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
FINLAND
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
FINLAND
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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Undefined
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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
• C H A P T E R 1 • OV E RV I E W
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
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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
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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
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Petrozavodsk
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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
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Kianta
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Suomu Rovaniemi
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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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• ARCHEAN
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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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• ARCHEAN
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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
2
• ARCHEAN
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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
CHAPTER
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• ARCHEAN
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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
• ARCHEAN
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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.
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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).
72
• 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
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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,
• CHAPTER 2 • ARCHEAN ROCKS
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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Chapter 3
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
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Proterozoic granitoids
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Portimo layered igneous complex Tornio
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Koillismaa layered igneNäränkävaara ous complex
Pirivaara
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 (<20 m), as in the Tornio, Kemi (virtually absent), Penikat and Narkaus (Portimo complex) intrusions, the other one much thicker (<150 m), as in the Suhanko and Konttijärvi intrusions (Portimo complex) and in the Western intrusion (Koillismaa complex). The thin, poorly developed marginal series is composed of a fine-grained chilled margin, subophitic, noncumulate-textured gabbroic rocks, and bronzititic cumulates. Felsic material (feldspar/quartz) present in these rocks has partly been extracted from floor rocks during a contamination process. In contrast, the thicker type has better-developed cumulate layers with an olivine cumulate layer (with poikilitic pyroxenes), even tens of meters thick, at the top. This olivine cumulate is separated by a pyroxene cumulate (<10 m) from an underlying gabbroic rock, which is either plagioclase-two pyroxene cumulate or a pyroxene cumulate in which felsic material (related to floor-rock contamination) is 106
present as an unmixed melt component. Details of the layered series differ between the different intrusions (Figure 3.2). Characteristically, the intrusions in the western and central parts of the Tornio–Näränkävaara belt display a repetition of ultramafic and mafic rocks that constitute megacyclic units (MCU) (Figure 3.2), five at maximum in Penikat. These have been interpreted to indicate major replenishments of the magma chamber with new magma batches (Alapieti and Lahtinen, 1986). The lower MCUs (I–III, at the Penikat intrusion and MCU I–II at the Narkaus intrusion) are relatively thin, have a higher proportion of ultramafic cumulates and show limited fractionation indicating frequent but pulsic inflow of magma. The stratigrafically higher cycles crystallized in a regime characterized by fractional crystallization and less frequent influxes of magma.
3.2. Kemi intrusion The Kemi intrusion is economically important due to the presence of huge chromitite reserves. This intrusion, described by Alapieti et al. (1989a), has a lenticular shape and is ~15 km long and 2 km wide in the middle. It lacks the basal reversal typical of the other Tornio–Näränkävaara belt intrusions. The lower part of the layered body comprises up to 500 m of peridotites (olivine±bronzite cumulates), overlain first by bronzitites and then websterites and diallage-augite cumulates. The relative volume of ultramafic rocks is clearly greater than in the other intrusions of the Tornio–Näränkävaara belt, except for the Näränkävaara intrusion. The ultramafic zone of the Kemi intrusion is characterized by the presence of at least 15 chromitite interlayers that, apart from the exceptionally thick main chromitite, are a few centimeters to 2.5 m thick. The main chromitite is discussed in more detail in section 5.2. The upper half of the ~2 km total section is composed of gabbros, gabbronorites, leucogabbros, and anorthosites.
• 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
Koillismaa layered igneous complex
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Portimo layered igneous complex
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Konttijärvi intrusion
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Narkaus intrusion
MCU III RK
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Magma intrusion MCU III pulses MCU II higher in Cr
Lower Zone
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AV
KJ
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MCU II Näränkävaara intrusion
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Cr
Cr
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MS Leucogabbro, anorthosite, diorite
Peridotite Chromitite
Magnetite gabbro Norite, gabbronorite, gabbro
Cr
Chromitite layer
Olivine norite Cr Cr Cr
Websterite Bronzitite, olivine bronzitite
Fig. 3.2. Simplified stratigraphic columns of layered intrusions in the Tornio–Näränkävaara belt and their correlation based on magma types (modified after Lahtinen et al., 1989). Also shown are occurrences of the principal PGE reefs (SJ–Sompujärvi, AP–Ala-Penikka, PV–Paasivaara, SK–Siika-Kämä, RK–Rytikangas), marginal series PGM-sulfide mineral deposits (KJ–Konttijärvi, AV–Ahmavaara), chromitite layers (Cr), and vanadium ore. MCU–megacyclic unit, MS–marginal series.
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Penikat intrusion Sompujärvi block
55°
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Kilkka block
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Keski-Penikat block
PV reef AP reef SJ reef
40°
Fault
Ala-Penikat block
Layering
Fig. 3.3. Simplified geological map of the Penikat intrusion showing the areal distribution of the megacyclic units and the Sompujärvi (SJ), Ala-Penikka (AP), and Paasivaara (PV) PGE reefs (after Alapieti et al., 1990).
3.3. Penikat intrusion The first layered intrusion to the east of Kemi is the PGE-potential Penikat intrusion, which has a present-day surface area of 1.5–3.5 km by 23 km (Figure 3.3). The ultramafic rocks are most often bronzite±chromite cumulates, while olivine or bronzite-augite cumulates 108
are less common. Minor chromitite interbeds from a few centimeters to 0.5 meters thick occur within the ultramafic rocks. The mafic rocks are mostly plagioclase-bronzite and plagioclase-bronzite-augite cumulates, more rarely plagioclase-augite cumulates. A common feature with the Western intrusion of the Koillismaa complex is the several hundred
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bCa pCa pCba p±bCa pegm pCa pa(b)Ca pbCa pCa(b) pC paCa bpCa
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6.1 1.8 0
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Leucogabbro, anorthosite Transition zone Plagioclase-augite-bronzite cumulate Plagioclase-bronzite cumulate Ultramafic cumulates
Cr wt.%
Fig. 3.4. Simplified stratigraphy and variation of whole-rock Cr across the Penikat intrusion. The positions of the three principal PGE reefs (SJ–Sompujärvi, AP–Ala-Penikka, PV–Paasivaara) and some other PGE showings (p) are marked in the stratigraphic column. Data for interval 0–2000 m taken from Halkoaho (1993) and data for 2000 m upwards taken from this study. Also shown is the variation of precious metal contents through the transition zone (modified after Huhtelin et al., 1990). Abbreviations for rock types in the transition zone: p–plagioclase, b–bronzite, a–augite, pegm–mafic/ultramafic pegmatite; letters before C denote cumulus phases, those after C intercumulus phases. CHAPTER
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N
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Fault Siika-Kämä reef Offset mineralization
MCU III
MCU II
MCU I
Layered series Plagioclase-augite-bronzite cumulate Bronzite, bronzite-augite, and olivine cumulate Plagioclase-augite-bronzite cumulate Bronzite, bronzite-augite, and olivine cumulate
Paleoproterozoic spracrustal rocks Mafic metavolcanic rocks and sills Quartzite Conglomerate, mica schist
Plagioclase-augite-bronzite cumulate Bronzite cumulate
Archean granitoids
Fig. 3.5. Geological map of the Kilvenjärvi block of the Narkaus intrusion, showing megacyclic units (MCU) and their mafic and ultramafic layers. The Siika-Kämä PGE reef and offset PGE deposit are also shown (modified after Huhtelin et al., 1989b).
meters thick sequence of leucogabbros and anorthositic rocks in the upper part of the stratigraphy. In addition, the topmost of part of the sequence (>2520 m) is rich in quartz and biotite (quartz diorite in Figure 3.4), resembling the granophyres overlying the Western intrusion. The Penikat intrusion, together with the 110
Narkaus intrusion (section 3.4.), can be divided into five stratigraphic units which are assigned to megacyclic units I to V and abbreviated as MCU I to V (Figures 3.2–3.4). The fourth and fifth are the thickest, occupying at least 3/4 of the ~3-km-thick igneous section remaining after Paleoproterozoic erosion. The three lowermost megacyclic units
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(MCU I–III) differ from MCU IV and V in that the extent of differentiation is considerably less well-developed. This implies that MCU I–III had less time to undergo fractional crystallization before recharge of the magma chamber with a pulse of primitive magma and initiation of a new MCU. This also implies that MCU IV and V contained additional residual liquid derived from the lower MCU I–III, allowing more advanced differentiation in the upper two cycles. As Figure 3.4 shows, the fifth megacyclic unit has produced the most fractionated rocks in terms of Cr contents. An abrupt drop in Cr content occurs immediately above the highest low-grade PGE reef, where it falls from a level of 100–200 ppm to 30–70 ppm, and again at the stratigraphic height of 2520 m down to ~10 ppm, with concomitantly rapid increase of quartz. The Penikat intrusion is mineralogically sufficiently well-preserved to allow the cryptic variation of the primary ferro-magnesian silicates (pyroxenes) to be studied throughout the layered sequence. Alapieti and Halkoaho (1995) reported microprobe data on augite from the intrusion base to the lower part of MCU V. Despite the great thickness of the studied section (~2 km), the overall range of augite compositions is rather limited with Mg# between 0.90 and 0.76. There is a distinct compositional reversal between MCU I and MCU II, indicative of a new pulse of magma, while Mg# remains nearly constant (~0.77) from the base of MCU III upwards.
3.4. Portimo layered igneous complex The Portimo layered igneous complex (in short Portimo complex; Figure 3.1) is composed of four principal structural units: • the Narkaus intrusion (Figures 3.1 and 3.5) • the Suhanko intrusion (Figure 3.6) • the Konttijärvi intrusion (Figure 3.6) • the Portimo dikes (Figure 3.6) CHAPTER
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Each intrusion contains a marginal series and an overlaying layered series. The marginal series of the Suhanko and Konttijärvi intrusions differ from that of the Narkaus intrusion in thickness and prevailing rock types. The Narkaus marginal series generally varies from 10 to 20 m in thickness, while the Suhanko and Konttijärvi marginal series may reach several tens of meters. The Narkaus marginal series is mainly composed of pyroxenite with some plagioclase-bearing rocks in its lower parts, whereas olivine cumulates commonly constitute the upper half of the Suhanko and Konttijärvi marginal series. A striking difference between the layered series of the intrusions is the presence of marked reversals in the Narkaus intrusion, as shown by the thick ultramafic olivine-rich cumulate layers, whereas crystallization in the Suhanko and Konttijärvi intrusions continued without notable reversals (Figure 3.2). The Suhanko layered series commences with plagioclase-bronzite orthocumulates (with poikilitic augite) that also contain some bronzite cumulate interlayers. This poikilitic rock is separated from the overlying, rather monotonous plagioclase-bronzite-augite adcumulates by a few meters thick pyroxenite. About midway in the stratigraphy, bronzite disappears as a cumulus mineral, but returns higher up in the Suhanko sequence. Four poikilitic anorthosite layers also occur in the upper Suhanko layered series (Figure 3.8A). Granophyric material is limited to discontinuous patches and crosscutting dikes in the upper Suhanko and Konttijärvi layered series. The major reversals in the Narkaus layered series resemble those of the Penikat intrusion and enable its layered series to be divided into three megacyclic units (Figures 3.2 and 3.5). The lowermost (MCU I) commences with a thick (~80 m) bronzite cumulate layer with a massive chromitite layer close to its top. The rest of MCU I as well as the gabbroic parts of MCU II and MCU III are mainly composed of plagioclase-bronzite-augite adcumulates
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Layered Series Granophyre Gabbronorite, gabbro Pyroxenite Poikilitic gabbro Marginal Series Gabbronorite, pyroxenite, peridotite, subophitic gabbro Portimo dike A Sections in Figure 3.7
B Konttijärvi
Fine-grained autolith Ultramafic pipe Rytikangas reef Fault Layering
1 km
Rytikangas
Ahmavaara
Suhanko
Fig. 3.6. Geological map of the Suhanko and Konttijärvi intrusion, depicting the general setting of the Portimo dikes and marginal and layered series. Also shown is the location of the Rytikangas PGE reef (modified after Iljina, 1994). Sites of the cross and longitudinal sections depicted in Figure 3.7 are marked on the Konttijärvi intrusion.
with the exception of a poikilitic plagioclase cumulate layer above the ultramafic basal part of MCU III. MCU unit II, however, is found only in the Kilvenjärvi block (Figure 3.5) and fades away eastwards. Mafic and ultramafic dikes, known as the Portimo dikes, are found in the basement below the Konttijärvi intrusion and in the Ahmavaara area of the Suhanko intrusion (Figure 3.6). They have also been found as fragments in the marginal series of the Konttijärvi intrusion (Figure 3.8B). The dikes have not been dated and their association to the main intrusions is based on geochemical observations as discussed later. The dikes are subparallel to the basal contact of the intrusion and merge 112
with it locally, so that a dike may actually form the basement of the intrusion. In view of their location and original subhorizontal orientation, these dikes were previously referred to as ‘sublayer sills’ (Iljina et al., 1989). The term ‘pre-intrusion dikes’ has also been used by Iljina et al. (1992) to allude to the nominal age difference between them and the intrusions. The cumulus sequences in the layered series of the small Konttijärvi intrusion and the western end of the Suhanko intrusion resemble each other. Pyroxenite, which separates the lowermost poikilitic orthocumulate from the overlying gabbroic adcumulate, attains a thickness of some tens of meters in both sections. The gabbroic rocks of the marginal
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100 m
Site of longitudinal section shown in Figure B
B
100 m Site of cross-section shown in Figure A
Layered series
Marginal series
Gabbroic cumulates
Olivine/pyroxene cumulates
Portimo dikes
Pyroxene cumulates
Gabbroic rocks
Archean granitoids
Heavily contaminated marginal series
Fault
Fig. 3.7. Cross-section (A) and longitudinal section (B) of the Konttijärvi intrusion (cf. Figure 3.6).
series and the heavily contaminated lower part of the marginal series of the Konttijärvi intrusion, indicated in Figure 3.7, are mostly pyroxene cumulates with variable portions of felsic material introduced by floor-rock contamination. This and the thick layered series pyroxenite make the present-day Konttijärvi stratigraphy largely ultramafic. The lower contact of the Konttijärvi intrusion (Figure 3.7) is also unique among the Tornio–Näränkävaara belt intrusions. It has ridges and troughs and its location with respect to the undisturbed baseCHAPTER
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ment granitoid is almost impossible to define. At the center of the igneous body the contact seems to dive to form a pothole, some tens of meters in diameter (Figure 3.7). The above described features of the Konttijärvi intrusion have led to the recognition of two separate intrusions instead of a single Suhanko–Konttijärvi intrusion (e.g., Alapieti et al., 1989b and Iljina, 1994). Fine-grained, non-cumulate-textured autoliths up to a few tens of meters thick and several hundred meters long occur in the
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B
A
Fig. 3.8. (A) Contact between poikilitic anorthosite and gabbroic adcumulate in the upper Suhanko layered series. (B) Partly corroded websterite xenoliths of the Portimo dike in olivine cumulate of the Konttijärvi marginal series. The rusty color of the cumulate is due to secondary carbonate and sulfides. The xenoliths are surrounded by a black magnetite rim. Photos: Markku Iljina.
Suhanko marginal series in many places. The chemical composition of these plagioclasetwo pyroxene rocks (see below), is similar to the mean composition of the Suhanko intrusion. They have thus been interpreted as early chilled margin rocks that were disrupted and entrained by subsequent magma pulses (Iljina, 1994). The autoliths in Ahmavaara area show distinctly high chromium contents (Iljina, 2005), in the order of ~1000 ppm, and they are chemically related to the Portimo dikes. An iron-rich ultramafic pegmatoidal mass two hundred meters in diameter is found in the western limb of the Suhanko intrusion, (the Ahmavaara block, Figure 3.6) where it is located above the ‘anticline’ of the base of the intrusion. Magnetic measurements and drill hole evidence suggest that the pegmatite forms a pipe-like body, which dips perpendicularly to the magmatic layering of the hosting cumulates. Comparable pegmatites have also been reported by Mutanen (1997) in the Koitelainen intrusion, where ultramafic pegmatoids (clinopyroxenites rich in magnetite and ilmenite) occur as pipes and veins.
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3.5. Koillismaa layered igneous complex The Koillismaa layered igneous complex (in short, Koillismaa complex), which was described in detail by Alapieti (1982), is an assemblage of three main structural units: the Western intrusion, the Näränkävaara intrusion and an unexposed connecting “dike” between the two intrusions (Figures. 3.1 and 3.9). The Western intrusion originally crystallized as a large, 3-km-thick magma chamber, but was later disrupted by tectonic movements and is now represented by several separate, tilted blocks. From south to north, these blocks are Pirivaara, Syöte, Porttivaara, Pyhitys, Kuusijärvi, Tilsa, Lipeävaara, Kaukua, and Murtolampi. Their present-day surface area varies from a few km2 up to ~100 km2 and the total volume of the magma that produced the Koillismaa complex has been estimated to be at least 2000 km3 (Alapieti, 1982). In addition to gabbroic bodies, temporally and spatially related, relatively undeformed Atype quartz syenite and granite plutons have been reported (Luukkonen, 1988; Lauri and Mänttäri, 2002). The individual fragments of the Western intrusion possess slightly differing lithological
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Murtolampi
Kaukua Lipeävaara
Kynsijärvi
Tilsa Kuusijärvi
Pyhitys Porttivaara
Syöte
N Pirivaara
10 km
Paleoproterozoic Granite (1.8 Ga) Syenite (2.44 Ga) Diabase Gabbro, serpentinite Mafic metavolcanic rock Felsic metavolcanic rock Metasediment Albite-quartz rock Granophyre Koillismaa layered igneous complex Magnetite gabbro Leucogabbro, anorthosite Olivine gabbronorite Gabbronorite Marginal series Archean Basement complex
Fig. 3.9. Geological map of the fragmented Western intrusion, Koillismaa layered igneous complex. The map shows the general lithology and distribution of the intrusion blocks (Pirivaara, Syöte, Porttivaara, Pyhitys, Kuusijärvi, Tilsa, Lipeävaara, Kaukua, and Murtolampi). Note the coeval Kynsijärvi syenite on the NE side of the Kuusijärvi block.
