbiditic graywackes Mg-rich tholeiitic basalts

Vig

Granitoid intrusions

Sivakkajoki Formation Polymictic conglomerates and cross-bedded feldspathic arenites Basaltic and andesitic volcanic rocks and volcaniclastic deposits Thin-bedded sulfide-bearing metapelites and graywackes Feldspathic epiclastic and pyroclastic deposits

Tat

Tasanvaara tonalite

Kot

Korpivaara tonalite

Vig

Viluvaara tonalite

Pat

Pampalonuurro porphyritic tonalite dike complex

Fig. 2.6. Geological map of part of the Pampalo structural domain within the Hattu supracrustal belt of the Ilomantsi terrain (after Sorjonen-Ward, 1993). CHAPTER

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31

Granitoid intrusions

A

Silvevaara granodiorite

Medium-grained biotite tonalite and leucotonalite

U-Pb(zircon)2757±4 Ma

Potassium feldspar-porphyritic hornblende granodiorite

εNd(2750)–0.4 to –2.1

Hattu schist belt Geophysically responsive, mainly thin-bedded metapelites Graywackes, mica schists and hydrothermally altered schists

Kuittila tonalite

Basaltic to intermediate volcanic and volcaniclastic intercalations

U-Pb(zircon)2746±9 Ma

Polymictic conglomerates and feldspathic turbiditic graywackes

εNd(2750)+0.4 to +2.3

Highly strained metabasalts and banded iron-formations Depositional younging direction High strain zones associated with progressive folding Trace of F2 antiformal hinge

0

1 km

qz

gar

po

B

C

Fig. 2.7. Rock types and their distribution in the southern part of the Hattu supracrustal belt of the Ilomantsi terrain. (A) Geological and structural map of part of the Kuittila structural domain (after Sorjonen-Ward, 1993), emphasizing relationship between deformation and emplacement of the Kuittila tonalite. (B) Detail of contact between turbiditic graywacke and the Kuittila tonalite. Note brittle–ductile fracture zones transecting tonalite and more ductile strain recorded in metasediments. Scale bar is approximately 1 dm in length. (C) Photomicrograph in plane-polarized light showing strain partitioning around garnet porphyroblast (gar), recorded by dynamically recrystallized intergrowth of quartz (qz), biotite, and pyrrhotite (po). Scale bar is approximately 1 mm in length. Photos: Peter Sorjonen-Ward.

32

• CHAPTER 2 • ARCHEAN ROCKS

and complex deformation (Figures 2.8A, B, C), the Hattu schist belt has retained a high degree of stratigraphical coherence (Figures 2.6 and 2.7A), which has enabled the delineation of two distinct and partially overlapping felsic volcanic complexes, developed within sporadically emergent but generally turbiditedominated basins. Sorjonen-Ward (1993) defined a number of formations within the Hattu supersequence. The lowest is the Sivakkojoki Formation, which consists mainly of feldspathic graywackes, representing resedimentation of penecontemporaneous volcanics and older felsic crust. Conglomerate interbeds within the formation contain clasts from many sources and exhibit local cross-bedding indicating deposition in shallow water, suggesting that volcanic edifices may have locally become subaerial. The Hosko Formation in the northern part of the area appears to show an upwards transition from turbidites to more proximal volcaniclastic rocks, with widespread sericitic alteration and replacement of plagioclase by microcline. The Hosko granite might be a synvolcanic shallow intrusion responsible for hydrothermal alteration, but structural relationships are not yet confirmed. The turbidites of the Kuljunki Formation (Figure 2.8B) may be correlative with the Hosko Formation, as it records the same distal to proximal transition. The Kuljunki Formation is overlain by the Tiittalanvaara Formation, which includes coarse polymictic conglomerates (Figures 2.6 and 2.8A), and records transgression, culminating in turbidites and sulfide-facies banded iron-formations. This marks the transition to a diverse but restricted phase of volcanism, defined as the Pampalo Formation, which contains tholeiitic basalts, doleritic sills, mafic to intermediate pyroclastic deposits and a single distinctive volcaniclastic komatiite unit. The structural architecture of the Hattu schist belt is characterized by upward-facing, generally steeply dipping structures. It is possible to establish a close, sequential

relationship between tightening of folds, attenuation of fold limbs, development of shear zones with strike-slip displacements, and the propagation of new folds due to strain incompatibilities between shear zones. Refold interference patterns or attenutation and excision of certain units at outcrop and map scale therefore most likely represent progressive deformation of initially upright structures with strain becoming more partitioned within discrete narrow zones. The kinematic histories of these zones suggest the importance of regionally coaxial and vertical constrictional strains, although evidence for more localized local strike-slip deformation is certainly present (Figure 2.8A). Deformation of the Hattu schist belt was closely associated with granitoid emplacement. In particular a distinctive suite of biotite tonalites intruded the sequence, initially as tabular semiconcordant (though not necessarily subhorizontal) sheets during the early stages of fold propagation, and were subsequently deformed along with their host rocks (Figures 2.6 and 2.7). This was accompanied by structurally controlled hydrothermal alteration and gold mineralization, which was subsequently recrystallized and deformed under upper greenschist to lower amphibolite facies conditions. Microstructural evidence clearly indicates dynamic recrystallization of hydrothermally altered assemblages. Annealed textures with garnet and locally, staurolite and kyanite porphyroblasts indicate that the thermal metamorphic peak was synchronous with or outlasted deformation (Figures 2.7C and 2.8E, F). A progressive, rather than episodic interpretation of deformation is therefore preferred, due to the geometrical congruence of overprinting phases and the relatively short time span between volcanism, deformation, and granitoid emplacement. Thus, younger structures appear to represent the partitioning of deformation into more discrete, highstrain zones, with an increasing component of vertical constrictional strain accompanying

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33

A

B

C

D

ky

and ky

E

F

G

H

34

• CHAPTER 2 • ARCHEAN ROCKS

emplacement and deformation of syntectonic plutons. This progressive evolution can result in either distinct overprinting fabrics, or else transposition and recrystallization of early fabrics, thus requiring care in defining sequential deformational phases. The mechanical behavior of various lithological units has clearly had a significant influence on deformation style. Major displacements were localized adjacent

to banded iron-formations (Figure 2.8C), while shortening of mafic units has in some instances been accommodated by the development of strike-slip duplexes, as shown by the imbrication of the Pampalo Formation in Figure 2.6. The resultant regional geometry suggests a transpressional regime with plutons being emplaced into dilatant sites within a N–NE-trending dextral shear system.

Fig. 2.8. (facing page) Rock types and microstructures from the Ilomantsi terrain. (A) Polymictic conglomerate typical of the Tiittalanvaara Formation, in the upper part of the preserved stratigraphic succession. Tiittalanvaara, northern part of Hattu supracrustal belt. Clasts and matrix consist predominantly of reworked intrabasinal volcanic and volcaniclastic deposits. Clasts are highly elongate perpendicular to outcrop surface, and sinistral folds record a regional strain path involving vertical constriction combined with transpression. Scale bar is approximately 1 dm in length. (B) Intense differentiated crenulation cleavage development (subhorizontal in photograph) and associated volume loss by solution transfer in thin-bedded laminated turbidite package within mesoscopic fold hinge zone. Kuljunki, northern part of Hattu supracrustal belt. Scale bar is approximately 1 dm in length. (C) Complex, inferred progressive deformation in quartz-grunerite-magnetite banded iron-formation intercalated with metaturbidites at northeastern margin of Kuittila tonalite; quartz-vein in axial planar orientation with respect to sinistral minor fold appears to be superimposed on tight to isoclinal dextral folds. Apparent superimposed fold generations and transition from ductile flow to semibrittle displacements may nevertheless represent a combination of strain rate control on rock behavior and local rotation as larger scale fold limbs amplify and need not have regional tectonic significance. Scale bar is approximately 1 dm in length. (D) Typical deformed pillow basalts, possibly representing the substrate upon which the sedimentary and felsic volcaniclastic sequence of the Ilomantsi terrain were deposited. Utrio, southeastern part of Hattu supracrustal belt. Scale bar is approximately 1 dm in length. (E) Continuity of sericite and biotite alignment in pseudomorphed, inferred andalusite porphyroblast with deflected external fabric, from hydrothermally altered schist along western margin of Kuittila tonalite. Kyanite (ky) clearly post-dates sericite crystallization but is confined to pseudomorphs. Therefore, relative timing with respect to continued deformation of the matrix remains unresolved, and a Proterozoic metamorphic origin for the kyanite is possible. Crossed nicols. Scale bar is approximately 1 mm in length. (F) Hydrothermally altered mica schist from near margin of Kuittila tonalite, showing optically contiguous relicts of andalusite porphyroblast (and) partly replaced by sericite aligned parallel to external foliation, with subsequent anoriented growth of kyanite (ky). Crossed nicols. Scale bar is approximately 1 mm in length. (G) Typical banding, defined by variations in tourmaline abundance, within the Naarva leucogranite, which intruded within the boundary zone between the Hattu supracrustal belt and the Lieksa complex. Scale bar is approximately 1 dm in length. (H) Pink potassic granite leucosomes discordant across highly strained migmatites, possibly representing progressive emplacement within the same magmatic system, and very characteristic of the southwestern part of the Ilomantsi terrain. Such migmatites are nevertheless relatively late, as they are associated with contractional shear zones that overprint the earlier tectonic and metamorphic fabrics recorded in Ilomantsi terrain supracrustal rocks. Scale bar is approximately 1 dm in length. Photos: Peter Sorjonen-Ward.

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35

Kovero supracrustal belt The Kovero schist belt (Nykänen, 1971; Tuukki et al., 1987) appears to be contiguous with the Hattu schist belt (Figure 2.5), though the intervening terrain is poorly exposed and attempts to date the sedimentary and volcanic rocks, as well as intrusive granitoids have so far been unsuccessful. The Kovero belt has also been affected by a younger phase of deformation, metamorphism, and granitoid emplacment that makes correlation more difficult. The most common rock types within the greenstone belt are Fe-rich tholeiitic basalts, which appear to be lowest in the stratigraphy. In many places they are associated with Mgrich tholeiitic basalts and komatiitic olivine (±pyroxene) cumulates, now altered into serpentinites and tremolite-chlorite rocks (Tuukki et al., 1987; Tuukki, 1991). Together they probably represent a deeply eroded remnant of a submarine lava complex. Distinctive felsic volcanic rocks, with abundant hydrothermal pyrite deposits are closely associated with the ultramafic and mafic lavas. It is therefore quite reasonable to correlate this stratigraphic horizon with the Pampalo Formation in the Hattu supracrustal belt, allowing for lateral facies variations.

Nunnanlahti and Ipatti supracrustal belts Archean supracrustal rocks also occur along the western margin of the Ilomantsi terrain, and could correlate with either the Kovero supracrustal belt, or the Tipasjärvi greenstone belt of the Kianta terrain (Figure 2.5). The Ipatti belt is exposed discontinuously beneath the Paleoproterozic unconformity and in places it can be shown that rocks have been leached during weathering and now comprise part of the Paleoproterozoic Hokkalampi paleoregolith (Kohonen and Marmo, 1992). The Ipatti supracrustal belt has also been affected by Svecofennian deformation, particularly in its type area near Koli, where it is folded around in a tight SW-plunging syncline. Lithologi36

cally, the sequence is rather diverse, including thin-bedded turbidites and mafic volcaniclastic deposits, concordant felsic porphyritic sills and sporadic basalts (Rossi, 1975). The nearby Nunnanlahti greenstone belt (Figure 2.5) is approximately 15 km long and 2–3 km wide. Although it may originally have been contiguous with the Ipatti belt, it now represents an almost allochthonous tectonically imbricated remnant amongst both Proterozoic and Archean rock units. This intense Svecofennian tectonic reworking was already recognized by the earliest geologists to work in the region (Frosterus and Wilkman, 1920). The main Nunnanlahti shear zone has had a complex history, being interpreted as an early thrust, or steep frontal ramp within a thrust system, which emplaced the Nunnanlahti greenstones over Proterozoic turbidites. The deformation zone was then reactivated as an oblique sinistral shear zone, with kinematics deduced from rotated porphyroclasts, cleavage duplexes, fold asymmetry and truncations of lithological units and magnetic anomalies at map scale. In proximity to the Nunnanlahti shear zone, regional structural trends in both basement and cover rocks are progressively transposed into NW-orientations with an intense moderately dipping foliation and Splunging lineation. In the most highly strained domains the foliations in the protomylonitic Archean granitoids, as well as the Nunnanlahti greenstones are essentially congruent with those in the Proterozoic sediments (Kohonen et al., 1991). Primary stratigraphic relationships within the Nunnanlahti greenstone belt are difficult to establish, due to the complex deformation history. Rock types range from massive and pillowed tholeiitic basalts, ultramafic rocks and felsic volcanic rocks, with some pelitic schists and chert (Kohonen et al., 1989). Ultramafic rocks include massive serpentinites, some of which are very homogeneous and appear to retain textures suggesting a dunitic cumulate origin. Serpentinites have also been

• CHAPTER 2 • ARCHEAN ROCKS

extensively altered to talc-magnesite rocks, including the economically significant Kärenvaara soapstone deposits, and associated biotite, tremolite, and chlorite schists. None of the rocks types present in the area have been amenable to isotope dating and lithological boundaries are generally so highly strained that any evidence for intrusive or truncating relationships has been destroyed. Highly strained granitic dikes also truncate the Nunnanlahti greenstones, and although these have not been dated, the absence of Proterozoic granitoids elsewhere in the region provides another argument supporting an Archean age for the Nunnanlahti greenstones. Likewise, Proterozoic mafic dikes have been observed to truncate serpentinite and foliated metabasalt, further demonstrating that the Nunnanlahti greenstones are indeed Archean in age (Sorjonen-Ward and Rossi, 1997). However, the timing of talc-carbonate alteration and soapstone formation is not fully constrained, despite its significance to regional metamorphic and tectonic studies. Only one example of a tabular, Proterozoic mafic dike truncating the talc-carbonate has been documented, from drill core (Tapio Kuivasaari, pers. comm., 2002). The dike shows a pronounced reaction selvage, although such features could also be expected where a tholeiitic dike within talccarbonate rock was subjected to static regional metamorphism. Therefore, the extent to which the talc-carbonate alteration represents a Proterozoic retrogression and fluid influx, as opposed to an entirely Archean phenomenon remains unresolved.

Lieksa complex – granitoids and high-grade gneisses The Lieksa complex is defined here as a NNEENE trending zone dominated by porphyritic potassium feldspar granodiorites that commonly contain pyroxene (Figure 2.5). Areas of mafic granulite with orthopyroxene and clinopyroxene indicate that these porphyritic granitoids may have crystallized, or were at

least metamorphosed under granulite facies conditions. If the latter were the case, the results of Pb isotope studies on potassium feldspar would suggest that very little time would have elapsed between emplacement and metamorphism, as the U-Pb zircon age of 2730 ± 20 Ma accords very well with the Pb whole-rock–potassium feldspar isochron age of 2728 Ma (Halla, 1998). The lead isotope studies of Halla (1998, 2002) further suggest that the Lieksa complex granitoids could have been derived from a mixture of juvenile, mantle-derived material and reworked older continental crust. The Silvevaara granodiorite, which intrudes the Hattu schist belt to the east of the Lieksa complex resembles the porphyritic granites of the latter, and shares the same prominent magnetic signature, wherever mafic minerals and magnetite have avoided hydration and retrogression (Sorjonen-Ward, 1993). Although this intrusion is evidently somewhat older than the Lieksa complex, with an interpreted magmatic SHRIMP age of 2757 ± 4 Ma (Sorjonen-Ward and Claoué-Long, 1993), it also indicates the existence of substantially older material in the deep crust. Neither the western boundary, nor the internal features of the Lieksa complex have been studied in detail. Because, however, the porphyritic granitoids are exposed across an area greater than normal crustal thickness, the nature of their geometry at depth has considerable relevance to understanding the potential source and volume requirements for magmatism and hence thermal and tectonic regime. Limited field data indicate that while a regionally expressed steep foliation is defined by biotite, gently dipping and folded compositional banding is common, sometimes with alignment of tabular potassium feldspar phenocrysts in a weakly strained groundmass. This can be interpreted as a high-temperature, emplacement related fabric, suggesting that the granitoids might represent coalesced tabular sheets that were emplaced and crystallized under granulite facies conditions during

CHAPTER

2

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•

37

regional compressive deformation, and were progressively deformed during cooling and uplift. The eastern boundary towards the Hattu schist belt coincides with a relatively abrupt transition to a zone of felsic granite sheets, sometimes muscovite bearing, and migmatites that clearly represent injection of felsic melt into amphibolite facies metasediments that probably represent higher grade and more recrystallized equivalents of Hattu schist belt rock types (Sorjonen-Ward, 1993). The presence of muscovite-bearing granites, and also the distinctive Naarva leucogranite, which has tourmaline as a major phase, together with biotite, muscovite and garnet, means that the boundary zone between the Lieksa complex and Hattu schist belt may be of fundamental significance. As well as the strongly magnetic character of the Lieksa complex, the regional gravity data show a distinct negative anomaly coinciding with the Hattu supracrustal belt. The later stages of deformation throughout the Ilomantsi terrain are interpreted as a consequence of NE–E-directed compression, resolved as a combination of thrusting with a top to the east sense, and dominantly dextral transpression (Sorjonen-Ward, 1993; Luukkonen and Sorjonen-Ward, 1998). Therefore a plausible geodynamic scenario would be oblique emplacement of the Lieksa complex over the Hattu schist belt, and hence coupling of the exhumation of the granulite facies rocks of the Lieksa complex with partial melting of underthrust Hattu supracrustal belt sediments at depth. As yet, there are no isotope or chemical data from the peraluminous leucogranites that would constrain the timing of this event or allow characterization of age and composition of potential source material. However, based on the U-Pb zircon and Pb-Pb model ages from the Lieksa complex (Halla, 1998) and the U-Pb results for zircon, monazite, and titanite from granitoids intruding the Hattu supracrustal belt (Vaasjoki et al., 1993), the 38

two units would most likely have been juxtaposed between 2.73–2.69 Ga. Late kyanite has also been observed to locally overprint relict and sericitized andalusite porphyroblasts (Figure 2.8E and F), belonging to the typical biotite-garnet-staurolite peak metamorphic assemblages of the Hattu supracrustal belt (Sorjonen-Ward, 1993). Although this might be seen as evidence of an increase in pressure related to Archean late orogenic thrusting, the isolated nature of these kyanite occurrences and the convincing documentation of Svecofennian thermal overprinting that produced kyanite in the Kianta terrain (Pajunen and Poutiainen, 1999) makes this interpretation less likely.

Granitoids intruding the Hattu and Kovero supracrustal rocks In contrast to many granitoid-greenstone terrains in Finland, it has proven possible to map discrete plutons within and around the Hattu supracrustal belt, and in some cases, to document contact relationships (Sorjonen-Ward, 1993). Although geochemical studies so far have emphasized lithogeochemical exploration aspects rather than petrogenesis (Nurmi and Sorjonen-Ward, 1993), the available data are sufficient for some general conclusions to be made (O’Brien et al., 1993a). The Kuittila suite of enclave-poor biotite tonalites has been studied in most detail and is particularly important in relation to constraining the timing and style of deformation and orogenic hydrothermal alteration processes in the Hattu supracrustal belt. This is because the Kuittila tonalite and Tasanvaara tonalite both form elongate intrusions some 50 km2 in extent (Figures 2.6 and 2.7A), with marginal apophyses that clearly truncate lithic units, while sharing an overall similar strain history (Sorjonen-Ward, 1993). The western margin of the Kuittila tonalite in particular is characterized by plagioclase-phyric and quartzplagioclase-phyric dike swarms that are highly strained and concordant with the axial planar

• CHAPTER 2 • ARCHEAN ROCKS

foliation to mesoscopic and regional folds in the country rocks (Figure 2.7A and B). The overall geometry suggests emplacement into a releasing bend within a steeply dipping dextral reverse sense shear system. The current shape of the pluton is considered to be primary, even though considerable post-emplacement strain may be recorded by the enclosing sediments (Figure 2.7). This is deduced from arrays of molybdenite-bearing quartz veins attributed to late- to post-magmatic processes in the Kuittila tonalite (Sorjonen-Ward, 1993), which show little evidence of buckling or substantial shear displacement. Similarly, it is difficult to interpret the northern and eastern margins of the Tasanvaara tonalite in any other way than fracture controlled propagation of dikes and eventually more coherent batches of tonalitic magma into the east-younging eastern limb of the regional N-plunging Pihlajavaara anticline, disrupting the stratigraphic sequence, quite possibly as the fold was amplifying (Figure 2.6). Several kilometers to the north of the Tasanvaara tonalite, the Korpivaara tonalite also seems to be intimately associated with the localization of deformation in the country rocks, in particular the Juttuhuhta oblique-sinistral duplex and the related Pampalo shear system (Figure 2.6). Indeed, the orientation and kinematic indicators in this area would be compatible with emplacement of the Korpivaara tonalite occupying the hanging wall above an oblique normal shear system, especially as the eastern margin of the pluton is also inward dipping, resulting in a funnel-shaped cross-sectional geometry. The Kuittila tonalite has a U-Pb zircon age of 2745 ± 10 Ma, compared to 2748 ± 6 Ma for the Tasanvaara tonalite and Sm-Nd data from both plutons produce a spread of TDM model ages, mostly in excess of the zircon age, including several over 2.85 Ga. Molybdenite extracted from the magmatic-related W-Mo mineralization in the Kuittila tonalite also initially yielded an age of near 2.85 Ga (Stein

et al., 1998), but revision of these analyses now produces ages that coincide remarkably well with the zircon dates. Although derivation from a crustal precursor of age greater than 2.8 Ga is therefore permissible, and such crustal material is indeed widespread in migmatites of the Kianta terrain, calculated εNd(at 2750 Ma) values of +0.9 to +2.1 also preclude source material from being significantly older. This is in contrast to the 3.0–3.2 zircon xenocrysts and TDM model ages and εNd(at 2750 Ma) values of –0.4 to –2.1 obtained from the nearby Silvevaara granodiorite (O’Brien et al., 1993a; Sorjonen-Ward and Claoué-Long, 1993; Vaasjoki et al., 1993). The magmatic age determined from SHRIMP studies of the Silvevaara granodiorite may be marginally older than that for the Kuittila tonalite and the Tasanvaara tonalite, though the two results overlap statistically at 2σ confidence levels; field relationships are not conclusive though it is inferred from aeromagnetic data that the Tasanvaara tonalite postdates the Silvevaara granodiorite. Because the Silvevaara granodiorite is also mineralogically very different, containing hornblende and potassium feldspar phenocrysts, it is interesting to speculate whether the two intrusive types were derived from different crustal sources and depths, or record different degrees of interaction with an as yet undefined penecontemporaneous mantle-derived magmatic component. The Kuittila suite shows REE profiles similar to those of other Archean TTG plutons, although HREE depletion is perhaps not so marked (O’Brien et al., 1993a). The Kuittila pluton can be subdivided into a tonalitic outer phase and a leucotrondhjemitic interior, in addition to the marginal porphyritic dikes swarms, all of which appear to be geochemically consanguineous. There is a systematic decrease in TiO2, Fe2O3, MgO, CaO, P2O5 , and Hf, Zr, and LREE with increasing SiO2, which would be expected for preferential source retention of mafic minerals, and accessory phases, including apatite, zircon, and monazite. On

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the other hand, the porphyry dikes, unlike the tonalite and trondhjemite, do not show Eu depletion, which is consistent with the abundance of plagioclase phenocrysts. Unless they represent cumulate enrichment, this would be compatible with derivation of tonalite from a tholeiitic basaltic source, given that experimental data indicate a broader stability field for plagiclase-biotite-quartz assemblages in tonalitic rocks compared to tholeiitic basalt (Huang and Wyllie, 1986). Pitkäjärvi (1988) also attempted petrogenetic modeling of REE chemistry of trondhjemitic granites in the southern part of the Ilomantsi terrain. He concluded that the most appropriate source material would be of quartz dioritic to tonalitic composition, with a calculated modal mineralogy of plagioclase (56%), quartz (13%), biotite (12%), amphibole (17%) and accessory apatite and zircon. The process was modeled with a high degree of melting of plagioclase and quartz, and conversely small amounts of mafic and accessory minerals, resulting in the observed HREE depletion, high Sr and Ba values, and lack of Eu depletion. A higher degree of melting of a tholeiitic source would also be plausible. Although attempts to accurately date the granitic rocks in this area have so far been unsuccessful (Vaasjoki et al., 1993), melting of an older tonalitic to dioritic basement can easily be integrated with constraints from field mapping. The later stage of deformation of the Kovero schist belt includes NE–E-directed folding and thrusting that appears to control the emplacement of discordant felsic granitic sheets within older quartz dioritic-tonalitic-granodioritic plutons, including the Pogosta granodiorite, which has a zircon age of 2724 ± 5 Ma (Vaasjoki et al., 1993). However, it should be noted that this later magmatism is not exclusively trondhjemitic. Potassic monzogranites are prominent and widespread, as stromatic highly strained migmatites and discrete plutons along the western margin of the Kovero supracrustal belt (Figure 2.8H). They are also distinctly 40

magnetic and appear to pass transitionally northwards into the Lieksa complex, although the nature, significance, and age of this transition is obscure.