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successions. This is probably because they represent different parts of the originally sill-like magma chamber and because the chamber had a variable thickness preventing the development of a completely uniform stratigraphy along the strike. Above a 50–100-m-thick, reversely zoned marginal series, topped by harzburgites and bronzitites, is a ~3000-mthick layered series that can be divided into three mineralogically distinct units: the lower, middle, and upper zones (Figure 3.2). The layered series contains cumulus plagioclase throughout. Olivine is liquidus phase in the lower zone but is replaced by augite and inverted pigeonite in the middle zone. The leucogabbroic and anorthositic upper zone contains a plagioclase-magnetite-augite cumulate unit in its middle, hosting a vanadium deposit. A conformable granophyre sheet typically one kilometer in thickness covers the intrusion. At the base of the intrusion, the footwall granite gneisses have been partially melted and pervasively metasomatically altered to an albite-quartz rock (Figure 3.9). Peculiar fine-grained, equigranular, noncumulate-textured gabbroic rocks are found within the layered series of the Pirivaara, Syöte, Porttivaara, and Kuusijärvi blocks. Based on their mode of occurrence and texture, these bodies have been termed microgabbronorite xenoliths (Iljina et al., 2001). They resemble the fine-grained rocks described from the Lukkulaisvaara intrusion (the Olanga complex, eastern Karelia; Glebovitsky et al., 2001). The Koillismaa microgabbronorite xenoliths occur in a ~20-m-thick zone, which parallels the igneous layering and is also characterized by felsic (or anorthositic) patches and gabbro pegmatite bodies. The size of the mafic xenoliths ranges from a decimeter to a few meters and they are composed of orthopyroxene, clinopyroxene, and plagioclase. Besides being higher in Cr and lower in TiO2, the Koillismaa microgabbronorite xenoliths differ from the surrounding cumulates in their REE patterns (see below), precluding any simple, 116
direct genetic link with the magma of the host intrusion. The genesis of the capping granophyric rocks has been problematic for a long time. Piirainen et al. (1978) regarded this biotitealbite rock as a quartz-keratophyric lava. Subsequently, Alapieti (1982) considered the granophyre cogenetic with the intrusions, a hybrid rock formed by mixing of magma with partial melts of roof rocks. Iljina et al. (2001), Karinen and Salmirinne (2001), and Lauri et al. (in press) proposed that the granophyre represents a unit that includes a large proportion of felsic metavolcanic material. These volcanic rocks are thought to have erupted before the emplacement of the Western intrusion and thus subsequently subjected to thermal effects of the mafic magma injected between the volcanic rocks and the granitoid basement. The Näränkävaara intrusion, with a maximum width of ~5 km and a length of more than 20 km on the Finnish side, differs from the other intrusions of the Tornio–Näränkävaara belt in being largely ultramafic (~80% of the surface section). Somewhat surprisingly, chromite only occurs as an accessory cumulus phase and no notable chromite concentrations have been found. This is despite the preponderance of ultramafic rocks at Näränkävaara and the precence of chromitites with ultramafic cumulates in the Kemi, Penikat, and Narkaus intrusions in the Tornio–Näränkävaara belt, the Akanvaara and Koitelainen intrusions in Lapland (Mutanen, 1997), and several Russian intrusions (e.g., Sharkov et al., 1995; Torokhov et al., 1997). The total stratigraphic thickness of the sequence is uncertain: the body has a funnelshaped cross-section and extends to a depth that may exceed 10 km (Elo, 1992). Nevertheless, the general structure of the exposed portion, formed by the lower ultramafic and upper mafic zone, is well established (Alapieti et al., 1979; Alapieti, 1982). The former is composed of olivine cumulates in its lower part followed by olivine-bronzite, bronzite, bronzite-augite, and bronzite cumulates and again by bronzite-au-
• 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
Western intrusion blocks Porttivaara Kuusijärvi
Syöte
Depth (m)
Gravity Bouguer (mgal)
SW
NE Pyhitys
20
Lipeävaara
Measured Calculated
10 0 -10
Line: Näränkä 2
-20
1000
RMS=1.96 Gabbro d=2900 kg/m3 Pyroxenite d=3300 kg/m3 Peridotite d=2950 kg/m3
Serpentinite d=2590 kg/m3
2000 3000
Basement d=2670 kg/m3
4000
A
5
10
15
Distance (km)
“Feeder Dike”
B
A
N B
Näränkävaara intrusion
Fig. 3.10. Gravity-based block model of the Näränkävaara intrusion, Western intrusion blocks, and the connecting “Feeder Dike” with a view towards W with an inclination of 25°. Also shown is a gravity profile across the Näränkävaara intrusion and its interpretation. Modified after Salmirinne and Iljina (2003).
gite cumulates. The bronzite-augite±olivine cumulates with poikilitic plagioclase are overlain by gabbros (plagioclase-hypersthene-augite cumulates) and dioritic rocks (plagioclaseaugite cumulates). The connecting “dike” is thought to form a hidden, long and narrow differentiated intrusion (Figure 3.10) mimicking the Great Dyke CHAPTER
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or the Jimberlana intrusion. Magnetic and gravity data suggest that the upper surface of the intrusive body lies at a depth of 0.3–1.5 km with its maximum width and vertical thickness being approximately 12 km and 5.5 km, respectively (Elo, 1992; Salmirinne and Iljina, 2003).
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4. Parental magmas and isotope studies 4.1. Parental magmas The geochemical character of the parental magma of the layered intrusions has been assessed using lower chilled margin whole-rock analyses, weighted average compositions of cumulate sequences, and chemical compositions of dike rocks that are thought to be genetically related to the intrusions (Alapieti et al., 1990; Iljina, 1994; Saini-Eidukat et al., 1997; Vogel et al., 1998). Based on the mineralogy and chemical composition of the cumulates, it has previously been concluded that the igneous bodies in the western and central part of the Tornio–Näränkävaara belt were fed with two kinds of parental magmas, a higher-Cr and a lower-Cr magma with estimated Cr contents of >1000 ppm and <600 ppm, respectively. As shown in Figure 3.2, some of the intrusions were generated either from the higher-Cr (Kemi and Tornio) or the lower-Cr magma (Suhanko and Konttijärvi), but others (Penikat and Narkaus) experienced influx of both magma types, first the higher-Cr and then the lower-Cr. It has been suggested that, compared with the lower-Cr parental magma, the higher-Cr magma had slightly higher SiO2 and was more enriched in LREE relative to HREE, thus having a stronger boninitic or siliceous high-magnesian basaltic (SHMB) nature. The TiO2 values of the weighted averages of the higher-Cr megacyclic units differ from those of the higher-Cr dikes, the latter being somewhat higher in TiO2. Iljina (1994) and SainiEidukat et al. (1997) showed that this magma was close in composition to the B1 magma of the Bushveld Complex (Figure 3.12). On the other hand, the lower-Cr parental magma was reported to correspond to the Bushveld B2/B3 magma composition in having a lower and flatter chondrite-normalized REE pattern (Iljina, 1994). In this respect, it also 118
resembles the Koillismaa and Lukkulaisvaara microgabbronorites. The Cr content (210–460 ppm) of the Koillismaa chilled margin appears representative of the lower-Cr magma type. However, the higher TiO2 content (0.5–0.8 wt.% TiO2) and the crystallization path of the Western intrusion that produced a thick magnetite gabbro suggest a third type of magma: a more evolved tholeiitic basalt. A compilation of chemical compositions of representative rock samples and weighted average compositions of the megacyclic units is shown in Table 3.1. Reported chilled margin analyses have an MgO content between 6.5 and 11.4 wt.%, while MgO in the dike rocks is usually higher, exceeding 20 wt.% in rocks enriched in cumulus phases. The most magnesian primary silicates in the Kemi (Fo83) and Penikat (En82) intrusions indicate that the magma that generated these layered series was not very primitive, with MgO probably below 10 wt.%. This was, however, considerably more MgO-rich than the melt that produced sections of the lower-Cr intrusions (Porttivaara Fo77; Syöte Fo73). In common with the Koitelainen and Akanvaara intrusions (cf. Hanski et al., 2001), the magma that produced the Kemi and Penikat intrusion was not a primary mantle melt but had experienced sialic contamination and accompanying crystal fractionation in a lower level, staging magma chamber in the crust, a process also shared by some other mafic layered intrusions (e.g., the Bushveld Complex; Arndt et al., 1997). This kind of AFC model was also proposed by Saini-Eidukat et al. (1997). Thus all of the chromitite-bearing 2440 Ma layered intrusions in Finland seem to have crystallized from a relatively fractionated basaltic magma in terms of Mg# and MgO. Figure 3.11 shows chondrite-normalized REE patterns for various rock types from the Tornio–Näränkävaara belt and, for comparison, from the Koitelainen and Lukkulaisvaara intrusions. The microgabbronorite xenoliths from Koillismaa, low-Cr autoliths from Su-
• 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
Table 3.1. Major and trace element analyses of mafic dikes and microgabbronorites and weighted averages of certain megacyclic units from the Tornio–Näränkävaara belt. Also shown are the compositions of a chilled margin from Koitelainen and microgabbronorite from Lukkulaisvaara. wt.% SiO2 TiO2 Al2O3 FeOtot MnO MgO CaO Na2O K2O P2O5 ppm V Cr Ni Zn Sc Rb Sr Y Zr Nb Ba
1 51.74 0.61 13.17 9.27 0.15 8.40 10.10 2.80 0.72 0.06
2 52.78 0.15 12.42 8.82 0.17 15.74 8.18 1.03 0.36 <0.00
199 466 104 65 31 20 337 13.6 51 2.0 188
102 1026 343 76 27 10 158 2.7 6
La Ce Nd Sm Eu Tb Yb Lu
8.63 17.2 9.63 2.21 0.73 0.38 1.23 0.20
2.5 2.26 1.12 0.28 0.15 <0.1 0.34 <0.1
Th U
1.3 0.24
3 52.06 0.08 17.34 4.01 0.12 11.29 13.02 1.54 0.06 0.05
4 52.00 0.68 12.30 10.20 0.18 13.00 9.26 1.74 0.71 0.10
5 51.40 0.40 8.00 12.00 0.31 13.80 11.00 0.62 0.45 0.07
6 51.53 0.62 11.88 10.29 0.19 13.63 8.32 1.40 0.92 0.07
460 166
1100 360 130 28 28 200 <2 70 <2 210
230 2300 250 160 37 17 20 7 40 2 47
200 1326 408 83 32 40 133 13 74 3.0 225
11.1 24 12 2.6 0.78 0.5 1.61 0.26
10.9 25 12 2.5 0.61 0.3 1.08 0.16
10.2 19.5 9.6 2.2 0.80 0.40 1.30 0.20
2.5 7 4 0.75 0.32 0.2 0.59 0.09
2.3 0.4
2 1.2
2.1 0.50
0.3 <0.1
97 0.5 2 1 0.22 0.16 0.1 0.23 0.03
7 50.70 0.22 14.20 7.06 0.14 14.70 11.60 1.22 0.22 0.02
8 52.40 0.24 13.20 8.17 0.16 14.10 9.01 1.81 0.53 0.03
3330 270
3290 230
150
190
9 51.30 0.12 17.10 5.89 0.17 10.60 12.20 2.26 0.18 0.01 150 290 180 63 33 <10
10 52.70 0.33 16.90 7.27 0.16 9.03 10.70 2.38 0.57 0.04
11 50.10 0.23 18.80 6.95 0.13 10.10 11.50 1.97 0.31 0.03
12 51.70 0.31 17.00 6.54 0.13 9.91 11.30 2.55 0.56 0.03
280 280
260 340
217 24
210
260
220
47
31
27
180
93
14
2.5 25 75
27 120
2.7 110
1. 2. 3. 4. 5. 6. 7, 8.
Koitelainen chilled margin. Additional elements (ppm): Pr 2.27, Gd 2.37, Dy 2.28, Ho 0.49, Er 1.34, Tm 0.21. Koillismaa microgabbronorite xenolith, an example from the Syöte block. Lukkulaisvaara microgabbronorite (Glebovitsky et al., 2001). Loljunmaa dike (major elements from Alapieti et al., 1990 and trace elements from Iljina, 1994). Portimo dike (Iljina, 1994). Viianki dike (Vogel et al., 1998). Additional elements (ppm): Pr 2.6, Gd 2.9, Dy 2.4, Ho 0.50, Er 1.40, Tm 0.20. Weighted averages of the lowermost megacyclic units (MCU I) of the Narkaus and Penikat intrusions, respectively (Alapieti et al., 1990). 9. Suhanko low-Cr autolith (Iljina, 1994). 10–12. Weighted averages of the Suhanko body, megacyclic unit III of Narkaus intrusion and megacyclic unit IV of Penikat intrusion (Alapieti et al., 1990).
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hanko, and microgabbronorite bodies from Lukkulaisvaara (all noncumulate, basaltic rocks) display surprisingly low concentrations of REE with only a slight enrichment of LREE over HREE. The cumulate rocks from Koillismaa and Koitelainen have very similar REE patterns with sloping LREE and relatively flat HREE, suggesting derivation from similar, LREE-enriched magma types that differed significantly from those parental to the above mentioned noncumulate rocks. The chilled margin samples from Koillismaa also have relatively high LREE/HREE but in detail they differ from each other in having different REE patterns. Figure 3.12 compares the REE characteristics of dike rocks from three locations – Loljunmaa, Portimo, and Viianki – with those of the Bushveld B1 and B3 magmas. Of these dikes, Loljunmaa is located on the southeastern side of the Penikat intrusion (see Figure 3.1) and Viianki at the FinnishRussian border 100 km to the south of the Näränkävaara intrusion. The REE patterns of the dikes mimic closely that of B1, except that the level is higher in the latter – this can be at last partly explained by the more primitive nature of the dikes. A similar type of relationship is observed between the Bushveld B3 and Portimo lower-Cr magma type compositions (Figures 3.11C and 3.12). Both have distinctly lower LREE/HREE compared to the dikes and B1, and they are lower in total REE.
4.2. Isotope studies Perttunen and Vaasjoki (2001) reported a UPb zircon age of 2430 ± 4 Ma for the Kemi intrusion and 2430 ± 36 Ma for the Penikat intrusion. Previously, Alapieti (1982) had published an age of 2436 ± 5 Ma for a combination of samples from several bodies of the Koillismaa complex. These results agree well with the age data obtained for the Akanvaara and Koitelainen intrusions (Mutanen and Huhma, 2001) and suggest a short period of 120
time over which the ~2440 Ma intrusions were emplaced. Using whole-rock samples and mineral separates, Huhma et al. (1990) obtained a Sm-Nd isochron of 2410 ± 64 Ma and an initial εNd of –1.6 for the Penikat intrusion; these resemble the preliminary data from the Koillismaa complex (2452 ± 160 Ma and εNd –0.6; O. Thalhammer, pers. comm., 1994) and they are similar to those from other layered intrusions of this age group, indicating a large contribution of REE from Archean enriched lithosphere. Hf isotope analyses on zircons from the Näränkävaara intrusion yielded a negative εHf value (–1.6), which is consistent with the general Nd isotope systematics of the ~2440 Ma magmatism, whereas the Kemi intrusion produced an unexpectedly high εHf (at 2440 Ma) of +3.7 (Patchett et al., 1981). Pb isotope data are available on silicate whole rocks from the Kemi intrusion (Manhes et al., 1980) and on galenas from a sulfide mineralization in the Suhanko intrusion (Alapieti et al., 1989b). The high 207Pb/204Pb ratio in galenas suggests that Pb originated mainly from the Archean upper crust.
5. Mineral deposits 5.1. Ore types Several types of ore mineralization have been discovered in the Tornio–Näränkävaara belt and two economic oxide deposits have so far been mined: the Cr deposit in the Kemi intrusion and the Ti-V deposit in the Porttivaara block of the Western intrusion of the Koillismaa complex (Figure 3.9). PGE and chalcophile element mineralization occurs in the Penikat, Suhanko, Konttijärvi, and Narkaus intrusions and the Koillismaa complex. These can be classified into three main types: (1) disseminated and massive PGE-enriched Cu-NiFe sulfides of the marginal series, (2) reef-type PGE deposits, and (3) offset base metal and
• 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
100
Rock/Chondrite
100
Microgabbronorite bodies Lukkulaisvaara
Microgabbronorite xenoliths Koillismaa 10
10
1
A La Ce Pr Nd
1
La Ce Pr Nd
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
100
B Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
100
Chilled margin Koillismaa
Rock/Chondrite
Suhanko autoliths
10
1
10
C La Ce Pr Nd
1 Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
100
D La Ce Pr Nd
100
Koitelainen cumulates
Rock/Chondrite
Koillismaa cumulates
10
1
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
10
E La Ce Pr Nd
1 Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
F La Ce Pr Nd
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Fig. 3.11. Chondrite-normalized REE patterns. (A) Koillismaa microgabbronorite xenoliths (Iljina et al., 2001). (B) Lukkulaisvaara microgabbronorite bodies (Olanga complex, Glebovitsky et al., 2001). (C) Suhanko low-Cr autoliths (Iljina, 1994). (D) Koillismaa chilled margin (Alapieti, 1982; Iljina et al., 2001). (E) Koillismaa cumulates surrounding the microgabbronorite xenoliths (Iljina et al., 2001). (F) Gabbro and pyroxenite cumulates from the lower part of the Koitelainen intrusion (Hanski et al., 2001).
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5.2. Kemi chromite deposit
100
Bushveld B1 Rock/Chondrite
Bushveld B3 10
1
Dikes: Loljunmaa Viianki Portimo La Ce Pr Nd
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Fig. 3.12. Chondrite-normalized REE patterns of Loljunmaa, Portimo, and Viianki dikes (Vogel et al., 1998; Iljina, 1994) and Bushveld B1 and B3 magmas (Hatton and Sharpe, 1989; Maier et al., 2000).
PGE deposits in the footwall rocks. The first type is almost exclusively confined to welldeveloped, thicker marginal series as found in the Suhanko and Konttijärvi intrusions of the Portimo complex and the Western intrusion of the Koillismaa complex. In Suhanko, both disseminated and massive concentrations of sulfides have been found, while the two other intrusions have only disseminated sulfides. The reef-type PGE enrichments found in the Penikat intrusion and Portimo complex can be divided into the major or principal enrichments, which are laterally continuous and have PGE concentrations of at least several ppm, and PGE showings that are less continuous and grade rarely above one ppm. The principal PGE reefs include the Sompujärvi (SJ), AlaPenikka (AP), and Paasivaara (PV) reefs in the Penikat intrusion, and Siika-Kämä (SK) and Rytikangas (RK) reefs in the Portimo complex (Figure 3.2). SJ, PV, and SK are considered highly viable for economic exploitation. These and other reef-type PGE deposits have low, barely visible concentrations of Cu, Ni, and Fe sulfides, wheras chromite is often present, as in the Sompujärvi and Siika-Kämä reefs. 122
The Kemi main chromitite layer located in the lower, ultramafic part of the intrusion, can be followed along strike over the entire length of the intrusion for more than 15 km (Alapieti et al., 1989a). Laterally, it varies in thickness from a few mm up to ~90 m at the thickest part of the intrusion, where the mine is located. The whole length of the mineable portion of the layer is 4.5 km. The estimated ore reserves in 1998 were 70 million tons and mineral resources 143 million tons, giving a total of 213 million tons (Leinonen, 1998). The main chromitite is separated stratigraphically from the lower contact of the intrusion by 50 m to 100 m of altered bronzitite, but the lowermost chromite-enriched layer can be found as low as 0.5 m above the chilled margin. Furthermore, a 10-cm-thick, fine-grained, ultramafic marginal rock sample was reported to contain 15 vol.% of chromite phenocrysts. This fact, together with a recently discovered chromite-bearing feeder dike beneath the thickest part of the intrusion (Alapieti, 1996), indicates that the magma carried chromite in suspension and ore-forming processes had begun already at depth before the emplacement of the magma into the final magma chamber. While chromites in the relatively thin chromitites of the Akanvaara and Koitelainen intrusions are peculiar in that they have very low MgO contents (Mutanen, 1997), chromites in the main ore body at Kemi exhibit high MgO contents frequently exceeding 10 wt.% (Alapieti et al., 1989a; Gornostayev et al., 2000). Compared with chromitite deposits in other layered intrusions (e.g., Bushveld) or podiform deposits in ophiolites, the average Cr content and Cr/Fe ratio of the Kemi ore are rather low, 34 wt.% and 1.50 wt.%, respectively.
• 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
5.3. Mustavaara Fe-Ti-V oxide deposit, Koillismaa complex Fractional crystallization in the Western intrusion of the Koillismaa complex has led to a significant accumulation of titanomagnetite in the upper layered series (Figure 3.2). The presence of magnetite enrichment has been verified in the southern intrusion blocks (Pirivaara, Syöte, Porttivaara, and Kuusijärvi), while only limited outcrop and geophysical observations from the Lipeävaara block suggest its existence in the northern Western intrusion as well. The Mustavaara vanadium mine, that produced 13.4 Mt of ore in 1975–1985, was located in the upper zone of the Porttivaara block (Juopperi, 1977). The host rock is a magnetite gabbro (typically plagioclase-pyroxene-titanomagnetite cumulate) that forms a coherent stratum more than 20 km in length and ~200 m in thickness between two anorthositic gabbro layers. The magnetite gabbro is divided into four layers, of which three have been mined at Mustavaara. In these three layers, the mean weight percentage of oxides ranges between 14% and 33%. The whole-rock V2O5 concentration varies between 0.40 wt.% and 0.70 wt.%, while the mean V2O5 content of oxide concentrates from each layer is rather constant, 1.55 wt.% to 1.68 wt.%. Compared with the average whole-rock V2O5 values of 1.5 wt.% to 2.0 wt.% in the exploited vanadiferous magnetite gabbros of the Bushveld Complex (Cawthorn and Molyneux, 1986), the Mustavaara ore was significantly lower in grade.