2.2. Kianta terrain The Kianta terrain is critical to understanding the nature and origin of Archean greenstone and granite terrains in Finland and has long been the subject of mapping programs by the Geological Survey of Finland (Wilkman, 1924; Matisto, 1958; Hyppönen, 1983; Luukkonen, 1986, 1987, 1992, and 1993) and thematic investigations by research groups from the universities of Oulu (Piirainen, 1988), Rennes (Martin et al., 1984), and Turku (Halkoaho et al., 1996; Papunen et al., 1989, 2001). The Kianta terrain is bisected from north to south by several greenstone belts, more than 200 km in length, but generally less than 10 km in width (Figures 2.2, 2.9, and 2.10). From north to south these are known as the Suomussalmi, Kuhmo, and Tipasjärvi greenstone belts, which are mutually similar in terms of stratigraphy and tectonic events. Intense deformation has obliterated primary structures and textures in many places, but well-preserved low strain domains include the Siivikkovaara–Kellojärvi area in the Kuhmo belt (Papunen, 1960; Hanski, 1980; Halkoaho et al., 2000), the Taivaljärvi area in the Tipasjärvi belt (Taipale, 1988; Papunen et al., 1989), and the Saarikylä (Engel and Dietz, 1989) and Kiannanniemi areas in the Suomussalmi belt where, because of the potential for komatiitehosted nickel mineralization, detailed field studies have been undertaken to better characterize volcanic facies and eruptive processes (Papunen et al., 2001). Although there is no reason to correlate the supracrustal belts of the Ilomantsi and Kianta terrains, the general similarities in structural architecture, and age and lithological characteristics of granitoids suggests that by at least 2.74 Ga, they were developing as a single coherent terrane in

• CHAPTER 2 • ARCHEAN ROCKS

30°00’E

29°00’E 0

5 km

10

RUSSIA

Jumaliskylä

Moisiovaara

64°30’N

Härmänkylä

Koskenmäki

Kuhmo

Kuhmo greenstone belt Polymictic conglomerates, turbiditic graywackes, and sericitic quartzites

Kianta terrain

Intermediate and felsic volcanic rocks and volcaniclastic deposits Komatiites and komatiitic olivine (± pyroxene) cumulates Mg-rich tholeiitic basalts and komatiitic basalts Fe-rich tholeiitic basalts Layered hornblende gabbros and uralite porphyry intrusives

Post-Archean porphyritic monzogranite (2.45–2.39 Ga) Tonalite and granodiorite intrusive into greenstone belt rock units Monzogranite intrusions post-dating greenstone belt rock units Highly strained to cataclastic leucotonalite and leucotonalites Tonalitic-trondhjemitic-granodioritic migmatites and pelitic gneisses Tonalitic-trondhjemitic-granodioritic migmatites and gneisses, including granulite facies domains

Banded amphibolite, typically derived from tholeiitic basalts

Fig. 2.9. Principal geological features of the Kuhmo greenstone belt and surrounding Kianta terrain (after Luukkonen and Sorjonen-Ward, 1998). Semitransparent gray texturing in Kianta terrain to the east of the Kuhmo greenstone belt relates to total magnetic intensity recorded by regional airborne surveys (reproduced from Geological Survey of Finland databases). Stronger patterning generally correlates with either higher metamorphic grade or less intense Paleoproterozoic hydration and retrogression. CHAPTER

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29°00’E 0

5 km

10

Post-Archean rock units Iivaara alkaline intrusive complex (Late Devonian) Näränkävaara layered mafic intrusive complex (2.45–2.39 Ga)

Archean rock units Younger granitoid intrusions Tonalite and granodiorite intrusive into greenstone belt rock units Monzogranite intrusions post-dating greenstone belt rock units Peranka

“Younger greenstones” (Saarikylä Group and correlatives)

Selkoskylä

Komatiites and komatiitic olivine (± pyroxene) cumulates Saarikylä

Mafic, intermediate, and felsic volcanic rocks and volcaniclastic deposits

Tormua

RUSSIA Juntusranta

“Older greenstones” (Luoma Group and correlatives) Mafic, intermediate, and felsic volcanic rocks and volcaniclastic deposits

Kiannanniemi

Banded amphibolite, typically as enclaves in older TTG migmatites

Older granitoids

65°00’N

Tonalitic-trondhjemitic-granodioritic migmatites and gneisses

Fig. 2.10. Principal geological features of the Suomussalmi greenstone belt of the Kianta terrain (after Luukkonen and Sorjonen-Ward, 1998).

response to the same craton-wide tectonomagmatic processes. The relationships between the greenstone belts of the Kianta terrain and surrounding granite-gneiss terrains have also been studied in detail. However, the earliest attempts to date magmatic processes and tectonic events in a 42

coherent way led to considerable controversy, due to the use of different isotope techniques. This was principally manifest in Rb-Sr wholerock ages being systematically younger than zircon ages from the same plutonic rocks (Martin et al., 1984; Luukkonen, 1985; Halliday et al., 1988; Martin and Barbey, 1988; Martin,

• CHAPTER 2 • ARCHEAN ROCKS

1989; Vaasjoki, 1989). Those who favored the zircon dates as recording plutonic ages argued that there was widespread resetting of the RbSr system during a Paleoproterozoic thermal event. In contrast, those who claimed that the Rb-Sr data were robust records of igneous cooling interpreted the older zircon ages as representing xenocrysts inherited from the melting of older crust. Although this particular controversy has been conclusively settled in favor of magmatic zircon reflecting emplacement (Martin, 1989; Vaasjoki, 1989; Vaasjoki et al., 1999), other studies have demonstrated the inheritance of older zircon in some parts of the Karelian domain (Sorjonen-Ward and Claoué-Long, 1993; Vaasjoki et al., 1993) and the issue is still highly relevant, especially for felsic volcanic rocks. Moreover, all of these studies have effectively drawn attention to the complexity of the region, and demonstrated that crustal evolution involved at least two major stages, with considerable tectonic and thermal reworking. Thus, the earliest welldocumented event in the Kianta terrain was amphibolite to granulite facies metamorphism and formation of tonalite–trondhjemite migmatites during pervasive deformation at 2843 ± 18 Ma (D2 event of Luukkonen, 1985, 1988a). Some evidence exists for eruption of mafic lavas (now greenstones) on this older continental substrate and has led to the currently preferred model of the greenstone belt as essentially an ensialic rift (Luukkonen, 1988a, 1992). Luukkonen (1988b) dated a differentiated mafic sill, which is believed to be cogenetic with the greenstone sequence at 2790 ± 12 Ma, while Tulenheimo (1999) reported an age of 2757 ± 20 Ma from an ultramafic cumulate complex that has assimilated granitic wallrocks. These relationships are of fundamental significance to any interpretation of the evolution of the region as they imply that the migmatites do in some sense form a basement to the Kuhmo greenstone belt. The second major phase of granite generation and intrusion accompanied the deformation of the

greenstone belts after 2.74 Ga (Luukkonen, 1988a, 1992; Sorjonen-Ward et al., 1997). The petrogenetic studies of Martin et al. (1983a,b) and Martin (1986, 1987a,b) have also provided the basis for generic comparisons of calc-alkaline plutonic magmatism in Archean and younger convergent regimes.

Suomussalmi greenstone belt The Suomussalmi greenstone belt is located at the northern end of the Kianta terrain, where northerly trends abruptly change to an easterly trend, marking the boundary with the Koillismaa terrain (Figure 2.10). Two distinct geological units have been recognized in the Suomussalmi greenstone belt, the Luoma Group and Saarikylä Group, separated by a mylonitic zone with intense albite-sericite alteration. Isotope age determinations indicate that the Luoma Group may be the oldest well-preserved supracrustal unit documented from Finland (Vaasjoki et al., 1999). However, although Engel and Dietz (1989) proposed that an angular discordance existed between the Luoma and Saarikylä groups, no information is available concerning the structural and metamorphic history of the Luoma Group prior to deposition of the Saarikylä Group. Both units were, however, affected by the main pervasive deformation recorded throughout the Kuhmo and Suomussalmi greenstone belts, and were intruded by granitoids around 2.7 Ga (Patchett et al., 1981). The Luoma Group consists of mafic, intermediate, and felsic lavas and pyroclastic rocks which were deposited in shallow water or possibly even in a subaerial environment, and include sporadic stratiform Ag-Zn-Pb mineralization (Kopperoinen and Tuokko, 1988). The presence of andesitic compositions is rather unusual for Archean rocks in Finland, as is the U-Pb zircon age of 2966 ± 9 Ma. Whole-rock Rb-Sr results (Martin and Querré, 1984) and Pb-Pb isotope data (Vidal et al., 1980) nevertheless indicate that the rocks of the Luoma Group were subjected to some

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kind of thermal disturbance at 2500 ± 100 Ma, which has not been recorded in other parts of the Suomussalmi greenstone belt (Vaasjoki et al., 1999). Given that there is evidence for zircon inheritance and heterogeneity in other felsic sequences in Finland (Vaasjoki et al., 1993), it is also conceivable that the eruptive age of the Luoma Group age is considerably younger, especially in view of the fact that these rocks do not record the 2.86–2.83 Ga migmatite event documented throughout the Kianta terrain, but instead display typical D3 structures (Luukkonen, 1985, 1988a). The Saarikylä Group in the eastern and central part the Suomussalmi greenstone belt is dominated by komatiitic olivine (±pyroxene) cumulates and komatiitic and tholeiitic basalts (Figure 2.11D). These rocks represent the deeply eroded remnants of shield volcanoes or lava ridges formed by submarine fissure eruptions. The komatiitic olivine (±pyroxene) cumulates probably represent deeply eroded parts of the lava flows and/or the lava channels of this large lava complex. The komatiitic and tholeiitic basalts are massive, with pillow structures, but because of the intense deformation primary structures have often been destroyed. Layered mafic sills up to tens of meters thick intrude the lavas. Intermediate and felsic volcanic rocks, volcaniclastic rocks, and graphitic schists overlie the mafic lavas. A number of nickel prospects have been identified in association with the komatiitic and tholeiitic cumulates, such as at Hietaharju and Peura-aho (Kojonen, 1981) and the region is currently under active investigation because of its gold prospectivity (Papunen et al., 2001).

Kuhmo greenstone belt The Kuhmo greenstone belt has been the subject of detailed study from the point of Archean crustal evolution and komatiitic magmatic processes, as well as for nickel and gold exploration (Papunen, 1960; Hanski, 44

1980; Jahn et al., 1980; Martin et al., 1984; Piirainen and Taipale, 1985; Luukkonen, 1988a; Halkoaho and Pietikäinen, 1999; Papunen et al., 2001). Stratigraphic relationships are relatively well understood, and many units can be traced along the entire length of the belt, and indeed correlated with equivalent stratigraphic levels in the Tipasjärvi supracrustal belt to the south, and the Suomussalmi greenstone belt to the north. In general, the greenstone belt defines a synclinorial structure, formed during regional D3 deformation (Luukkonen, 1988a, 1992). Even though strain is locally intense, primary depositional and eruptive features are widely preserved (Figure 2.11A). The stratigraphic sequence appears to begin with felsic volcanic rocks, but is dominated by mafic rocks. The former are found as several isolated occurrences along the eastern and western marginal areas of the Kuhmo greenstone belt and are correlated stratigraphically with the more extensive Koivumäki Formation in the Tipasjärvi supracrustal belt (see below). This is consistent with U-Pb zircon ages of 2798 ± 15 Ma from the Juurikkaniemi Group in the Ontojärvi area and 2810 ± 48 Ma from the felsic unit at Vuosanka (Luukkonen, 1992). However, primary stratigraphical transitions between felsic and inferred overlying mafic volcanic rocks have not been observed. In the northern part of the Kuhmo greenstone belt, at Moisiovaara, mafic sills have been dated to 2790 ± 18 Ma (Luukkonen, 1988b). These provide important constraints to the geodynamic setting and timing relationships between various elements of the Kianta terrain, as komatiitic dikes evidently truncate tholeiitic banded gneisses that had already been affected by one or more deformation events (Figure 2.11B). In the southern part of the belt, however, the only age determination available from within the mafic to ultramafic sequence is 2757 ± 20 Ma (Tulenheimo, 1999). Therefore, it is possible that stratigraphic or structural breaks are present. The type stratigraphic sections have been

• CHAPTER 2 • ARCHEAN ROCKS

A

B

C

D

Fig. 2.11. Representative rock types in the Kianta terrain. (A) Polygonal jointing in komatiite flow at Näätäniemi, in relatively weakly strained domain at southern end of Kuhmo greenstone belt. Scale bar is approximately 1 dm in length. (B) Intrusive relationships between fine-grained komatiitic dikes related to the main greenstone sequence in the Kuhmo greenstone belt and older banded tholeiitic amphibolites. Such relationships are critical in demonstrating that greenstone belt magmatism occurred at least partly within an older continental crustal context. Deformation of komatiitic dikes relates to the principal tectonomagmatic event recorded through the Kianta and Ilomantsi terrains. Repolampi, northern end of Kuhmo greenstone belt. Scale bar is approximately 1 dm in length. (C) Complex relationships between deformation, anatexis, and melt migration in multiply deformed migmatites characteristic in particular of the eastern part of the Kianta terrain. Kelkkakangas, compass diameter is nearly 7 cm. (D) Pillow basalts in low-strain domain at Peura-aho in the Suomussalmi greenstone belt. Note hyaloclastic breccia in interstices and amygdales with radiate orientation, typical for shallow eruption depths. Scale bar is approximately 1 dm in length. Photos: Peter SorjonenWard.

defined in the Siivikkovaara area in the southern part of the belt, where primary features are best preserved, although nomenclature remains to be formalized (Papunen, 1960; Hanski, 1980; Hyppönen, 1983). It should also be noted that regional metamorphism and hydrothermal alteration have led to textural replacement and loss of primary mineral-

ogy. Ultramafic rocks (Figure 2.11A) thus have serpentine-talc-magnesite in cumulus layers, and tremolite-chlorite-albite-chromite-carbonate in former spinifix layers and have commonly lost their original magnetic character. Mafic rocks contain garnet-hornblende-plagioclase-chlorite. The lowermost and thickest unit exposed is the Pahakangas

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“Formation” (Papunen, 1960; Hanski, 1980; Halkoaho et al., 2000). This comprises a thick succession of submarine massive and pillowed tholeiitic basalt flows, commonly separated by magnetite-grunerite-quartz BIF horizons that show considerable variations in thickness, suggesting that the basalts also accumulated in a regional topographic depression. Individual flows may attain a thickness of 70 m, and the total sequence exceeds 1000 m. Gruau et al. (1992) discussed the effects of Proterozoic metamorphism with respect to REE mobility and Sm-Nd isotope resetting in the Siivikkovaara area, so that geochemical data need to be selected and evaluated carefully. Pahakangas tholeiitic basalts were also subject to primary or diagenetic interaction with intercalated BIF but in general display flat REE patterns and lack of Eu anomalies, as would be expected from the primitive character of the melt. The basaltic Pahakangas phase of volcanism terminated with the deposition of a sulfide-facies iron-formation, followed by eruption of komatiitic lavas that represent both distal and proximal flow regimes. The Siivikko komatiitic volcanism includes abundant relatively thin lava flows, especially towards the base but also a significant proportion of cumulates, in particular the Kellojärvi cumulate complex. Four zircon fractions from the Niittylahti gabbro, which belongs to the Kellojärvi cumulate complex, yielded a U-Pb age of 2757 ± 20 Ma, which is so far the only direct date obtained from the mafic and ultramafic sequence (Tulenheimo, 1999). The base of the Siivikko volcanic phase is marked by tremolite rock interpreted as a product of assimilation of underlying BIF by komatiitic magma. The first three flows recorded are fractionated with orthocumulate and spinifex textures, overlain by a large number of flows characterized by orthocumulates, polygonal jointing and flow top breccias, suggesting a progressively increasing supply of magma or proximity to vent. The lower flows 46

also show subtle differences in chemistry, including LREE enrichment, that are attributed to assimilation of felsic material. Further evidence of assimilation, and confirmation of a rifted continental crustal setting is indicated by the presence of granitic enclaves in serpentinite cumulate bodies. The Kellojärvi cumulate complex is 24 km2 in extent and consists of serpentinites and talc-magnesite rocks derived principally from olivine adcumulates and mesocumulates and minor olivine-clinopyroxene adcumulates. An original thickness of 1.5–2.5 km has been inferred. Sheared mylonitic talc-carbonate rocks occur at tectonic contacts between the cumulates and the granodioritic country rocks. However, erosional and flow structures are well preserved within the complex, while detailed mapping of the marginal zones has demonstrated partial melting and assimilation in several areas, thus providing evidence for eruption on an older granitic substrate (Halkoaho et al., 1996; Tulenheimo, 1999). As a result, the margins of the complex are characterized by hybrid cumulates of pyroxenitic composition, varying from 20 to 50 m thick. There is therefore abundant evidence for interaction between the komatiitic cumulates and a felsic substrate. Enclaves of Pahakangastype tholeiites moreover indicate the complex extruded through the earlier lavas and to the surface. Mobilization of granitic country rock in this way can clearly lead to potential confusion when attempting to determine timing relationships between greenstones and granitoid magmatism. Although significantly disrupted by smallscale faulting, the upper part of the section at Siivikkovaara shows a transition from komatiite to pillowed and variolitic komatiitic high-Mg basalts and eventually to a distinctive suite of Cr-rich basalts (Halkoaho et al., 2000). The high-Cr basalts are also associated with sporadic komatiite flows but are distinguished from the underlying high-Mg komatiitic basalts by Cr values ranging from 1300 to 4500

• CHAPTER 2 • ARCHEAN ROCKS

ppm. Quartz-filled drainage cavities in some pillows indicate shallow water depths for eruption and flows range from 0.5 to 5 m in thickness. No evidence of relict chromite or chromian magnetite has been found, despite the preservation of these minerals in other, equally metamorphosed and recrystallized rock types in the Kuhmo greenstone belt. Therefore Halkoaho et al. (2000) considered that Cr was originally in clinopyroxene rather than spinel. Because there is no evidence for cumulate concentrations of clinopyroxene in these flows, nor hydrothermal alteration, Halkoaho et al. (2000) concluded that the Cr enrichment was a primary magmatic feature related to a relatively low oxygen fugacity in the source region, but were unable to establish whether this was an inherent feature of the Archean mantle or a relatively local phenomenon, possibly related to fractionation of olivine from the associated komatiites. The Cr basalt and komatiite sequence at Siivikkovaara area are evidently overlain discordantly by a poorly exposed sequence of graded and current-bedded mafic to felsic pyroclastic and epiclastic deposits, including lahar breccias that contain clasts of komatiite and high-Cr basalt; no basement granitoid clasts have been found (Nieminen, 1998). Various tectonic and magmatic models have been presented to explain the origin of the mafic and ultramafic volcanism of the Kuhmo greenstone belt, including gravitational instability on a continental substrate (Barbey et al., 1984) and arc volcanism above a subduction zone that generated TTG magmatism, presumably during the regional D3 event (Piirainen, 1988). Luukkonen (1992) did not couple the mafic and ultramafic volcanism to granitoid generation and compressive deformation but proposed that the greenstone belt was initiated by rifting of a continental substrate due to the impingement of a mantle plume at the base of the lithosphere (cf. Campbell and Griffiths, 1992). This concept is consistent with the evidence for komatiitic and mafic sills truncating

previously deformed banded amphibolites in the northern part of the Kuhmo greenstone belt, and also the evidence for assimilation of crustal material recorded in the Kellojärvi cumulate complex. There is little doubt that plume impingement beneath late Archean continental crust could also trigger partial melting of the lower crust, given fertile rock compositions, in which case bimodal magmatism could also be explained. If this were the case, there ought to be geochemical and mineralogical evidence for melting at relatively low pressures, compared to slab or mantle wedge melting in subduction zones.

Tipasjärvi greenstone belt The Tipasjärvi greenstone belt is considered to be a southwards continuation of the Kuhmo and Suomussalmi greenstone belts (Figures 2.2 and 2.5). It forms two narrow branching belts of predominantly mafic and felsic volcanic rocks, each just under 30 km long, with a maximum width of 4 km. Depositional younging directions are sporadically preserved in the western branch (Taipale, 1988; Taipale et al., 1993) and indicate a tight synformal structure, possibly with tectonic repetition of stratigraphy as well. Intense sericitic and kyanite hydrothermal alteration is associated with the felsic pyroclastic deposits, including quartzphyric crystal tuffs, that host the Taivaljärvi Ag-Pb-Zn deposit (Kopperoinen and Tuokko, 1988; Papunen et al., 1989). Metamorphism of altered rocks has locally produced kyanitequartz and plagioclase-cordierite assemblages. According to Taipale (1983, 1988) and Kopperoinen and Tuokko (1988), the felsic volcanic rocks form the lower part of the stratigraphic sequence, defined as the Koivumäki Formation. The transition to the overlying Vuoriniemi Formation is marked by the onset of sporadic mafic volcanism, though a hiatus is indicated by the presence of a persistent horizon of sulfidic and graphitic siliceous pelites and magnetite facies BIF. These are overlain by tholeiitic basalts, basaltic tuffs, distinctive

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Cr-rich basalts and komatiitic lavas with spinifex structure, assigned to the Kallio Formation and finally by mica schists. Zircons from felsic volcanic rocks within the ore zone have been dated at 2791 ± 8 Ma, which is considered to be one of the more reliable age constraints on volcanism in the Kianta terrain (Vaasjoki et al., 1999). Galena, pyrite, and sphalerite from the Taivaljärvi deposit have also been analysed and are relatively homogeneous with respect to Pb characteristics. Moreover, the galena isotope composition lies along the same chord as whole-rock Pb-Pb results from unaltered host rocks, suggesting a close relationship between volcanism and mineralization (Vaasjoki et al., 1999). If correlation with the Kuhmo greenstone belt is attempted, the Tipasjärvi greenstone belt corresponds to the upper part of the sequence at Kuhmo. The Hattu schist belt shows many lithological similarities with the Tipasjärvi greenstone belt, and there too, mafic to ultramafic volcanic rocks occur towards the top of a substantial felsic volcanic and epiclastic succession. However, current age constraints preclude direct correlation, the lowest exposed rock units in the Hattu schist belt being some 40 Ma younger than at Tipasjärvi. An appraisal of the possibility of variable degrees of zircon heterogeneity, either inherited from source regions, or by wall-rock contamination during ascent and eruption, may be warranted, but is beyond the scope of this review.

Granitoids, gneisses, and crustal evolution in the Kianta terrain The granitoids and migmatites of the Kianta terrain have been studied intensively with respect to structural evolution (Luukkonen, 1985, 1988a, 1992) and petrogenesis (Martin et al., 1983a,b, 1984; Martin, 1987a,b). Resolution of apparent contradictions in isotope dating (Martin, 1989; Vaasjoki, 1989) has now led to a consensus from which it is clear that three stages of granitic magmatism are represented. The last of these is Paleoprote48

rozoic in age and consists of porphyritic granite plutons and dike swarms, in places with rapakivi feldspar texture, and is clearly discordant with respect to Archean orogenic structures (Figure 2.9). The U-Pb zircon date of 2435 ± 12 Ma (Luukkonen, 1988a) shows that these represent a bimodal aspect to the extensive mafic layered intrusive complexes, which can be traced across the northern end of the Ranua terrain, and along the boundary between the Koillismaa and Kianta terrains (Figures 2.10 and 2.12; Chapter 3). Regional mapping to the east of the Kuhmo greenstone belt has shown that this area is dominated by complex and diverse stromatic and nebulitic migmatites, typically tonalitic to trondhjemitic in composition (Figure 2.11C), with abundant enclaves of banded mafic amphibolites (Luukkonen, 1986, 1987, 1993). Compositionally, the banded amphibolites were Fe-rich tholeiites, of uncertain age, and were regarded by Luukkonen (1992) as disrupted remnants of an earlier mafic crust that was isotopically homogenized during a major melting event, accompanying amphibolite facies metamorphism and pervasive ductile deformation. This event was classified as D2 in the regional structural framework established by Luukkonen (1985, 1988b). Enveloping surfaces to D2 structures are typically gently to moderately dipping, though locally steeper, in contrast to the generally steep structures associated with the younger stages of deformation. The composite S1-S2 fabric and differentiated banding in the amphibolites is defined by dimensional alignment of plagioclase and hornblende or actinolite. This pervasive and widepsread tectonic and magmatic event has been dated by several isotope methods, including a Rb-Sr whole-rock isochron of 2.86 ± 0.09 Ga for tonalitic gneisses, corroborated by Sm-Nd studies (Martin et al., 1983a). The paleosome from banded migmatites at Lylyvaara also yielded a zircon U-Pb age of 2843 ± 18 Ma (Luukkonen, 1985). Martin et al. (1983a,b) and Martin (1987a,

• CHAPTER 2 • ARCHEAN ROCKS

b) referred to these banded migmatitic rocks as the Kivijärvi gray gneisses and have attempted to model their origin and crystallization. Compositionally they represent tonalites, trondhjemites and granodiorites, in which banding is due to variations in mafic mineral abundances, notably biotite and hornblende, with felsic minerals being plagioclase and quartz; potassium feldspar is present but rare. Chemically these rocks are typical for Archean TTG series granitoids, though notably peraluminous. Initial Sr isotope ratios are low (0.7023), close to the mantle evolution trend for the late Archean; Sm-Nd results and common lead data (Vidal et al., 1980) also militate against derivation of these rocks from a source dominated by isotopically evolved old continental crust. This places some constraints on petrogenetic models that require two-stage melting via a crustal source, rather than direct derivation from mantle rocks. Martin (1987a, b) used fractionation of REE as the basis for ascertaining likely source compositions for the TTG magmas, for degrees of partial melting considered realistic under Archean geothermal gradients. Some fractional crystallization of plagioclase and hornblende has evidently occurred, to explain the presence of granodioritic compositions, but has not had a significant effect on overall REE patterns, which tend to show pronounced HREE depletion, but no negative Eu anomaly. Martin (1987a,b) concluded that direct derivation from mantle rocks, whether modeled with spinel lherzolite or garnet lherzolite compositions failed to produce REE fractionation consistent with observed data, nor was high degree of melting of tholeiitic basalt under eclogite facies conditions appropriate. On the other hand, tholeiitic amphibolites containing 10–25% residual garnet, at 10–45% melting produced REE fractionation patterns, high La/Yb ratios and depleted Yb values within the range observed for the Kivijärvi gneisses. After concluding that the TTG magmas were extracted from a relatively young tholeiitic

source that underwent melting in the stability field of garnet and hornblende, Martin (1986, 1987a,b) then argued that the Archean orogenic geothermal gradient permitted melting of subducted ocean crust, before the subducted slab was completely dehydrated. This older generation of TTG intrusions apparently formed the basement to the supracrustal magmatism of the Kuhmo greenstone belt (Luukkonen, 1988a, 1992; Lukkonen and Sorjonen-Ward, 1998, Piirainen, 1988). This has been deduced primarily from truncation of the pre-D2 foliated mafic amphibolites by ultramafic dikes at Repolampi (Figure 2.11B), and an extensive differentiated mafic sill at Moisiovaara, which includes a pegmatoid gabbro phase with zircon dated at 2790 ± 18 Ma (Luukkonen, 1988a). No granitoid intrusions coeval with the felsic volcanism within the Kuhmo greenstone sequences have been specifically identified, whereas deformation of the greenstone belt during the regional D3 event of Luukkonen (1985, 1988a) was accompanied by widespread granitic magmatism (Figure 2.9). Some of these tonalitic to granodioritic plutons demonstrably intrude supracrustal units of the Kuhmo greenstone belt, particularly in the south and east of the region (Horneman et al., 1988), and have ages of 2739 ± 8 Ma and 2694 ± 13 Ma (Hyppönen, 1983). As well as discrete plutons, structurally controlled magmatism is characteristic of D3, represented by agmatites and neosomes intruded within axial surfaces of F3 folds, both within the Kuhmo greenstone belt and in the surrounding older migmatite terrain (Luukkonen, 1985, 1988a). Chemically and petrographically, the D3 intrusions resemble the earlier generation of TTG magmas (Martin et al., 1983a,b; Martin, 1987b; Horneman et al., 1988) and also have a low intial Sr isotope ratio (0.7024), suggesting similar melting conditions and sources. In general, contacts between the greenstones and the older TTG migmatites were

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Post-Archean rock units Kuopio (K-Kuo) and Kaavi (K-Kaa) kimberlite clusters Neoproterozoic redbed sequence

A

hja

o

o äp

Svecofennian granites with partial derivation from Archean crust Paleoproterozoic supracrustal units

am

Ranua terrain

r Pe

us Ku

Posio

Koillismaa terrain

Ranua Simo Kemi

Paleoproterozoic (2.45–2.39 Ga) mafic layered intrusions

Simo

Taivalkoski

Oijärvi Siurua

Kianta terrain

Pudasjärvi

Significant Svecofennian deformation zones Boundaries between Archean rock units

Simo granitic gneiss complex Oijärvi greenstone belt Siurua granulite and granite complex Highly strained supracrustal gneisses

Puolanka

Sii Siilinjärvi carbonatite complex Var Varpaisjärvi granulite complex Rau Rautavaara gneiss complex Man Manamansalo granitic gneiss complex Pir Pirttimäki granitic gneiss complex Kaj Kajaani granitic gneiss complex

Man

Ämmänsaari

u inu a K

Kajaani Kaj

Pir

Iisalmi terrain

Rau Iisalmi

Var

Rautavaara

Varpaisjärvi Sii

B

Siilinjärvi vo Sa

K-Kuo

K-Kaa

Kuopio Outokumpu

0

50 km

C

50

• CHAPTER 2 • ARCHEAN ROCKS

obscured, or tectonically modified during D3 such that the boundary zone is typically marked by leucocratic medium-grained foliated and nebulitic cataclastic tonalite (Luukkonen, 1988a,b). The D3 event has been responsible for imparting the presently observed structural architecture of the Kuhmo greenstone belt, in which N–NE dextral transpression and E–W compression has produced a combination of fold interference patterns and dextral brittle– ductile shear zones (Figure 2.9). Neosomes are also present in generally NW-trending D4 shears and have been dated at 2657 ± 32 Ma (Luukkonen, 1985). Horneman and Hyvärinen (1989) also distinguished a diverse range of plutonic rock types surrounding the Tipasjärvi greenstone belt, including a continuum from stromatic and nebulitic migmatic gneisses, to discrete plutons, varying in composition from tonalite to monzogranite. The nature of contact relationships between supracrustal rocks and the granitoids is generally equivocal or unknown, although deformation events are shared by both the greenstones and granitoids (Horneman and Hyvärinen, 1989). This is unfortunate, as for example, the Haasianvaara tonalite along the northwestern margin of the greenstone belt has a concordant and rather precise age of 2830 ± 2 Ma (Horneman and Hyvärinen, 1989; Vaasjoki et al., 1999). A somewhat younger age of 2826 ± 14 Ma has been obtained for lithologically similar tonalites along the southeastern margin of the greenstone belt, at Huuskonvaara (Vaasjoki et al., 1999). If these dates relate to the emplacement age, then it is quite significant, being a potential example of older basement, in either depositional or tectonic contact with the Tipasjärvi supracrustal sequence. Horneman and Hyvärinen (1989) speculated that the

Haasianvaara tonalite and its amphibolitic and felsic supracrustal enclaves may correlate with the 2.86–2.83 Ga amphibolite facies deformation and migmatite event in the northern part of the Kianta terrain (Martin et al., 1983a; Luukkonen, 1985, 1988a, 1992). However, if a basement – cover interface is preserved in this area, the resultant structural pattern becomes quite complex, and considerable tectonic displacement might be invoked. Each of the tonalitic rock types defined by Horneman and Hyvärinen (1989) show major and trace element trends typical for the late Archean tonalite–trondhjemite association, despite the significant differences in age. Horneman (1990) attributed the REE characteristics of the tonalitic magmatism to melting of mafic lower crust under amphibolite facies conditions such that garnet and hornblende were retained in the source (cf. Martin et al., 1983b; Martin, 1986, 1987a). The assumed mafic source is believed to have been enriched in incompatible elements over MORB tholeiites, due to hydrothermal alteration or metasomatism and interaction with overlying crust during slab dehydration and melt migration (Condie, 1986; Martin, 1987b). More nebulitic and deformed tonalitic intrusions, that appear to be transitional into the supracrustal Nurmes gneiss complex, occur to the south of the Tipasjärvi greenstone belt. For these rocks, a metasomatized mantle composition enriched in LREE would be also be an appropriate source, although modeled chemistry would better match melting at somewhat greater depth, with garnet dominating in the residual phase. This Halmejärvi-type of tonalitic magmatism (Horneman and Hyvärinen, 1989) has been dated to 2745 ± 8 Ma, which is very similar to that of the tonalites intruding

Fig. 2.12. (facing page) Principal features of the Ranua and Iisalmi terrains of the Karelian domain. (A) Distribution of major crustal units described in this review. (B) Folded stromatic migmatite of the Simo terrain. Kuivaniemi, near Simo. (C) Typical examples of complex interaction between deformation and magmatic processes within migmatites of the Simo terrain. Kuivaniemi, near Simo. Compass diameter is nearly 7 cm. Photos: Peter Sorjonen-Ward. CHAPTER

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the Hattu schist belt (Vaasjoki et al., 1993). Horneman (1989) also recognized two distinct groups of felsic granitoids, one of which can be seen as part of a compositional continuum from tonalite to trondhjemite, derived from the same amphibolitic source, or by melting of tonalite during a later event; similarities in zircon ages (Horneman and Hyvärinen, 1989; Vaasjoki et al., 1999) would favor the former alternative. The second group of felsic granitoids ranges from granodiorite to monzogranite, typically with more fractionated REE patterns and Eu depletion, attributed to melting of the tonalite–trondhjemite series granitoids during a later event, with preferential retention of plagioclase in the source. This intepretation is supported by their late position in the sequence, heterogeneity of zircon populations and the considerable scatter in Pb isotope data from feldspars (Halla, 1998, 2002), although the effect of Svecofennian thermal and tectonic events in this area needs to be considered as well.

Nurmes gneiss complex Kontinen (1991) proposed that metasedimentary gneisses and stromatic to nebulitic migmatites form a major component of the western and southern Kianta terrain (Figure 2.5), and that they can be distinguished both chemically and texturally from migmatites and gneisses of plutonic origin. At outcrop scale, lithological and compositional banding is evident, suggestive of relict depositional layering, despite the presence of concordant leucotonalitic leucosomes, which might be an expression of in situ partial melting as well as externally derived melt injection. Granoblastic biotite-plagioclase gneisses are predominant, with additional alternations between quartzplagioclase and thinner garnet-biotite-plagioclase layers, the latter commonly containing relatively abundant graphite and sulfides. Chemical data from paleosomes were interpreted by Kontinen (1991) to reflect a combination of degree of hydraulic sorting, 52

weathering, and provenance composition. The positive correlation between Mg+Fe and K is taken to indicate the separation of clay and sand during sedimentary processes, whereas negative correlation between Mg+Fe and Ca is seen as evidence for chemical weathering and as a useful criterion for confirming a sedimentary rather than igneous origin for the gneisses. High Cr, Ni, and V contents at elevated (67–68 wt.%) SiO2 contents are also an indication of both a sedimentary origin, and a mixed felsic and mafic provenance, characteristic of many Archean gneiss terrains (cf. Taylor and McLennan, 1985; Sawyer, 1986). When data are normalized against the PostArchean Australian Shale (PAAS; Taylor and McLennan, 1985), the Nurmes gneisses are seen to be impoverished in large ion lithophile elements, suggesting a more primitive source. On the other hand, data are very similar to those published from the Quetico belt, which has been interpreted as a fore-arc sequence derived from a juvenile calc-alkaline arc (Sawyer, 1986). The Nurmes gneisses also compare well with the metasediments of the Hattu schist belt (O’Brien et al., 1993a). There are no constraints on polarity of deformation, nor have structural relationships with other elements of the Kianta and Ilomantsi terrains been established, so that it is not yet possibly to determine whether the Nurmes gneiss complex could represent an accretionary prism related to the Ilomantsi terrain. Provisional age data, indicating a minimum depositional age of 2720 Ma and evidence that 2.68 Ga Konivaara-type granodiorites truncate gneissic banding (Asko Kontinen, pers. comm., 2002) are at least consistent with such a hypothesis. Vaasjoki et al. (1999) dated pelitic gneiss enclaves from within migmatitc tonalites located close to the boundary zone between the Nurmes gneiss terrain and the Tipasjärvi greenstone belt. Ages obtained, although not precise (2748 ± 10 Ma or 2715 ± 20 Ma, depending on which fractions are assigned greater significance),

• CHAPTER 2 • ARCHEAN ROCKS

nevertheless fall within a range appropriate for Ilomantsi terrain provenance, rather than Kianta terrain volcanism. Reconnaissance data on enveloping surfaces of structures and lithological layering (Taipale et al., 1993; Asko Kontinen and Erkki Luukkonen, unpublished data) support the concept of combined N–NEthrusting and transpression, both within and along the boundary zone between the Kianta and Ilomantsi terrains (Figure 2.5).

2.3. Iisalmi terrain The Iisalmi terrain (Figures 2.2, 2.5, and 2.12) is particularly important in that the Varpaisjärvi granulite complex, in the western part of the terrain, includes the best documented Archean granulite facies rocks in Finland (Paavola, 1984; Hölttä, 1997; Hölttä and Paavola, 2000). In addition, the terrain records both some of the oldest and youngest Archean events in the Fennoscandian Shield, namely paleosomes of magmatic gneisses dated to nearly 3.2 Ga (Paavola, 1986; Hölttä et al., 2000a; Mänttäri and Hölttä, 2002), and the 2.6 Ga Siilinjärvi carbonatite complex (Puustinen, 1971; Patchett et al., 1981; Lukkarinen, 2000a). The existence of two distinct Paleozoic kimberlite provinces within the Iisalmi terrain also provides an excellent opportunity for investigating the composition and thermal evolution of the deep crust over time (Kukkonen and Peltonen, 1999; Peltonen et al., 1999; Hölttä et al., 2000b). Crustal-scale seismic refraction studies have also shown that the Moho beneath the Iisalmi terrain is unusually deep, with an estimated present crustal thickness of 55–60 km, compared to more typical values of around 40 km beneath the Kianta terrain (Korja et al., 1993; Korsman et al., 1999). This anomalous crustal thickness is likely a consequence of several processes, including thrust stacking during Svecofennian collision and post-collisional underplating, the latter being inferred from by U-Pb zircon ages obtained from mafic lower crustal xenoliths in kimberlites (Hölttä

et al., 2000b). However, the negative initial εNd values recorded by Svecofennian granitoids intruding the Iisalmi terrain (Huhma, 1986; Ruotoistenmäki et al., 2001) also indicate partial melting of deep Archean crust during the later stages of the Svecofennian orogeny, at 1.86–1.85 Ga. In the western part of the Iisalmi terrain, these intrusions appear to show brittle intrusive features, while Svecofennian resetting of K-Ar biotite ages (Kontinen et al., 1992) and epidote-albite assemblages in retrograde shear zones suggest a greenschist facies overprint at the present erosion level (Figure 2.13E and F). Therefore, even with a modest late orogenic geotherm, partial melting of fertile Archean rocks at around 800 ºC could have occurred at depths of 15–20 km below the present erosion level. It is also important to emphasize that Paleoproterozoic dike swarms demonstrably truncate metamorphic boundaries within the Iisalmi terrain, while recording the Svecofennian greenschist facies overprint. This clearly demonstrates not only that the granulite facies metamorphism recorded in parts of the Iisalmi terrain is of Archean age, but also that exhumation of the granulites was not merely a consequence of Svecofennian collisional processes.

Proterozoic reworking and the boundaries of the Iisalmi terrain The boundaries of the Iisalmi terrain at the present erosion level nevertheless substantially reflect Proterozoic events (Figure 2.12). It is indeed possible that it has been displaced in its entirety with respect to other terrains of the Karelian domain, as first suggested by Väyrynen (1939). The Iisalmi terrain is separated in the northwest from the Ranua terrain by the N–NE-trending Oulujärvi shear zone (Kärki et al., 1993). Within this broad deformation zone, Archean rocks have been tectonically reworked and emplaced over Proterozoic rocks of the Kainuu schist belt and intruded by Proterozoic granites, such that they form several isolated units (Kärki et

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al., 1993; Kontinen, 1993; Lukkarinen, 2000a; Vaasjoki et al., 2001), notably the Pirttimäki, Manamansalo, and Kajaani migmatitic granitic complexes (Figure 2.12). Likewise, the southern and western margin of the Iisalmi terrain at the present erosion level can only be defined in terms of the effects of Proterozoic tectonic reworking. In this region, the Archean–Proterozoic interface, commonly marked by a recognizable depositional unconformity (Korkiakoski and Laajoki, 1988; Pietikäinen and Vaasjoki, 1999; Lukkarinen, 2000a), has been deformed into complex domal interference patterns. This pattern has been variously attributed to diapiric instability (Eskola, 1949; Brun, 1980), or fold interference (Park, 1981); irrespective of origin the internal structures of these domes commonly retain coherent structures of doubtless Archean origin (Park, 1981). More specifically, Proterozoic overprinting can be interpreted as a consequence of thrusting, followed by dextral transpression within a network of shear zones that juxtapose Svecofennian assemblages directly against Archean gneisses, with locally intense transposition of Archean structures (cf. Park and Bowes, 1983; Ward, 1984; Kohonen et al., 1991; Paavola, 1991; Kärki et al., 1993; Lukkarinen, 2000a,b). In proximity to some shear zones, strong epidotization and albitization has transformed the banded migmatites into almost massive pale-colored mylonitic rocks, many with folding and displacements showing a consistent dextral shear sense (Paavola, 1991). The effect of Proterozoic tectonic and thermal overprinting increases eastwards through the Iisalmi terrain such that the prominent south-plunging linear fabric observed within much of the Rautavaara complex (Paavola, 1980, 1997, 1999) appears to be Svecofennian in origin. Supracrustal gneisses of the Rautavaara complex record a multiphase history, resolved as earlier medium-pressure assemblages superimposed by an amphibolite facies retrogressive event (Hölttä and Paavola, 54

Fig. 2.13. (facing page) Representative rock types from the Iisalmi terrain. (A) Two-pyroxene garnet amphibolite from Kumisevanmäki, near Sonkajärvi, within the Varpaisjärvi granulite complex. Scale bar is approximately 1 dm in length. Photo: Jorma Paavola. (B) Quartz-chlorite-cordierite assemblage in hydrothermally altered schists derived from an inferred mafic protolith, at Lumimäki, within the Rautavaara gneiss complex. These rocks typically also record a Proterozoic “retrograde” amphibolite facies history, superimposed upon medium-pressure late Archean metamorphism. Scale bar is approximately 1 dm in length. Photo: Jorma Paavola. (C) Coarse potassium feldspar phenocrysts are typical of granodioritic to quartz dioritic intrusions within the Rautavaara complex and Ilomantsi terrain. Scale bar is approximately 1 m in length. Photo: Jorma Paavola. (D) Archean megacrystic granitoid deformed to highly strained mylonite during the Paleoproterozoic Svecofennian orogeny, with feldspar porphyroclasts representing relict phenocrysts. This strain state is typical over much of the eastern part of the Iisalmi terrain. Scale bar is approximately 1 dm in length. Photo: Peter Sorjonen-Ward. (E) Glimmerite–carbonatite within open pit at Kemira Oy Siilinjärvi apatite mine. Note Proterozoic mafic dikes truncating vertical banded fabric, indicating limited Proterozoic tectonic overprint in western part of Iisalmi terrain. Photo: Peter Sorjonen-Ward. (F) Detail of brittle–ductile carbonatite veins intruding glimmerite zone of the Siilinjärvi carbonatite complex, derived from alteration of Archean granitoid gneisses. Note sharply truncated Proterozoic mafic dike, indicating that Proterozoic tectonic and thermal overprint was limited. Scale bar is approximately 1 dm in length. Photo: Peter Sorjonen-Ward.

• CHAPTER 2 • ARCHEAN ROCKS

A

B

C

D

E

F

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2000). As Proterozoic mafic dikes are typically highly strained and have been recrystallized under amphibolite facies conditions, it is therefore reasonable to conclude that the enclosing Rautavaara complex gneisses also record this Proterozoic amphibolite facies event. Along the eastern margin of the Rautavaara complex (Figure 2.5), Archean migmatitic gneisses have been tectonically emplaced over inverted Proterozoic sequences during the Svecofennian orogeny (Frosterus and Wilkman, 1920; Väyrynen, 1939; Park and Bowes, 1983; Ward and Kohonen, 1989), followed by emplacement of granitic sheets with Sm-Nd attributes indicative of an Archean provenance (Huhma, 1986). Archean porphyritic granitoids (Figure 2.11C) have been deformed into mylonitic augen gneisses (Figure 2.11D) in which potassium feldspar shows Pb-Pb characteristics consistent with Proterozoic lead loss from an evolved radiogenic precursor (Halla, 1998). A zone of intense sinistral transpressive deformation (Ward and Kohonen, 1989; Kohonen et al., 1991) towards the eastern margin of the terrain makes it difficult to define the nature and location of the boundary with the Kianta terrain. Hence the Nunnanlahti greenstone belt, being structurally allochthonous could be assigned to the Iisalmi terrain as well as the Kianta and Ilomantsi terrains (Figure 2.5).

Origin of the present metamorphic zonation pattern The position of the Iisalmi terrain at the boundary zone between the Svecofennnian and Karelian domains also means that Proterozoic thermal and tectonic effects superimposed on the Archean bedrock can be studied in detail (Paavola, 1986; Toivala et al., 1991; Kontinen et al., 1992). However, this has made relationships with other Archean terrains more difficult to establish. Similarly, the nature and timing of juxtaposition of granulite facies units with lower grade rocks within the terrain itself is not entirely clear. At the present 56

erosion level the granulites are bounded by discrete faults for which variable strike-slip, oblique-slip and dip-slip displacements have been documented (Paavola, 1984; Hölttä, 1997). There is nevertheless compelling evidence from which an Archean rather than Proterozoic origin for the observed pattern of metamorphic zonation may be inferred. Firstly, primary magmatic minerals such as orthopyroxene, and brittle intrusive features and chilled margins are commonly observed in Proterozoic dikes, dated from 2.3–2.1 Ga (Toivala et al., 1991), in both granulite facies and lower grade rocks in the western part of the Iisalmi terrain. Secondly, Archean K-Ar ages are recorded for hornblendes and some biotites from the Varpaisjärvi granulites. In the adjacent amphibolite facies magmatic gneisses Archean hornblende ages are preserved, whereas biotite ages were reset by a Proterozoic thermal event (Paavola, 1986; Kontinen et al., 1992). This suggests that the western part of the Iisalmi terrain cooled coherently below the biotite blocking temperature during the late Archean; the preservation of Archean biotite ages in the granulites is attributed to their relatively anhydrous nature (Kontinen et al., 1992). Thirdly, the Siilinjärvi carbonatite complex was intruded into granitic gneisses of the Iisalmi terrain at 2.61–2.58 Ga (Puustinen, 1971; Patchett et al., 1981; Lukkarinen, 2000a). Calcite-dolomite equilibria (Puustinen, 1974) and fluid inclusion data from apatite and zircon (Poutiainen, 1995) suggest final emplacement and equilibration in a greenschist facies environment, which is consistent with the brittle–ductile deformation style recorded by the apatite bodies and glimmerite (Figure 2.13E and F). Neither the carbonatite itself, nor its fenitic alteration aureole are in direct contact with granulites. However, because the north-south trend of the carbonatite complex is oblique to the metamorphic zone boundaries (Figure 2.12), it is likely that the carbonatite magmatism was related to a separate deformation phase, post-dating the

• CHAPTER 2 • ARCHEAN ROCKS

event that produced the presently observed distribution of metamorphic domains. This is also consistent with the lack of evidence for earliest Proterozoic (2.5–2.0 Ga) magmatism or metamorphic cooling events in U-Pb data from both the Varpaisjärvi granulites (Hölttä et al., 2000a), and mafic lower crustal xenoliths extracted from kimberlites (Hölttä et al., 2000b). All of the above evidence suggests that the rocks at the present erosion level in the western part of the Iisalmi terrain experienced a rather modest greenschist facies overprint during the Svecofennian orogeny. If Proterozoic deformation had exhumed the Varpaisjärvi granulites from deeper levels – whether by thrusting or extensional processes – then the orogenic geotherm ought to be recorded in resetting of K-Ar system in hornblende or UPb in titanite (cf. Bibikova et al., 2001). This is clearly not the case (Kontinen et al., 1992; Hölttä et al., 2000a). Therefore, exhumation of the Varpaisjärvi granulites by listric faulting and attenuation of the Karelian domain during Paleoproterozoic rifting (Ward and Kohonen, 1989) is unlikely. In that case, the observation by Paavola (1991), that mafic dikes tend to be more abundant in granulite facies rocks, must be attributed to rheological contrasts, rather than implying that dike abundance relates to crustal depth at the time of emplacement.