5.4. PGE reefs of the Penikat intrusion The lower part of the Penikat intrusion is characterized by repetition of four thick ultramafic cumulate layers all defining the base of the megacyclic units (Figures 3.2–3.4). Of these units, MCU IV hosts high-PGE-grade reefs. Well-documented examples, from lowest to highest, are the Sompujärvi, Ala-Penikka, and CHAPTER
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Paasivaara PGE reefs. The Sompujärvi reef is found at the base of MCU IV, Ala-Penikka reef well within this unit, and the Paasivaara reef in the uppermost part of the unity. The part is called the transition zone (Halkoaho, 1993; Huhtelin et al., 1990). In addition to these major reefs, there are also a number of PGE showings especially with the ultramafic and chromitite layers of MCU I–III (Figure 3.4). A PGE showing above the Paasivaara reef is hosted by a melagabbro–anorthosite sequence (12 m thick) within a thick succession of monotonous leucogabbros. The Penikat PGE reefs are correlated with the PGE enrichments of the Portimo complex in Figure 3.2. Reefs are persistent features at certain stratigraphic levels (Figure 3.3); the Sompujärvi and Ala-Penikka reefs have been traced virtually over the entire 23 km intrusion length. These reefs vary in thickness from a couple of decimeters to several meters and may locally reach 20 m. The average thicknesses of the reefs are 1.0 m and ~0.3 m, respectively. The Sompujärvi reef is most commonly hosted by a bronzite±chromite cumulate or metasomatic chlorite schist. More rarely, high PGE concentrations are found in the underlying gabbroic cumulates or overlying olivine cumulates. The Ala-Penikka reef is hosted by a leucocratic plagioclase cumulate with poikilitic augite, which is overlain by a plagioclase-bronzite cumulate some tens of meters thick, with nonpoikilitic intercumulus augite. This package is underlain and overlain by thick sequences of fairly monotonous gabbronorite adcumulates (see also the discussion of the Rytikangas reef of the Portimo complex). The stratigraphically highest Paasivaara reef is more erratic in nature, and base metal sulfides and PGM are hosted by a peculiar transition zone (Figure 3.4) that forms the topmost part (up to 40 m) of MCU IV. The transition zone is composed of plagioclase±augite±bronzite ad- and orthocumulates and a mixed rock, which itself is composed of irregular patches of poikilitic gabbro embedded in a plagioclase
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Layer ed int rusion Floor rocks
R358
R359 Overburden Diabase
Ul cumtrama ula fic tes
La Ma yere rgi d s na eri l s es er ies
Gabbroic cumulates
Albite-quartz rock pm
4p
0
u
+A
10 20 30 40 50 m
d +P
Pt
Fig. 3.13. Drill hole profile R358–R359 through the Kuusijärvi marginal series of the Koillismaa layered igneous complex, showing down hole lithology and Pt+Pd+Au concentrations (modified after Iljina et al., 2001).
matrix. Platinum-group elements in the reefs commonly occur in Pd-Te-Bi minerals, Pd-As-Sb minerals, and sperrylite (Törmänen, 1993). In the Sompujärvi reef, PGE-alloys and sulfides dominate locally. The grade of the PGE reefs varies but is roughly the same as in the Merensky reef or UG2 chromitite of the Bushveld Complex, i.e., on the order of several ppm. Locally, especially in the Sompujärvi reef, the concentration may rise up to several tens or hundreds of ppm.
124
5.5. Marginal series Cu-Ni-PGE and reef-type mineralization of the Koillismaa complex A classic example of Cu-Ni-PGE mineralization is found in the marginal series of the Western intrusion of the Koillismaa complex (Figure 3.9). This mineralized zone stretches sporadically along the total strike length of ~100 km of the marginal series. Disseminated sulfides, mainly pyrrhotite, chalcopyrite and pentlandite, are found in the lower half of the marginal series over a thickness of 20 m to 100 m. The grade commonly ranges between
• 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
Konttijärvi area
2b
Ko-3 0
4
8
4
ppm
Pd Ko-10
8 12 ppm
0
Dikes
Gabbroic cumulates
Archean granitoids
Pd
100 m
Ultramafic cumulates
Fig. 3.14. Geological cross-section of the Konttijärvi intrusion showing the geological setting of the basement granitoids, Portimo dikes and the main intrusion. Drill hole intersection Ko-10 represents mineralized Portimo dikes and adjacent granitoids, and Ko-3 represents an intersection through the marginal series and into the granitoids below. The horizontal and vertical scales are the same. Circled 2b refers to Figure 3.19. Modified after Iljina et al. (1992).
0.2 wt.% and 0.4 wt.% Cu and 0.2 wt.% and 0.3 wt.% Ni, but occasionally exceeds 1.0 wt.% for both metals. The average Ni and Cu contents of the sulfide fraction are 6 wt.% and 10 wt.%, respectively, while the wholerock sulfur concentration is typically 0.5–1.0 wt.%. Precious metal contents roughly follow those of the base metal sulfides and in many drillholes the grade is on the order of 0.5–1.0 ppm of combined Pt+Pd+Au. In certain places, the maximum combined PGE-Au contents reach 2–3 ppm (Figure 3.13), while in the massive Cu-sulfide veins the Pd contents reach several ppm. The platinum-group mineral species resemble those found in the Penikat intrusion and Portimo complex, as antimonides, tellurides, and bismuthides dominate the Pd mineralogy and an arsenide (sperrylite) dominates the Pt mineralogy (KoCHAPTER
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jonen and Iljina, 2001). A highly irregular PGE enrichment exists in the zone of the microgabbronorite xenoliths. This mineralization type can be regarded as reef-type due to the lateral continuity of the host rock unit. These microgabbronorites occur within a several tens of meters thick sequence that also contains gabbroic adcumulates, gabbro pegmatites, and felsic patches. The gabbro pegmatites and felsic patches in particular are enriched in base metal sulfides and platinum-group elements with grades reaching 1 wt.% of Cu and Ni and 1 ppm of Pt+Pd+Au.
5.6. Diverse Cu-Ni-PGE mineralizations in the Portimo complex Among the layered intrusions in Fennoscandia, the Portimo layered mafic complex
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Ahmavaara Pd (ppm) 4 12
S (wt. %) 10 30 Yp-60
Pd (ppm) 4
50 m
30 m
0m
0m
3
S (wt. %) 10 30
Basement
Marginal series
Yp-143
Suhanko
2a
Gabbroic cumulates
Ultramafic cumulates
Massive sulfides
Basement
Fig. 3.15. Two representative drill hole profiles through the Suhanko marginal series with disseminated and massive sulfide deposits. Drill hole Yp-143 is an example from the high-grade PGE Ahmavaara marginal series and Yp-60 shows an example of the other, more poorly PGE-mineralized marginal series type, represented here by the Suhanko massive pyrrhotite deposit. Circled 2a and 3 refer to Figure 3.19. Modified after Iljina et al. (1992).
is exceptional in hosting a variety of PGE mineralizations (Figures 3.2 and 3.14–3.19). The principal mineralization types are: • PGE-bearing Cu-Ni–Fe sulfide disseminations in the marginal series of the Suhanko and Konttijärvi intrusions • predominantly massive pyrrhotite deposits located close to the basal contact of the Suhanko intrusion • the Rytikangas PGE reef in the layered series of the Suhanko intrusion • the Siika-Kämä PGE reef in the Narkaus layered series • the offset Cu-PGE mineralization below the Narkaus intrusion Four other PGE enrichments are also depicted in Figure 3.19. These are (1) the PGE enrichment in the Portimo dikes below the Konttijärvi and Ahmavaara marginal series, (2) the PGE concentrations near the roof of 126
the Suhanko intrusions, associated mostly with pegmatites, (3) a Pt-anomalous pyroxenitic pegmatite pipe in the western limb of the Suhanko intrusion and (4) chromite and silicateassociated PGE enrichments in the lower parts of the Narkaus intrusion and MCU II. Disseminated PGE-bearing base metal sulfide showings, normally 10–30 m in thickness, occur throughout the marginal series of the Suhanko and Konttijärvi intrusions. Their distribution is erratic and they generally extend from the lower peridotitic layer downwards for some 30 m into the basement. The PGE contents vary from only weakly anomalous values to 2 ppm in most places in the marginal series of the Suhanko intrusion, but rise to >10 ppm in several places in the Konttijärvi and Ahmavaara areas. Highly PGE-enriched marginal series of this kind are rare in layered intrusions; another well-known occurrence is the Platreef in the northern Bushveld Complex. In the case of the Suhanko and Konttijärvi intru-
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Rytikangas Height (m)
Pt + Pd + Au (g/t) 0 4 8 12 0
Ala-Penikka
5 Cr (wt.%) Height 0.02 0.04 0.06 (m)
210
Pt + Pd + Au Cr (g/t) (wt.%) 0 1 2 3 4 5 0 0.02 0.06
Metacyclic unit IV
840
190
170
150
780 Yp-43
Ki-17
CUMULATES pbCa
pCa*
pbaC
Fig. 3.16. Comparison of the Rytikangas and Ala-Penikka PGE reefs from the Suhanko and Penikat intrusions, respectively, in terms of cumulus stratigraphy and variation in whole-rock precious metal and Cr concentrations as shown by drillholes Yp-43 (Rytikangas) and Ki-17 (Ala-Penikka). Circled 5 refers to Figure 3.19. Modified after Iljina and Lahtinen (1991). Abbreviations: p–plagioclase, b–bronzite, a–augite; letters before C denote cumulus phases, those after C intercumulus phases. The asterix (*) indicates that the mineral is poikilitic.
sions, the PGE grade and Cu and Ni contents of the sulfide fraction seem to correlate with the presence of the Portimo dikes underneath the intrusion (Figures 3.14–3.15). Massive sulfide mineralization is characteristic of the marginal series of the Suhanko intrusion. They form individual plate-like bodies generally varying in thickness from 20 cm to 20 m, separated from each other by more silicate-rich cumulate layers or granitoids. The mineralized bodies also vary in location from 30 m below the basal contact of the intrusion to a position 20 m above it and range in size from less than 1 million tons to more than 10 million tons. The sulfide paragenesis is composed almost exclusively of pyrrhotite, except in the Ahmavaara deposit, which also contains chalcopyrite and pentlandite. The massive CHAPTER
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pyrrhotite deposits show relatively low PGE values with the maximum Pt + Pd normally reaching a few ppm (circled 3 in Figures 3.15 and 3.19). However, similarly to the marginal series, disseminated sulfide mineralizations of the same intrusion (see above), the PGE concentrations are generally much higher in the Ahmavaara deposit, attaining a level of 20 ppm (circled 2a in Figures 3.15 and 3.19). The Rytikangas PGE reef represents the main PGE occurrence in the layered series of the Suhanko intrusion (Figures 3.2, 3.16, and 3.19), located in the middle of the western limb ~170 m above the base of the intrusion (Figure 3.6). Its position is known over a distance of 1.5 km. The Rytikangas reef is hosted by poikilitic plagioclase, plagioclase-bronzite, and plagioclase-bronzite-augite orthocumu-
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Siika-Kämä reef Knn-98
0 Height (m)
5 10 15
0
ppm
0.4
1 0.8
0
0.1
wt.%
wt.%
MCU III
210
190
MCU II
170
150 Gabbro pegmatite
Pt + Pd + Au
Gabbroic
S
Cr
Ultramafic
Fig. 3.17. Stratigraphic sequence across the Siika-Kämä PGE reef showing variations in bulk Pt+Pd+Au, S, and Cr in drill hole Knn-98. The figure represents a typical situation in which the PGE peak is at the base of MCU III.Vertical scale denotes structural height in meters. Circled 1 refers to Figure 3.19. Modified after Huhtelin et al. (1989b).
lates, all containing augite oikocrysts. This cumulate series overlies a 70-m sequence of monotonous plagioclase-bronzite-augite adcumulates and underlies 10 m of homogeneous plagioclase-bronzite mesocumulates with nonpoikilitic intercumulus augite. The orthocumulate layer varies in thickness from 30 cm to 10 m. The thickness of the reef itself is 30–50 cm and it typically occurs on top of the poikilitic orthocumulate layer. The cumulus stratigraphy and the drop in the whole-rock Cr content across the Rytikangas reef are practically identical to those of the Ala-Penikka reef of the Penikat intrusion (see Figures 3.2 and 3.16). The Siika-Kämä PGE reef of the Narkaus intrusion is most commonly located at the base of MCU III (Figures 3.6 and 3.17), but can also be found somewhat below this or in the middle of the olivine cumulate layer of MCU III. Chlorite-amphibole schist similar 128
to that at the Sompujärvi PGE reef in the Penikat intrusion commonly hosts the SiikaKämä reef. In some parts of the reef, the PGE mineralization is accompanied by chromite seams or disseminations. The thickness of the reef varies from less than one meter to several meters, and many drillholes penetrate a number of mineralized layers separated by PGE-poor layers, which can be several meters thick. The PGE concentration varies from anomalous values of several hundred ppb to tens of ppm. Some gabbroic pegmatites, abundant in the uppermost gabbroic adcumulates tens of meters below MCU III, are also mineralized and may contain several ppm of Pd and Pt. The Siika-Kämä deposit is one of the most sulfide-deficient PGE occurrences in the Portimo complex, in some places containing no visible sulfides and rarely exceeding a whole-rock sulfur content of 1 wt.%. The offset mineralization is sporadically
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Cu-PGE offset below the Narkaus intrusion
4
Knn-48 Overburden
Drillhole Knn-48 0 5 10 15 20 0 5 10 15 20 0 8 16 24 0 8 16 0 m
Narkaus intrusion -50
50
50 m Marginal series
-100
27.8
31.8
25.3
26.7
100
PGE in massive chalcopyrite PGE in disseminated chalcopyrite
Cu(wt.%) S(wt.%) Pd(ppm) Pt(ppm)
Archean basement complex
Layered intrusion
Overburden
Basement
Fig. 3.18. Schematic cross-section depicting the location of the offset PGE deposits beneath the Kilvenjärvi block of the Narkaus intrusion.Vertical variation of Cu, S, Pd, and Pt concentrations in drill hole Knn-48 is shown on the right. Circled 4 refers to Figure 3.19. Data from Huhtelin et al. (1989b).
distributed in the basement gneisses and granites below the Narkaus intrusion. The largest deposit is situated below the Kilvenjärvi block (Figure 3.5). This deposit is composed of a cluster of closely grouped smaller occurrences and is located in and near a N-trending major fault zone some tens of meters wide, against which the Kilvenjärvi block terminates. The offset mineralization represents the richest PGE deposit type within the Portimo area with Pt+Pd contents reaching 100 ppm (Figure 3.18). The offset mineralization is predominantly a Pd deposit, as it has a much higher Pd/Pt ratio than the other Portimo deposits (or any, for that matter, other Fennoscandian PGE deposit) and is extremely low in Os, Ir, Ru, and Rh. It is quite irregular in form, containing disseminated sulfide-PGM ‘clouds’, massive sulfide veins or bodies, and breccias, in which sulfide veins brecciate granitoids. The CHAPTER
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proportions of base metal sulfides and PGM are highly variable, but the massive sulfide bodies are always rich in PGE, while some samples containing almost no visible sulfides can carry several tens of ppm of Pd. In general, the more sulfide-rich occurrences are situated closer to the basal contact of the intrusion and those poorer in sulfides in a wider zone below the intrusion (Figure 3.18). Figure 3.19 shows a structural model for the Portimo complex and the positions of the deposits described above, as interpreted by Iljina (1994). Taking the boundary of the two parental magmas as a reference level, it can be seen that the Siika-Kämä reef, the highly mineralized Ahmavaara and Konttijärvi marginal series, and the mineralized Portimo dikes are located in the same position in magmatic stratigraphy. Accordingly, as a group they were referred to by Iljina (1994) as the Portimo
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LOCATION OF PGE MINERALIZATIONS
Higher-Cr magma Lower-Cr magma
PGE in gabbro pegmatites Suhanko massive sulfide deposit
Rytikangas reef
MCU III
PGE-anomalous ultramafic pipe
5
3
Portimo reef
2a, b
1
MCU II
Ahmavaara and Konttijärvi marginal series
MCU I PGE-bearing chromite and ultramafic layers
4
Offset PGE deposit
= Major PGE reef
Fig. 3.19. Schematic structural model for location of various PGE enrichments found in the Portimo layered igneous complex. The circled numbers refer to detailed illustrations in other figures as follows: 1 to Figure 3.17, 2 to Figures 3.14–3.15, 3 to Figure 3.15, 4 to Figure 3.18, and 5 to Fig. 3.16. Modified after Iljina (1994).
reef. Pulses of the earlier, higher- Cr parental magma are represented in the Konttijärvi and Ahmavaara areas by the Portimo dikes lying immediately below the marginal series. A marked decrease in PGE values and Ni and Cu contents of the sulfide fraction in the Suhanko marginal series occurs wherever the Portimo dikes are absent immediately below the marginal series. The Portimo and Rytikangas reefs were also tentatively noted to have higher whole-rock Se/S ratios than the other mineralization types in the Portimo complex (Iljina, 1994; Iljina and Lee, 2005).
5.7. PGE geochemistry The Paasivaara reef is Pt-dominated (Pd/Pt 0.4), while all the others (the ones discussed above) are Pd-dominated (Pd/Pt 2–8), except 130
the Sompujärvi and Siika-Kämä PGE reefs of the Penikat and Narkaus intrusions, in which the Pd/Pt ratio varies from ~0.8 to 5.0. Regardless of the quantity of base metal sulfides present, the PGE ratios and chondritenormalized PGE patterns (Figure 3.20) are practically constant throughout the Portimo reef. For example, in the Ahmavaara marginal series of the Portimo reef, the variation from disseminated to massive sulfides has only a minimal effect on the PGE patterns. A characteristic feature of the Rytikangas and Portimo reefs is a pronounced negative Ru anomaly, which has also been documented for the Sompujärvi, Ala-Penikka, and Paasivaara PGE reefs (although a less deep one in the PV). Additionally, the lower-grade PGE enrichments (mainly associated with chromitites) in MCU I–III of the Penikat and MCU I–II of
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100
A
Reefs: Ala-Penikka, Rytikangas, Siika-Kämä and Konttijärvi and Ahmavaara high-PGE 100 B grade disseminated and massive sulfides
C1 Chondrite normalised
Koillismaa marginal series 10
10
1
1
0.1
0.1 Suhanko lowPGE grade marginal series
0.01
0.01
0.001
0.001 Os
Ir
Ru
Rh
Pt
Pd
Au
Os
Ir
Ru
Rh
Pt
Pd
Au
Fig. 3.20. Chondrite-normalized whole-rock PGE and Au patterns. (A) Siika-Kämä, Rytikangas and Ala-Penikka PGE reefs, Konttijärvi and Ahmavaara high-grade PGE disseminated and massive sulfide deposits and lower-grade PGE massive sulfide deposits of the Suhanko marginal series (data from Iljina, 1994). (B) Marginal series disseminated sulfides from Koillismaa.
the Narkaus intrusions possess minor to deep negative Ru anomalies (Huhtelin et al., 1989a; Iljina, 1994). Such negative Ru anomalies are a rare phenomenon in PGE deposits worldwide. Among the platinum-group elements, Ru tends to fractionate with chromite crystallization: laurite (RuS2) is commonly found as inclusions within chromite. In contrast to the PGE behavior in the above mentioned reefs, the lower-grade PGE Suhanko massive sulfide deposits differ in having positive Rh and negative Pt anomalies whereas the Koillismaa marginal series PGE patterns show a steady positive slope (Figure 3.20).
6. Summary and discussion The Tornio–Näränkävaara belt crosses Finland CHAPTER
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almost along the Arctic Circle and extends also to the neighboring Sweden and Russia. It forms an essential part of the Fennoscandian 2.4–2.5 Ga layered intrusion complexes that were formed in a mantle-plume related intracratonic breakup of the Archean craton. 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. The individual intrusions and intrusion complexes of the Tornio–Näränkävaara belt are commonly located between the Archean complexes in the south and the Paleoproterozoic Peräpohja and Kuusamo schist belts in the north, with the exception of the Näränkävaara intrusion and its “connective dike”, which are exclusively surrounded by Archean rocks. The structural, lithological, and geochemical similarity of
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individual layered bodies of the Western intrusion of the Koillismaa complex suggest that they are tectonically dismembered fragments of a large intrusion. The stratigraphy of the layered intrusions in the Tornio–Näränkävaara belt can be divided into a reversely zoned marginal series and an overlying layered series. The latter has a total present-day thickness varying from a few hundred meters to nearly 3 km. The marginal series successions can be thin (<20 m), as in the Kemi and Penikat intrusions and the Narkaus intrusion of the Portimo complex, or much thicker, as in the Suhanko and Konttijärvi intrusions (Portimo complex) and the Western intrusion (Koillismaa complex). Significant differences exist among the layered series of the various intrusions. Characteristically, the intrusions in the western and central parts of the Tornio–Näränkävaara belt display a repetition of ultramafic and mafic rocks that constitute megacyclic units, at maximum, five in Penikat. This cyclic distribution of rock types coupled with the compositional variation of primary minerals indicates opensystem fractionation involving multiple injections of basaltic magma into high-level magma chambers. In contrast, the Western intrusion of the Koillismaa complex lacks the analogous megacyclic structure and instead fractionated to the point of producing significant Fe-Ti-V oxide enrichment now represented by magnetite gabbro, essentially unknown elsewhere in the Tornio–Näränkävaara belt. The cumulate sequences of the Tornio–Näränkävaara belt intrusions were generated by various parental magmas, which nevertheless shared some common features including enrichment in LREE over HREE and depletion of Ti and other HFSE. The igneous bodies in the western and central part of the belt were fed with two kinds of magma in terms of their Cr content, one higher in Cr and the other lower in Cr. Some of the intrusions were formed from one or the other magma type, while others expecienced injection of first the 132
higher-Cr and then lower-Cr magma type. It has also been concluded that a third parental magma type, a more evolved tholeiitic basalt, was involved in the Western intrusion of the Koillismaa complex. The mafic intrusive bodies and their immediate country rocks show a number of different kinds of mineralization. These are basal disseminated and massive Cu-Ni-Fe sulfides enriched in PGE, offset base metal sulfidePGM lodes in footwall rocks, reef-type PGE enrichments within the layered series, and Cr- and Fe-Ti-V oxide accumulations either at the base of the intrusion or higher up in the sequence. A world-class chrome deposit is located at the base of the Kemi intrusion and the magnetite gabbro layer of the Western intrusion of the Koillismaa complex has been exploited for vanadium at Mustavaara. The PGE and chalcophile element enrichment can be classified into three main types: (1) disseminated and massive PGE bearing Cu-Ni-Fe sulfides in the marginal series, (2) reef-type PGE deposits, and (3) offset deposits in footwall rocks. The first type is almost exclusively confined to well-developed, thicker marginal series as found in the Suhanko and Konttijärvi intrusions of the Portimo complex and the Western intrusion of the Koillismaa complex. In the Portimo complex, both disseminated and massive concentrations of sulfides have been discovered, while the Koillismaa complex contains only disseminated sulfides. Potentially world-class PGE reefs, five in number, have been found in the Penikat intrusion and in the Portimo complex. As a natural test laboratory for the genesis of PGE deposits, the Tornio–Näränkävaara belt offers a unique opportunity due to the variety of magma compositions emplaced within a short period of time. Whatever model is proposed for the genesis of the PGE deposits found in the Portimo complex and Penikat intrusion, it must be able to account for their high number within a relatively restricted area, high grade, distribution in separate intrusive
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bodies, and location in diverse host rocks. Stratigraphically, the first pulse of the lowerCr magma type marks the level of the lowermost high-grade PGE reef. In the case where the intrusions have a higher-grade marginal series, the early higher-Cr magma injection is only represented by a set of dikes beneath the intrusion. The next reef in the sequence is always found some hundred meters above the lowermost enrichment regardless of the host intrusion. The highest reef is located at the top of the lowermost lower-Cr megacyclic unit. None of the generally applied theories of PGE ore-forming processes taking place only within a single intrusion chamber can account for all of the observed characteristics. The fact that the megacyclic units, and not a particular intrusion or stratigraphic height, seem to control reef-type PGE mineralization requires a genetic model in which the intrusions with a megacyclic structure were fed from a common auxiliary magma chamber lower in the crust. All the above mentioned reef types and high-grade PGE marginal series have a moderate to deep negative Ru anomaly, which is a rare phenomenon worldwide. It is known that chromite accumulation fractionates Ru. Since the chromite reefs underlying the PGE reefs also have negative Ru anomalies, the depletion of Ru can be attributed to fractionation processes that took place in the proposed auxiliary magma chamber. This may link the Kemi intrusion and its huge chromite mass to the same system, as the feeder dike of the Kemi intrusion is noted for containing suspended chromite.
Acknowledgments We are grateful to The South Atlantic Resources Ltd. for allowing us to log their exploration drillholes from the upper Penikat intrusion, and to The Arctic Platinum Partnership Ay for permission to publish the Konttijärvi cross and longitudinal sections. Jorma Räsänen is CHAPTER
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thanked for his contribution to the geological map of the KLIC. Constructive comments on the manuscript by Hugh O’Brien are gratefully acknowledged.
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lia: Petrological implications. Can. Mineral. 39, 607–637. Halkoaho, T., 1993. The Sompujärvi and AlaPenikka PGE reefs in the Penikat layered intrusion, northern Finland. Acta Univ. Ouluensis. Ser. A, Sci. Rer. Natur. 249. Halkoaho, T.A.A., Alapieti, T.T., Lahtinen, J.J., 1990a. The Sompujärvi PGE reef in the Penikat layered intrusion, northern Finland. In: E. F. Stumpfl, H. Papunen (Eds.), Proceedings of the Fifth International Platinum Symposium, Helsinki 1989. Mineral. Petrol. 42, 39–55. Halkoaho, T.A.A., Alapieti, T.T., Lahtinen, J.J., Lerssi, J.M., 1990b. The Ala-Penikka PGE reefs in the Penikat layered intrusion, northern Finland. In: E.F. Stumpfl, H. Papunen, (Eds.), Proceedings of the Fifth International Platinum Symposium, Helsinki 1989. Mineral. Petrol. 42, 23–38. Hanski, E., Walker, R.J., Huhma, H., Suominen, I., 2001. The Os and Nd isotopic systematics of the c. 2.44 Ga Akanvaara and Koitelainen mafic layered intrusions in northern Finland. Precambrian Res. 109, 73–102. Hatton, C., Sharpe, M., 1989. Significance and origin of boninite-like rocks associated with the Bushveld Complex. In: A. Crawford (Ed.), Boninites. University Press, Cambridge, pp. 174–207. Huhma, H., Cliff, R.A., Perttunen, V., Sakko, M., 1990. Sm-Nd and Pb isotopic study of mafic rocks associated with early Proterozoic continental rifting: the Peräpohja Schist Belt in northern Finland. Contrib. Mineral. Petrol. 104, 369–379. Huhtelin, T.A., Alapieti, T.T., Lahtinen, J.J., Lerssi, J., 1989a. Megacyclic units I, II and III in the Penikat layered intrusion. In: T. Alapieti (Ed.), 5th International Platinum Symposium. Guide to the post-symposium field trip, August 4–11, 1989. Geol. Surv. Finland, Guide 29, 59–69. Huhtelin, T.A., Lahtinen, J.J., Alapieti, T.T., Korvuo, E., Sotka, P., 1989b. The Narkaus intrusion and related PGE and sulphide mineralizations. In: T. Alapieti (Ed.), 5th International
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Platinum Symposium. Guide to the postsymposium field trip, August 4–11, 1989. Geol. Surv. Finland, Guide 29, 145–161. Huhtelin, T.A., Alapieti, T.T., Lahtinen, J.J., 1990. The Paasivaara PGE reef in the Penikat layered intrusion, northern Finland. In: E.F. Stumpfl, H. Papunen (Eds.), Proceedings of the Fifth International Platinum Symposium, Helsinki 1989. Mineral. Petrol. 42, 57–70. Iljina, M., 1994. The Portimo Layered Igneous Complex with emphasis on diverge sulphide and platinum-group element deposits. Acta Univ. Ouluensis. Ser. A, Sci. Rer. Natur. 258. Iljina, M., 2005. Portimo layered igneous complex, In: T. Alapieti, T. Kärki (Eds.), 10th International Platinum Symposium. Field trip guidebook. Geol. Surv. Finland, Guide 51a, 77−100. Iljina, M., Lahtinen, J., 1991. Geokemian käyttö kerrosintruusioiden platinatutkimuksissa. In: R. Salminen, S. Roos (Eds.), Geokemian päivät Oulussa 28–29.11.1990. Publication of The Finnish Association of Mining and Metallurgical Engineers, Ser. B, 50, 106–117. (in Finnish) Iljina, M., Lee, C., 2005. PGE deposits in the marginal series of layered intrusions. In: J. Mungall (Ed.), Exploration for platinumgroup element deposits. Mineral. Assoc. Canada, Short Course Ser. 35, 75−96. Iljina, M., Alapieti, T.T., Lahtinen, J.J., Lerssi, J.M., 1989. The Suhanko-Konttijärvi intrusion and related sulphide and PGE mineralizations. In: T. Alapieti (Ed.), 5th International Platinum Symposium. Guide to the postsymposium field trip, August 4–11, Geol. Surv. Finland, Guide 29, 163–187. Iljina, M., Alapieti, T., McElduff, B., 1992. Platinum-group element mineralization in the Suhanko-Konttijärvi intrusion, Finland. Australian J. Earth Sci. 39, 303–313. Iljina, M.J., Karinen, T., Räsänen, J., 2001. The Koillismaa Layered Complex: general geology, structural development and related sulphide and platinum-group element min-
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eralization. In: A. Piestrzynski et al. (Eds.), Mineral deposits at the beginning of the 21st century. A.A. Balkema, Lisse, 649–652. Juopperi, A., 1977. The magnetite gabbro and related Mustavaara vanadium ore deposit in the Porttivaara layered intrusion, northeastern Finland. Geol. Surv. Finland, Bull. 288. 1–68. Karinen T., Salmirinne, H., 2001. Koillismaan kerrosintruusiokompleksin läntisen osan geologinen evoluutiomalli. Unpubl. report, M19/3543/2001/2, Geol. Surv. Finland. (in Finnish) Kojonen, K., Iljina, M., 2001. Platinum-Group Minerals in the Early Proterozoic Kuusijärvi Marginal Series, Koillismaa Layered Igneous Complex, Northeastern Finland. In: A. Piestrzynski et al. (Eds.), Mineral deposits at the beginning of the 21st century. A.A. Balkema, Lisse, pp. 653–656. Korsman, K., Koistinen, T., Kohonen, J., Wennerström, M., Ekdahl, E., Honkamo, M., Idman,H., Pekkala, Y. (Comps.), 1997. Bedrock Map of Finland, 1 : 1 000 000. Geol. Surv. Finland, Espoo. Lahtinen, J., 1985. PGE-bearing copper-nickel occurrences in the marginal series of the Early Proterozoic Koillismaa layered intrusion, northern Finland. In: H. Papunen, G. I. Gorbunov (Eds.), Nickel-copper deposits of the Baltic Shield and Scandinavian Caledonides. Geol. Surv. Finland, Bull. 333, 161–178. Lahtinen, J. J., Alapieti, T. T., Halkoaho, T. A. A., Huhtelin, T. A., Iljina, M. J., 1989. PGE mineralization in the Tornio–Näränkävaara layered intrusion belt. In: T. Alapieti (Ed.), 5th International Platinum Symposium. Guide to the post-symposium field trip, August 4–11, 1989. Geol. Surv. Finland, Guide 29, 43–58. Lauri, L.S., Mänttäri, I., 2002. The Kynsijärvi quartz alkali feldspar syenite, Koillismaa, eastern Finland–silicic magmatism associated with 2.44 Ga continental rifting. Precambrian Res. 119, 121–140. Lauri, L.S., Rämö, O.T., Huhma, H., Mänttäri, I.,
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Räsänen. J., 2005. Petrogenesis of silicic magmatism related to the ~2.44 Ga rifting of Archean crust in Koillismaa, eastern Finland. Lithos, in press. Leinonen, O., 1998. Use of chromitite microstructure image analysis to estimate concentration characteristics in the Kemi chrome ore. Acta Univ. Ouluensis. Ser. A, Sci. Rer. Natur. 305. Luukkonen, E.J., 1988. Moisiovaaran ja AlaVuokin kartta-alueiden kallioperä. English summary: Pre-Quaternary rocks of the Moisiovaara and Ala-Vuokki map-sheet areas. Geological map of Finland 1:100 000. Explanation to the maps of Prequaternary rocks, sheet 2421 and 4423 + 4441. Geol. Survey Finland, Espoo. Maier, W.D., Arndt, N.T., Curl, E.A., 2000. Progressive crustal contamination of the Bushveld Complex; evidence from Nd isotopic analyses of the cumulate rocks. Contrib. Mineral. Petrol. 140, 316–327. Manhes, G., Allégre, C.J., Dupré, B., Hamelin. B., 1980. Lead isotopic study of basic–ultrabasic layered complexes: Speculations about the age of the earth and primitive mantle characteristics. Earth Planet. Sci. Lett. 47, 370–382. Mutanen, T., 1997. Geology and ore petrology of the Akanvaara and Koitelainen mafic layered intrusions and the Keivitsa-Satovaara layered complex, northern Finland. Geol. Surv. Finland, Bull. 395, 1–233. Mutanen, T., Huhma, H., 2001. U-Pb geochronology of the Koitelainen, Akanvaara and Keivitsa mafic layered intrusions and related rocks. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Paper 33, 229–246. Patchett, J., Kouvo, O., Hedge, C., Tatsumoto, M., 1981. Evolution of continental crust and mantle heterogeneity: evidence from Hf isotopes. Contrib. Mineral. Petrol. 78, 279–297.
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Perttunen, V., 1991. Kemin, Karungin, Simon ja Runkauksen kartta-alueiden kallioperä. Summary: Pre-Quaternary rocks of the Kemi, Karunki, Simo and Runkaus mapsheet areas. Geological Map of Finland 1 : 100 000, Explanation to the Map of Rocks, Sheets 2541, 2542+2524, 2543 and 2544. Geol. Surv. Finland, Espoo. Perttunen, V., Vaasjoki, M., 2001. U-Pb geochronology of the Peräpohja Schist Belt. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Paper 33, 45–84. Piirainen, T., Hugg, R., Isohanni, M., Juopperi, A., 1974. On the geotectonics and ore forming processes in the basic intrusive belts of Kemi-Suhanko, and Syöte-Näränkävaara, northern Finland. Bull. Geol. Soc. Finland 46, 93–104. Piirainen, T., Hugg, R., Aario, R., Forsström, L., Ruot sa lainen, A., Koivumaa, S., 1978. Koillismaan malmikriittisten alueiden tutkimusprojektin loppuraportti 1976. Summary: The report of the Koillismaa research project. Geol. Surv. Finland, Rep. Invest. 18, 1–51. (in Finnish) Saini-Eidukat, B., Alapieti, T.T., Thalhammer, O.A.R., Iljina, M.J., 1997. Siliceous, highmagnesian parental magma compositions of the PGE-rich Early Paleoproterozoic layered intrusion belt of northern Finland. In: P. Rongfu (Ed.), Proceedings of the 30th International Geological Congress, Beijing, China, 4–14 August, 1996. Vol. 9, Energy and Mineral Resources for the 21st Century, Geology of Mineral Deposits, Mineral Economics, VSP, Utrecht, pp. 177–197. Salmirinne, H., Iljina, M., 2003. Koillismaan kerrosintruusiokompleksin tulokanavamuodostuman painovoimatulkinta ja alueen malmimahdollisuudet (osa 1). Unpubl. report, Q21/2003/1, Geol. Surv. Finland. (in Finnish) Sharkov, E.V., Bogatikov, O.A., Grokhovskaya, T.L., Chistyakov, A.V., Ganin, V.A., Grinevich,
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N.G., Snyder, G.A., Taylor, L.A., 1995. Petrology and Ni-Cu-Cr-PGE mineralization of the largest mafic pluton in Europe: the Early Proterozoic Burakovsky layered intrusion, Karelia, Russia. Internat. Geol. Rev. 37, 509–525. Törmänen, T., 1993. Penikkain kerrosintruusion Sompujärvi-, Ala-Penikka- ja Paasivaaramineralisaatioiden platinamineralogia. Unpubl. M.Sc.Thesis, Dept. of Geol. Univ. of Oulu. (in Finnish) Torokhov, M.P., Fedotov, G.A., Karzhavin, V.K., Sholokhnev, V.V., 1997. The Imandra Layered Complex and related mineralization.
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In: F.P. Mitrofanov, M.P. Torokhov, M. Iljina (Eds.), Ore deposits in the Kola Peninsula, Northwestern Russia. Research and exploration – where do they meet? 4th Biennial SGA Meeting, August 11–13, 1997, Turku, Finland. Excursion guidebook B4. Geol. Surv. Finland, Guide 45, 33–38. Vogel, D.C., Vuollo, J.I., Alapieti, T.T., James, R.S., 1998. Tectonic, stratigraphic, and geochemical comparisons between ca. 2500–2440 Ma mafic igneous events in the Canadian and Fennoscandian Shields. Precambrian Res. 92, 89–116.
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CENTRAL LAPLAND GREENSTONE BELT
E. Hanski, H. Huhma
Cover page: Kumpu Group conglomerate, Kumputunturi, Kittilä. Diameter of the largest clasts is ~10 cm. Photo: Eero Hanski.
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Hanski, E., Huhma, H., 2005. Central Lapland greenstone 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. 139–194. © 2005 Elsevier B.V. All rights reserved.
The Central Lapland greenstone belt records Paleoproterozoic depositional evolution for almost 600 Ma, beginning at ~2450 Ma with the eruption of mantle plume-related, komatiitic to rhyolitic lavas on the Archean cratonic basement. This magmatic phase also included emplacement of large layered mafic intrusions. During the following 300–400 Ma, deposition of a thick, transgressive quartzite-dolomite-basalt-pelite succession took place and was followed by komatiitic to picritic volcanism. Mafic magmas intermittently formed layered sill-like intrusions within the sediments, most notably at ~2220 Ma and ~2050 Ma. This prolonged extensional regime was interrupted by a collisional event, which led to thrusting of a ~2000 Ma slab of ancient oceanic lithosphere (the Kittilä Group) onto older cratonic rocks. This took place at ~1920 Ma, as deduced from felsic dikes and granodioritic plutons of this age within the Kittilä Group. Roughly simultaneously, a juvenile calc-alkaline arc complex was formed farther north (western part of the Inari area) and was shortly followed by the upthrust of the Lapland granulite belt. The supracrustal rock sequence in central Lapland was completed with the deposition of molasse-like, coarseclastic sediments in a fore-arc basin soon after ~1880 Ma synorogenic felsic plutonism and associated minor volcanism.