Varpaisjärvi granulite complex Much of the Iisalmi terrain consists of tonalitic and trondhjemitic granitoids and migmatites, with variable amounts of concordant enclaves of amphibolite. These amphibolite zones may be hundreds of meters wide, either homogeneous or banded and contain ultramafic enclaves (Paavola, 1988, 1991). Quartz dioritic paleosomes in some of these banded granitoid gneisses have yielded U-Pb zircon ages of 3136 ± 20 Ma and 3095 ± 18 Ma (Paavola, 1986), making them some of the oldest rocks exposed in the Fennoscandian Shield. These ancient rocks occur in close

proximity to the enderbites of the Varpaisjärvi granulite complex, which have U-Pb zircon ages around 2.7 Ga (Paavola, 1986). The granulite facies rocks are conspicuous in regional aeromagnetic data, and tend to define discrete fault bounded blocks. Recent Sm-Nd isotope studies (Hölttä et al., 2000a) have confirmed that there are distinct crustal subdivisions within the granulites themselves, while also demonstrating that the protoliths to at least some of the enderbitic granulites have ages up to 3.2 Ga. Nevertheless, some enderbites have U-Pb zircon ages, inferred to represent crystallization, as young as 2.68 Ga (Paavola, 1986), while U-Pb ages from zircons and monazites are both interpreted to constrain the peak of granulite metamorphism to around 2.63 Ga. In addition to the predominant hypersthene-bearing enderbites, which range in composition from diorite to tonalite, two-pyroxene amphibolites, possibly of volcanogenic origin are present (Figure 2.13A), as well as garnet-cordierite-sillimanite and quartzcordierite rocks (Paavola, 1984, 1988, 1991; Hölttä, 1997). The Varpaisjärvi granulites do not appear to represent typical restitic and depleted lower crustal compositions following melt extraction. The mafic granulites intercalated among the enderbites were classified into two distinct geochemical types by Hölttä (1997), and this distinction seems to be reflected isotopically as well (Hölttä et al., 2000a). Mafic rocks in the Jonsa block, which has a younger Sm-Nd model age than other Varpaisjärvi granulites, show a greater degree of compositional variation. Moreover, they are associated with quartz-cordierite and cordierite-orthoamphibole-orthopyroxene rocks that apparently represent the metamorphic derivatives of basalts and andesites that were hydrothermally altered by interaction with seawater. Locally, these distinctive rock compositions have resulted in unusual mineral assemblages, including sapphirine and kornerupine. Hölttä and Paavola (2000) documented

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a two-stage Archean metamorphic history, suggesting isothermal decompression and speculate that this relates to terrane accretion and crustal thickening. The first metamorphic event was defined at 9–11 kbar and 800–900 ºC, accompanied by partial melting, such that the present garnet-plagioclase-pyroxene assemblages are considered restitic. Melting progress was evidently dependent upon compositional differences, with a greater abundance of neosomes in Fe-rich intermediate rocks than in more mafic Mg-rich rocks. Melting reactions may have been promoted by decompression, with equilibration at 7 kbar and 700 ºC. Hölttä and Paavola (2000) also considered the possibility that emplacement of the enderbites might have been responsible for the regional scale contact metamorphism in the lower crust. However, they concluded that the enderbites were intruded up to 50 Ma prior to the granulite facies metamorphic peak and attribute the irregular distribution of granulite facies assemblages to fluid availability and infiltration, rather than lateral variations in heat distribution.

Rautavaara complex The Rautavaara complex forms the eastern part of the Iisalmi terrain and records intense Proterozoic tectonic and thermal reworking, which becomes progressively stronger eastwards, where Archean rocks are demonstrably allochthonous and have been emplaced over Proterozoic sediments (Frosterus and Wilkman, 1920; Park and Bowes, 1983). Archean structures may therefore be difficult to distinguish from Proterozoic overprinting, especially given that Proterozoic mafic dikes have commonly been sheared and transposed into near concordance with gneissic banding. Proterozoic thermal overprinting is also recorded in Pb isotope compositions of potassium feldspar augen in deformed megacrystic granites (Halla, 1998), which are widespread in the Rautavaara complex (cf. Frosterus and Wilkman, 1920). 58

While tonalitic-trondhjemitic migmatites (orthogneisses) are also typical of the Rautavaara complex, the most distinctive feature is the relative abundance of metasedimentary and metavolcanic paragneisses, many of which have been hydrothermally altered (Paavola, 1999). This is expressed mineralogically as assemblages containing kyanite (locally also andalusite and sillimanite), cordierite, amphibole, staurolite, and tourmaline. Quartz-chlorite assemblages are also widespread (Figure 2.13B). The intense Proterozoic overprint has made it difficult to document the initial Archean metamorphic regime, although it appears likely that there has been a significant retrograde equilibration from high-grade Archean assemblages. While expressing concerns about the possibility of Proterozoic disturbance to U-Pb isotope systems, Paavola (1999) conceded that ages obtained from the Rautavaara complex are significantly younger than those from the Iisalmi terrain to the west and the Kianta and Ilomantsi terrains to the east. Tonalite yielded a zircon age of 2677 ± 10 Ma and a porphyritic granite 2657 ± 15 Ma; these results do not differ greatly from the younger granite and neosomes ages reported from the Kianta terrain (Martin et al., 1983a,b; Luukkonen, 1985; Vaasjoki et al., 1999). On the other hand, xenotime from a quartzite provided a concordant age of 2616 Ma, which might still be interpreted as a cooling age, especially given the 2.63 Ga estimates for peak granulite facies metamorphism in the adjacent Varpaisjärvi granulites. The most surprising and anomalous data came, however, from an altered metasediment, with an age of 2657 ± 20 Ma. If this result represents a mixed detrital population, and does not record metamorphic or magmatic zircon growth, then these hydrothermally altered rocks are the youngest Archean supracrustal rocks yet found in Finland and have some significance in interpreting the timing of juxtaposition of the Iisalmi and Kianta and Ilomantsi terrains.

• CHAPTER 2 • ARCHEAN ROCKS

2.4. Ranua terrain The Ranua terrain is an essentially triangular Archean block at the northwestern margin of the Karelian domain (Figures 2.2 and 2.12). To the east, it is separated from the Koillismaa and Kianta terrains and the intervening Paleoproterozoic Kainuu schist belt by the complex Paleoproterozoic Hirvaskoski and Oulujärvi shear zones (cf. Kärki et al., 1993), while along its northwestern and southwestern margins, it is unconformably overlain by Paleoproterozoic sedimentary and volcanic sequences. The northwestern margin of the terrain, beneath the unconformity with the Peräpohja supracrustal belt, was also intruded at 2.4 Ga by the Kemi–Penikat–Portimo suite of mafic layered intrusions (Figure 2.12; Chapter 3). This extensive region remains one of the least understood areas in Finland, largely due to a combination of poor exposure, and the relatively monotonous nature of the predominantly tonalitic to granitic gneisses. There have been very few geological studies since reconnaissance mapping documented by Wilkman (1931) and Enkovaara et al. (1953), and these have concentrated mostly on the relationships with surrounding Proterozoic rock units (Perttunen, 1991). Mention should also be made of the eastern and southeastern margin of the Ranua terrain, where Proterozoic tectonic reworking and granitic magmatism within the Hirvaskoski and Oulujärvi shear zones has been substantial. The heterogeneous nature of strain in this zone has made it very difficult to unequivocally separate Archean and Svecofennian structural events and neosomes. Although an increasing number of isotope age determinations indicate that both ages are represented (Pietikäinen and Vaasjoki, 1999; Vaasjoki et al., 2001), the problem of zircon inheritance in felsic rocks remains. The same applies to attempts to resolve depositional ages of some contentious siliciclastic sediments in this region, as even if they were deposited during

the Proterozoic, they are likely to have zircon age spectra identical to that of their late Archean source area. This question surrounds an extensive and distinctive sequence of metasedimentary gneisses, known collectively as the West Puolanka gneisses and the Central Puolanka Group (Laajoki, 1986). Isotope studies have derived Archean Sm-Nd model ages for pelitic sediments (Kontinen et al., 1996), which accords with dating of zircons from siliciclastic sediments and inferred felsic volcaniclastic deposits (Huhma et al., 2000). If this is correct, then the West Puolanka gneisses and Central Puolanka Group form a distinct supracrustal unit along the eastern margin of the Ranua terrain. An Archean affinity would also be consistent with evidence accruing from a number of potentially correlative felsic volcanic and sedimentary units farther to the north (Räsänen and Vaasjoki, 2001; Räsänen and Huhma, 2001; Evins et al., 2000, 2002). Alternatively, Laajoki (Chapter 7) provides an evaluation of the evidence in favor of a Proterozoic depositional and eruptive age for the Central Puolanka Group. High-resolution aeromagnetic data became available for the Ranua terrain relatively recently, leading to the delineation of a discrete supracrustal belt in the western part of the terrain. This is now known as the Oijärvi greenstone belt and is reviewed in more detail below. Little progress has been made in subdividing and classifying the remainder of the Ranua terrain, although it is suggested here that the migmatitic tonalitic gneisses and granites to the west of the Oijärvi belt be designated as the Simo complex (Figure 2.12A). Rock types range from stromatic migmatites showing complex deformation and multiple stages of leucosome development (Figure 2.12B and C) to discrete granodioritic and tonalitic plutons that appear to be intimately associated with deformation of the Oijärvi greenstone belt. There are at present no constraints on the relative – or absolute – ages of the migmatitic and discrete plutonic units.

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Oijärvi greenstone belt As with much of the Ranua terrain, this region is poorly exposed and was one of the last areas in Finland to be covered by comprehensive airborne geophysical surveys. When data became available, the Oijärvi greenstone belt was clearly discernible magnetically as a narrow, anastomosing feature than could be traced for over 80 kilometers along strike, before it is obscured beneath unconformably overlying Paleoproterozoic sediments (Figure 2.12). Because of the obvious analogy with the Archean greenstone and schist belts in the Kianta and Ilomantsi terrains, the Geological Survey of Finland commenced a reconnaissance mapping and drilling program, which has provided some insights into the nature and distribution of rock units. The central part of the greenstone belt, in an area known as Karakkalehto, has been considered more prospective for gold and is consequently better understood than elsewhere (Tolppi, 1999). In this area, the greenstone belt appears to bifurcate, anastomosing around a large granitic intrusion (Figure 2.12A). The eastern contact of the greenstones is highly strained, but is likely to have been defined by an intrusive granitoid, rather than depositional basement. Small tonalitic intrusions and porphyritic dikes also clearly intrude the greenstones. There is a distinct lithological asymmetry to the belt, in that pillowed mafic and massive ultramafic volcanic rocks, with minor graphitic interflow sediments are more abundant in the east, whereas pelitic and graphitic schists and turbidites characterize the western part. However, a regional stratigraphic framework has yet to be defined, even though primary depositional and eruptive features are locally preserved; paucity of outcrop is a greater impediment to regional mapping than intensity of deformation. Localized zones of higher strain have been delineated, with particular prominent linear fabrics and are closely associated with hydrothermal alteration and quartz-carbonate brecciation. 60

Tolppi (1999) classified the mafic and ultramafic volcanic rocks into chemically distinct groups, including Fe- and Mg-tholeiites, Cr-rich basalts, basaltic komatiites, and ultramafic komatiites. The Cr-rich basalts generally resemble Mg-tholeiites, except that they are considerably enriched in Cr (450–4200 ppm), Ni (around 500 ppm) and have higher Al2O3/TiO2. The presence of these rocks in the Oijärvi greenstones is of interest from the perspective of regional correlation with the Kuhmo greenstone belt, where similar rocks have been described (Halko-aho et al., 2000). Tolppi (1999) considered that gold mineralization and alteration took place below the biotite isograd near peak metamorphism. Tolppi (1999) also observed a static porphyroblastic amphibolite facies overprint, with hydrothermal sericite and chlorite partially replaced by muscovite and biotite, almandine replacing chlorite and quartz in Fe-rich rocks, and tremolite and cummingtonite overprinting ultramafic talc-bearing assemblages. Because dolerite dikes are seen to truncate alteration fabrics, but also record the amphibolite facies metamorphism, Tolppi (1999) considered this latter metamorphic event as Svecofennian. This conclusion is clearly of regional significance, both in understanding Proterozoic geodynamic history, as well as Archean tectonic and thermal evolution.

Siurua granulite complex Detailed characterization and subdivision of the Ranua terrain to the east of the Oijärvi greenstone has not yet been attempted. However, in the Siurua area (Figure 2.12), Enkovaara et al. (1953) described a number of narrow zones, a kilometer or less in width and up to 10 in length, consisting of various granulite facies assemblages. These include mafic, Fe-rich granulites with hypersthene and hedenbergite and more felsic gneisses with cordierite-plagioclase-quartz assemblages, intruded by granitic neosomes of garnet-plagioclasequartz containing abundant magnetite, zircon,

• CHAPTER 2 • ARCHEAN ROCKS

and apatite. This is suggestive of sediments that may have been previously hydrothermally altered or compositionally modified by melt extraction, as well as local partial melting. There are therefore similarities to with the granulites of the Iisalmi terrain (Hölttä, 1997; Hölttä and Paavola, 2000). Mutanen and Huhma (2003) dated a trondhjemitic gneiss from Siurua and obtained an age of 3500 Ma, from a somewhat discordant and heterogeneous zircon population. This is currently one of the oldest rocks identified in the Fennoscandian Shield. The whole-rock Sm-Nd model age (TDM) is 3.48 Ga, which clearly supports the inference from the zircon data. The Siurua granulites as mapped by Enkovaara et al. (1953) define a narrow northerly trending zone nearly 50 long, subparallel to and some 20 east of the Oijärvi greenstone belt. As noted earlier, Tolppi (1999) found that the late Archean metamorphic peak in the Oijärvi greenstone belt was at greenschist facies conditions (250–400 ºC and 1.5–2.5 kbar). Hence there is either a tilted crustal section, exposing deeper crustal levels to the east, or there has been a late orogenic tectonic juxtaposition of different crustal units. The former interpretation is difficult to reconcile with the vertical foliations and southerly plunges documented for the Siurua granulite complex (Enkovaara et al., 1953), though data are admittedly few, while the latter interpretation would be reminiscent of current interpretations from the Iisalmi terrain (Hölttä and Paavola, 2000).

3. The Karelian domain in northern Finland In contrast to the situation in the central part of the country, the Archean rocks of northern Finland do not form extensive, coherent terrains, but are exposed as isolated basement windows, or have been substantially modified and disrupted by Proterozoic magmatism

and deformation. It is therefore difficult to discuss Archean geology in isolation from the superimposed effects of various Proterozoic processes and events. For the purposes of this review, all Archean rocks to the south and west of the main frontal thrust of the Paleoproterozoic Lapland granulite belt are assigned to the Karelian domain, with the exception of the Ropi terrain (Figures 2.1A and 2.3). The Lapland granulite belt represents the consequences of collision between the Kola and Karelian domains (Hörmann et al., 1980; Barbey et al., 1984; Marker, 1985; Gaál et al., 1989). Deep seismic and electrotelluric studies (Behrens et al., 1989; Korja et al., 1989; Luosto et al., 1989) indicate that the Lapland granulite belt was emplaced over the Karelian domain along a gently dipping detachment that can be traced at least to middle crustal depths. Isotope age data from metaigneous and metasedimentary granulites (Meriläinen, 1976; Huhma, 1986; Sorjonen-Ward et al., 1994) and cross-cutting plutons constrain this collisional event to between 1.91 Ga and 1.78 Ga. Seismic and gravity studies have also been used to infer the presence of felsic Archean basement at relatively shallow depths beneath the Lapland greenstone belt (Elo et al., 1989; Gaál et al., 1989). This is consistent with the presence of Archean basement windows surrounded and intruded by Proterozoic rocks, as in the Pomokaira terrain (Mikkola, 1941; Räsänen et al., 1989), and the abundant evidence for Archean isotope inheritance in Proterozoic granitic rocks in southern and western Lapland (Huhma, 1986; Öhlander and Skiöld, 1994; Perttunen and Vaasjoki, 2001; Väänänen and Lehtonen, 2001). The present distribution of Archean rocks in northern Finland thus represents complex reworking during the Svecofennian orogeny and collision between the Kola and Karelian domains. Some of the geographical terrains described here have been defined principally because the original late Archean relationships are not demonstrable, and may therefore also coincide

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with Proterozoic tectonic boundaries.

3.1. Koillismaa terrain This terrain is bordered to the north and west by the Paleoproterozoic Kuusamo supracrustal belt, while to the south, a discontinuous arcuate zone of tectonically disrupted, differentiated layered mafic intrusions (Alapieti, 1982) separates it from the Kianta terrain. Although the extent of Proterozoic tectonic reworking within the Koillismaa terrain is not clear, the prominent change in structural trend, from NNNE in the northern end of the Kianta terrain, to ESE in the Koillismaa terrain is considered sufficient justification for classification as a separate structural unit. This change in foliation trend coincides with prominent gravity and magnetic anomalies, at least part of which can be attributed to the layered intrusions or their subsurface continuations (Alapieti, 1982; Elo, 1992; Airo, 1999). Basal sedimentary units of the Kuusamo supracrustal belt were deposited unconformably upon quartz dioritic to trondhjemitic orthogneisses and mafic to pelitic paragneisses of the Koillismaa terrain (Silvennoinen, 1972, 1973, 1989, 1991). In addition to basement clasts, mafic and felsic detritus derived from the Paleoproterozoic layered intrusions and associated bimodal volcanic rocks have been found. The current structural geometry of the layered intrusions requires substantial tectonic disruption along the boundary zone between the Koillismaa and Kianta terrains, interpreted by Ward et al. (1989) as a consequence of listric extensional faulting during deposition of the Kuusamo supracrustal sequences and subsequent inversion during the Svecofennian and Kola–Karelian collisional events. Silvennoinen (1991) noted that the Archean gneisses are more intensely fractured and foliated and commonly chloritic in proximity to the unconformity, which is attributed to Proterozoic deformation. However, some constraints suggesting a rather modest amount of Svecofennian deformation are 62

provided by the mafic dike swarms truncating Archean foliations; this is also clearly apparent in interpretations of regional magnetic data, which furthermore indicate that Proterozoic hydrothermal overprinting over much of the Koillismaa terrain was relatively limited (Airo, 1999). The Koillismaa terrain terminates westwards against a complex deformation zone variously referred to as the Hirvaskoski shear zone (Kärki et al., 1993) or Posio shear system (Sorjonen-Ward et al., 1997). This deformation zone has been repeatedly activated throughout the Paleoproterozoic so that it is difficult to establish whether it was initiated and active in the Archean (Sorjonen-Ward et al., 1992, 1997; Kärki et al., 1993; Vaasjoki et al., 2001). However, it effectively bisects the Karelian domain in northeastern Finland, with the Pudasjärvi and Napapiiri terrains to the west and the Koillismaa and Kianta terrains to the east. Räsänen and Vaasjoki (2001) have recently identified a zone of metasedimentary gneisses and inferred rhyolitic volcanic rocks and pyroclastic deposits within this highly deformed zone, for which a U-Pb zircon age of 2796 ± 10 Ma was obtained. As elsewhere, the possibility of inherited Archean detrital zircon in metasediments, or retention of restitic zircon during partial melting of an Archean source needs to be evaluated. Nevertheless, these results are consistent with emerging evidence for the existence of a discontinuous zone of late Archean supracrustal rocks, extending along the eastern margin of the Pudasjärvi Terrain and northwards into Lapland, substantially disrupted by Svecofennian deformation and magmatism (Huhma et al., 2000; Räsänen and Huhma, 2001; Evins et al., 2000, 2002).

3.2. Napapiiri terrain The Napapiiri terrain encompasses a diverse and poorly understood assemblage of supracrustal gneissic and granitic rock units, ex-

• CHAPTER 2 • ARCHEAN ROCKS

tending across southern and central Finnish Lapland, from the Swedish border to Russia (Figures 2.1 and 2.3). The Napapiiri terrain is partly synonymous with the terms Central Lapland granitoid complex or Kemijärvi complex (Ahtonen and Melqvist, 1997), but is used here to emphasize its Archean aspect. This is because several recent studies, particularly in the eastern part of the terrain (Räsänen and Huhma, 2001; Räsänen and Vaasjoki, 2001; Vaasjoki et al., 2001; Evins et al., 2000, 2002) have revealed that Archean rocks are more widespread than previously thought, while an appreciation of the substantial Proterozoic thermal and tectonic overprint requires some reassessment of the geological evolution of the northern part of the country (Vaasjoki et al., 1999; Corfu and Evins, 2002). Much of the western and central part of the Napapiiri terrain is characterized by a prominent NNE-trending magnetic fabric. Väänänen (1998) described migmatitic metasedimentary gneisses and granites at the western edge of the terrain and defined them as the Venejärvi complex (Figure 2.3). Attempts to date this complex have not been successful, with evidence for both Archean inheritance and Proterozoic ages for U-Pb zircon and SmNd dating of both neosomes and paleosomes. In the eastern part of the terrain, Proterozoic granitic magmatism is manifested as sheets and networks of equigranular to porphyritic monzogranite, which are slightly peraluminous and strongly enriched in LREE, but with low Nb and Y (Ahtonen and Melqvist, 1997). Rastas et al. (2001) have reported Archean ages from hydrothermally altered felsic rocks along the northern margin of the Napapiiri terrain, at Honkavaara (Figure 2.3), near the contact with the Lapland greenstone belt. Given the intensity of hydrothermal alteration and proximity to the greenstone belt it is also possible that some of these felsic rocks represent Paleoproterozoic sediments with Archean detrital zircons. On the other hand, the nature of the regional magnetic pattern

would be consistent with a lithologically diverse Archean rock package, variably affected by Proterozoic anatexis. Such an interpretation finds further support from the recognition of Archean supracrustal rocks, including felsic lava, some distance away, along the eastern margin of the Napapiiri terrain (Räsänen and Huhma, 2001). Zircon fractions from dacitic to andesitic felsic volcanic rocks of the Loviselkä Formation (Figure 2.3) have been dated at 2775 ± 25 Ma and are intercalated with quartzofeldspathic gneisses that, despite isoclinal folding and metamorphism to assemblages containing staurolite-garnet-cordierite and andalusite, still preserve evidence of a thick-bedded graded turbiditic origin. When these results are combined with the data and interpretations of Räsänen and Vaasjoki (2001) from the western margin of the Koillismaa terrain and Evins et al. (2002), they acquire still greater significance in regard to regional correlations between Archean rock units, as will be discussed later. A number of studies have attempted to correlate magnetic characteristics with granite chemistry and mineralogy, particularly in the southeastern part of the terrain (Airo; 1999; Airo and Ahtonen, 1999). Puranen (1989) also found that the Proterozoic granites intruding the terrain have relatively high abundances of ferrimagnetic magnetite, even though they are rather poor in iron compared to other Svecofennian granites. This was interpreted as a consequence of derivation of Proterozoic monzogranites from a highly oxidized Archean source terrain.

Suomu terrain The Suomu terrain covers some 1000 km2 in area and occupies a transitional position between the Napapiiri and Ranua terrains (Figures 2.3 and 2.14). It is has been strongly affected by Proterozoic thermal and tectonic events (Corfu and Evins, 2002), but the Archean age and character of much of the complex has recently been demonstrated and docu-

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mented (Evins et al., 1997, 2000, 2002; Airo, 1999). The Suomu terrain has been subdivided into biotite-bearing tonalitic to granodioritic gneisses, which comprise over 80% of the complex, and the Aholanvaara supracrustal complex, near the southeastern margin of the terrain. The two units are readily distinguishable in aeromagnetic data, with the tonalitic gneisses being rather subdued magnetically, in contrast to the more intense and variable anomaly patterns associated with supracrustal rocks, particularly magnetite-biotite pelitic schists (Airo, 1999; Evins et al., 2002). Evins et al. (2002) have dated several samples of the tonalitic gneisses using the NORDSIM ion microprobe and quote a pooled zircon age for crystallization at 2823 ± 10 Ma and 2815 ± 21 Ma. Several zircon cores yielded ages up to 2.87 Ga, suggesting derivation of the tonalites from an older source; this is also consistent with the presence of discrete, variably sized mafic and ultramafic supracrustal enclaves that have an internal structural history discordant with respect to the host gneisses. Biotite-amphibole dioritic gneisses are also sporadically present and concordant with respect to the banding in the tonalitic gneisses (Evins et al., 2002). They have, however, yielded younger ages, around 2555 ± 16 Ma, suggesting a two-stage magmatic evolution for the Suomu terrain. Sillimanite-grade quartzites and arkosites are the most characteristic rock types of the Aholanvaara supracrustal package, although metapelitic and calc-silicate gneisses are also present, and intruded by amphibolite sills, resulting in distinctive garnet-gedrite-sulfide contact skarns. This lithological association and metamorphic style is very reminiscent of the Paleoproterozoic Kuusamo schist belt (cf. Evins and Laajoki, 2001), which raises further questions about its Archean affinity. Nor have zircon provenance studies by ion microprobe helped to dispel this uncertainty; a detrital zircon age spectrum from Aholanvaara quartzite, taken from the contact with the tonalitic gneiss, 64

included eight nearly concordant grains with 207 Pb/206Pb ages between 2706 Ma and 2744 Ma, while zircons with ages corresponding to the Suomujärvi complex tonalites (2.85–2.80 Ga) were conspicuously absent. The contact between the Aholanvaara quartzites and the tonalitic gneisses is actually exposed: Evins et al. (2002) noted that there is no obvious difference in structural and metamorphic history on either side of the contact, and that the contact itself does not appear highly strained, nor is there obvious evidence for a weathered unconformity. The absence of detrital zircons representing the age range of the Suomujärvi tonalitic gneisses is another reason why Evins et al. (2002) considered the contact to be tectonic in nature. However, the extent to which original Archean structures are preserved remains uncertain – if the mafic sill intruding the Aholanvaara supracrustal units is correctely correlated with the 2.21 Ga Tokkalehto gabbro (Evins and Laajoki, 2001), then much of the intense SW-plunging and NE-trending folding and foliation in the Suomu terrain must clearly be Proterozoic in age.