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1. Introduction The Central Lapland greenstone belt in northern Finland, together with its continuations in northern Norway, Sweden, and Russian Kare lia, forms one of the largest known Paleoproterozoic greenstone belts. Its age and stratigraphic relations have been controversial for a long time. Recent geochemical and isotope studies coupled with extensive field investigations carried out by the Lapland Volcanite Project (Geological Survey of Finland) in the late 1980’s have considerably improved our understanding of the geochemical characteristics of the belt and the stratigraphic and age relations of its individual formations (Räsänen et al., 1995; Lehtonen et al., 1998). A large quantity of previously unpublished geochronologic data from the Central Lapland greenstone belt have recently been reported (Vaasjoki, 2001). The supracrustal evolution of the Central Lapland greenstone belt lasted for several hundreds of m.y., resulting in thick sequences of sedimentary and volcanic rocks generated in various geotectonic settings. In this paper, we present an overview of this depositional evolution on the basis of earlier studies and some new Nd isotope data. In addition to supracrustal rocks, we consider three episodes of intrusive basic magmatism that produced mafic–ultramafic layered intrusions at 2440, 2220, and 2050 Ma, some of them with notable ore reserves.
2. Main geologic units of northern Finland The bedrock of northern Finland is part of the Fennoscandian Shield and is composed mainly of Archean granite gneiss-greenstone terrains and Paleoproterozoic schist belts, which were mostly deposited on or, in some cases, tectonically juxtaposed on the Archean basement. Archean TTG and quartzofeldspathic gneisses, granites, and supracrustal 142
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belts, dominated by pelitic metasediments or mafic metavolcanic rocks, are found in several large domains such as the Archean areas of eastern and western Lapland and the eastern part of the Inari area (Figure 4.1). In the south, the Pudasjärvi and eastern Finland basement complexes extend northwards to southernmost Lapland. In addition, small Archean gneiss domes surrounded by Proterozoic rocks are found in central Lapland. The focus of this article is on the Paleoproterozoic rocks, particularly the Central Lapland greenstone belt, which runs as an almost uninterrupted zone from northern Norway through central Finnish Lapland where it bifurcates to reach both the western and eastern borders of Finland (Figure 4.1). The Central Lapland greenstone belt contains several different areas in which either volcanic or sedimentary rocks prevail. Volcanic terrains in Kittilä and Salla are assigned to the Kittilä and Salla greenstone areas, respectively. These are separated by the Sodankylä schist area, which is composed predominantly of metasedimentary rocks. Lehtonen et al. (1998) regarded the Salla greenstone area as the easternmost part of the Central Lapland greenstone belt, whereas Nironen et al. (2002) considered it the northernmost part of the Kuusamo belt, a supracrustal belt in southeast Lapland running across the Finnish-Russian border. In southwestern Lapland there is the Peräpohja belt that is separated from the Central Lapland greenstone belt by the Central Lapland granitoid complex. Smaller, less intensively studied Paleoproterozoic schist belts are found in the most northwestern corner of Finland (in the “right hand of the Finnish Lady”) and in northernmost Finland north of the Lapland granulite belt (Figure 4.1). The granulite zone can also be regarded as belonging to the Paleoproterozoic formations on the basis of whole-rock Sm-Nd analyses and U-Pb analyses of detrital zircons; these indicate a Paleoproterozoic age for at least part of the protoliths (Huhma and Meriläinen, 1991; Tuisku and Huhma, 1998a,
LAPLAND
GREENSTONE
B E LT
Caledonides
NORWAY d lan Lap
Inari area
Carbonatite Granitoids
it nul gra elt
nt Ce
eb
Layered mafic–ultramafic intrusions, anorthosite Granulite
l ra
Juvenile arc complex
pla La
Ki
nd
Paleoproterozoic supracrustal belts Archean supracrustal belts
ns ee gr ne
to
Archean gneiss and granite
So
be
Sa
lt
SWEDEN
Central Lapland granitoid complex
RUSSIA
Pe
Ku
100 km
Fig. 4.1. Main geologic units of northern Finland (modified after Korsman et al., 1997). Abbreviations: Pe–Peräpohja schist belt, Ku–Kuusamo schist belt, So–Sodankylä schist area, Sa–Salla greenstone area, Ki–Kittilä greenstone area.
b; Daly et al., 2001). Adjacent to the Lapland granulite belt on its northeastern side are small Paleoproterozoic schist belts and ~1950–1900 Ma juvenile calc-alkaline plutonic rocks (Barling et al., 1997; Huhma, 2001). Some of the largest layered mafic–ultramafic intrusions in northern Finland are shown in black in Figure 4.1. Felsic plutonic rocks are found widely south of the Central Lapland greenstone belt, in an area designated the Central Lapland granitoid complex. It is a heterogeneous entity containing, in addition to Paleoproterozoic granites and migmatites, remnants of Archean granitoids and schists (Evins et al., 2002). Rocks of the Central Lapland greenstone belt are cut by synorogenic monzodiorites and granodiorites of CHAPTER
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the ~1880 Ma Haaparanta (Haparanda) suite occurring mainly in the western part of Finnish Lapland, and by postorogenic, ~1800 Ma, “Nattanen-type” granites forming relatively small, scattered stocks throughout Lapland. Restricted felsic magmatism with an age of ~1920 Ma is also found in central Lapland. The youngest rocks in northern Finland are represented by some Neoproterozoic mafic dikes and Phanerozoic alkaline complexes in eastern Lapland, such as the Sokli carbonatite, and a small piece of the Caledonides in the northwestern corner of Finland (Figure 4.1).
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3. Central Lapland greenstone belt 3.1. General features The Paleoproterozoic Central Lapland greenstone belt is exposed over an area of 100 km by 200 km and is part of a larger greenstone belt extending up to the Norwegian coast. In northern Finnish Lapland, it runs parallel to and plunges beneath the southwestern contact of the Lapland granulite belt, while in the west and east it is bordered by Archean granitegneiss terrains. In the south and northwest it is intruded by plutons of granitic rocks (Figures 4.1 and 4.2). The Paleoproterozoic greenstone belt is underlain by Archean basement gneisses, and the lower parts of the supracrustal sequence were deposited unconformably on these 3.1 Ga to 2.6 Ga gneisses (Kröner and Compston, 1990; Meriläinen, 1976). The Central Lapland greenstone belt records a complex geologic history of more than 500 Ma culminating in orogenic deformation at ~1.9 Ga. Affected by the Svecofennian collisional event in the southwest and thrusting of the Lapland granulite belt in the northeast, it acted as a foreland fold-and-thrust belt to two, nearly coeval convergent systems with op posing polarity (Sorjonen-Ward et al., 1997). Structures in central Lapland have been interpreted with progressive compression and regionally imposed dextral torque (Ward et al., 1989). Hanski (1997) suggested that the Kittilä greenstone area represents an oceanic allochthon, but the most part of the belt seems to be autochthonous or parautochthonous. In general, apart from some recumbent folding close to the contact of the granulite belt, major overturned nappes have not been easy to recognize (Sorjonen-Ward et al., 1997). Recently, Evins and Laajoki (2002) reported evidence for south-vergent nappe structures in the Sodankylä schist area 100–150 km south of the basal contact of the granulite belt. Later tectonic activation is manifested 144
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by some notable shear zones. Among these are the “Sirkka-line” at the southern margin of the Kittilä greenstone area and the northtrending Kolari shear zone running close to the Finnish-Swedish border. Central Lapland displays a relatively complex pattern of metamorphic zonation, but a simplified picture is as follows. Within the core zone of the belt, comprising the major part of the Kittilä greenstone area and the northwestern part of the Sodankylä schist area, greenschist facies metamorphism prevails. Towards the Central Lapland granitoid complex in the south and Svecofennian granites in the west, the degree of regional metamorphism increases rather abruptly to amphibolite facies. The eastern part of the belt has mostly undergone middle-amphibolite facies metamorphism. There is a transition from lower grades to upper amphibolite and granulite facies towards the northeast before reaching the granulite belt proper.
3.2. Lithostratigraphy Although the supracrustal rocks of the Central Lapland greenstone belt have most often been regarded as belonging to the Karelian formations (e.g., Eskola, 1963; Simonen, 1986), their correlation with the traditional Karelian formations (Sariolian, Jatulian, and Kalevian; see Laajoki, 1986; Chapters 1 and 7) has been controversial for decades, partly due to the lack of reliable age determinations. Nevertheless, the discovery made by Hackman (1927) indicating the presence of two principal stratigraphic entities separated by a major unconformity has generally been accepted. For the younger unit containing exclusively epiclastic metasedimentary rocks, Hackman (1927) coined the term ‘Kumpu quartzites’, while Sederholm (1932) introduced the term ‘Lapponian’ for the widely distributed volcano-sedimentary sequence of the lower unit (see also Mikkola, 1941). For a more detailed account of the history of the stratigraphical
LAPLAND
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27°
M Kittilä
S Sodankylä
Kolari
Pyhätunturi
67°
K
SWEDEN
RUSSIA
0
RUSSIA
Salla
40 km
Lithostratigraphic groups of central Lapland greenstone belt
Other rock units
Lainio and Kumpu Groups
Paleozoic carbonatite
Kittilä Group
Paleoproterozoic intrusive rocks
Savukoski Group
Lapland granulite
Sodankylä Group
Archean supracrustal rocks
Onkamo Group
Archean gneisses and granites
Salla Group
Fig. 4.2. The areal distribution of the main lithostratigraphic units (groups) in central Finnish Lapland (simplified after Räsänen et al., 1995). M denotes the location of the Möykkelmä dome, and S and K mark the major komatiite occurrences, Sattasvaara and Kummitsoiva, respectively.
research carried out in northern Finland, the reader is referred to Hanski (2001). A premise for the adoption of the term ‘Lapponian’ in the first place (Sederholm, CHAPTER
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1932) and implicit in subsequent stratigraphic schemes (e.g., Silvennoinen, 1985) was the notion that the Lapponian rocks are older (>2300 Ma) than the Jatulian rocks. However, geologi-
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cal, geochemical, and geochronological investigations carried out in the 1980’s and 1990’s have revealed that both the Kumpu quartzites and a large part of Sederholm’s Lapponian volcanic rocks are younger than the Jatulian and do not have any obvious counterparts in the conventional Karelian stratigraphic units (Lehtonen et al., 1998; Rastas et al., 2001). This contradiction and the obvious necessity to update the stratigraphic terminology led to the rejection of the traditional stratigraphic nomenclature (Räsänen et al., 1995; Lehtonen et al., 1998). Recent geochronological studies have shown that the Paleoproterozoic supracrustal rocks earlier assigned to the Lapponian rocks were deposited over ~450 Ma (see Hanski et al., 2001b). Lehtonen et al. (1998) divided them into five lithostratigraphic groups which, from oldest to youngest, are the Salla, Onkamo, Sodankylä, Savukoski, and Kittilä Groups. In addition, Lehtonen et al. assigned the younger, traditional Kumpu formations to two units, the Lainio and Kumpu Groups. The distribution of these stratigraphic units in the Central Lapland greenstone belt is shown in Figure 4.2. For simplicity, the Kumpu and Lainio Groups are combined on that map. In the following, the Paleoproterozoic supracrustal evolution of the Central Lapland greenstone belt is discussed by describing each of the above-mentioned groups in geochronologic order. While Lehtonen et al. (1998) provided information on the type of formations of each group, we will maintain our focus at the group level.
3.3. Salla Group The Salla Group represents the lowermost Paleoproterozoic lithostratigraphic unit of the Central Lapland greenstone belt, occurring most abundantly in the Salla greenstone area in the southeastern part of the Central Lapland greenstone belt (previously called the Salla Formation by Manninen, 1991). Together with their extensions on the Russian side of 146
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the border, the Salla Group and the overlying Onkamo Group metavolcanic rocks occupy here an approximate area of 40 km by 100 km. Due to the gentle dips of the volcanic deposits, the total thickness of the group is difficult to estimate but may reach two kilometers (Tuomo Manninen, pers. comm., 2001). Other notable occurrences of the Salla Group are found in the vicinity of the Akanvaara and Koitelainen layered intrusions in the central part of the Central Lapland greenstone belt, whereas Salla Group rocks are not found in the western part of the belt. The volcanic rocks were erupted onto the rifted Archean basement. In the Salla area, however, the basement rocks are exposed only on the Russian side and the contact of the Salla greenstone area with other rock units on its western side is tectonic (Manninen, 1991). In the Peurasuvanto area north of the Koitelainen intrusion, the Salla Group metavolcanic rocks surround the 3100 Ma Tojottamanselkä basement gneiss dome. The succession begins with a basal volcaniclastic conglomerate with a variety of rock clasts including rounded Archean gneiss cobbles (Peltonen et al., 1988). Both the Salla Group and the overlying Onkamo Group rocks are overwhelmingly volcanic in origin. The Salla Group metavolcanic rocks vary from intermediate to felsic with the more evolved varieties generally higher in stratigraphy. Even though the primary minerals of the Salla Group metavolcanic rocks have generally been obliterated by greenschist facies regional metamorphism, the original volcanic structures of the extrusive rocks are well-preserved in many places, especially in the Salla area (Figure 4.4A). Felsic metavolcanic rocks are characterized by welded ash flow tuffs and crystal tuffs. The field characteristics of the metalavas coupled with the scarcity of intervening epiclastic metasedimentary rocks indicate that eruptions occurred rapidly under subaerial conditions.
LAPLAND
GREENSTONE
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A
Ruppapalo granodiorite Nyssäkoski Sotkaselkä 2.2 2.6 1
Levi
Sirkka
2.9
3.4
Kumputunturi
4 2
Sovasjoki Dome
3.7
” ine al
Ylläs
Mantovaara conglomerate k irk “S
4.4 3.3 3.9
Nuttio
Jeesiörova
7
5
1.4
-1.1 -1.6
Kittilä
-0.8
Latvajärvi
6 3
Kaarestunturi
Kallo monzonite
Paleoproterozoic supracrustal rocks
10 km
Other rock units
----------------
Kumpu Group
Onkamo Group
.. . . . . . . . . . . . . . . . . . .
Lainio Group metasediments Lainio Group metavolcanic rocks Kittilä Group
Paleoproterozoic intrusive rocks Serpentinite
Savukoski Group
Archean rocks
Sodankylä Group
Fault
B
Salla Group
3.7
2
Initial εNd value Location of studied conglomerate clasts: 1 = Mantovaara 2 = Linkupalo 3 = Vesikkovaara 4 = Sätkänävaara 5 = Hangasoja 6 = Kellostapuli 7 = Aakenustunturi
Ruppapalo granodiorite
Sirkka Sovasjoki Dome
Kittilä
Fig. 4.3. (A) Geological map of the Kittilä area. Numbers in white boxes in the Kittilä Group area denote εNd (at 2015 Ma) values for mafic metavolcanic rocks. Circled numbers mark locations where felsic porphyry and granitoid clasts of the Lainio and Kumpu Group conglomerates have been studied. (B) Gray-tone aeromagnetic map of low-altitude total intensity of the Kittilä area.
Kallo monzonite
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147
A
B
C
D
E
F
G
H
Fig. 4.4. (A) Amygdaloidal mafic lava, Salla Group. Tag length 15 cm. (B) Partly assimilated granitoid fragments from basement in mafic lava, Onkamo Group. Tag length 10 cm. (C) Fuchsite-bearing, crossbedded quartzite, Sodankylä Group. Alteration features in the lower part of the figure. Tag length 5 cm. (D) Ultramafic (komatiitic) volcaniclastic metavolcanic rock, Savukoski Group. Coin diameter 25 mm. (E) Photomicrograph of komatiitic hyaloclastite, Savukoski Group. Crossed nicols, bar length 1 mm. (F) Mafic pillow lava, Kittilä Group. Tag length 15 cm. (G) Photomicrograph of felsic porphyry, Kittilä Group, crossed nicols. Bar length 3 mm. (H) Lainio Group conglomerate containing felsic porphyry and granitoid pebbles. Bar length 3 cm. Photos: (A) and (B) by Tuomas Manninen, (C) by Pentti Rastas, (D) and (E) by Jorma Räsänen, (G) and (H) by Reijo Lampela.
148
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Sample/Primitive Mantle
300
100
Salla Group metavolcanic rocks
10
2 Rb
Th Nb K Ce Sr Nd Zr Eu Gd Dy Er Lu Ba U Ta La Pr P Sm Hf Ti Tb Y Yb
Fig. 4.5. Primitive mantle-normalized trace element patterns for average compositions of basaltic andesite (yellow), andesite (green), dacite (blue), and rhyolite (red) from the Salla Group, Salla area. Data from Manninen (1991).
Geochemistry and Nd isotopes The chemical composition of the Salla Group metavolcanic rocks varies from basaltic andesite to rhyolite with the bulk of the SiO2 contents falling in the range of 54–70 wt.%. Diagrams such as the AFM or Jensen cation plot (not shown) place andesitic rocks in the calc-alkaline field probably due to secondary processes, while dacites and more evolved rocks display tholeiitic affinity (Manninen, 1991). All the members of the observed unimodal rock series are low in TiO2 and were probably derived from a common parental magma through fractional crystallization, as deduced from incompatible trace elements distributions. In Figure 4.5, primitive mantlenormalized trace element patterns are shown for the average compositions of basaltic andesite, andesite, dacite, and rhyolite from the Salla area. The patterns are similar for the whole series apart from the relatively low levels of Sr and Ti in the most evolved rocks, most likely resulting from plagioclase and CHAPTER
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oxide fractionation, respectively. An evident crustal signature is manifested by the strong LREE enrichment and deep negative Ta anomalies (Figure 4.5). Furthermore, these anomalies suggest that a significant amount of contamination with crustal material took place before the extensive fractional crystallization that produced the basalt–rhyolite series. Two samples of felsic metavolcanic rocks from the Salla area yielded εNd(at 2440 Ma) values of –4.4 and –3.1 (Table 4.1, Figure 4.7) which are consistent with the contaminated character of the magma as indicated by the geochemistry.
Geochronology Confident dating results have recently been obtained for felsic metavolcanic rocks from the eastern and western side of the Koitelainen intrusion, yielding U-Pb zircon ages of 2438 ± 14 and 2438 ± 11 Ma, respectively (Räsänen and Huhma, 2001; Manninen et al., 2001). These zircon ages are within the error of the ages reported for the large mafic layered intru-
LAPLAND
GREENSTONE
B E LT
•
149
MÖYKKELMÄ 250 m
Komatiitic basalt Andesite Komatiitic basalt
Basaltic andesite N
Basalt S
Drillhole R6
Quartzite (Sodankylä Group)
Archean granite gneiss basement
Fig. 4.6. Block diagram showing mafic to ultramafic metavolcanic rocks of the Onkamo Group on Archean basement at the southwestern corner of the Möykkelmä dome, ~30 km north of Sodankylä (M in Figure 4.2).
sions, Koitelainen and Akanvaara (Mutanen and Huhma, 2001). However, felsic metavolcanic rocks of the Salla Group are also found among the country rocks of these intrusions, demonstrating that at least part of these metavolcanic rocks are older than the intrusions (Räsänen and Huhma, 2001). On the Russian side, the equivalent rocks of the Salla greenstone area have been regarded as belonging to the Ludikovian deposits and this would have a post-Jatulian age (e.g., Kulikov et al. 1980; Radchenko et al., 1994). However, a recently obtained U-Pb zircon age for a dike rock (Onkamonlehto dike) demonstrates that the Salla Group is pre-Jatulian. Although not very robust, this age determination gives a minimum age of ~2380 Ma for the intermediate to felsic volcanic country rocks of the dike (Manninen and Huhma, 2001).
3.4. Onkamo Group Supracrustal rocks of the Onkamo Group are widely distributed, yet discontinuous in central Lapland (Figure 4.2). Our information on the Onkamo Group comes primarily from the meta150
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volcanic rocks studied in the Salla area at the Finnish-Russian border (previously called the Mäntyvaara Formation by Manninen, 1991), at Möykkelmä in the Sodankylä area (Räsänen et al., 1989), and at Teuravuoma in the Kolari area (Väänänen, 1989). The Onkamo Group rocks were deposited either on the lavas of the Salla Group or directly on Archean gneisses and granites. The former is true in the Salla area and allowed Manninen (1991) to study the contact relationship between the Salla and Onkamo Groups. He observed an Al-rich paleoweathering crust a few meters thick on the rhyolitic rocks of the Salla Group, grading upwards to a ~10 m thick sericite schist unit with a conglomeratic lower part. The Onkamo Group rocks lie directly on Archean basement at Möykkelmä in the municipality of Sodankylä (Räsänen et al., 1989). Here a 250-m-thick section of metavolcanic rocks has been penetrated by a drill hole (Figure 4.6). The metavolcanic rocks flank a small Archean basement dome (Möykkelmä dome) and they are overlain by quartzites of the next Paleoproterozoic unit, the Sodankylä Group (Figure 4.6). The Onkamo Group rocks mostly comprise subaerial, amygdaloidal mafic to intermediate lavas and fragmental ultramafic rocks, originally agglomerates and tuffs. In the Salla area, variolitic pillow lavas and pillow breccias are also found. At Möykkelmä, mafic and ultramafic units alternate (Figure 4.6). The sequence starts with pyroclastic komatiitic eruptions and andesitic lavas containing partly digested gneiss xenoliths from the basement (Figures 4.4B). These are followed by another pulse of ultramafic metavolcanic rocks overlain first by highly amygdaloidal basaltic andesites and then basalts. The mafic metavolcanic rocks thus become more primitive upwards in the section.
Geochemistry and Nd isotopes The Onkamo Group rocks are komatiites, komatiitic basalts, high-magnesian basalts,
LAPLAND
GREENSTONE
B E LT
Table 4.1. Sm-Nd isotope data for volcanic and plutonic rocks from Central Finnish Lapland. Sm Nd (ppm) (ppm)
147
Sm/144Nd (±0.4%)
143
Nd/144Nd (±2)
T* (Ma)
εNd(T) TDM (Ma)
Salla Group felsic metavolcanic rocks A1404 Koutoiva Felsic tuff A1454 Mäntyvaara Felsic tuff
7.89 6.16
37.00 30.57
0.1272 0.1221
0.511300±10 0.511285±10
2440 2440
-4.4 -3.1
3101 2944
Onkamo Group mafic metavolcanic rocks 11735 Mäntyvaara Pillow lava 11736 Mäntyvaara Pillow lava 11739 Mäntyvaara Pillow lava 12781 Mäntyvaara Komat. basalt 12782 Mäntyvaara Komat. basalt 12788 Mäntyvaara Komat. basalt 12789 Mäntyvaara Komat. basalt 12790 Mäntyvaara Komat. basalt 12920 Mäntyvaara Komatiite A1435 Möykkelmä Andesitic lava
3.03 2.04 2.17 1.72 1.37 1.67 2.18 1.91 1.21 3.04
15.45 9.99 10.76 7.01 4.25 8.00 10.67 9.42 5.29 16.49
0.1185 0.1235 0.1219 0.1481 0.1943 0.1265 0.1233 0.1226 0.1381 0.1115
0.511285±11 0.511327±10 0.511270±11 0.511800±10 0.512147±30 0.511414±10 0.511311±11 0.511327±10 0.511690±12 0.511106±10
2440 2440 2440 2440 2440 2440 2440 2440 2440 2440
-1.9 -2.7 -3.3 -1.2 -9.0 -1.9 -2.9 -2.4 -0.2 -3.2
2831 2922 2967 2925 2869 2941 2891 2747 2902
Kittilä Group, Kautoselkä Formation Tarv I Tarvasenvaara Plag-porphyrite Tarv I Tarvasenvaara Plagioclase 6579 Tarvasenvaara Amygd. lava 6583 Tarvasenvaara Mafic lava 5963 Kivipurnuvaara Mafic lava A1449 Karjakko-oja Mafic dike 11774 Nuttio Mafic lava
8.29 0.40 7.99 8.76 6.33 5.34 6.08
43.89 2.12 41.05 44.90 31.80 25.79 24.93
0.1142 0.1126 0.1177 0.1179 0.1204 0.1253 0.1473
0.511463±10 0.511449±21 0.511535±10 0.511536±13 0.511588±13 0.511655±10 0.512056±10
2015 2015 2015 2015 2015 2015 2015
-1.6 -1.5 -1.1 -1.2 -0.8 -0.8 1.4
2423 2406 2396 2399 2381 2397 2265
Kittilä Group,Vesmajärvi Formation 5155 Järvikäinen Mafic lava 13-17-LVP-84 Järvikäinen Mafic lava 13-17-LVP-84 Järvikäinen Mafic lava 13-17-LVP-84 Järvikäinen Clinopyroxene 13-17-LVP-84 Järvikäinen Clinopyroxene 13-17-LVP-84 Järvikäinen Clinopyroxene 13-18-LVP-84 Järvikäinen Mafic lava 13-18-LVP-84 Järvikäinen Mafic lava 13-18-LVP-84 Järvikäinen Clinopyroxene** 13-18-LVP-84 Järvikäinen Clinopyroxene** 5489 Järvikäinen Pillow lava A1200 Kiuasautonoja Mafic dike 5333 Penikkajärvi Pillow lava 84/20-1 Rajala Mafic lava 84/21-1 Rajala Mafic lava
3.34 1.95 2.15 1.10 1.10 1.11 2.42 2.19 2.76 1.75 2.67 1.81 2.35 3.34 4.39
10.34 5.99 7.12 2.17 2.20 2.22 7.41 6.14 9.41 4.75 7.89 7.06 6.68 10.13 13.67
0.1954 0.1971 0.1822 0.3056 0.3043 0.3032 0.1972 0.2159 0.1773 0.2234 0.2043 0.1547 0.2130 0.1992 0.1944
0.512836±12 0.512840±11 0.512617±10 0.514245±13 0.514267±85 0.514221±10 0.512838±15 0.513076±10 0.512608±12 0.513154±21 0.512949±10 0.512277±12 0.513046±11 0.512856±10 0.512761±10
2015 2015 2015 2015 2015 2015 2015 2015 2015 2015 2015 2015 2015 2015 2015
4.2 3.8 3.3 3.1 3.9 3.3 3.7 3.5 4.4 3.1 4.1 3.8 3.7 3.6 2.9
Kittilä Group,Veikasenmaa Formation 7412 Kiimarova Pillow lava 7472 Kiimarova Pillow lava 9771 Veikasenmaa Pillow lava 9782 Veikasenmaa Mafic porphyrite 9701 Veikasenmaa Mafic lava A1563 Selkäsenvuoma Mafic dike
3.33 3.15 6.12 5.25 2.50 1.14
10.77 9.92 19.47 19.03 8.56 3.99
0.1867 0.1921 0.1899 0.1667 0.1767 0.1731
0.512684±11 0.512730±11 0.512776±10 0.512417±12 0.512578±11 0.512443±19
2015 2015 2015 2015 2015 2015
3.4 2.9 4.4 3.4 4.0 2.2
Kittilä Group, Köngäs Formation 10220 Köngäs Pillow lava 10225 Köngäs Pillow lava
9.46 43.11 12.66 54.37
0.1327 0.1408
0.511910±11 0.512035±12
2015 2015
2.3 2.7
Sample
Location
CHAPTER
Rock type/ Mineral
4
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GREENSTONE
B E LT
1978
2133 2109
•
151
Sample
Location
Rock type/ Mineral
Sm (ppm)
Nd (ppm)
147
Sm/144Nd (±0.4%)
143
Nd/144Nd (±2)
T* (Ma)
εNd(T) TDM (Ma)
Calc-alkaline dikes in Nuttio serpentinites R4/9.40 Nolppio Mafic dike R4/18.90 Nolppio Mafic dike Kittilä Group felsic rocks A280 Kapsajoki Felsic porphyry A581 Veikasenmaa Felsic porphyry A893 Kiimaselkä Felsic porphyry A887 Yräjärvi Felsic porphyry
5.76 5.64
33.09 32.63
0.1053 0.1046
0.511553±11 0.511525±10
2015 2015
2.4 2.1
2093 2119
28.71 20.51 6.26 10.95
121.0 93.74 26.79 48.70
0.1434 0.1323 0.1412 0.1359
0.512126±11 0.511982±12 0.512094±10 0.512028±12
2015 2015 2015 2015
3.8 3.8 3.7 3.8
1987 1985 1994 1983
Felsic dikes and plutons (<2000 Ma) A246 Nyssäkoski Felsic porphyry A1262 Kuotko Felsic porphyry A1206 Ruoppapalo Granodiorite A746b Haisuvuoma Monzonite
2.79 3.07 1.29 5.71
16.57 18.98 7.11 30.48
0.1016 0.0978 0.1097 0.1132
0.511395±25 0.511420±10 0.511468±14 0.511446±10
1919 1920 1914 1880
-0.9 0.5 -1.5 -3.2
2240 2125 2309 2419
Lainio Group felsic metavolcanic rocks A513 Latvajärvi Felsic metavolc.
6.69
47.35
0.0854
0.511087±14
1880
-3.5
2322
Clasts in Lainio and Kumpu Group conglomerates A939 Linkujoki Felsic porphyry A1470 Mantovaara Felsic porphyry 105.2-PPR-94 Hangasoja Felsic porphyry A1416a Vesikkovaara Felsic porphyry A1416b Vesikkovaara Felsic porphyry A635b Kellostapuli Granite
12.08 0.95 6.71 1.93 1.58 0.70
57.13 5.30 42.65 13.34 9.56 2.42
0.1278 0.1080 0.0950 0.0875 0.1000 0.1747
0.511901±10 0.511582±12 0.511199±10 0.511035±11 0.511190±38 0.512202±21
2015 1928 1880 1880 1880 1880
3.4 1.3 -3.6 -5.0 -5.0 -3.3
2019 2105 2364 2420 2482
For methods see Hanski et al. (2001a). Measurements were made on VG Sector 54 mass spectrometer at the Geological Survey of Finland. 143Nd/144Nd ratio is normalized to 146Nd/144Nd=0.7219 and error is 2 standard error of the mean in the last significant digits. Measurements on the La Jolla standard yielded a 143Nd/144Nd ratio of 0.511851±6 (standard deviation, n=48). Typical error in εNd is 0.4 units. TDM is calculated according to DePaolo (1981) for rocks with significant LREE enriched pattern. * The age used for calculating εwd is based on U-Pb zircon dating (underlined) and geological correlations (see text). ** Clinopyroxene fraction altered and turbid, not included in age calculation (see Figure 4.13).
basalts, basaltic andesites, and andesites and are thus, on average, more varying and more primitive than the Salla Group rocks. The mafic to intermediate rocks are either tholeiitic or calc-alkaline with their calc-alkaline nature most probably caused by interaction with sialic crust and postmagmatic hydrothermal alteration. The MgO content of the komatiitic rocks rarely exceeds 20 wt.%, typically falling between 14 wt.% and 18 wt.%. Primitive basaltic rocks have a relatively high SiO2 and can be called siliceous high-magnesian basalts (SHMB). In common with the Salla Group rocks, the Onkamo Group metavolcanic rocks display a strong signature of crustal contamination: high LREE/HREE and negative HFSE 152
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anomalies in primitive mantle-normalized trace element plots (Räsänen et al., 1989). In komatiitic rocks, (La/Yb)N approaches 5, and they differ in this respect from typical Archean komatiites in Finland or younger komatiitic rocks of the Savukoski Group in central Lapland (see below). In contrast to the Salla Group, in which the mafic to felsic volcanic rocks form a cogenetic suite (Figure 4.5), at Möykkelmä the basaltic member seems to have evolved from a different parental magma than the andesites and dacites (Räsänen et al., 1989). A special feature of the Möykkelmä basalts is their exceptionally high Cr content (700–1100 ppm). A major contribution from Archean LREE-
LAPLAND
GREENSTONE
B E LT
Vesmajärvi Formation mafic rocks Veikasenmaa porphyries
Jouttiaapa Jeesiörova
4
Depleted mantle
Pechenga tholeites
Nu
ttio
2
Jeesiörova
Link u Jormua joki po rph EMORB yry Mantovaara
dik es
Pechenga ferropicrites 0
0
Runkausvaara
εNd
Jormua OIB
Kau
tos
Penikat Koitelainen Akanvaara
Gro
afic
Gro
up f elsic
äF
m.
Nyssäkoski porphyry Ruoppapalo
ma
fic
-2
roc k
s
up m
Salla
elk
Onk
amo
-4
2
Haaskalehto
CHUR
-2
4
roc Keivitsa ks
Haisuvuorna Kellostapuli Latvajärvi Hangasoja
roc k
s
-4
Vesikkovaara -6
-6 Granites Field of Archean granitoids
-8
-8 2450
2350
2250
2150
2050
1950
1850
1750
Age (Ma) Komatiite, picrite
Postorogenic granite
Mafic volcanic rock
Monzonite, granite (1880 Ma)
Layered intrusion (2050 Ma)
Granodiorite
Gabbro–wehrlite association intrusion (2200 Ma)
Felsic porphyry (2015–1880 Ma)
Layered intrusion (2440 Ma) Fig. 4.7. Nd isotope evolution diagram for Paleoproterozoic plutonic and volcanic rocks of northern Finland. Initial εNd values are shown for rocks that have been dated by the U-Pb or Sm-Nd method. Error bars are based on several analyses on each rock association. Evolution lines are shown for rocks lacking direct isotope datings: mafic to ultramafic metavolcanic rocks from the Onkamo Group (green), felsic metavolcanic rocks from the Salla Group (purple), mafic metavolcanic rocks from the Kautoselkä Formation (Kittilä Group, blue), dikes cutting Nuttio serpentinites (mustard green), and a felsic porphyry clast (Linkupalo) in Lainio Group conglomerate (red). Also shown are data on metavolcanic rocks from Jormua, Pechenga, and Peräpohja (Jouttiaapa) for comparison. Data from this study, Huhma (1986), Huhma et al. (1990, 1995), Hanski (1992), Hanski et al. (1990, 2001a, c), and Peltonen et al. (1996, 1998). The field formed by trajectories of Archean granitoids is based on Huhma (1986), O’Brien et al. (1993), Hölttä et al. (2000), and Hanski et al. (2001c). Evolution of depleted model upper mantle after DePaolo (1981). CHAPTER
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enriched lithosphere is also evident in the Nd isotope composition of the magnesian lavas of the Onkamo Group (Table 4.1, Figure 4.7). Eight samples analysed from the Salla area have an average Nd (at 2440 Ma) value of –2.2 and a TDM of ~2900 Ma (one analysis gave an Nd of –9 and suggests strong metamorphic fractionation at ~1800–1900 Ma). A single analysis of an andesite from the Möykkelmä area is consistent with the data from Salla, yielding an Nd (at 2440 Ma) value of –3.2.
Geochronology The metalavas rocks of the Onkamo Group have so far not been directly dated. However, on the basis of their geological position and geochemical composition, they can be correlated with rocks in the Vetreny belt, southeastern Russian Karelia. The Vetreny suite komatiites have yielded Sm-Nd isochron ages of 2449 ± 35 and 2410 ± 34 Ma and a U-Pb zircon age of 2437 ± 3 Ma has been obtained for a dacite (Puchtel et al., 1997). The Salla Group is cut by ~2440 Ma, chromitite-bearing mafic layered intrusions at Akanvaara and Koitelainen, but such a relationship has not been observed for the Onkamo Group, thus permitting contemporaneity of these intrusions and the Onkamo Group metavolcanic rocks. In any case, it seems that the Salla and Onkamo Groups do not deviate much in age from each other and, broadly speaking, belong to the same, short-lived magmatic episode that produced large volumes of igneous rocks, layered intrusions, dike swarms, subaerial or shallow-water lavas, over large areas. They thus exhibit the characteristic features of magmatism related to an incipient mantle plume (cf. Ernst and Buchan, 1997). In the Russian part of the shield, corresponding rocks have been assigned to the Sumi–Sariolian stage of magmatism (e.g., Gaskelberg et al., 1986; Zagorodny et al., 1986).
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3.5. Sodankylä Group Active volcanism of the Onkamo Group was followed by a more tranquil period resulting in a thick, epiclastic sedimentary sequence. These sediments are represented by the quartzites and mica schists of the Sodankylä Group, which are found mainly on the eastern and southern side of the Kittilä greenstone area (Figure 4.2). The Sodankylä Group metasediments were deposited either on Archean gneissose basement or volcanic rocks of the Salla or Onkamo Groups. Besides orthoquartzites, sericite quartzites, and mica schists, the Sodankylä Group comprises minor carbonate rocks with occasional stromatolitic structures and mafic metavolcanic rocks. Dolomites from the Sodankylä Group have an anomalous ␦13C value of +6, which is typical of the Jatulian carbonate rocks (Karhu, 1993). Cr-bearing mica, fuchsite, is a characteristic mineral of the sericite quartzites, often staining the rocks with different tints of green (Figure 4.4C). The abundance and distribution of quartzites suggest that the depositional basin was markedly widened from a relatively narrow rift basin after cessation of the Salla and Onkamo Group volcanism. Primary structures in the metasediments, well exposed at Virttiövaara for example, include cross-bedding, graded bedding, herringbone structures, and mud cracks, indicating a tidal environment (Nikula, 1988). Volcanic rocks of the Sodankylä Group are mainly represented by amygdaloidal mafic lavas forming units some tens of meters thick. One of them is the northwestern extension of the Greenstone III Formation of Silvennoinen (1972), which can be followed from the Kuusamo belt to the Pyhätunturi area (Figure 4.2) in the eastern part of the Central Lapland greenstone belt (Räsänen and Huhma, 2001).
LAPLAND
GREENSTONE
B E LT
10
. Kittilä Group
Th/TiO2 = 3 8
Savukoski Group Sodankylä Group
6
Th (ppm)
Onkamo Group Salla Group
4
2
0 0
1
2
3
4
TiO2 (wt. %) Fig. 4.8. Th vs. TiO2 diagram for mafic to intermediate metavolcanic rocks from the Central Lapland greenstone belt.
Geochemistry and geochronology The mafic metavolcanic rocks of the Sodankylä Group are tholeiitic basalts and basaltic andesites with slightly LREE-enriched chondrite-normalized REE patterns. They differ from the older metavolcanic rocks of the Onkamo and Salla Groups in having a less pronounced crustal signature as demonstrated by the Th vs. TiO2 plot (Figure 4.8). Also, felsic rocks interpreted as volcanic in origin have been encountered within the Sodankylä Group around the Central Lapland granitoid complex, but their interpretation is problematic as they have yielded Archean zircon ages (Rastas et al., 2001; Räsänen and Huhma, 2001). Mafic magmatism is also manifested by ~2200 Ma hypabyssal intrusions forming concordant, differentiated mafic–ultramafic sills CHAPTER
4
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within the quartzites of the Sodankylä Group and thus yield a minimum age for the metasediments. Intrusions of this age group have not been seen to cut the overlying lithostratigraphic groups (Savukoski, Kittilä, Lainio, Kumpu), which is consistent with isotopic and other evidence indicating that these units did not exist at the time of the intrusion of the ~2200 Ma sills (see below). Isotope analyses of bulk detrital zircon samples from Sodankylä Group quartzites indicate a predominant Archean provenance for these rocks (Rastas et al., 2001). In this respect, the Sodankylä Group quartzites are similar to the “Jatulian” Kivalo Group quartzites in the Peräpohja belt (Perttunen and Vaasjoki, 2001), but differ clearly from the sedimentary rocks of the Lainio and Kumpu
LAPLAND
GREENSTONE
B E LT
•
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Groups; the latter have a significant Paleoproterozoic source component (Rastas et al., 2001; Hanski et al., 2001b).