3.3. Tuntsa terrain Mikkola (1941) first defined a distinctive suite of medium- to high-grade supracrustal gneisses, intruded by granites and trending in a northeasterly direction from Savukoski towards the Russian border, as the Tuntsa–Savukoski series. Similar rock types are widespread in the Belomorian terrain in the Kola Peninsula and Russian Karelia (Gaál and Gorbatschev, 1987; Stenar, 1988) and it is clear that the Tuntsa terrain can be traced into Russia, coinciding with a progressive change in lithological and structural trends from NE to NW (Figures 2.3, 2.14 and 2.15). This arcuate change in regional trend is evident in regional aeromagnetic data (Korhonen et al., 2001a,b) as well as from geological mapping (Koistinen et al., 2001). From the Finnish perspective, the Tuntsa ter-

• CHAPTER 2 • ARCHEAN ROCKS

rain appears to lie well within the Karelian domain, with its northern boundary defined by the basal thrust of the Lapland granulite belt, and the Pomokaira and south Lapland terrains occurring to the west. Relationships between the Tuntsa domain and Archean rocks to the south are unfortunately obscured by the Paleoproterozoic Salla and Kuusamo supracrustal belts. In Russia, however, the Belomorian terrain forms a broad zone separating the Kola and Karelian domains, and its age, origin and tectonic significance have long been a source of controversy (Stenar, 1988). Recent isotope studies have nevertheless provided a framework for integrating structural, metamorphic and petrogenetic studies in the Belomorian terrain; the preferred interpretation is that the Belomorian terrain collided with other elements of the Karelian domain during the late Archean, and that this boundary zone was the locus for renewed deformation when the Kola domain collided with the Karelian domain during the early Proterozoic (Bibikova et al., 2001; Daly et al., 2001). The effect of Proterozoic deformation within the Tuntsa terrain may be more difficult to discern, especially given that many structures and rock units have gently dipping enveloping surfaces, subparallel to those in the Lapland granulite belt (Figure 2.16). Some tectonic reactivation is therefore likely, particularly as that deformation in the Salla and Kuusamo supracrustal belts to the south have been interpreted as a consequence of foreland deformation during emplacement of the Lapland granulites (Ward et al., 1989). Moreover, Proterozoic thermal overprinting on Archean rocks is widely documented from the Southern Lapland terrain (Evins and Corfu, 2002). Despite these potential problems, and the generally poor exposure, it has been possible to subdivide the rocks of the Tuntsa terrain into five distinct units, namely the Naruska, Ahmatunturi and Vintilänkaira–Kemihaara granitoid complexes, and the Tuntsa and Tulppio supracrustal belts (Juopperi and

Vaasjoki, 2001).

Granitoid complexes The granitoid complexes consist of tonalitic, granodioritic, and granitic gneisses, commonly containing gneiss and amphibolite inclusions, sometimes of considerable extent. Deformation is particularly intense in proximity to the Lapland granulite belt (Mikkola, 1941). The Naruska granitoid complex, in the southern part of the terrain, tends to show gradational transitions with the Tuntsa paragneisses, suggesting a deeper erosional level within a single lithotectonic unit (Juopperi and Vaasjoki, 2001). The results of U-Pb zircon analyses nevertheless suggest protracted evolution for the Naruska granitic magmatism, with granitic to tonalitic gneiss samples from the transition zone yielding ages ranging from 2744 ± 25 Ma to 2705 ± 5 Ma. However, one sample, from a partially retrogressed granite within the Tuntsa paragneiss complex provided a significantly younger age of 2636 ± 11 Ma. Titanite ages are also close to zircon ages, providing some constraints on the degree of Proterozoic thermal overprinting (Juopperi and Vaasjoki, 2001). The granitoids of the Kemihaara–Vintilänkaira (Figure 2.15) complex are poorly exposed and have not been mapped in detail. Some intrusions in the southern part of the complex are likely to be Paleoproterozoic rather than Archean in age, although no contacts with the Tulppio supracrustal belt have been observed (Juopperi and Vaasjoki, 2001). The most reliable U-Pb zircon age estimate obtained so far is from a tonalitic rock, dated at 2805 ± 4 Ma (Juopperi and Vaasjoki, 2001). However, of particular interest is the presence of syenitic intrusions with zircon ages of 2795 ± 20 Ma. This would be an unusual age for alkali magmatism, the only other Archean syenites and carbonatites in the Fennoscandian Shield being in the Kola Peninsula (Zozulya et al., 2001) and the Siilinjärvi carbonatite in the Iisalmi terrain (Puustinen, 1971). Because

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Muonio

g

Pomokaira

g g g g

Muonio

Mö

To

Kittilä

Sodankylä

Kolari

Tuntsa

Savukoski

Napapiiri Salla Kemijärvi Suomu 0

40

Rovaniemi

Archean

Proterozoic

km

Paleozoic carbonatite

gg

Lapland granulite belt

Svecofennian (1.9–1.86 Ga) orogenic granitoids Younger (1.86–1.82 Ga) granitoids with Archean crustal inheritance

Paleoproterozoic (2.5–2.0 Ga) Lapland greenstone belt Paleoproterozoic (2.5–1.9 Ga) supracrustal rocks

Paleoproterozoic granites and thermal reworking in Napapiiri terrain

Tuntsa supracrustal gneiss terrain

Granitoids, migmatites, and gneisses

Mafic and ultramafic metavolcanic rocks

Fig. 2.14. Terrains defined within the Karelian domain in northern Finland, including basement windows exposed within the Paleoproterozoic Lapland greenstone belt. The Napapiiri and Suomu terrains record a complex Proterozoic thermal overprint, intruded by extensive granitic bodies. The Tuntsa terrain, which is contiguous with the Belomorian terrain in Russia, shows less thermal overprinting, but the extent of structural disruption, associated with emplacement of the Lapland granulite belt from the north is unclear. Mö–Möykkelmä, To–Tojottamanselkä.

66

• CHAPTER 2 • ARCHEAN ROCKS

Post-Archean rock units

A

Paleozoic Sokli carbonatite Lapland granulite belt (1.9 Ga) Paleoproterozoic (2.5–2.0 Ga) Lapland greenstone belt Kemihaara

Archean rock units in Tuntsa terrain Naruska granitoid complex (2.74–2.70 Ga) Granitic compositions dominant

Tulppio

Tonalitic compositions dominant

Tuntsa supracrustal gneiss belt Pelitic, psammitic, and quartzitic gneisses

Vintilänkaira

Tuntsa

Ahmatunturi

Mafic and ultramafic metavolcanic rocks

Ahmatunturi and Vintilänkaira–Kemihaara granitoid gneiss complexes (>2.80 Ga)

Savukoski

Granitoids, migmatites, and gneisses Granitic compositions dominant Naruska

Tuntsa supracrustal belt Metasedimentary gneisses

0

Mafic and ultramafic metavolcanic rocks

B

10 km

20

C

Fig. 2.15. Tuntsa terrain. (A) Principal geological units, based on Juopperi and Vaasjoki (2001). (B) Cliff section showing gently dipping structural architecture, typical of Tuntsa terrain gneisses and granitoids. John Ridley is approximately 1.8 m in height. Near Naruskajoki, in southern part of terrain. (C) Detail of highly strained stromatic migmatites with felsic leucosomes at same locality as (B). Photos: Peter Sorjonen-Ward.

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the Devonian Sokli carbonatite complex is situated very close to these syenites, the possibility of zircons being xenocrystic ought to be considered, although titanite from the same intrusion records an age of 2683 ± 1 Ma (Juopperi and Vaasjoki, 2001). The Ahmatunturi granitoid complex has also yielded zircons of age 2833 ± 22 Ma, which is significant in that it provides a minimum age constraint on deposition in the Tulppio supracrustal belt (Juopperi and Vaasjoki, 2001). When compared with the results for the Naruska granitoids, it appears that the Tuntsa terrain records the juxtaposition of two crustal units of different age, or alternatively, two stages of magmatism and deformation. The presence of a polymictic conglomerate at Nuolusvaara, near the Russian border, containing clasts of mafic schist, pelitic gneiss and tourmaline-bearing pegmatite, within a matrix metamorphosed to lower amphibolite facies (Juopperi and Veki, 1988), is consistent with both of these scenarios.

Tuntsa and Tulppio supracrustal belts The relatively high degreee of deformation and metamorphism has precluded mapping of primary rock facies or stratigraphical relationships in the Tuntsa terrain (Juopperi and Vaasjoki, 2001). The paragneisses of the Tuntsa supracrustal belt form a coherent unit 15–25 km across strike and consist almost entirely of medium-grade metamorphic sedimentary rocks. In contrast, the Tulppio supracrustal belt comprises only discontinuous schist and gneiss remnants, although they are lithologically more diverse than those of the Tuntsa belt. The most extensive of these remnants is characterized by medium-grade metamorphic ultramafic and mafic volcanic rocks, the former being interpreted as cumulates of Archean komatiitic lavas, the latter as strongly altered submarine Mg-rich and Fe-rich tholeiitic lavas. Locally the metavolcanic rocks are associated with quartz-feldspar 68

schists, amphibole- and garnet-rich aluminous schists as well as quartzites and cherty rocks (Juopperi, 1994).

3.4. Pomokaira terrain Archean granodioritic gneisses and lesser quartzofeldspathic metasediments are exposed throughout the foreland immediately adjacent to the main frontal thrust of the Lapland granulite belt, and in several basement windows, notably at Möykkelmä and Tojottomanselkä, where they are unconformably overlain by basal units of the Lapland greenstone belt (Figure 2.15). Negative εNd values from plutons of the Nattanen granite suite (Huhma, 1986), some of which intrude the boundary between the Lapland granulite belt and the Pomokaira terrain indicate that Archean rocks are present at depth for a considerable distance behind the thrust front. The Pomokaira terrain appears to be contiguous with the northern part of the Tuntsa terrain, although the boundary zone is largely obscured by the 2.44 Ga Koitelainen layered intrusive complex and supracrustal rocks of the Lapland greenstone belt. The tonalitic gneisses exposed in the small (4 km 2 ) Tojottamanselkä basement inlier provided the first evidence for rocks older than 3.0 Ga in the Fennoscandian Shield, with a multigrain zircon population yielding a U-Pb age of 3110 ± 34 Ma (Kröner et al., 1981). A whole-rock Rb-Sr isochron of 2729 ± 244 Ma was obtained for the same sample and was attributed to resetting during a metamorphic disturbance. These results were later corroborated by SHRIMP analysis, with an estimated intrusive age of 3115 ± 29 Ma and a subsequent thermal reworking at 2836 ± 30 Ma (Kröner and Compston, 1990). Jahn et al. (1984) also interpreted Pb isotope data as recording a metamorphic resetting during the late Archean or earliest Proterozoic, based on a whole-rock isochron of 2640 ± 240 Ma. A whole-rock Sm-Nd isochron of 3060 ± 123 Ma was considered to be consistent with the

• CHAPTER 2 • ARCHEAN ROCKS

zircon data and was interpreted as the age of emplacement of the tonalitic precursor to the gneisses. Jahn et al. (1984) also found that the Tojottamanselkä gneisses had an εNd value (with respect to CHUR) of –3.7 ± 1.8, implying derivation from a protolith that was already enriched in LREE. On this basis they proposed a multistage evolution commencing with extraction of basalt from mantle, and melting of basalt to produce a tonalitic to trondhjemitic crust, several hundred million years prior to the 3.1 Ga event recorded by the Tojottamanselkä zircons. For comparison, note that the oldest inherited zircon found by Kröner and Compston (1990) was 3248 ± 10 Ma.

3.5. Muonio terrain Lehtonen (1984) identified three separate areas of migmatitic biotite-plagioclase gneisses in the Muonio district, near the Swedish border. Granodioritic to tonalitic compositions predominate, with some hornblende gneiss intercalations. These gneiss occurrences are up to 10 in length and several kilometers in width, and form fault-bounded anticlinal features surrounded by sillimanite-grade arkosic gneisses and metavolcanics rocks, which are correlated with the Paleoproterozoic Lapland greenstone belt. An Archean age for these gneisses is also supported by U-Pb zircon studies, which yielded ages of 2444 ± 96 Ma and 2591 ± 16 Ma (Lehtonen, 1984; Väänänen and Lehtonen, 2001). Titanite ages of 1845 Ma are consistent with the intense tectonic and thermal reworking associated with the Svecofennian orogeny, suggesting that these inliers of Archean basement are analogous to the classic basement gneiss domes described along the margin of the Karelian domain in southeastern Finland (Eskola, 1949). They are considered separately here, because it is unclear whether they represent part of a contiguous region of Archean basement, extending eastwards and southwards beneath the Lapland greenstone belt towards the Pomokaira and southern

Lapland terrains, or alternatively, correlate with the Archean of the Ropi terrain and northern Sweden (Figures 2.3 and 2.14). An Archean provenance is also indicated for the so-called Hetta granites, which occur in the area between Muonio and the Norwegian border (Figure 2.3), based on Pb isotope studies of feldspars (Meriläinen, 1976) and heterogeneous and xenocrystic zircon populations (Lehtonen, 1984; Mänttäri, 1995). These lithological units continue into adjacent Norway as the Jer’gul gneiss complex (Siedlecka et al., 1985), which has been subdivided into two major units based on lithology and chemical composition (Olsen and Nilsen, 1985) – the Ak’kanasvarri gneisses, which are typically quartz dioritic to tonalitic hornblende gneisses, and the Biennaroavvi gneisses, which are more evolved trondhjemitic magmas in origin. Olsen and Nilsen (1985) used trace element modeling to infer melting of an amphibolite source under garnet-stable conditions for the former, and partial melting of the Ak’kanasvarri gneisses for the latter. Combined data from both units produced a Rb-Sr isochron of 2993 ± 195 Ma.

3.6. Ropi terrain Archean rocks are exposed in northwestern Norway as windows beneath Caledonian nappes and as the Raisædno gneiss complex (Siedlecka et al., 1985), which can be traced into the extreme northwestern part of Finnish Lapland (Figures 2.1 and 2.3). This area is referred to here as the Ropi terrain and consists mainly of banded migmatitic granitoids and gneisses, their composition varying in composition from tonalite to granodiorite (Lehtovaara, 1995). In addition, the Ruossakero–Sarvisoaivi–Ropi tun turi greenstone belts can be traced as remnant supracrustal units several kilometers in width and more than 10 km along strike (Lehtovaara, 1995). These consist principally of amphibolites derived from basaltic lavas, overlain by a

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paleoeregolith and schists inferred to have been volcaniclastic in origin, with sporadic sericite quartzites, mica schists, and mica gneisses (Hannu Idman, pers. comm., 1995). Some ultramafic rocks, which may have been originally cumulates of ultrabasic lavas or intrusions, have also been recognized, containing a low-grade nickel mineralization. The Ropi terrain is contiguous southwards with the Råstojaur gneiss complex, which forms part of an extensive region of Archean crust in northern Sweden, widely overlain and intruded by Proterozoic rocks (Skiöld and Öhlander, 1989; Öhlander et al., 1993; Martinsson et al., 1999). The Paleoproterozoic sequences in northern Sweden and adjacent Finnish Lapland share a number of features in common, suggesting that they both record Paleoproterozoic rifting and fragmentation along the southwestern margin of the Karelian domain. However, there is a major NE-vergent Svecofennian deformation zone, characterized by high strain and metamorphic grade, and a distinctive suite of 1.89-1.86 Ga synorogenic calc-alkaline to post-collisional potassic granitoids, separating the Archean of northern Sweden, and the Ropi terrain, from the Muonio terrain and Karelian domain of Finnish Lapland (Figures 2.1, 2.3, and 2.14). It is therefore possible that the Ropi terrain is exotic with respect to the Karelian domain, or at least represents part of a continental fragment ribbon rifted from and translated along the Karelian continental margin (cf. SorjonenWard et al., 2001).

4. The Kola domain in Finland The Kola domain is a complex mosaic of Archean and Paleoproterozoic terrains that were amalgamated and accreted to the Karelian domain between 2.0 Ga and 1.8 Ga (Hörmann et al., 1980; Barbey et al., 1984; Berthelsen and Marker, 1986a; Gaál et al., 1989). The Kola domain in Finland, as in 70

adjacent Norway, can be further divided into three tectonic units which can be traced for several hundred kilometres, showing a NEdirected tectonic polarity – the late Archean Inari terrain in the southwest and Sørvaranger terrain in the northeast, separated by the Paleoproterozoic Polmak–Pasvik–Pechenga belt (Gaál et al., 1989) (Figure 2.16). An unconformable relationship between the basal units of the Polmak–Pasvik–Pechenga belt and the Sørvaranger terrain has been demonstrated in several places, but much of the sequence is allochthonous; likewise, the Inari terrain has been thrust over the Pechenga belt, with considerable tectonic reworking in the contact zone (Marker, 1985; Gaál et al., 1989). The Finnish segment of the Pechenga–Polmak–Pasvik belt has been mapped as the Opukasjärvi Group (Kesola, 1991, 1995). Although basal units of this sequence have been shown to unconformably overlie gneisses of the Sørvaranger terrain, contact relationships with the Inari terrain gneisses to the southwest are more complicated. As is typical for the Sørvaranger terrain, dips are gentle to moderate. Kesola (1991) interpreted the main foliation parallel to the enveloping surface to the Opukasjärvi Group schists as regional S3, which postdates garnet-staurolite porphyroblast growth. This implies intense Proterozoic tectonic reworking of at least the northeastern margin of the Inari terrain, although Kesola (1995) also considered that the Inari terrain and Sørvaranger terrain are sufficiently similar in terms of lithology that they may have originally formed part of a single crustal unit. In Finland the Kola domain and Karelian domain are separated by the Lapland granulite belt. This is a zone more than 50 km wide, consisting predominantly of highly strained and anatectic peraluminous metasedimentary granulites, for which SHRIMP zircon studies indicate mainly Proterozoic provenance ages (Sorjonen-Ward et al., 1994). Granulite-facies enderbitic pyroxene-bearing intrusions and anorthosites dated at 1.95 Ga and 1.90

• CHAPTER 2 • ARCHEAN ROCKS

Ga, respectively (Meriläinen, 1976; BernardGriffiths et al., 1984), provide a maximum age constraint for emplacement over the Pomokaira and Tuntsa terrains of the Karelian domain. Thermobarometry indicates that maximum pressures in the marginal zone attained nearly 12 kbar (Tuisku and Makkonen, 1999), although 6–7 kbar is more typical (Raith and Raase, 1986). Interpretations of deep crustal seismic reflection and refraction data (Behrens et al., 1989; Luosto et al., 1989), gravity surveys (Elo et al., 1989), and electromagnetic data (Korja et al., 1989) are all consistent with surface observations indicating that the granulites were emplaced southwards along a basal detachment zone that can be traced at least into the middle crust. The contact between the Lapland granulite belt and southern margin of the Kola domain appears to be steeper and possibly of opposite dip (Gaál et al., 1989), suggesting a large scale pop-up structure or retrowedge (cf. Beaumont et al., 1994), in which the Archean gneisses of the Inari terrain are imbricated with rocks of the granulite belt and have themselves locally been metamorphosed to granulite grade (Hörmann et al., 1980; Raith and Raase, 1986). Because the Inari terrain has also been intruded by quartz diorites and gabbros dated at 1.95–1.93 Ga (Meriläinen, 1976), a number of authors have integrated the above features into a model involving the formation of a continental margin magmatic arc in the Inari terrain, which was eventually terminated by collision and emplacement of the Lapland granulite belt over the Karelian domain (Hörmann et al., 1980; Barbey et al., 1984). Berthelsen and Marker (1986) proposed an alternative polarity, attributing the 1.95 Ga calc-alkaline magmatism in the Inari terrain to south-directed subduction and underthrusting of the Sørvaranger terrain. In both cases, the implications are that the Kola and Karelian domains might have developed in quite different settings during the Archean.

4.1. Inari terrain The Inari terrain (Figures 2.3 and 2.16) consists predominantly of migmatitic biotite- and biotite-hornblende orthogneisses ranging in composition from tonalite to monzogranite, with U-Pb zircon ages between 2.73 Ga and 2.50 Ga (Meriläinen, 1976; Gaál et al., 1989). Kesola (1995) defined two separate gneiss complexes based on lithological differences. The Suorre–Tievjan complex consists of intensely migmatized gneisses and augen gneisses of granitic composition, with a U-Pb zircon age of 2502 ± 8 Ma, which is appreciably younger than other Archean ages from comparable rock types in Finland. However, because the titanite age from the same rock is concordant at 1997 Ma (Kesola, 1995), and Proterozoic ages have been obtained for titanite throughout the Inari terrain (Meriläinen, 1976), the U-Pb systems of zircons may also have been affected by Proterozoic events. The Moresveijohjkan complex in the northwestern part of the Inari terrain is distinctly more mafic, consisting of pyroxene-bearing quartz diorites with abundant enclaves of hornblende-biotite gneiss. Remnant supracrustal units are also present in the Inari terrain, the most significant being the Kuorboaivi schist belt (Meriläinen, 1976; Gaál et al., 1989). The affinities of these rocks are controversial, and Kesola (1991) correlates them with the Paleoproterozoic Opukasjärvi Group, implying complex Proterozoic tectonic imbrication of Archean rock units.

4.2. Sørvaranger terrain The Sørvaranger terrain is developed most extensively in northern Norway, where the NW-trending Garsjø and Bjørnevatn supracrustal belts are tectonically juxtaposed against tonalitic and trondhjemitic migmatitic gneisses (Siedlecka et al., 1985). The Garsjø and Bjørnevatn belts are lithologically diverse, including banded iron-formations, invari-

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Vainospää granite (1780 Ma) Tšuomasvarri ultramafic intrusion Luossajavri gabbro (1731 ± 2 Ma)

Nuorgam

NORWAY

Utsjoki

Näätämö

Inari terrain Suorre–Tievja gneiss complex granites (2520 ± 8 Ma), migmatites, and paragneisses

Sevettijärvi

Opukasjärvi Group Karigasniemi

Silisjoki gneiss complex Metabasalt, meta-andesite

Kola domain

Metarhyolites and pelitic schists Metaconglomerate and metaarkose

Inari

Opukasjärvi Group Pirivaara granite (2604 ± 21 Ma) Garsjøen gneiss complex

RUSSIA Lapland granulite belt

Pdl Karelian domain

Pdl

Pgv

0

10 km

Pgv

Pgv

Fig. 2.16. The Kola domain in Finland. Figure at upper right shows regional relationship between Karelian domain, Lapland granulite belt, and Kola domain. Larger scale figure at lower right shows the Inari terrain in the southwest, separated from the Sørvaranger terrain to the northeast by the supracrustal rocks and highly strained gneisses of the Opukasjärvi Group; the latter are most likely of Paleoproterozoic age, although relationships and age determinations are contentious. Pdl–Luossajavri gabbro, Pgv–Vainospää granite. Based on Kesola (1991, 1995).