3.6. Savukoski Group Fine-grained metasedimentary rocks of the Savukoski Group, including phyllites and black schists, and some mafic tuffites, overlie the clastic metasedimentary rocks of the Sodankylä Group. No significant discordance or hiatus has been observed at the contact of these two groups, indicating gradual deepening of the depositional basin. The Savukoski Group contains the first manifestations of graphite- and sulfide-bearing schists in the Paleoproterozoic stratigraphy of the Central Lapland greenstone belt, and provide an important key horizon. A few mafic tuff and tuffite interbeds with locally abundant concretions are found within phyllites. They are strongly magnetic and hence form useful marker horizons on geophysical maps. Also mafic lavas are locally present within pelitic metasediments. These are well-exposed at Linkupalo, east of Kittilä, where variolitic pillow lavas, massive lavas, volcanic breccias, agglomerates, tuffites, and associated concordant diabases and gabbros have been described (Lehtonen et al., 1998). The pelitic metasediments are overlain by primitive volcanic rocks with komatiitic and picritic chemical affinities. These rocks have been assigned to the komatiite–picrite association and can be followed for more than 350 km from the northern margin of the Salla greenstone area in the east towards the northwest across the border to northern Norway (Saverikko, 1983, 1985; Hanski et al., 2001a; Henriksen, 1983; Barnes and Often, 1990). The komatiitic and picritic metavolcanic rocks cannot be distinguished from each other in the field, but chemical analyses show that there are areas in which either of these two rock types predominates. Abundant komatiitic rocks are present at Sattasvaara in Sodankylä, 156
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Jeesiörova in Kittilä and Kummitsoiva in Savukoski, while picritic rocks characterize the Sotkaselkä area (see Figures 4.2 and 4.3). Also more fractionated basaltic metalavas, commonly with pillow structures, are associated with the highly magnesian varieties. Though mineralogically heavily altered, the magnesian metavolcanic rocks exhibit well-preserved primary volcanological structures, particularly in the Sattasvaara area (Figure 4.4D, E). Both the komatiitic and picritic rocks occur as pillowed and massive lavas, pillow breccias, and various kinds of volcaniclastic rocks including agglomerates, lapilli tuffs, and reworked tuffs. Thick, differentiated lava flows are rare (Räsänen, 1996). The subaqueous komatiite–picrite association of the Savukoski Group differs from the older subaerial komatiites of the Onkamo Group in their environment of eruption, though fragmental komatiitic rocks are common in both rock suites. In most strongly altered regions, especially close to the Sirkka tectonic line at the southern margin of the Kittilä greenstone area, komatiitic rocks have been affected by intense carbonatization and are now deep-green chromian marbles. These are good targets for gold prospecting (Sorjonen-Ward et al., 1992) and colorful enough for decorative purposes.
Geochemistry The Savukoski Group ultramafic and mafic metavolcanic rocks form geochemically a diverse group of rocks. They can be classified as belonging to two series both of which have evolved to basaltic rocks from a magnesian parental magma with MgO close to or higher than 20 wt.%. On an [Al2O3] vs. [TiO2] diagram, the primitive magmas and moderately evolved members of these series fall in the fields of Ti-enriched komatiites and picrites (see Figure 4.9). Elevated TiO2 contents in picrites result in low Al2O3/TiO2 ratios of 4 to 5. In komatiites this ratio normally ranges from 12 to 14 and is still relatively low compared
LAPLAND
GREENSTONE
B E LT
0.05
40 Picrite Al-depleted picrite
Picrite
0.03 t% )
[TiO2]
Ti-enriched, Al-depleted komatiite
10
20
0.00 0.00
0.05
Ti-enriched komatiite
Al-undepleted
Al-depleted komatiite
3
0.01
Al 2O
3 /T iO 2 = 5
(W
0.02
Sample/Chondrite
0.04
10
Komatiite
Ti-depleted komatiite, boninite
0.10
0.15
0.20
0.25
[Al2O3]
1
La Ce Pr Nd
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Fig. 4.9. Analyses of komatiitic and picritic metavolcanic rocks of the Savukoski Group plotted on an [Al2O3] vs. [TiO2] diagram of Hanski et al. (2001a). [Al2O3] and [TiO2] are Al2O3 and TiO2 projected from olivine composition and are calculated in mole proportions (normalized to unity) using the following equations: [Al2O3] = Al2O3/(2/3 - MgO - FeO) and [TiO2] = TiO2/(2/3 - MgO - FeO) (see Hanski, 1992, p. 106).
Fig. 4.10. Examples of chondrite-normalized REE spectra for Savukoski Group ultramafic metavolcanic rocks that vary from LREE-depleted komatiites to LREE-enriched picrites.
with Al-undepleted komatiites as defined by Nesbitt et al. (1979). The two magma types differ markedly from each other in their rare earth element and high-field strength element contents. The komatiites are strongly LREEdepleted and moderately HREE-depleted displaying hump-shaped chondrite-normalized REE patterns, whereas the picrites are strongly enriched in LREE and all HFSE. There are also intermediate types having REE patterns between those of typical komatiites and picrites (Figure 4.10). Given the strong depletion of the komatiites in highly incompatible elements, their moderate enrichment in Ti is a peculiar feature and sets them apart from classic LREE-depleted komatiites, such as those in Munro Township, Canada (Hanski et al., 2001a).
of the Savukoski Group pelitic metasediments is provided by crosscutting intrusive bodies, such as the Keivitsa layered intrusion. Two samples from this intrusion have yielded UPb zircon ages of 2058 ± 4 Ma and 2054 ± 5 Ma (Mutanen and Huhma, 2001). Quartz porphyries ~10 km southwest of the Keivitsa intrusion have a comparable age (2048 ± 5 Ma; Mutanen and Huhma, 2001). Also diabase dikes that cut the Savukoski Group metasediments have yielded U-Pb zircon ages of 2060 Ma to 2050 Ma (Rastas et al., 2001; Räsänen and Huhma, 2001). Preservation of primary clinopyroxene in the komatiitic rocks at Jeesiörova in the Kittilä area allowed Hanski et al. (2001a) to utilize pyroxene–whole-rock pairs to determine the direct Sm-Nd isotope age for these rocks. The result, 2056 ± 25 Ma, is comparable with the Sm-Nd age of 2085 ± 85 Ma published earlier for komatiites from northern Norway (Krill
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et al., 1985). The εNd (at 2060 Ma) values for komatiitic rocks range from +2 to +4 (Figure 4.7), with the most LREE-depleted, least contaminated komatiites having the highest εNd. Picritic samples have also yielded positive εNd (at 2060 Ma) values around +4. The available geochemical and isotope data also indicate that the komatiites and picrites were generated from a geochemically heterogeneous but isotopically similarly depleted mantle source. Hanski et al. (2001a) suggested that the heterogeneities in the source region were created by complex depletion and enrichment processes shortly prior or related to dynamic melting in an adiabatically ascending mantle source. The isotope and chemical data show that the effect of sialic crustal contamination remained negligible in the komatiites and picrites of the Savukoski Group. In this respect, they differ profoundly from the komatiites of the Onkamo Group (Räsänen et al., 1989).
3.7. Kittilä Group The Kittilä Group represents one of the largest accumulations of mafic metavolcanic rocks in the Fennoscandian Shield. It forms a single terrane, often called the Kittilä greenstone complex, and covers an area of more than 2600 km2 in the central part of the Central Lapland greenstone belt (see Figure 4.2). On the basis of geophysical measurements, the present day vertical thickness of the folded volcanic pile has been estimated to reach 6 km, giving an indication of the magnitude of the volcanism of this complex, as opposed to the relatively thin volcanic intercalations in the Sodankylä Group, for example. When followed from the Kittilä area to the north, the wide, relatively flat-flying and weakly deformed volcanic complex becomes a narrow, steeply E-dipping, tectonized zone. This is probably related to the overthrusting of the Lapland granulite belt from the northeast. Consequently, the rocks of the Kittilä Group almost wedge out before reaching the Finnish–Norwegian border 158
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(Figure 4.1). The depositional basement of the Kittilä Group has not been confidently established. The contacts of the Kittilä Group rocks with the other lithostratigraphic units described above seem to be tectonic. For instance, a thrust zone runs close to the eastern margin of the complex separating the volcanic rocks of the Kittilä Group from the Archean Sovasjoki gneiss dome and its autochthonous Paleoproterozoic cover (Figure 4.3). The southern contact of the Kittilä greenstone area coincides with the roughly E–W-trending fault zone, the “Sirkka line” (Gaál et al., 1989). The Kittilä greenstone area is dominated by metavolcanic rocks, but there are also several types of sedimentary interbeds and larger sedimentary units. These comprise metagraywackes, phyllites, graphite- and sulfide-bearing schists, and minor carbonate rocks. Most notable are, however, banded iron-formations assigned to the Porkonen Formation. Oxide facies rocks predominate, but there are also sulfide- and silicate-carbonate facies precipitates including manganosiderites (Paakkola and Gehör, 1988).
Stable isotopes Unlike the carbonate rocks in the cratonic Sodankylä Group, those in the Kittilä Group do not have anomalously high ␦13C values (Karhu, 1993). Sulfur isotope data of the black schists of the Savukoski Group in the vicinity of the Keivitsa mafic intrusion display elevated ␦34S values mostly in the range of +15% to +20% (Hanski et al., 1996). As shown by Figure 4.11, isotope data of the black schists from one locality representing the Kittilä Group record very different, close to chondritic ␦34S values (–3 to +4; Heikki Pankka, unpublished data). The available ␦34S data obtained for the Kittilä Group metasediments are thus in line with the oceanic geotectonic interpretation suggested for this unit (see below).
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Frequency
8
0 -5
0
5
Black schists:
10
␦34S
Kittilä Group
15
20
25
30
Savukoski Group
Fig. 4.11. Sulfur isotope composition of sulfides in black schists from the Savukoski and Kittilä Groups (data from Hanski et al., 1996, and Heikki Pankka, unpublished).
Field characteristics and geochemistry of mafic metavolcanic rocks The Kittilä Group contains several generations of genetically unrelated metavolcanic rocks. Lehtonen et al. (1998) described two types of formations, the Vesmajärvi and Kautoselkä Formations, which differ in their mode of occurrence and geochemistry. The Vesmajärvi Formation comprises various kinds of submarine mafic metavolcanic rocks, including pillow lavas, pillow breccias, and hyaloclastites, and cogenetic dikes and gabbroic sills. Locally, narrow mafic dikes form subparallel sets with interdike screens of metagabbro. These are interpreted as incipient sheeted dike complexes, feeder channels for the associated pillow lavas (Lehtonen et al., 1998). Also minor felsic rocks are encountered amidst the mafic metavolcanic rocks. Mafic amygdaloidal lavas, tuffs, and tuffites are the main rock types of the Kautoselkä Formation, which is found in the southeastern part of the greenstone complex. The volcanic structures of the Vesmajärvi Formation indicate submarine eruptions, which is consistent with the rocks having a chemical affinity of oceanic basalts. Geochemically, they are tholeiitic basalts with Mg# [atomic Fe/(Mg+Fe)] between 0.72 and 0.43. CHAPTER
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Rare examples of high-Mg lavas have been encountered, but they are basalts enriched in mafic phenocrysts rather than rocks crystallized from a komatiitic magma. The Vesmajärvi metavolcanic rocks have a moderate TiO2 content and their chondrite-normalized REE patterns are usually slightly LREE-enriched, but also flat or slightly LREE-depleted patterns occur (Figure 4.12A). The closest modern analogues are EMORBs, while NMORB types are subordinate (Figure 4.12B). No indication of sialic contamination is apparent in the chemistry of these lavas. The mafic metavolcanic rocks of the Kautoselkä Formation are relatively evolved, Fe-rich tholeiitic basalts and andesites with FeOtot and Mg# averaging 13.5 wt.% and 0.43 wt.%, respectively (Lehtonen et al., 1998). They have high LREE/HREE (red dots in Figure 4.12A) with (La/Yb)N ~10 and are enriched in incompatible elements including TiO2. They plot in the field of within plate basalts in trace element discrimination diagrams such as Zr/Y vs. Zr, Zr-Ti-Y, and Zr-Nb-Y. However, Nb and Ta are less enriched compared with LREE, Th, and U and, consequently, the rocks display negative Nb-Ta anomalies in primitive mantle-normalized spidergrams and straddle the transition zone between within plate basalts, E-MORBs, and volcanic arc basalts in
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Kittilä Group mafic metavolcanic rocks B
A
Hf/3
300
ic
NMORB
Vo Ca l lcalk canic alin arc e bas alt s Th ole iit
Sample/Chondrite
100
10
EMORB + Thol. WPB
Alk. WPB
3 La Ce Pr Nd
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Th
Nb/16
Fig. 4.12. (A) Chondrite-normalized REE patterns and (B) discrimination diagram of Wood et al. (1979) for mafic metavolcanic rocks from the Kittilä Group.
the Th-Hf-Ta and Th-Hf-Nb diagrams (red dots in Figure 4.12B). A subset of samples has a high P2O5 content (0.5–1.1 wt.%) as opposed to values less than 0.4 wt.% in most samples. The high-P samples also have high Ti/V ratios exceeding 100 in several cases. Such values have been observed in some oceanic alkali basalts (Shervais, 1982). There are also pillow lavas (see Figure 4.4F) in the western part of the complex (as exemplified by the Köngäs Formation) with high incompatible trace element concentrations. They display high LREE/HREE [(La/ Yb)N 3–6] and plot in the fields of within plate basalts or OIB (dark blue diamonds in Figure 4.12B). In addition, basaltic lavas chemically akin to present-day island arc tholeiites (IAT) have been found close to the eastern margin of the Kittilä greenstone area in association of serpentinite lenses (Hanski, 1997). They possess a low TiO2 (<0.70 wt.%) content and low LREE/HREE (dark blue dots in Figure 4.12A).
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Geochronology of mafic rocks We have analysed the Nd isotope composition of tholeiitic lava samples and mineral separates from the Vesmajärvi Formation. The samples were collected from outcrops on the western side of the Kumputunturi fell (see Figure 4.3A), containing pillow lavas and genetically related dikes. The analytical results are listed in Table 4.1 and plotted in Figure 4.13. Seven whole-rock and three pyroxene analyses yielded an isochron with an age of 1987 ± 36 Ma and an initial εNd value of +3.8 ± 0.3. The latter is close to the value of the contemporaneous model depleted mantle of DePaolo (1981). This and elemental geochemical observations demonstrate that the magma had insignificant or no interaction with an old sialic crust during its ascent or emplacement (Figure 4.7). We also analysed mafic metalavas from the western part of the Kittilä greenstone area, including EMORB- and OIB-like rocks of the Veikasenmaa and Köngäs Formations, respectively. At 2015 Ma, the former have positive εNd values (+2.9 to +4.4) similar to those obtained for the Vesmajärvi Formation,
LAPLAND
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143
Nd/144Nd
0.5140 0.5136
Vesmajärvi Fm. tholeiites T = 1990 ± 35 Ma εNd = +3.7 ± 0.2
0.5132 0.5128 0.5124 0.5120 .14
0.18
0.22
0.26
0.30
0.34
147
Sm/144Nd Clinopyroxene Whole rock Fig. 4.13. 143Nd/144Nd vs. 147Sm/144Nd diagram for whole-rock and mineral separates from tholeiitic basalts of the Vesmajärvi Formation (Kittilä Group).
while the Köngäs alkali basalts have slightly lower εNd values (+2.3 to +2.7) (Table 4.1, Figure 4.3). In contrast, the Kautoselkä Formation provides mostly negative εNd (at 2015 Ma) values (Figure 4.6). The origin of these LREE-enriched rocks is clearly distinct from the origin of the basalts of the Vesmajärvi Formation. We have also analysed one sample that is chemically and isotopically intermediate (εNd +1.4 at 2015 Ma) between the lavas of the Vesmäjärvi and Kautoselkä Formations (Table 4.1).
Nuttio serpentinites and related dikes At the eastern edge of the Kittilä volcanic complex (Figure 4.3), there is a N-trending chain of serpentinites and dunites designated as the Nuttio serpentinite belt (Hanski, 1997). Drilling and field observations and geophysical measurements have shown that the ultramafic bodies are in tectonic contact with their country rocks. They have usually been altered to serpentinites or soapstones, but locally CHAPTER
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well-preserved dunites are found containing strained olivines with high Fo contents, even up to 97%. Chemically, the ultramafic rocks are highly refractory with the Al2O3 and TiO2 contents commonly falling below 0.1 wt.% and 0.06 wt.%, respectively. They contain small chromitite lenses with chromite having a metallurgical composition (high Cr/Al). Preliminary analyses show that the Os isotope composition of the Nuttio chromitites falls within the narrow range measured for the chromitites from the Outokumpu ophiolite complex, yielding close to chondritic initial ratios at 2000 Ma (Hanski et al., 1995; Walker et al., 1996). Another feature common with the Outokumpu serpentinites is the presence of metasomatic Cr-rich quartz rocks adjacent to some of the Nuttio belt serpentinites (cf. Kontinen, 1998). The ultramafic bodies are cut by numerous mafic to ultramafic dikes with island arc tholeiitic, calc-alkaline, and boninitic affinities. The boninites have MgO contents up to 25 wt.% (in volatile-free analyses) and are extremely low in the incompatible elements Ti, Zr, REE, and Y. They are thus among the most magnesian and incompatible elementdepleted boninites ever discovered. The trace element characteristics of the three dike types differ clearly from each other (Figure 4.14), but all of them display a marked decoupling of LREE and HFSE abundances, a feature shared with modern island arc lavas. Analogous dikes have not been observed in the volcanic and sedimentary rocks outside the serpentinites, indicating that the intrusion of the dikes took place before the postulated tectonic emplacement of the serpentinites. All these features are consistent with the interpretation that the Nuttio serpentinites represent dismembered ophiolitic ultramafic rocks. Moreover, the chemical composition of chromites and dike rocks suggest a suprasubduction zone environment. At present, no direct age data are available for the rocks of the Nuttio serpentinite belt. Because the serpentinite lenses were emplaced
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200 100
Dikes cutting Nuttio serpentines
Sample/Primitive Mantle
Calc-alkaline
10
Tholeiitic 1
Boninitic 0.1 Th
Nb
La
Ce
Pr
Nd Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Fig. 4.14. Primitive mantle-normalized, extended REE diagram for dike rocks cutting Nuttio belt serpentinites.
tectonically, they can, in principle, be older, coeval or younger than their country rocks. We have determined the Nd isotope composition of two calc-alkaline dikes hosted by the serpentinites (see Table 4.1). Although precise dating cannot be achieved by these data, some inferences can be made due to the low Sm/Nd ratio of the samples and resultant steep growth curves on an εNd vs. age diagram (Figure 4.7). The obtained TDM ages are ~2100 Ma for both dike samples (Table 4.1) and can be taken as a maximum age of crystallization. This suggests that the host ophiolitic ultramafic rocks are lower Paleoproterozoic in age and thus consistent with the interpretation that the Nuttio serpentinites are penecontemporaneous with the metavolcanic rocks of the ~2000 Ma Kittilä Group. Hence, the age of the Nuttio belt ultramafic rocks could be close to, but not necessarily the same as, the age of the gabbros from the Jormua and Outokumpu ophiolite 162
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complexes in eastern Finland (1970–1950 Ma; Huhma, 1986; Peltonen et al., 1998). As shown by Figure 4.7, the trajectories of the Nuttio dike samples pass close to the initial εNd values of the E-MORBs from Jormua and the Pechenga ferropicrites.