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• CHAPTER 2 • ARCHEAN ROCKS

ably associated with mafic volcanic rocks, and local ultramafic and quartzitic layers, within a dominantly psammitic to semipelitic sequence (Gaál et al., 1989). Proterozoic tectonic reworking of the Sørvaranger terrain appears to be restricted to the contact zone with the Polmak–Pechenga–Pasvik belt, such that Marker (1985) and Gaál et al. (1989) considered the generally gentle to moderate NE–ENE-dipping enveloping surface to rock units and thrusts to be a relict of the original Archean architecture. The Sørvaranger terrain in Finland has been designated as the Garsjø complex (Kesola, 1995), which includes gneisses that are obviously supracrustal in character and highly strained quartzofeldspathic gneisses whose origin is less clear; both types are closely associated and appear to share a common deformation history. Recognizable remnants of magnetite-grunerite banded iron-formations and tholeiitic mafic volcanic rocks are preserved in particular in the Näätämö and Vätsäri areas (Kesola, 1991; Figure 2.16) and closely resemble those described from the Garsjø and Bjørnevatn belts in adjacent Norway (Siedlecka et al., 1985). No depositional basement to the supracrustal rocks has been found, although polymicitic conglomerates have been described (Gaál et al., 1989). On the other hand, the Garsjø complex was intruded by the relatively homogeneous plutons of the Pirivaara granite suite (Figure 2.16), equivalent to the Neiden granites in Norway, with a U-Pb zircon age of 2604 ± 21 Ma. This is anomalously young when compared to latest Archean granitic magmatism elsewhere in Finland, so that the possibility of Proterozoic disturbance should also be considered.

5. Insights into the deeper Archean crust in Finland Information concerning the structure, composition, age, and thickness of the deep crust, and the degree to which lower crustal and lithospheric mantle coupling has evolved with time can be obtained directly and indirectly through • Studying deep crustal sections tectonically exhumed during later events; • Evaluating source compositions from chemical and isotope characteristics of granitoids; • Constraining P-T-t histories of xenolith suites, to determine the age, depth distribution and petrophysical characteristics of different rock types – an approach known as 4D lithospheric mapping (O’Reilly and Griffin, 1996); and • Deep seismic refraction and reflection surveys, ideally in combination with gravity and magnetotelluric investigations.

5.1. Exhumed deep crustal sections in Finland? Exposed sections of the deep crust, such as the Kapuskasing zone in Canada and the Vredefort dome in the Kaapvaal craton, have been important in providing insights into the composition, thermal properties, and density structure of Archean lithosphere and the nature of tectonic and thermal reworking, all of which can be used to constrain interpretations of crustal sections based on geophysical data (Percival et al., 1992; Rudnick and Fountain, 1995). It is of great interest to understand whether granulite metamorphic events are directly coupled with the tectonic processes that exhume a terrain, or whether uplift and exhumation is due to some younger event (Sandiford, 1989). In Finland, exhumation of rocks from depths corresponding to pres-

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sures of 10–12 kbar occurred during the emplacement of the Paleoproterozoic Lapland granulite belt southwards over the Pomokaira terrain and Lapland greenstone belt after 1.9 Ga; this also resulted in medium-pressure metamorphism within the foreland (Raith and Raase, 1986; Gaál et al., 1989). The highest pressure assemblages recorded from the Lapland granulite belt are nearly 12 kbar, from ultramafic rocks and anorthosites, which are, however, likely to represent Paleoproterozoic rather than Archean magmatic cumulates (Tuisku and Makkonen, 1999). Kyanite-bearing assemblages in Paleoproterozoic sediments within the Karelian domain in Russia and the results of isotopically constrained thermobarometric studies also indicate that late Archean crust of the Belomorian terrain was exhumed from middle crustal levels at around 1.80 Ga (Bibikova et al., 2001). The Tuntsa terrain might also record tectonic juxtaposition of different levels of Archean crust during the early Proterozoic, although at present there are no P-T constraints from this region. There is in addition the remarkable possibility that the serpentinized harzburgites of the Jormua ophiolite complex represent Archean subcontinental lithospheric mantle (Peltonen et al. 2003; Chapter 6), in which case these rocks would be the only known example of Archean lithosphere extensively exposed at the surface of the Earth. This is evidently a consequence of attenuation of the rifted continental margin at 2.0 Ga, followed by tectonic obduction back onto the Karelian domain during the Svecofennian orogeny. It is therefore important to appreciate that the lateral variations in metamorphic grade observed throughout the Archean of Finland do not necessarily represent late Archean postorogenic stabilization and erosion, but may instead relate to Proterozoic tectonic and thermal reworking. Where such uncertainty exists concerning the timing of juxtaposition of terrains of varying metamorphic grade, or the uplift and exhumation of Archean complexes, useful constraints can be 74

provided by consideration of the distribution of Proterozic unconformity surfaces and the degree of recrystallization and strain recorded by Proterozoic mafic dikes. Observations of the strain state and metamorphic grade of mafic dikes and of Svecofennian granitic intrusions in the Iisalmi terrain and Ranua terrain (Paavola, 1984, 1986) indicate that the medium-pressure granulites described by Hölttä et al. (2000a) were already exposed at high crustal levels and juxtaposed against lower grade terrain prior to the Svecofennian orogeny. In the Rautavaara area, there is a stronger Svecofennian overprint, and metamorphic re-equilibration under amphibolite facies conditions, suggesting differential Svecofennian uplift of a tilted crustal section. The lithological diversity and evidence for hydrothermal alteration of supracrustal rock units prior to granulite facies metamorphism in the Iisalmi terrain and Rautavaara terrain (Hölttä et al., 2000) indicate that the deep crust in at least this part of the Karelian domain is likely to be very heterogeneous in composition. Greenschist to lower amphibolite facies metamorphism and locally intense foliation development is characteristic of Proterozoic sedimentary and volcanic cover sequences, as well as mafic dike swarms throughout the Karelian domain and is attributed to burial during overthrusting (Kontinen et al., 1992; Sorjonen-Ward, 1993). This is consistent with the clockwise P-T-t history recorded for the Svecofennian orogeny in eastern Finland (Ward, 1987; Pajunen and Poutiainen, 1999) and the widespread resetting of Archean basement isotope systems in areas currently devoid of Proterozoic cover (Kontinen et al., 1992; O’Brien et al., 1993). Biotite generally yields Proterozoic K-Ar ages throughout eastern Finland, whereas hornblende has been more robust, generally retaining Archean ages (Kontinen et al., 1992; O’Brien et al., 1993b). Titanite U-Pb ages have also been reset in the Belomorian province, with the youngest ages

• CHAPTER 2 • ARCHEAN ROCKS

of 1.78–1.75 Ga being from titanite within late hydrothermal alteration parageneses (Bibikova et al., 2001). Pajunen and Poutiainen (1999) determined metamorphic conditions and hydrothermal fluid activity in Proterozoic shear zones within Archean basement in the Kuhmo and Nurmes terrains, recognizing a prograde event accompanied by saline water-rich fluids and decompression associated with a more typical late orogenic metamorphic CO2–H2O fluids; hydrothermal xenotime relating to the latter mineral assemblage was dated to 1852 ± 2 Ma. This suggests that between 1.9–1.8 Ga the present erosion level of the Karelian province experienced temperatures between 400–500 ºC during burial to maximum depths of around 15 km, with metamorphic dehydration reactions producing at least localized fluid–rock interaction. This is par ticularly evident in the structural control on magnetic signatures in the Archean of eastern Finland (Sorjonen-Ward, 1993; Airo, 1999), although it is uncertain whether fluids were derived from underlying Archean rocks, or the overlying Proterozoic allochthon.

5.2. Distribution and composition of buried Archean crust Some further inferences concerning the distribution and composition of the deep Archean crust in the Karelian domain can be deduced from the regional responses to deformation, burial and heating during the Svecofennian and Kola–Lapland orogenies. For example, the P-T-t history defined by Pajunen and Poutiainen (1999) and Bibikova et al. (2001) in principle allows the possibility of decompression melting within the deep Archean crust during late Svecofennian orogenic stabilization, between 1.85 and 1.80 Ga. The role of magmatic underplating in modifying the thermal regime of the lower crust also needs to be considered, in view of the 1.80 Ga ages obtained from lower crustal mafic granulite and mantle xenoliths (Hölttä et al., 2000b;

Peltonen and Mänttäri, 2001) and some of the deep crustal high-velocity layers observed in seismic profiles (Korja et al., 1993). Although there are no exposed Proterozoic granitoids to indicate partial melting of Archean crust in the Ilomantsi and Kianta terrains, monzogranitic and pegmatitic intrusions dated at 1.82–1.80 Ga intrude allochthonous and autochthonous Archean basement along the eastern edge of the Ranua terrain and throughout the Southern Lapland terrain (Vaasjoki et al., 2001). This indicates that Archean lower crust was not too refractory for melting, but that the solidus for fertile rocks was attained only in areas that were sufficiently thickened, or where the Karelian province was underthrust beneath the Svecofennian province. For example, systematic Sm-Nd studies by Huhma (1986) show that some granitoids along the boundary zone between the Svecofennian and Karelian domains have initial εNd values of –1 to –4, indicating that Svecofennian collision had led to partial melting and assimilation of underthrust Archean crust by 1860 Ma. Ruotoistenmäki et al. (2001) also attributed variations in initial εNd within a suite of 1.86–1.85 Ga gabbroic to granitic plutons in the western part of the Iisalmi terrain to variable degrees of derivation from an evolved Archean lower crustal source – the lowest εNd value being –6.5 – and a juvenile enriched lithosphere. The shift to values closer to the depleted mantle Sm-Nd evolutionary trend away from the Karelian boundary zone has indeed been one of the strongest lines of evidence for arguing that the Svecofennian domain is not underlain by Archean crust (Huhma, 1986; Patchett and Kouvo, 1986). As well as implying a relatively enriched late Archean reservoir in the lower crust, the Archean-derived granitoids intruding the margin of the Karelian domain typically show weakly radiogenic lead in potassium feldspar, and low 208 Pb/206Pb ratios in zircon. These results are consistent with an Archean deep crustal source enriched in residual phases such as garnet and

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pyroxene, and with high Th/U and low U/Pb – in other words compositions complementary to much of the currently exposed Archean crust in this region. Widespread Paleoproterozoic melting of Archean crust in northern Finland is also evident from Lu-Hf studies (Patchett et al., 1981), Pb-Pb whole rock and potassium feldspar data (Meriläinen, 1976), and heterogeneity in zircon populations (Lauerma, 1982; Huhma, 1986; Mänttäri, 1995) as well as the Sm-Nd survey by Huhma (1986). Rämö (1991), in seeking appropriate Archean crustal Sm-Nd compositions for modeling partial melting of Archean crust during the formation of the Salmi rapakivi granite batholith in the southeastern part of the Karelian domain, noted that the εNd for the Nattanen granites, which intrude the Pomokaira terrain, is less negative than comparative values published for exposed Archean rocks, suggesting a more radiogenic source composition. The petrological and geochemical characteristics of the Nattanen granites is also consistent with derivation from an igneous source, and implies that Archean crust lies beneath the allochthonous Lapland granulite belt (Haapala et al., 1987). Öhlander and Skiöld (1994) conducted a similar Sm-Nd survey in adjacent northern Sweden and found that both 1.90–1.87 Ga calc-alkaline granitoids and 1.80 Ga felsic weakly peraluminous minimum-melt monzogranites record variable degrees of derivation from Archean crust; the latter, so-called Lina-type granites have εNd values as low as –9.3 at 1.80 Ga, compared to a mean value of –12.4 (Öhlander and Skiöld, 1994). Despite their minimum-melt features, the weakly peraluminous to metaluminous character and δ18O values between +5 and +8% mean that the Lina-type granites do not qualify as collisional S-type granites, irrespective of whether they were directly derived from Archean basement or indirectly through remelting of 1.9 Ga granitoids (Öhlander et al., 1987b; Öhlander and Skiöld, 1994). This suggests that the lower crustal Archean com76

ponent in their source material would have been predominantly of metaigneous rather than pelitic sedimentary character. Finally, Huhma (1986) and Huhma et al. (1990) noted that the Sm-Nd data for some basalts formed during early rifting of the Karelian domain imply relatively LREE-enriched compositions, which is difficult to reconcile with the established depleted mantle reservoir beneath the Fennoscandian Shield. It is not clear whether these features result from crustal contamination, as is evidently the case for komatiites erupted through Archean crust in Lapland (Räsänen et al., 1989), or metasomatism of the Archean subcontinental lithosphere during the various Paleoproterozoic rifting events (cf. Peltonen et al., 1998).

5.3. Xenoliths and deep seismic studies Xenoliths entrained by kimberlite diatremes and basalts are widely used to obtain information about the composition, age, and thermal evolution of the lower crust and lithospheric mantle (Rudnick, 1992; O’Reilly et al., 2001). It appears that there is not only a strong coupling between crustal formation and stabilization of underlying lithosphere, but also that there are subtle secular changes in mantle composition, which make Archean lithosphere inherently more buoyant and resilient to subduction and destruction (Griffin et al., 1999; O’Reilly et al., 2001). The recent recognition of two kimberlite clusters within the Iisalmi terrain (Griffin et al., 1995; Tyni, 1997) has provided an ideal opportunity for investigating the nature of the deep crust and lithosphere near the margin of the Karelian domain, and the extent to which Archean lithosphere has been modified by Svecofennian and younger processes (Peltonen et al., 1999; Hölttä et al., 2000b). Interpretations of results have also been complemented by heat flow data (Kukkonen and Peltonen, 1999) and comparison with the deep crustal density structure inferred from the SVEKA seismic refraction profile,

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which also passes through the Iisalmi terrain (Korja et al., 1993; Korsman et al., 1999). Studies of mantle xenolith populations (Peltonen et al., 1999) have revealed that the lithospheric mantle beneath the Iisalmi terrain is stratified, comprising two compositionally distinct layers. The upper layer is characterized by depleted harzburgite xenoliths in the garnet-spinel facies, derived from depths of 100–150 km. These rocks appear to have undergone metasomatic enrichment, probably in association with the kimberlitic magmatic event, since Nd and Sr isotope studies of xenoliths produce isochrons matching the ages obtained from the kimberlites (Tyni, 1997; Peltonen et al., 1999). The lower layer, at depths from 170–230 km, consists of garnet facies harzburgite-lherzolite-wehrlite, and also includes eclogites. The upper layer is regarded as Archean subcontinental lithospheric mantle, whereas there are three potential alternatives for the formation of the lower lithospheric layer. The first involves attenuation of Archean continental lithosphere during rifting and formation of a passive margin, most likely between 1.97 and 1.95 Ga (Peltonen et al., 1998), with accretion of a mafic underplate beneath what is now the western margin of the Karelian domain. The second alternative would be underthrusting of deep Svecofennian oceanic lithosphere, which may well have the appropriate residual cumulate-like geochemical signatures (Peltonen et al., 1999). This is more difficult to reconcile with the polarity of Svecofennian collision, except possibly at the later stages, around 1.86 Ga, when bimodal plutons were emplaced into the Iisalmi terrain (Paavola, 1991; Ruotoistenmäki et al., 2001). The third possibility is in relation to plume impingement and underplating at around 1.80 Ga, which is consistent with xenolith data in other parts of the Fennoscandian Shield and isotope characteristics of mafic magmatism (Eklund et al., 1998), as well as the abundant postorogenic 1.80–1.78 Ga granitic magmatism. Even if

isotope studies are unable to discriminate between these three alternatives, the geometry of the lower accreted lithospheric layer ought to differ, being more uniform over a wider area in the case of plume impingement compared to eastwards tapering for the rifting and attenuation scenario. In summary, while Kukkonen and Peltonen (1999) concluded from the absence of sheared fabrics in xenoliths that the petrologically defined lithosphere–asthenosphere boundary is at least 230 km deep, the composition of xenoliths (Peltonen et al., 1999) indicates that the lower part of the mantle lithosphere beneath the Karelian domain may be of Proterozoic rather than late Archean age. On the other hand, the upper part of the subcontinental lithospheric mantle is likely to be Archean, although it has evidently been modified during the latest Proterozoic or early Paleozoic. Studies of lower crustal xenoliths sampled by the kimberlites lend further support to the postorogenic 1.8 Ga underplating scenario. A considerable scatter in isotope results obtained by Hölttä et al. (2000b) from mafic granulites representing crystallization at depths corresponding to the present middle crust. Peltonen et al. (1999) also considered that residual cumulate composition in the lower part of the mantle lithosphere could well be complementary to mafic magmas emplaced in the lower crust. Therefore it is apparent that the Archean crust, as well as the mantle lithosphere includes a considerable component of Proterozoic material, and by implication, Proterozoic thermal reworking. If the Archean lower crust has indeed been modified and reworked by magmatic intrusion and underplating, this ought to be evident in deep seismic data. In Russia, the Moho depth may also have been modified by Paleoproterozoic rifting, magmatic underplating, and convergent tectonics. However, there is a relatively well-defined Moho from the Belomorian terrain westwards to the Finnish border, which steadily

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increases to about 40 km, before jumping to 37 km (Systra et al., 2001). The Europrobe BABEL reflection surveys have not transected the Archean crust, except where it is underthrust beneath Proterozoic rocks in northern Sweden, while the results of the FIRE reflection seismic surveys were not available at the time of writing. Refraction data are however available from the SVEKA profiles (Luosto et al., 1990; Korja et al., 1993; Korsman et al., 1999), which provide information on densities and depths of crustal layers. Results are complementary to the xenolith data in that thick crust, with relatively high density lower crust, is present beneath the Iisalmi terrain and the western margin of the Karelian domain generally. There appears to be a marked decrease in depth to the Moho eastwards, coinciding approximately with the position of the Kuhmo greenstone belt (Yliniemi et al., 1996; Korsman et al., 1999) and the zone of Proterozoic tectonic reworking along the boundary between the Rautavaara complex and the Ilomantsi terrain (Luosto et al., 1990). Although this is mostly attributable to the Svecofennian collision (Kohonen et al., 1991; Korsman et al., 1999), it is by no means clear to what extent this was controlled ultimately by the inherited late Archean lithospheric architecture. Farther east, in Russian Karelia, similar controversies in interpretation relate to discriminating between Proterozoic and Archean structures, across the boundary between the Belomorian terrain and the western part of the Karelian domain (Berzin et al., 2001; Samsonov et al., 2001; Slabunov and Bibikova, 2001).

6. Discussion and synthesis 6.1. Archean thermal regimes and tectonic consequences We can recognize several distinct magmatic and tectonic events in the Karelian domain 78

in Finland, but as yet there are insufficient constraints on large scale crustal architecture and timing for developing a robust and testable tectonic model. This is due as much to the lack of information and the effects of Proterozoic disruption as to concerns about fundamental differences between Archean and modern earth processes (Sleep and Windley, 1982; Hamilton, 1998). The issue is not simply whether or not the Archean thermal regime inhibited or allowed Phanerozoic-style plate tectonics, for even in the absence of convincing criteria such as ophiolites, or blueschist facies accretionary complexes, the existence of extensive strike-slip shear zones demonstrates that the continental lithosphere in the late Archean was sufficiently rigid to record large scale horizontal compression (Sleep, 1992). Of equal importance is the extent to which the Archean thermal regime influenced the degree of melting, lithospheric rheology, and post-collisional responses to thermal and gravitational disequilibrium within the crust. Studies of Archean high-grade terrains suggest that late orogenic geotherms in continental crust were not distinguishable from those in later orogens (Bickle, 1978; Griffin et al., 1980; Pollack, 1997). In contrast, considerations of the efficiency of heat loss from the Earth have led to propositions that Archean lithospheric plates would have been smaller, and spreading ridge length accordingly greater, than in the modern Earth (Bickle, 1978; de Wit et al., 1992), and that the frequency and consequences of plume-plate interaction were greater in the late Archean (Campbell and Griffiths, 1992; Wyman et al., 1999). The absence of documented blueschist facies terrains and estimates of the ambient mantle temperatures in the Archean being somewhere between 50–100 °C (Arndt, 2001) or 100–200 °C (Campbell and Jarvis, 1984) greater than in the younger Earth would be consistent with higher geothermal gradients in convergent tectonic settings (Martin, 1987a,b), such that hydrated oceanic lithosphere might have com-

• CHAPTER 2 • ARCHEAN ROCKS

menced melting beneath subduction zones, at depths less than 70 km and temperatures of 600–700 °C (Wyllie, 1979). In addition, greater degrees of melting would result in thicker and more buoyant oceanic lithosphere, that would be more resistant to subduction, even more so if extensively hydrated, which might promote crustal growth by lateral accretion of oceanic plateaux (de Wit et al., 1992; Abbot and Mooney, 1995). An elevated Archean thermal regime would have significant consequences for magmatism during collision, and thermal evolution of the crust following collision. Based to some extent on studies from the Kianta terrain, Martin (1987a,b) concluded that many trace element characteristics of tonalite–trondhjemite magmatism are consistent with extensive melting of subducted hydrated oceanic lithosphere, at relatively shallow depths, in contrast to Phanerozoic terrains, where arc magmatism is attributed to melting in the lithospheric wedge above the subduction zone. Ridley (1992) proposed that under an orogenic geotherm, much of the lower crust, of tonalitic composition, would be partially molten, which would regulate crustal strength and buoyancy. When the effects of radiogenic heat production are considered, as crustal anatexis enriches the middle crust in K, U, and Th, there is potentially an even greater effect on crustal strength (Sandiford and McLaren, 2002). For example, Jamieson et al. (1998) have conducted numerical modeling of the thermal evolution of collisional fold belts with temperature dependent rheologies. They found that the location of crustal units with high heat production has a significant effect on temperature distribution. Concentration of radiogenic elements is expected firstly in tonalitic to granodioritic magma derived from partial melting of contemporaneous or older mafic crust, and secondly from sediments buried within the accretionary prism. Thus some first order correlation might be expected between duration of an orogenic event, rate of

sediment supply, and the abundance and timing of late orogenic granitic magmatism and accompanying metamorphism. Is the buoyancy and strength, and hence equilibrium thickness of Archean lithosphere therefore a two-stage self-organizing phenomenon (Bak, 1996; Hodges, 1998), regulated firstly by lithospheric composition and heat flow and secondly by the temperature-dependent rheology of crustal rocks? Ultimately, the degree of interaction between anomalous thermal regimes related to plume activity and the rates of convergence and extension at plate boundaries, would determine whether crustal growth would take place as rapidly formed oceanic plateaux or magmatic arc complexes. Was there a critical crustal thickness under an elevated Archean geotherm that modulated intracrustal melting and differentiation of the crust into a mafic lower crust and felsic upper crust, analogous to the onset of felsic volcanism in modern Iceland (cf. Marsh et al., 1987), or was subduction of hydrated oceanic lithosphere – or even crustal duplexing – always required to attain appropriate P-T conditions for melting mafic lower crust? In either case, internal crustal differentiation, leading to upwards concentration of lithophile, radiogenic elements in tonalitic to granitic magmas could then significantly influence deformation style and metamorphic evolution at higher crustal levels, including perhaps the depth of the brittle–ductile transition. When the effects of perturbed orogenic geotherms and radiogenic heat production decay to a critical threshold for crustal strength, a given terrain may be considered stabilized, at least until subjected to some later anomalous event. According to O’Reilly et al. (2001), and Poudjorn–Domani et al. (2001), the Archean lithosphere is distinctive in terms of composition and buoyancy, which makes it inherently stable, unless metasomatized and infiltrated by younger magma to such an extent that it forms isolate relict domains within an essentially younger lithosphere, as in the south China

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craton. Lenardic et al. (1999) also argued that destruction of Archean lithosphere is inherently unlikely when surrounded by younger fold belts, which preferentially accommodate strain during subsequent collisional events. A general conclusion from the Karelian domain in Finland is that greenschist to lower amphibolite facies greenstone sequences tend to be steeply dipping and tightly folded, and intruded by discrete homogeneous plutons. In contrast, migmatite terrains, with relict supracrustal components (not merely mafic intrusives and cumulates) commonly have more gently dipping enveloping surfaces. This dichotomy in structural style and metamorphic grade suggests a thermal (and lithostatic loading) control on crustal rheology and mode of deformation. Similar relationships are apparent in other Archean terrains, notably the Yilgarn craton where a general trend of decreasing depth of exhumation from east to west can be inferred. In the western part of the craton, granulite facies gneiss terrains are intruded by monzogranites of similar age (Nemchin et al., 1994), which often form gently dipping sheets. Farther east, in the Southern Cross province, greenschist to amphibolite facies greenstone belts are steeply dipping and associated with domal plutonic complexes and large scale transpressive shear zones (Dalstra et al., 2000; Greenfield and Chen, 1999); this architecture closely resembles the structural relationships and metamorphism in the Kianta and Ilomantsi terrains of the Karelian domain. Still farther east, the extensive greenschist facies supracrustal sequences of the Eastern Goldfields province include extensive lowstrain domains, local tectonic imbrication and repetition of stratigraphy, and discordant plutons with zones of higher strain and metamorphic grade exposed in antiformal culminations. Interpretations of seismic reflection data (Swager et al., 1997; Drummond et al., 2000) are consistent with this variation in structural style with depth, showing an upper crustal layer with open to tight folding and duplex80

ing, terminating downwards at subhorizontal detachment zone, inferred to represent the base of the greenstone sequence. Below this, prominent reflectivity indicates asymmetric imbrication within the middle crust, with a more homogeneous lower crust, consistent with late orogenic lower crustal melting and magma transfer into the middle and upper crust. The detachment zone is interpreted as a fundamental rheological boundary, along which granitic sheets were emplaced, and feeding plutons emplaced as discrete intrusion into the upper crust (Drummond et al., 2000; Sorjonen-Ward et al., 2002). This seems to represent dynamic feedback between crustal strength, thickening and degree of melting. An inherent aspect of this process is that anatexis can occur in a contractional deformation regime and does not necessarily require or cause regional-scale extensional collapse or crustal thinning. However, a consequence of extraction of a volume of magma from one particular level in the crust and transfer to higher levels would effectively be equivalent to imposing a flattening strain on the original melt layer. Is it possible therefore that there is a coupling between melt production and crustal rheology that reinforces the transition between upper and lower crust, again in the manner of self-organizing systems (Bak, 1996; Hodges, 1998)? If Archean orogenic processes were regulated by internal responses as much as external factors, then the growth and reworking of Archean cratons could be argued as diachronous on a global scale, instead of viewing the end of the Archean as an abrupt global transition triggered by mantle cooling below a particular temperature threshold. For example, despite the global prevalence of Archean cratons stabilized around 2.7–2.6 Ga, others, such as the Pilbara and Kaapvaal cratons record progressive growth, differentiation, and stabilization in the time interval from 3.5 Ga to 3.0 Ga (de Wit et al., 1992; Bickle et al., 1993); by 2.9–2.7 Ga, the Pilbara craton provided a

• CHAPTER 2 • ARCHEAN ROCKS

remarkably stable platform environment for Hamersley basin sedimentation (Blake and Barley, 1992; Krapez, 1993) and the latter formed an orogenic foreland setting in which the Witwatersrand basin accumulated (Coward et al., 1995). Stabilization of the Fennoscandian Shield follows a similar pattern except that it is clearly several hundred million years younger and similar in age to the Superior craton in Canada and Yilgarn craton in Australia. Thermal equilibrium in the deep crust of the Karelian domain, if we take the zircon ages from Iisalmi terrain granulites as an indication of cooling to a postorogenic geotherm, had been attained by 2.63 Ga, alkaline and carbonatite magmatism was manifest soon after, and deposition of the earliest unconformably overlying volcanic and sedimentary units in Lapland occurred some time later at 2.5 Ga. Another typical element of collisional and accretionary terrains, not yet recognized in the Karelian domain, is the presence of late orogenic to post-collisional sedimentary basins. The Siilinjärvi carbonatite complex does provides some indirect evidence for latest Archean alkaline magmatism and extension, but there is no record of related sedimentary basins or volcanism. This contrasts for example with the post-collisional Timiskaming phase of terrestrial sedimentation and alkali magmatism in the Abitibi belt (Sutecliffe et al., 1993; Jackson et al., 1994), or the Merougil and Kurrawang sequences in the Yilgarn craton (Krapez et al., 2000). It should be noted, however, that the Kurrawang sequence consists predominantly of submarine mass flow deposits and is devoid of volcanogenic intercalations. Neither is it strictly a postorogenic sequence, although it overlies volcanic rocks of the Kalgoorlie sequence with an erosional discordance, as the two sequences nevertheless share the same tectonic fabrics and metamorphic history. The coarse clastic resedimented deposits in the Ilomantsi and Kianta terrains (Sorjonen-Ward, 1993) are the only known candidates for this type of

depositional environment in the Archean of Finland, but it is to be emphasized that they are intimately associated with the volcanic evolution of the greenstone belts. There is at present no stratigraphic or environmental framework for interpreting other diverse supracrustal sequences, such as the mature siliciclastic metasediments in the Tuntsa terrain and hydrothermally altered rocks in the Rautavaara complex. If the paragneisses of the Western and Central Puolanka Groups were conclusively shown to be Archean, then the sedimentological studies of Laajoki (1986; Chapter 7) would provide insights into Archean processes and paleogeography, and would delineate a major sedimentary and felsic volcanic province separating the Napapiiri and Pudasjärvi terrains from the Koillismaa and Kianta terrains to the east. Whether this has any fundamental significance in terms of accretion of two quite separate crustal units remains to be seen.

6.2. Regional scenarios and correlations At this stage, our understanding of the isolated Archean terrains of the northern part of the Karelian domain is too fragmentary to permit any synthesis of tectonic evolution. The same applies to attempts to reconstruct the early history of those terrains that contain rock units older than 2.8 Ga, such as the 3.2 Ga migmatites of the Iisalmi terrain and the early generation of migmatites and their paleosomes in the Kianta terrain. A number of tectonic models have been presented to explain magmatism in the Kuhmo greenstone belt (Piirainen, 1998; Taipale, 1998; Jegouzo and Blais, 1993, 1995). Despite the advances in interpreting and reconstructing eruptive and depositional processes and environments, and inferring magmatic sources and settings of mafic and ultramafic rocks from petrogenetic studies (Halkoaho et al., 2000; Puchtel et al., 1999), the intensity of reworking during later Archean events makes it difficult to derive a coherent,

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robust, and testable tectonic model. The oceanic plateau model for Archean continental growth is an appealing one (de Wit et al., 1992), and characteristics of komatiitic volcanism has been used to invoke collision with an oceanic plateau in the Kostamuksha greenstone belt (Puchtel et al., 1998). There are nevertheless lithostratigraphic constraints on application of such models in Finland, including the evidence for eruption of komatiites and high-Mg basalts in an ensialic environment, or at least the almost ubiquitous bimodal aspect of ultramafic volcanism accompanied by felsic magmatism. For the later events in the Karelian domain, an obvious plate tectonic scenario could be devised as follows. The 2.75–2.73 Ga coeval volcanism and tonalitic plutons in the Hattu schist belt are reminiscent of arc magmatism, while there is isotope evidence for derivation of some granitic magmatic and sediments from older continental crust. Older migmatites are exposed to the north and west, in the Iisalmi and Kianta terrains, which would be appropriate source material. The intervening Nurmes gneiss complex, and potentially the silicilastic and volcanic precursors to supracrustal gneisses in the Rautavaara complex would then be ideally placed as an accretionary wedge overridden by the Ilomantsi terrain. A polarity of this kind would also seem to provide an explanation for medium-pressure metamorphism in the underthrust Iisalmi terrain. But such a scenario becomes less tenable under closer scrutiny, unless we argue that the polarity of this event has been substantially obscured by younger crustal reworking. For example, the available kinematic constraints from the Ilomantsi terrain suggest N–NE dextral transpression during later stages of deformation, which is at least locally resolved in moderately dipping terrain as thrusting with a top to the east component. Similarly, there is limited reconnaissance mapping to suggest that the southeastern part of the Kianta terrain was thrust southwards and eastwards 82

with respect to the Ilomantsi terrain, which would imply a different polarity and setting for the Nurmes gneiss complex. These issues, as well as the time difference between arc-like magmatism in the Ilomantsi terrain (2.75 Ga) and deep crustal metamorphism in the Iisalmi terrain (2.68–2.63 Ga) are difficult to reconcile with such a simple collisional model. Abrupt changes in the kinematic framework of evolving orogens are of course not unusual in the modern Earth, but it is equally probable that the observed structural patterns reflect responses to different, superimposed tectonic regimes. Still more complex scenarios can be envisaged if we attempt to integrate constraints and concepts from the Russian part of the Karelian domain. For example, regional mapping combined with thermochronological studies (Bibikova et al., 2001) and reflection seismic studies (Berzin et al., 2001; Samsonov, 2001) strongly support the idea of tectonic accretion of the Belomorian terrain by westward or southwestward emplacement over the Karelian domain at around 2.7 Ga. The scale of this event is such that it ought to have had significant consequences for the Karelian domain in Finland. At this stage, we would envisage that the Kianta and Ilomantsi terrains were deforming within a common kinematic framework, characterized by NNE dextral transpression, or E–W compression partitioned into a combination of thrusting and NE-directed simple shear (Sorjonen-Ward et al., 1997). This at least would provide a mechanism for exhumation of the granulites and pyroxene-bearing granodiorites of the Lieksa complex while emplacing the Kutsu monzogranites and Naarva leucogranites over the Ilomantsi terrain. Because this represents an opposite sense with respect to the Belomorian thrusting, the Karelian domain is potentially an example of a doubly vergent orogeny (Koons, 1990) or records the formation of a backthrust retrowedge (Beaumont et al., 1994), which has been described from many orogens (Cook and Varsek, 1994),

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including Archean terrains (Sorjonen-Ward et al., 2002). However, more detailed field work combined with chronological constraints are required before the relationship between the formation and exhumation of Karelian highgrade terrains and adjacent lower grade terrains is adequately understood.

6.3. Comparisons and contrasts between Archean and Svecofennian crustal processes There are some intriguing parallels in the tectonic and thermal evolution of the Archean of eastern Finland and the Svecofennian domain in southern Finland. These are of interest when considering whether the formation of Archean lithosphere in itself exerts deterministic control on subsequent crustal processes and responses (O’Reilly et al., 2001; Poudjom-Domani et al., 2001) or whether the Archean to Proterozoic transition indeed records secular changes in Earth processes, particularly thermal regimes, through radiogenic heat production (cf. Kukkonen and Lahtinen, 2001). In both areas, linear belts of greenschist to lower amphibolite grade metasediments and volcanic rocks, with steep enveloping surfaces and simple structural geometry are juxtaposed against migmatitic gneiss terrains of broadly coeval age. Neither area has preserved significant amounts of lateto postorogenic sedimentary basins, which may be an indication of either the failure to form extensive areas of topographically elevated terrain, or an erosional artefact relating to isostasy and crustal composition. A qualitative correlation between steep enveloping surfaces for foliations in low-grade greenstone belts and more gently dipping enveloping surfaces higher grade migmatite gneiss terrains may represent a fundamental thermal and rheological contrast and decoupling between the upper and lower crust. In both areas too, there is a distinct phase of thermal and tectonic reworking, some fifty million years after crustal formation, resulting

in widespread potassic granitic magmatism and granulite facies metamorphism, though barometric data indicate that the presently exposed Archean granulites record somewhat deeper crustal levels than in the Svecofennian (Väisänen and Hölttä, 1999; Hölttä and Paavola, 2000). Until better age constraints become available, it also seems that the late orogenic Archean felsic magmatism includes enderbites as well as monzogranites, which is not the case in the Proterozoic of southern Finland. A magmatic underplating event has been invoked throughout various parts of the Fennoscandian Shield at around 1.8 Ga (Eklund et al., 1990; Kempton et al., 2001; Markwick and Downes, 2000) but this postdates, and cannot be the cause of the 1.84–1.80 Ga Svecofennian granitic magmatism and metamorphism. Similarly, in the Archean of eastern Finland, there is as yet no record of late orogenic mafic magmatism that might have caused extensive intracrustal melting. Although modeling of seismic refraction and gravity data indicate potential mafic layers in the deep crust in eastern Finland (Korsman et al., 1999), it will be recalled that xenolith studies record Proterozoic rather than Archean underplating (Hölttä et al., 2000b). Clearly, both regions need to be examined carefully for evidence of a mantle magmatic input at this time, or alternatively, metamorphic evidence for rapid decompression associated with granite emplacement, such as might be expected if the heat source was a consequence of delamination of tectonically thickened lithosphere (cf. Houseman et al., 1981). The 50-Ma time lag between Svecofennian arc magmatism and the potassic granites, and the similar delay recorded in the Ilomantsi terrain is an important constraint on interpretation. A delay of this magnitude might be expected between impingement of a plume at the base of the lithosphere and the onset of partial melting in the middle crust, in the case of conductive heat transfer (Hobbs et al., 1998). However, from the point of view of lithospheric delamination following collision

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(England and Houseman, 1989) a 50-Ma gap between collision and uplift and anatexis in the middle crust is a rather long time frame. As an alternative, the effect of redistribution of radiogenic elements within the crust on late orogenic thermal evolution might be considered as a more viable explanation (Jamieson et al., 1998; Sandiford and McLaren, 2002).

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LAYERED MAFIC INTRUSIONS OF THE TORNIO– NÄRÄNKÄVAARA BELT

M. Iljina, E. Hanski

Cover page: Magmatic layering in ultramafic zone of Megacyclic unit I, Penikat intrusion. Tag width 5 cm. Photo: Vesa Perttunen.

Iljina M., Hanski E., 2005. Layered mafic intrusions of the Tornio–Näränkävaara belt. In: Lehtinen, M., Nurmi, P.A., Rämö, O.T. (Eds.), Precambrian Geology of Finland – Key to the Evolution of the Fennoscandian Shield. Elsevier B.V., Amsterdam, pp. 101–138. © 2005 Elsevier B.V. All rights reserved.

Most of the ~20 of Finland’s early Paleoproterozoic layered mafic–ultramafic intrusions are found in a roughly E–W trending, 300-km-long belt in northern Finland. Known as the Tornio–Näränkävaara belt, it represents a major failed rift system into which large volumes of mafic and minor A-type granite magma were intruded at ~2440 Ma. The mafic intrusions have late Archean felsic gneisses on their southern side and Paleoproterozoic volcano-sedimentary sequences on their northern side. Deposition of these supracrustal sequences took place on the unconformity that truncates the igneous layering of the mafic intrusions. This indicates relatively shallow depth of intrusion and rapid uplift and erosion. Composition of chilled margins, cumulates, and cogenetic dikes as well as established crystal fractionation sequences allow recognition of three different parental magmas. Two of these resemble siliceous high-Mg basalt (SHMB) and are thus akin to the Bushveld B1 and B3 magmas, the third is a more evolved tholeiitic basalt. The SHMB types are found in the western and central parts of the Tornio–Näränkävaara belt, the third in the eastern part of the belt. All the mineralization types characteristic of layered mafic intrusions are present. These include chromite and PGE-enriched base metal sulfides in the bottom parts of the intrusions, stratiform PGE, chromite, and magnetite enrichments higher in the cumulate sequences, and offset PGE-base metal deposits below the intrusions. A world-class chrome deposit is located at the base of the Kemi intrusion and a magnetite gabbro of the Koillismaa complex has been exploited for vanadium. Five potentially world-class reef-type PGE deposits are distributed among three separate intrusions: Penikat, Suhanko, and Narkaus. Sulfide mineralization in the marginal series shows, in places, high PGE concentrations relative to typical basal sulfide mineralizations. The location of the reefs and high-grade PGE marginal series seems to be controlled by the megacyclic structure of the intrusions. This, together with the compositional similarities of the intrusions (mineral, modal, whole-rock, PGE), suggests that the magmas that formed these three intrusions and the chromite-bearing Kemi intrusion had a common history in a lower-level auxiliary magma chamber before emplacement.

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1. Introduction The gabbroic and serpentinitic ultramafic rocks of the Tornio–Näränkävaara belt were already outlined and mapped in the first half of the 1900s. However, it was not until the 1960’s and 1970’s that these intrusions were commonly accepted as differentiated bodies of basic magma and the similarity between these and other layered mafic intrusions (e.g., the Rustenburg Layered Suite) was acknowledged. Still, at this stage, the internal structure and crystallization ages remained rather poorly known. Extensive exploration in the 1980’s and isotope age determinations accumulated evidence suggesting that the Tornio–Näränkävaara belt is part of a globally recognized episode of mafic igneous activity at ~2450 Ma. The results of these investigations, carried out mostly by the University of Oulu, Outokumpu Oyj, and the Geological Survey of Finland, have been reported in numerous publications and academic theses (e.g., Alapieti, 1982; Lahtinen, 1985; Alapieti and Lahtinen, 1986, 2002; Alapieti et al., 1989a; Lahtinen et al., 1989; Halkoaho et al., 1990a,b; Huhtelin et al., 1990; Huhma et al., 1990; Iljina et al., 1992; Iljina, 1994; Iljina et al., 2001). Mafic ~2450 Ma intrusions represent intracratonic plume-related igneous activity and are characterized by relatively high MgO and Cr contents and relatively high SiO2 compared with MgO. These magma types have been referred to either as basaltic komatiites, boninites, or siliceous high-magnesium basalts and have been found to be prospective for Ni-Cu-PGE sulfide, PGE, and Cr- and Fe-Ti-V oxide deposits. The South African Bushveld Complex, Zimbabwean Great Dyke, Chinese Jinchuan intrusion and Finnish Kemi intrusion host well-known examples of economic deposits of these types, the last mentioned being part of the Tornio–Näränkävaara belt. In this paper we summarize the main geological features of the Tornio–Näränkävaara belt, with an emphasis on related mineraliza104

tion. This information is supplemented by new data from the Penikat intrusion and Portimo and Koillismaa layered igneous complexes.

2. Geologic setting of the Tornio–Näränkävaara belt The Tornio–Näränkävaara belt is a discontinuous zone of layered intrusions crossing northern Finland almost along the Arctic Circle and extending some kilometers into Sweden (Tornio intrusion) and several tens of kilometers into Russia (the Olanga complex). The belt contains roughly half of the 2.4–2.5 Ga layered igneous complexes within the Fennoscandian Shield. This widespread pulse of mafic magmatism has been interpreted as representing the initial stage of continental rifting (e.g., Piirainen et al., 1974; Amelin et al., 1995). Some of the intrusions are located close to each other, thus forming igneous complexes in which the individual intrusions were probably connected by dikes or intermediate magma chambers at the time of emplacement. Some time after crystallization the intrusions were faulted by multistage deformation into several smaller blocks, some of which can be found today as independent igneous bodies. The deformation was also accompanied by greenschist to lower-amphibolite facies metamorphism resulting in replacement of the primary igneous minerals by low-temperature assemblages. Recrystallization was especially pervasive in the central part of the belt, leaving essentially no primary mafic silicates in the cumulate rocks; nevertheless, the original texture is usually still recognizable. On the other hand, in the eastern and western parts of the belt, magmatic minerals are well preserved in many places. The Tornio–Näränkävaara belt consists of the Tornio, Kemi, and Penikat intrusions in the west, the Portimo layered igneous complex (Portimo complex) in the middle, and the Koillismaa layered igneous complex (Koillismaa

• C H A P T E R 3 • L AY E R E D M A F I C I N T R U S I O N S

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Kuohunki Nutturalampi Kilvenjärvi Lihalampi Siika-Kämä Konttijärvi intrusion

Suhanko intrusion

Fig. 3.1. Mafic–ultramafic layered intrusions (black) in the Tornio–Näränkävaara belt (simplified after Korsman et al., 1997). Low altitude aeromagnetic map is shown from the area of the Koillismaa layered igneous complex. The “connecting dike” in the Koillismaa area refers to a strong magnetic and gravimetric anomaly joining the Näränkävaara intrusion to the Pyhitys and Kuusijärvi blocks of the Western intrusion. The Western intrusion also comprises the Pirivaara, Syöte, Porttivaara, Tilsa, Lipeävaara, Kaukua, and Murtolampi blocks.

complex) in the east (Figure 3.1). All but the last are found at the southern or southeastern margin of the Peräpohja schist belt or close to the margin within the adjacent Archean Pudasjärvi basement complex (Figure 3.1). In fact, all intrusions or intrusion fragments, even those surrounded by Archean gneisses, have a cap of supracrustal rocks (see inset of Figure 3.1). Unequivocal evidence has been obtained demonstrating that the lowermost supracrustal rocks of the Peräpohja schist belt, including polymictic conglomerates, are younger than the layered intrusions and were deposited unconformably on the tilted, uplifted and partly eroded layered intrusions (Perttunen, 1991). The deposition of these supracrustal rocks is interpreted to have taken place at >2.3 CHAPTER

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Ga, suggesting a relatively shallow depth of emplacement of the layered intrusions as well as rapid uplift and erosion. The contact between the Western intrusion of the Koillismaa complex and the greenstone belt on its western side is tectonic. However, there is a small supracrustal package between the Kuusijärvi and Lipeävaara blocks as well as between the Syöte and Porttivaara blocks (Figure 3.1). These rocks contain (in stratigraphic order) felsic–intermediate volcanic rocks, conglomerates, quartzites, and mafic metavolcanic rocks that can be correlated with the lowermost rocks of the Kuusamo schist belt. At least the conglomerates and overlying supracrustal rocks are younger than the intrusion. This accords with observations

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from the southeastern margin of the Peräpohja schist belt. Exploration has revealed a number of mineral deposits, of which the Kemi chrome ore is world-class in size and outstanding with respect to the economic cluster it has created. In addition to Cr oxide deposits, one titanian magnetite deposit has been exploited for vanadium. PGE-enriched base metal sulfides are found at the base of some intrusions and numerous PGE reefs grading to exploitable concentrations of PGE have been delineated.

3. Cumulus sequences 3.1. General characteristics The general stratigraphy of the Tornio–Näränkävaara belt intrusions can be divided into a marginal series and an overlying layered series. The marginal series represents a basal reversal, in which rocks become more primitive upwards. There are two kinds of marginal series successions in terms of thickness: one is thin ( 1

Basic hyaloclastic tuffite, carbonate rocks Pillow lava Massive lava Pillow breccia and hyaloclastite

Chromitites

Sheeted dikes, gabbro, and mantle peridotite screens Fe-Ti-gabbro and plagiogranite (1954 ± 11 Ma) Isotropic gabbro (1960 ± 12 Ma)

“Early” OIB-type dikes (~2.1 Ga) Deep dikes

Gabbroic dikes (1953 ± 2 Ma)

Gabbro pods Mantle tectonite (serpentinite)

Clinopyroxenitic mantle dikes (~2.1 Ga)

Mantle foliation

Hornblenditic mantle dikes (~2.1 Ga)

Fig. 6.3. Stratigraphic reconstruction of the Jormua ophiolite. The lowermost unit separated by a fault refers to the western block of the ophiolite (see Table 6.1 and Figure 6.2). The western block is lithologically distinct from the remaining ophiolite. Recent ion microprobe age determinations (Peltonen et al., 2003) suggest that the ~2.1 Ga clinopyroxenitic dikes from both the central and western blocks contain inherited Archean zircon grains and thus these blocks represent ancient subcontinental lithospheric mantle. Hornblenditic dikes within the western block and “early” OIB dikes at the central block are most likely related and Paleoproterozoic in age, being older than the main suite basalts and gabbros.

km in thickness and several square kilometers in extent and also as individual dikes intruding mantle tectonites deeper in the ophiolite stratigraphy. The presence of lava, gabbro, and mantle tectonite as interdike screens suggests that the contact between the main ophiolite units are transitional. Dikes in the sheeted dike complexes are generally 20–120

cm thick, aphyric or plagioclase-phyric with sharp chilled mutual contacts (Figure 6.4B). Half-split dikes and marginless septa are common, attesting to an extensional setting typical of ophiolitic dike-in-dike complexes. These dikes are generally well-preserved, whereas individual dikes deeper in the ophiolite stratigraphy are strongly altered because of the CHAPTER

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ser pentinization of the adjacent peridotites (Figure 6.4C). Two distinct types of basalts are present in the Jormua ophiolite, the “main suite” basalts and “early dikes.” The former include all the MORB-type lavas and sheeted dike complexes, whereas the latter are OIB-type and occur as subordinate dikes deep in the ophiolite stratigraphy (Figure 6.3). Field observations imply that the emplacement of the “early dikes” preceded that of the “main suite” basalts. The main suite basalts are subalkaline EMORB with flat chondrite-normalized REE patterns (Figure 6.5A) and only moderately depleted Nd composition [εNd (at 1950 Ma) ~ +1.9]. Most of the basalt samples, especially lavas, cannot be related to each other by fractional crystallization (see below) but instead represent distinct, rather primitive melt fractions directly fed from an asthenospheric diapir. This is consistent with the absence of large cumulate units (magma chambers), where pre-eruption fractionation would have occurred. The chemical composition of the OIB-type “early dikes” is truly distinct from that of the main suite basalts. They have high Nb/Y similar to alkali basalts or basanites. Their low Al2O3, high Cr and Ni, fractionated LREE and HREE (Figure 6.5A) together with non-depleted εNd (at 1950 Ma) close to zero are

compatible with an origin as ultramafic lamprophyre melts derived from a mantle domain in the stability field of garnet. The chemical and Nd isotope composition of the basalts implies that two distinct mantle sources were incolved. Peltonen et al. (1996b) modeled trace element abundances and came to the conclusion that the “early dikes” represent melts from a distinct OIB-like deep mantle source. The “early dikes” thus provide important evidence for the existence of OIB-type mantle sources already at 2 Ga. The modeling further suggested that the main suite lavas and sheeted dikes were not derived from a normal depleted mantle source either. Trace element ratios imply that they contain a small and rather uniform proportion of an OIB-like component and that their chemical composition is consistent with mixing of a NMORB end member with a small amount of an an OIB-like end member. Magma mixing was considered unlikely because of the complete absence of compositionally intermediate dikes between the MORB-like main suite basalts and OIB-type “early dikes.” Instead, Peltonen et al. (1996b) suggested that the OIBlike dikes were emplaced during the initial stages of continental rifting and oceanic basin formation. Meanwhile, they metasomatized the uppermost convective mantle from which

Fig. 6.4. (facing page) (A) Hydrothermally altered pillow lava. Note the concentrically zoned pillows with vuggy interiors and fine-grained pillow rims against the hyaloclastic interpillow matrix; Asko Kontinen for scale. (B) Outcrop of sheeted dike complex consisting of 100% of subparallel EMORB dikes. Plagioclase-phyric dikes (with drill holes) are being cut by slightly younger apphyric dikes with chilled margins. Diagonal light streaks are traces of late fractures. (C) Main suite EMORB dikes (“deep dikes”) intruded into mantle tectonites. The dark dike margins are due to postmagmatic dike–peridotite interaction during serpentinization and regional metamorphism. (D) Gabbroic feeder dike (dark) intruding mantle tectonite. Note the prominent concentration of plagioclase (now largely epidote) into the core of the dike. (E) “Knobby”-textured mantle peridotite with serpentine pseudomorphs after orthopyroxene standing up with higher relief. (F) Small massive chromitite pod (black) approximately 1by ≥5 m in size. (G) Clinopyroxenitic cumulate dike (brown weathering surface) intruding mantle peridotite. (H) Garnet-bearing hornblenditic mantle dike, garnet (white pseudomorphs) crystals define comb-layering. Photos by the author except (A) by Ari Linna, and (B), (C), and (F) by Asko Kontinen; (D), (E), (G), and (H) reprinted with the permission from Oxford University Press.

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Lavas and dikes

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Fig. 6.5. (A) Chondrite (Boynton, 1984) normalized rare earth element patterns for lavas and basaltic dikes from the Jormua and Outokumpu ophiolites. For similar patterns of Nuttio basalts the reader is referred to Chapter 3 of this volume. The Jormua ophiolite contains two distinct suites of basaltic rocks: EMORB type lavas and dikes with flat chondrite normalized patterns and less common OIBtype dikes with fractionated patterns. Note that the basalts spatially associated with Outokumpu-type ultramafic massifs have lower absolute REE abundances and LREE depleted patterns indicative of their derivation from more depleted sources than the Jormua EMORBs. (B) Chondrite-normalized rare earth element patterns for gabbro and plagiogranite samples from Outokumpu and Jormua. Note the generally lower REE abundances of the Outokumpu gabbros compared to those from Jormua consistent with their coeval formation with the associated basalts. Plagiogranites from Jormua are characterized by more fractionated patterns (accompanied by negative Eu-anomaly) than Jormua gabbros.

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the main suite basalts were soon to be generated. Alternatively, the source of the main suite obtained its OIB-like component through thermal erosion of the base of the old OIBmetasomatized subcontinental lithospheric mantle. Importantly, the absence of any kind of geochemical subduction signature in the basalts implies that the Jormua ophiolite did not form in an arc-related geotectonic setting.

Gabbros and plagiogranites Gabbros are a subordinate component of the Jormua ophiolite. Two main types are present: (a) relatively large, up to > 1 km2 size upper-level stocks spatially associated with volcanic rocks and (b) thin, only a couple of meters wide but tens to hundreds of meters long gabbro dikes intruding and brecciating mantle tectonites (Figure 6.4D). Most of the group (a) gabbro intrusions can be regarded as belonging to the oceanic crustal unit, but some are completely enclosed by mantle tectonites. They range in composition from high-Mg olivine gabbros to ilmenite-rich ferrogabbros with minor tonalite–trondhjemite segregations, which closely resemble oceanic plagiogranites of younger ophiolites and modern oceanic ridges (Kontinen, 1987). The group (b) gabbroic dikes are found stratigraphically beneath the upper-level gabbro stocks. Field evidence suggests that they represent feeder dikes for the upper-level gabbro bodies. Samples from the upper-level gabbro stocks, gabbroic feeder dikes, and plagiogranites have yielded a whole-rock + clinopyroxene Sm-Nd isochron of 1936 ± 43 Ma with an initial εNd (at 1950 Ma) of +2.0 ± 0.3. Importantly, the average main suite basalt plots exactly along this isochron implying that the lavas, sheeted dikes, plagiogranites, upper-level gabbros, and gabbroic feeder dikes are cogenetic and represent progressively deeper expressions of the oceanic crust-forming magmatism in Jormua (Peltonen et al., 1998). Originally, both types of gabbros consisted of low-pressure plagioclase+clinopyroxene±olivine cumulates. How-

ever, their internal structures and alteration of primary minerals are distinct. While the upper-level gabbros frequently underwent extensive closed-system fractionation, the gabbro dikes crystallized in dynamic conduits and developed mineral layering parallel to the conduit walls. Locally, large clinopyroxene phenocrysts occur aligned parallel to the dike margins. Some crystals show microtextures indicative of pervasive ductile deformation and they may represent “megacrysts” transported from deeper levels of the mantle (Peltonen et al., 1998). In such feeder dikes, the dike centers are composed of progressively more evolved cumulates (Figure 6.4D). The gabbro stocks and feeder dikes also underwent distinct types of alteration: while olivine and clinopyroxene were replaced by chlorite and amphibole in the high level gabbros, the feeder dike gabbros became rodingitized due to serpentinization of the enclosing mantle peridotites. In the AFM diagram of Irvine and Baragar (1971), samples from the upper-level gabbro stocks and feeder dike gabbros form separate groups. First, upper-level gabbro samples show extensive compositional range along the MgO–FeOtot join, indicative of extensive tholeiitic fractional crystallization of their parental magmas (Figure 6.6). These gabbros range from primitive Mg-gabbros to ferrogabbros that may contain up to 10 vol.% ilmenite. Low abundances of incompatible elements, such as REE, imply that the amount of intercumulus liquid in the gabbros is low and that postcumulus growth took place (Kontinen, 1987). Chondrite-normalized REE patterns (Figure 6.5B) remain subparallel through the crystallization sequence with all showing clear positive Eu-anomalies. Such patterns indicate that the accumulation of olivine (+spinel) and plagioclase have controlled the cumulate compositions, whereas clinopyroxene or amphibole fractionation was less important. The feeder dike gabbros have a similar range in MgO–FeOtot and similar REE patterns but are CHAPTER

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Fig. 6.6. The AFM diagram for Jormua and Outokumpu gabbros and plagiogranites. Boundary between tholeiitic and calc-alkaline series after Irvine and Baragar (1971). Note how the gabbroic feeder dikes are depleted in alkalies due to rodingitization reactions.

extremely depleted in alkalis. This is a typical compositional feature of gabbros that have been enclosed by peridotites undergoing serpentinization. Such gabbros typically become depleted in silica and enriched in calcium, and lose their alkalies due to interaction with serpentinizing hydrous fluids. Ultimately, they become transformed into grossular and diopside-bearing rodingites – “by-products of serpentinization” (e.g., O’Hanley, 1996). This implies that the feeder dike gabbros at Jormua that have the typical metarodingite mineral assemblage diopside-epidote-amphiboles-grossular garnet, were emplaced into the peridotite protoliths before extensive serpentinization of their host rocks. The plagiogranite analyses plot along the (Na2O+K2O)–FeOtot join in the AFM diagram and show extreme alkali (sodium) enrichment 252

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(Figure 6.6). Plagiogranites have equal Zr/Y with high-level gabbros and typically occur as segregations and dikes within highly fractionated gabbro pods (Kontinen, 1987). They have yielded a crystallization age equal to that of the gabbros (~1.95 Ga, Table 6.2), implying that their origin is intimately related to the oceanic crust-forming magmatism. The REE patterns of plagiogranites are more fractionated than those of the most evolved gabbros and show pronounced negative Eu-anomalies.

4.2. The mantle section The well-exposed mantle section makes the Jormua ophiolite unique among ancient ophiolites. It permits the direct study of processes that took place in the upper mantle during the early Proterozoic continental breakup and

formation of a new oceanic basin (Peltonen et al., 1998). In fact, mantle rocks cover approximately 70% of the total exposure of the Jormua ophiolite (i.e., > 30 km2). The mantle section consists of mantle peridotites and various types of intrusive rock types. Because of their intimate genetic relationship with the crustal unit, basaltic dikes and gabbroic feeder dikes, which also intrude mantle tectonites, were described already in the preceding section. In addition, the western block peridotites are veined by abundant clinopyroxenitic and hornblenditic dikes that do not have counterparts in the crustal unit. They are not coeval with the formation of the oceanic crust in Jormua and are therefore described separately below.

Serpentinites Most of the Jormua mantle sequence consists of thoroughly serpentinized and regionally metamorphosed lherzolites and harzburgites, which do not show any evidence for magmatic layering or cumulus textures. Instead, their textures and chemical compositions – discussed in more detail in Section 7 – are consistent with them representing mantle peridotites that have undergone variable degrees (~7–25%) of partial melting. The primary mineralogy of the peridotites has been nearly completely destroyed by multistage serpentinization and regional metamorphism with the exception of occasional chromite relicts. Still, the central parts of the larger serpentinite massifs display obvious mantle tectonite fabrics and foliation defined by bastite pseudomorphs after elongated/flattened orthopyroxene crystals (Figure 6.4E). This foliation is intersected at steep angles by 1950 Ma gabbroic feeder dikes (emplaced >50 Ma before the onset of the regional deformation), which clearly implies that this foliation must be of a mantle origin. Locally, some serpentinite domains are moderately enriched in altered chromite and the possibility remains that they represent small dunitic cumulate pods within the residual peridotites.

These dunites are not, however, comparable to the thick layered cumulate sequences common in many younger ophiolites. Chromite is the only primary mineral that has been preserved to some extent in Jormua metaperidotites. It occurs as discrete grains or is sometimes concentrated into thin seams. Most of the grains are thoroughly altered but occasionally translucent deep red chromite cores are present and surrounded by ferrian chromite and chromian magnetite. The present silicate mineralogy of the mantle tectonites is dominated by non-pseudomorphic antigorite. Such non-pseudomorphic textures form through recrystallization of pseudomorphic serpentine textures or directly through hydration of Fe-Mg silicates at elevated temperatures (O’Hanley, 1996). Bastite ovoids represent pseudomorphosed primary orthopyroxene and the intervening antigorite domains with some magnetite dust derive from mantle olivine. Stable prograde mineral parageneses vary according to the bulk-rock-composition of the serpentinites. The antigorite-olivine-tremolite assemblage, for example, belongs to the ideal prograde sequence of metamorphosed serpentinites equilibrated at the lowermost-amphibolite facies (Will et al., 1990). In some less calcic samples the stable mineral paragenesis is antigorite-olivine. Qualitative estimates for the metamorphic peak temperature at Jormua are 480 and 530 °C for pressures of 2 and 5 kb, respectively. Later, metamorphic olivine and tremolite became partly replaced by pseudomorphic lizardite. Talc-carbonate alteration is present as narrow marginal zones of serpentinite massifs. Talc-carbonate rocks consist of carbonate and talc in approximately equal proportions, together with some magnetite and sulfides (pyrite, pyrrhotite, pentlandite, gersdorffite, and trace chalcopyrite). The carbonate-talc and antigorite-carbonate-talc assemblages stabilized under the same prograde conditions but at significantly higher XCO2 than the carbonate-free mineral assemblages. CHAPTER

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Although the general lithological characteristics of the eastern block suggest that the peridotites associated with the gabbros and sheeted dikes represent oceanic lithospheric mantle stabilized at ~1.95 Ga, the Re-Os study of Tsuru et al. (2000) resulted in a different interpretation. They demonstrated that Re-Os isotope composition of chromite separated both from eastern block serpentinites and local chromitite boulders are consistent with closed-system behavior. Chromite from serpentinites yield very depleted present-day 187 Os/188Os with an average calculated initial γOs(at 1950 Ma) of –5.1 ± 0.8. Such a negative value requires that the peridotites were depleted in Re already approximately one billion years before the time of the formation of the Jormua ophiolite at 1.95 Ga and, therefore, they most likely represent old subcontinental lithospheric mantle (SCLM). This suggests that true oceanic mantle (asthenospheric diapir at 1.95 Ga) is probably not exposed at Jormua, but that all peridotites represent stretched slivers of the subcontinental lithospheric mantle. This does not contradict with the presence of ocean floor basaltic rocks (dikes, lavas, gabbro pods) within the eastern block peridotites. In a compatible scenario, listric faulting would have exposed SCLM at the incipient oceanic basin, which subsequently became intruded by basalts fed from the underlying asthenospheric diapir (for more details, see Section 8).

Clinopyroxenitic and hornblenditic mantle dikes of the western block As emphasized above, the western block peridotites of the Jormua ophiolite are unique in being intruded by clinopyroxenitic and hornblenditic cumulate dikes (Table 6.1; Figure 6.4G). Such dikes are not typical of oceanic mantle units but are a more typical feature of the subcontinental lithospheric mantle. Peltonen et al. (1998) stressed the similarity of these intrusive rocks with those found within orogenic lherzolite massifs of the French Pyrenees – particularly that of Lherz 254

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(Conquéré, 1971; Bodinier et al., 1987a, 1987b; Fabriés et al., 2001). Clinopyroxenite dikes are medium-grained ortho- and mesocumulates. Clinopyroxene is the only cumulus mineral and has been extensively replaced by secondary low-Al actinolitic amphibole. Hornblenditic dikes form a more heterogeneous group of dikes: they include mediumgrained hornblendite dikes and veins which may contain garnet, pegmatitic varieties, garnetite veins, and carbonatitic segregations (Figure 6.4H). The primitive mantle-normalized REE patterns are particularly informative in petrogenetic studies. Clinopyroxenites have low abundances of REE and slightly upwardconvex patterns consistent with clinopyroxene accumulation (Figure 6.7A). The mantle normalized pattern shapes for the hornblendites clearly reflect their mineralogical composition. Hornblenditic samples yield patterns similar to those expected for pure hornblende on the basis of published partitioning coefficients (Figure 6.7B). In addition, dikes with abundant garnet pseudomorphs yield HREE-enriched and LREE-depleted pattern shapes indicative of accumulation of garnet (Figure 6.7C). Transitional cumulates are dikes which contain both magmatic amphibole and garnet (now preudomorphosed) in varying amounts (Figure 6.7D). The presence of magmatic garnet in these dikes is indicative of their relatively high crystallization pressures of the order of 10–15 kb (Green, 1969; Vétil et al., 1988). The exact timing of the emplacement of clinopyroxenitic and hornblenditic dikes is not well-constrained. Two clinopyroxenite dikes, one from the western and one from the central block, so far dated by ion microprobe, contain two concordant zircon populations with distinct ages: Archean (~2.7–2.8 Ga) and Paleoproterozoic (~2.05 Ga and 1.95 Ga), (Peltonen et al., 2003). This age data imply that at least the western block peridotites must represent Archean subcontinental lithospheric mantle. Hornblenditic dikes (and related carbonatitic veins) have yielded crystallization ages equal

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B

50

50

Clinopyroxenitic mantle dikes 10

10

1

1

La Ce Pr Nd

La Ce Pr Nd

Sm Eu Gd Tb Dy Ho Er Tm Yb

C 50

D 50

10

10

1

Garnet-rich mantle veins

La Ce Pr Nd

Sm Eu Gd Tb Dy Ho Er Tm Yb

Hornblenditic mantle dikes

1

Sm Eu Gd Tb Dy Ho Er Tm Yb

Transitional mantle dikes

La Ce Pr Nd

Sm Eu Gd Tb Dy Ho Er Tm Yb

Fig. 6.7. Primitive mantle-normalized (McDonough and Sun, 1995) REE patterns for clinopyroxenitic and hornblenditic mantle dikes from the western block of the Jormua ophiolite. Clinopyroxenites are equigranular ortho- and mesocumulates (A), whereas hornblendites form a more heterogeneous suite, consisting of pure hornblendites (B), garnet-rich dikes (C), and transitional cumulates (D).

to or slightly older than those of gabbros and plagiogranites (Table 6.1) and could represent alkaline magmatism related to the initial stages of continental rifting. They do not have their counterparts in the crustal sequence of the Jormua ophiolite and therefore it is probable that the magmatism evolved from early OIBtype magmatism towards EMORB-type in the course of continental breakup. It is likely that the mantle peridotites of those ophiolitic blocks that contain either OIB-type, clinopyroxenitic or hornblenditic dikes represent the remnants of the Archean subcontinental lithospheric mantle. It is interpreted that the clinopyroxenites and OIB-type dikes and hornblendites were emplaced in the SCLM at ~2.1

Ga. The involvement of the ascending asthenospheric diapir and associated magmatism at 1.95 Ga inevitably led to intense heating of the adjacent streched remnants of the Archean SCLM, and resulted in strong recrystallization of primary 2.7 Ga and 2.1 Ga zircon crystals in these dikes into anhedral metamorphic grains, with ages close to 1.95 Ga (Table 6.2).

5. Outokumpu-type ultramafic massifs The second occurrence of ophiolitic rocks is found within the North Karelia schist belt, which is located at the junction of the CHAPTER

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N

Losomäki

Miihkali Luikonlahti KUOPIO

Kylylahti

Sola OUTOKUMPU Täilahti JOENSUU

Kivijärvi

ec Sv es

nid

en of

Petäinen

20 km

Puiroonmäki

Serpentinite massifs (~1.97 Ga) Allochthonous metaturbidites “upper Kaleva” Autochthonous metaturbidites “lower Kaleva”

Jatulian: mainly quartzites, minor metavolcanic and calc-silicate rocks 1.89–1.80 Ga granitoids + minor gabbros Archean Karelian craton

Fig. 6.8. Distribution of Outokumpu-type ultramafic massifs in the North Karelia schist belt. Note that some massifs (Täilahti, Puiroonmäki) are found in close vicinity to the westernmost (subsurface) margin of the Karelian craton. After Säntti et al. (in preparation).

Neoarchean Karelian craton in the east and the 1.93–1.80 Ga Svecofennian island arc complex in the west (Figure 6.1). Within this domain, several tens of ultramafic massifs of variable size are distributed over an area of more than 5000 km2 (Huhma and Huhma, 1970; Koistinen, 1981; Figure 6.8). The ultra256

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mafic massifs range from several kilometers long and several hundred meters thick tabular bodies to just a few tens of meters long and some meters thick lenses (Gaál et al., 1975; Koistinen, 1981). Their estimated total volume exceeds 200 km3 (Kontinen, 1998a). Examples of these massifs are illustrated in Figures 6.9

and 6.10. The chemical composition of the peridotites implies that they are refractory residual mantle peridotites, i.e., harzburgites and dunites (see Section 7). The ultramafic rocks have commonly been intruded by gabbroic and basaltic stocks and dikes which are absent in the enclosing metasediments. Importantly, some large massifs, e.g., Outokumpu, are practically devoid of all kinds of mafic rocks. These gabbro intrusions have yielded a U-Pb zircon age of ~1.97 Ga (Huhma, 1986; Table 6.2) and thus they intruded the peridotites some 70 Ma before their inferred ~1.90 Ga obduction. Koistinen (1981) proposed that these ultramafic massifs might represent fragments of ancient ophiolites and thus oceanic lithosphere. Soon after, this view was strengthened by the identification of the Jormua mafic–ultramafic complex in the northwestern extension of the belt as a well-preserved ophiolite; also, similar crystallization ages were obtained for the Outokumpu and Jormua gabbros (Huhma, 1986; Kontinen, 1987). Although some basaltic and gabbroic dikes intrude the peridotites, the ophiolitic sequence of Outokumpu is far from complete: an extensive sheeted dike complex is absent, a layered cumulate sequence has not been positively identified, and seafloor-type volcanic rocks are uncommon, being present only in the Losomäki area (Park and Bowes, 1981). However, the presence of chromitite bodies with high IPGE/PPGE ratios and mantle-like initial Os isotope compositions strengthens the ophiolite connection (Vuollo et al., 1995; Walker et al., 1996). The incomplete nature of the Outokumpu ophiolite is certainly partly due to tectonic dismembering and selective preservation. However, the nonfractionated composition of the basalts (see below) and extremely low Pb content of the sulfide ores, together with their intimate association with mantle tectonites, suggest that Outokumpu-type massifs more likely represent fragments of ancient peridotitic seafloor (Gaál and Parkkinen, 1993).

5.1. Ultramafic rocks Serpentinization, metasomatic alteration, and regional metamorphism of the peridotite massifs have resulted in complete replacement of the primary silicate minerals. The only remaining primary mineral is chromite, which is well-preserved within the chromitite bodies and may still yield information of the igneous evolution of the complex (Vuollo et al., 1995; Walker et al., 1996). The metamorphic equilibria of the ultramafic massifs has been studied in detail by Säntti (1996) who came to the conclusion that the ultramafic massifs were thoroughly serpentinized into lizardite before the onset of the regional metamorphism. The regional metamorphic isograds transect the Outokumpu nappe and thus individual ultramafic massifs record varying metamorphic grades (Figure 6.11A, B, C). According to Säntti et al. (in preparation) ultramafic massifs record four distinct mineral parageneses depending on the grade of the regional metamorphism (Table 6.3). The main constituent of the antigorite zone massifs is non-pseudomorphic antigorite found as a fine-grained mass of interpenetrating, randomly oriented to subparallel blades and flakes. Increase in the metamorphic grade has resulted in the appearance of olivine and tremolite porphyroblasts, which give the rocks a mottled appearance. Chromite (now largely chromian magnetite) schlierens represent banding inherited from the mantle tectonite protolith. The ultramafic bodies within the higher grade zones have massive, porphyroblastic or crystalloblastic textures without any preferred orientation. This implies crystallization of the metamorphic paragenesis in a late, postkinematic stage of the regional metamorphism. In addition to the serpentinization that thoroughly hydrated the ultramafic massifs, the outer margins of the peridotite massifs became metasomatically altered. Removal of Mg and addition of Ca and CO2 produced successive shells of carbonate rocks and silicified rocks around the massifs CHAPTER

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OUTOKUMPU MINE CROSS-SECTION Y=186.63 10A

23A

24A

703

461 718 112A

726

719

728

724

2 km

Mica schist

Calc-silicate rock

Black schist

Drill hole

Quartz rock

Serpentinite

Ore

Tectonic slide

Fig. 6.9. Vertical cross-section of an ultramafic massif associated with semimassive Cu-Zn-Co-Ni sulfide ore, Outokumpu. Note how the quartz and calc-silicate alteration shells, together with black schists, envelop the serpentinite bodies (modified from Koistinen, 1981).

(Haapala, 1936; Kontinen, 1998a). As Kontinen (1998a) pointed out, individual ultramafic massifs are often completely surrounded by such thin metasomatic alteration shells. This implies that the alteration of the peridotites into carbonate and quartz rocks took place after the obduction-related fragmentation of the ultramafic massifs. Sedimentary origin for the carbonate and quartz rocks can be discarded on the basis of the presence of abundant chromite and mantle-like abundances of the least mobile elements such as Ir, Cr, Ni, and Zr (Kontinen, 1998a). The metamorphism of the serpentinites in the Outokumpu region resulted in breakdown of the primary Cr-bearing phases (chromite, clinopyroxene) and subsequent redistribution of Cr by metamorphic fluids resulted in the formation of rare mineral species such as eskolaite (Cr2O3; Kouvo and Vuorelainen, 1958; Peltonen 258

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et al., 1996a) and extensive substitution of Cr in garnet, diopside, epidote, tremolite, muscovite, and staurolite (Figure 6.11D, E; Eskola, 1933; Treloar, 1987).

5.2. Basaltic rocks Mafic rocks are particularly common in the Losomäki, Miihkali, and Kylylahti serpentinite massifs. Small stocks and dikes of medium- to coarse-grained metagabbro are the most common variant, whereas fine-grained basaltic dikes are uncommon. The volume of mafic intrusions relative to these ultramafic hosts ranges from 5 vol.% to 25 vol.%. Field observations suggest that the gabbros represent intrusions into the mantle tectonites (Asko Kontinen, pers. comm., 2001). Many occurrences comprise clear dikes or small

N

200 m

Outokumpu association

Country rocks

Serpentinite

Mica schist

Carbonate rocks

Black schist

Tremolite/diopside skarn

Calc-silicate rocks

Sulfide ore

Metabasalt Granite

Fig. 6.10. Geological map of the ultramafic massif associated by the Luikonlahti Cu-Zn ore. Quartzrich alteration margins are absent but calc-silicate rocks (tremolite/diopside skarns) are abundant at the margins of the serpentinite massif and frequently are the host rock for the ore. The Luikonlahti body is extensively intruded by granitic dikes related to the younger Maarianvaara granite. Modified from the map of the Malmikaivos Ltd.

pods with apophyses and chilled margins against peridotite. Dike-in-dike intrusion structures are present in several gabbro occurrences suggesting emplacement in an extensional tectonic regime. All intrusions enclosed in the ultramafic massifs are severely tectonized, strongly schistose, and folded,

which attests to their pretectonic origin and emplacement (Figure 6.11F). Narrow (