Felsic rocks In the western part of the Kittilä greenstone area, volumetrically minor but geotectonically significant felsic porphyries are intimately associated with mafic lavas and dikes. These are plagioclase- and quartz-phyric rocks (see Figure 4.4G) and occur as subvolcanic dikes, lavas, and crystal tuff layers up to 10 m in thickness. Fragments of felsic porphyry are found in the associated mafic metavolcanic rocks and diabase dikes and some of the porphyries are cut by apophyses injected from the diabases. The reverse is also seen, as fragments of a fine-grained mafic rock are found within
LAPLAND
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Felsic porphyries from the Kittilä area 0.38
Latvajärvi Formation Samples A513, A745 Intercepts at 1880 ± 8 Ma & 123 ± 240 Ma MSWD = 2.2, n = 5
2020 1980
0.36
Kittilä Group porphyries Samples A280, A581, A887, A893 Intercepts at 2015 ± 2 Ma & 339 ± 50 Ma MSWD = 0.42, n = 13
1900 0.34 1860
206
Pb/238U
1940
Nyssäkoski dike Sample A246 Intercepts at 1919 ± 8 Ma & 148 ± 350 Ma MSWD = 2.7, n = 5
1820 0.32
0.30 4.8
5.2
5.6 207
6.0
6.4
Pb/235U
Fig. 4.15. Concordia diagram showing isotope data on zircons from felsic porphyries from the Kittilä area (for analytical data, see Rastas et al., 2001).
felsic porphyries. These field relationships have been interpreted to demonstrate that the felsic porphyries and associated mafic volcanics are products of contemporaneous basaltic and silicic magmas (Lehtonen et al., 1998). The felsic rocks can therefore be utilized to determine the age for the mafic rocks of the Kittilä Group. The porphyries dated by Rastas et al. (2001) at four localities (Veikasenmaa, Kapsajoki, Kiimarova, and Yräjärvi) have mutually consistent U-Pb zircon ages between 2012 ± 5 Ma to 2018 ± 7 Ma and yield a combined age of 2015 ± 2 Ma (Figure 4.15). Within error, these ages overlap with the Sm-Nd isochron age of the Vesmajärvi Formation mafic metavolcanic rocks reported in this study. Geochemically, the felsic porphyries range from low-K dacites to high-silica rhyolites with high contents of REE, Th, Ta, and Zr. In the most evolved rocks, light REE and heavy REE reach levels up to 400 and 100 times chondritic, respectively, with relatively CHAPTER
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low (La/Yb)N indicative of low-pressure melt equilibration. On trace element discrimination diagrams designed for felsic rocks (Pearce et al., 1984; Gorton and Schandl, 2000), the porphyries plot in the ocean ridge or within plate fields. We have determined the whole-rock Nd isotope composition of the same porphyry samples that were utilized for zircon dating by Rastas et al. (2001). Given the highly evolved nature of these silica-rich rocks, it is interesting that all samples yielded positive initial εNd values of +3.8 (Table 4.1); these are indistinguishable from the values obtained for the most depleted tholeiitic mafic metavolcanic rocks of the Kittilä Group (Figure 4.7). This suggests that no sialic crust was involved in the generation of these felsic porphyries or that the crustal residence time of the sialic crust was very short. Therefore, the Archean basement was not the source of the porphyries and probably did not even exist beneath the porphyries and associated mafic rocks at
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B
200
500
Sample/Primitive Mantle
200 100
10 10
1 0.7
1 Rb Th Nb K Ce Sr Nd Zr Eu Gd Dy Er Lu Ba U Ta La Pr P Sm Hf Ti Tb Y Yb
Rb Th Nb K Ce Sr Nd Zr Eu Gd Dy Er Lu Ba U Ta La Pr P Sm Hf Ti Tb Y Yb
Ruoppapalo granodiorite
Latvajärvi Formation trachyte
Nyssäkoski felsic dike
Clast from Lainio Group conglomerate
Pebble from Kumpu Group conglomerate
Fig. 4.16. Primitive mantle-normalized trace element diagrams for felsic rocks from the Kittilä area. (A) ~1920 Ma felsic porphyries and Ruoppapalo granodiorite. (B) ~1880 Ma Latvajärvi Formation trachyte and a porphyry clast from Lainio Group conglomerate.
the time of their generation. The two most likely mechanisms to produce silicic igneous rocks with juvenile isotope compositions are extensive fractional crystallization of basaltic magma or dehydration and partial melting of mafic, amphibolitic oceanic crust. We have identified another felsic magmatic pulse of dikes and plutons cutting metavolcanic rocks of the Kittilä Group. These are ~100 Ma younger than the dikes described above and have turned out to be significant in providing age constraints to the (tectonic) emplacement of the Kittilä Group and deposition of the overlying Kumpu Group rocks (see below). The dikes are up to 10 m thick and contain plagioclase, quartz, amphibole, and biotite phenocrysts; hence they differ from the ~2015 Ma felsic porphyries, which contain only felsic phenocrysts. Compositionally these younger dikes are more restricted than the ~2015 Ma porphyries, plotting mostly in the fields of dacite and rhyodacite. Isotopically 164
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they are less radiogenic as their initial εNd values are around zero (Table 4.1). Rastas et al. (2001) published a U-Pb zircon age of 1919 ± 8 Ma for a felsic dike at Nyssäkoski (see Figure 4.15). Geochemical studies have shown that the Ruoppapalo granodiorite close to the northeastern margin of the Kittilä greenstone area (see Figure 4.3) is very similar to these dikes (Figure 4.16A). This granodiorite has a U-Pb zircon age of 1914 ± 3 Ma, matching well the age of the Nyssäkoski dike (Rastas et al., 2001). The initial εNd values for these rocks are also similar (–1.5, –0.9; Figure 4.7).
3.8. Lainio and Kumpu Groups Metasediments A major stratigraphic break already recognized by Hackman (1927) is found in the supracrustal sequence of the Central Lapland greenstone belt between the coarse-clastic, molasse-like metasediments of the Lainio
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and Kumpu Groups and the underlying older formations. The Lainio Group is located in restricted areas in the western part of the greenstone belt, where it forms fells such as Ylläs and Aakenustunturi on the western side of the Kittilä granite (Figure 4.3A), raising from the relatively flat landscape in the area of the older formations. The Kumpu Group rocks also form high hills or fells, including Levi, Mantovaara and Kumputunturi in the Kittilä region, Kaarestunturi in the Sodankylä region, and Pyhätunturi in Pelkosenniemi (Figures 4.2, 4.3A) (Räsänen et al., 1995). Lehtonen et al. (1998) distinguished the Lainio and Kumpu Groups on the basis of their deformation histories: in contrast to the Lainio Group, the Kumpu Group does not display effects of the earliest deformation phases of the Svecofennian orogeny. This observation allowed Lehtonen et al. (1998) to place the Kumpu Group higher in their lithostratigraphic scheme. However, it has now become evident that the deposition of both units took place at a post-1880 Ma stage of the development of the greenstone belt (see below) and therefore they are treated together in this work. The Lainio and Kumpu Groups form 200–2000 m thick sedimentary units, which comprise meta-arkoses, quartzites, polymictic conglomerates with graywacke interbeds, and siltstones with desiccation cracks. The Kumpu Group metasediments commonly display a characteristic red-brown or purple tint due to hematite pigment. The clasts in the Lainio and Kumpu Group conglomerates are dominated by rock types found in the underlying formations. They include mafic metavolcanic rocks, tuffites, different kinds of schists (sericite schist, graywacke, phyllite), quartzite, arkose quartzite, vein quartz, chert, iron ore, albitite, rare carbonate rock as well as magmatic rocks such as metagabbro, diabase, granitoids, and felsic porphyries (see Figure 4.4H) (Hackman, 1927; Mäkelä, 1968; Kortelainen, 1983; Räsänen and Mäkelä, 1988). Among the most CHAPTER
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conspicuous clasts are red jasper fragments apparently derived from the Kittilä Group. Pebbles of older conglomerates have also been found. At Mantovaara, undeformed quartzite with a well-preserved clastic texture is found as a common clast type, thus differing from the older, schistose quartzites of the Sodankylä Group. The sedimentological features of the Lainio and Kumpu Group rocks suggest that the deposition took place in a fluvial environment. The rock successions have been related to alluvial fans and braided river deposits (Kortelainen, 1983; Räsänen and Mäkelä, 1988; Nikula, 1988).
Metavolcanic rocks Intermediate to felsic metavolcanic rocks, ~20 km west of Kittilä (Figure 4.3), have been assigned to the Latvajärvi and Tuulijoki Formations in the lower part of the Lainio Group (Lehtonen et al., 1998). They have been dated at 1880 ± 8 Ma by Rastas et al. (2001) using the zircon U-Pb method (see Figure 4.15). The metavolcanic rocks of the Latvajärvi Formation are mainly K-rich trachytes, wheras the Tuulijoki Formation comprises trachyandesites. Representative trace element patterns for felsic porphyries from the Latvajärvi Formation are shown in Figure 4.16B. As shown by the Th vs. Yb plot of Figure 4.17, the 1880 Ma metalavas of the Lainio Group clearly differ from the two older types of felsic porphyries within the Kittilä Group. The εNd (at 1880 Ma) value of –3.5 obtained for a Latvajärvi sample is also distinct from the initial Nd isotope composition of the older porphyries (Figure 4.7). On the other hand, it is equal to the value obtained for a coeval Haaparanta suite monzonite (sample A746b, Table 4.1, Figure 4.7) and suggests substantial involvement of Archean crustal material.
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Pebbles in conglomerates
2015 Ma felsic porphyries
Aakenustunturi Hanhioja Mantovaara
Yb (ppm)
10
Linkupalo
1880 Ma Latvajärvi Fm. metavolcanic rocks
1920 Ma felsic porphyries
0 0
10
20
30
40
Th (ppm) Fig. 4.17. Yb vs. Th plot for three age groups of felsic metavolcanic and dike rocks from the Kittilä region. Symbols with black rims mark porphyry pebbles from the Lainio and Kumpu Group conglomerates (for location, see Figure 4.3).
Isotope studies of conglomerate clasts and detrital minerals Useful time constraints on the deposition of the Lainio and Kumpu Group metasediments are provided by rock pebbles in conglomerates and detrital zircons in quartzites. The bulk zircon data indicate a substantial contribution of post-Archean material in the source region of the Lainio and Kumpu Group metasediments (Hanski et al., 2001b). Rastas et al. (2001) reported field observations and isotope data demonstrating that a Kumpu Group conglomerate at Sätkänävaara (Figure 4.3A) contains diabase clasts derived from ~2050 Ma diabase dikes that cut the underlying metasediments of the Savukoski Group. A Lainio Group conglomerate at Linkupalo (Figure 4.3A) contains felsic porphyry clasts that are geochemically indistinguishable from the ~2015 Ma felsic porphyries of the Kittilä Group (Figures 4.16 and 4.17). As 166
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the latter show the diagnostic radiogenic Nd isotope composition among felsic rocks in northern Finland, we analyzed one porphyry clast from the Linkupalo conglomerate for Nd isotopes. The initial εNd (at 2015 Ma) value of +3.4 is in good agreement with the Nd isotope data from the Kittilä Group porphyries and thus confirms the genetic link (Table 4.1). This result establishes the maximum time of deposition of the Linkupalo conglomerate at ~2015 Ma. Quartz-feldspar porphyry fragments with a still younger age are found in the Mantovaara conglomerate (Figure 4.3A), which overlies Kittilä Group metavolcanic rocks and belongs to the Kumpu Group. Zircons separated from one 20-cm-diameter porphyry pebble yielded an age of 1928 ± 6 Ma (Rastas et al., 2001). This age is similar to that obtained for the Nyssäkoski felsic dike (Figure 4.15) and the Ruoppapalo granodiorite crosscutting Kittilä
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Group metavolcanic rocks. Comparison of the major and trace element data also reveals the similarity of the Nyssäkoski dike, Mantovaara porphyry pebble, and Ruoppapalo granodiorite (Figure 4.16A). The initial εNd value of +1.3 obtained for the pebble is somewhat higher that that of the Nyssäkoski dike (–0.9) but clearly differs from the more radiogenic Nd isotope composition of the older porphyry group (Table 4.1, Figure 4.7). The porphyry clasts in the Mantovaara conglomerate were thus most likely derived from the felsic dikes that intruded the underlying Kittilä Group metalavas. This means that the sedimentation of the Kumpu Group took place later than ~1920 Ma. We have also studied felsic igneous clasts in some other conglomerates belonging to the Lainio Group. Geochemical and petrographic data on several felsic porphyry pebbles (Aakenustunturi, Vesikkovaara, Hangasoja; Figure 4.3A) show their affinity to the trachytic metavolcanic rocks of the ~1880 Ma Latvajärvi Formation (Figures 4.16 and 4.17). One of the clasts from Hangasoja was analysed for Nd isotopes. As expected, the εNd (at 1880 Ma) value of –3.6 is identical to those obtained for the Latvajärvi Formation trachytes (Table 4.1, Figure 4.7). The genetic link between the porphyry clasts and Latvajärvi Formation metavolcanic rocks is further corroborated by the U-Pb zircon date of 1873 ± 11 Ma obtained by Rastas et al. (2001) for a porphyry clast from the Vesikkovaara conglomerate. These authors also reported a U-Pb zircon age of 1888 ± 22 Ma for a granite pebble from a conglomerate at Kellostapuli (Figure 4.3A). This date fits well with the ages of the Haaparanta suite plutons (Väänänen and Lehtonen, 2001; Hiltunen, 1982; Skiöld and Öhlander, 1989) and thus suggests a source component in this igneous suite. Recent ion microprobe U-Pb zircon studies show that both the Lainio and Kumpu Groups include a ~1880 Ma detrital zircon population (Hanski et al., 2000). This matches well the CHAPTER
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age of the felsic to intermediate plutons of the Haaparanta suite in western Finnish Lapland. The above-listed field, petrological, and isotope observations provide strong evidence for a post-1880 Ma deposition of the sediments of the Lainio and Kumpu Groups.
4. Mafic plutonism 4.1. ~2440 Ma intrusions in Lapland Layered mafic intrusions with an age of ~2440 Ma are abundant in the eastern part of the Fennoscandian Shield both in Finland and Russia and are related to the early Paleoproterozoic rifting of the Archean craton (Alapieti et al., 1990; Chapter 3). Two of them, Akanvaara and Koitelainen, are located within the Central Lapland greenstone belt (Figure 4.18). U-Pb zircon ages of 2439 ± 3 Ma (Koitelainen) and 2436 ± 6 Ma (Akanvaara) suggest that these two intrusions are coeval (Mutanen and Huhma, 2001). We focus on these two intrusions, while the other economically important intrusions in the Tornio–Näränkävaara belt farther south (Figure 4.1) are described in Chapter 3 of this volume.
Akanvaara intrusion The Koitelainen and Akanvaara layered intrusions contain significant chromitite reserves and have been targets for exploration and petrological studies for several decades. Detailed descriptions of these intrusions were given by Mutanen (1997). The Akanvaara intrusion is situated in the eastern part of the Central Lapland greenstone belt at the junction of the Salla greenstone area and the Sodankylä schist area (Figures 4.2 and 4.18). It forms a 15-km-long, curved, roughly N-trending monocline with a dip of 25° to 40° and a total stratigraphic thickness of ~3.1 km. Country rocks are felsic to intermediate metavolcanic rocks of the Salla Group (Figure 4.2). Figure 4.19 portrays a stratigraphic col-
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Paleoproterozoic intrusive rocks 2050 Ma layered intrusions
Koitelainen
2220 Ma layered intrusions
Keivitsa-Satovaara
2440 Ma layered intrusions
Haaskalehto
Paleoproterozoic supracrustal rocks
Sodankylä Savukoski
Archean rocks
Akanvaara
40 km
Fig. 4.18. Map showing locations of ~2440 Ma (red), ~2220 Ma (yellow), and ~2050 Ma (blue) layered mafic–ultramafic intrusions in central Lapland.
umn of the Akanvaara intrusion and a detailed drill core section (R322) from the bottom part of the column. Briefly, the Akanvaara intrusion comprises three units: a basal, 65-m-thick, fine-grained gabbro, layered sequence, and >260-m-thick granophyre on top. The layered sequence was divided by Mutanen (1997) into three major units: the Lower, Main, and Upper Zones with respective approximate thicknesses of 640 m, 570 m, and 1900 m. The Lower Zone is composed mainly of bronzite cumulates and overlying noritic gabbros. The intrusion contains at least 23 separate chromitite layers with most of them (the Lower Chromitites) hosted by the bronzite cumulates of the Lower Zone as illustrated in Figure 4.19. The lowest of them is located close to the bottom of the layered sequence. The Main Zone begins with a chromitite (the Uppermost Lower Chromite according to 168
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the terminology of Mutanen, 1997) which is followed by a 25-m-thick peridotite unit and a thicker sequence of noritic gabbro cumulates with minor anorthositic interlayers. The Upper Zone also commences with a chromitite (the Upper Chromitite), which is the most regular and tectonically undisturbed of the chromitites and can be followed more than 8 km along strike. A 1–2-m-thick mottled anorthosite is found above the chromitite, and serves as a persistent marker horizon. This is overlain by noritic gabbros, which occupy almost one half of the Upper Zone, and is capped by an anorthositic unit. The appearance of magnetite as a cumulus phase marks the base of the magnetite gabbro. The uppermost part of the intrusion beneath the granophyres comprises ferrogabbro, apatite-ferrogabbro, and apatiteferrodiorite. The chromitite layers of the Akanvaara
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Akanvaara intrusion Drill core R322 Granophyre
100
Noritic gabbro
Diorite Ferrogabbro
160 γOs +2.7
Depth in meters
800 m 400
Pyroxenite Olivine pyroxenite Chromitite
Magnetic gabbro
UZ
0 Noritic gabbro
Pyroxenite Pegmatoid Plagioclasebearing pyroxenite
220
Chromitite 280
γOs +4.1
γOs +3.2
Upper Chromitite Orthopyroxenite
Anorthosite
MZ γOs +6.1
340
Peridotite/Pyroxenite Uppermost Lower Chromitite
Gabbro Chromitite
γOs +3.3
400
...... .. ... ...... .. ...
LZ Orthopyroxenite + Lower Chromitites Marginal Gabbro
Chromitite Microgabbro Gabbro
460
Acid volcanic rock
Fig. 4.19. Stratigraphy of the Akanvaara intrusion (modified after Mutanen, 1997) and a cross-section of the bottom part of the intrusion as revealed by drill core R322. Also shown are initial γ Os values for chromitites as determined by Hanski et al. (2001c). LZ–Lower Zone, MZ–Main Zone, UZ–Upper Zone.
intrusion are not restricted to the ultramafic part of the cumulate sequence, but the two uppermost chromitites are located in the broadly gabbroic middle part. The chromitites range in thickness from a few cm up to 3 m and their Cr2O3 contents are between 6 wt.% and 32 wt.% (see Figure 6 in Mutanen, 1997). The most peculiar feature of the chromites in the Akanvaara intrusion (and the Koitelainen intrusion as well) is their low MgO content (normally between 0.1 wt.% to 1.1 wt.%) and consequently extremely low Mg/(Mg+Fe2+).
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Koitelainen intrusion The Koitelainen mafic layered intrusion (Figure 4.18) is rounded, approximately 25 km by 30 km in size and has a maximum stratigraphic thickness of ~3.2 km. As in Akanvaara, the hanging wall comprises metavolcanic rocks of the Salla Group. The foot wall rocks, exposed in the anticlinal area in the middle of the intrusion, contain ~3100–2800 Ma Archean gneisses (Tojottamanselkä and Kiviaapa domes) and metavolcanic rocks of the Salla Group (Figures 4.2 and 4.18). It thus appears that the Koitelainen intrusion was injected into the contact zone of the Archean basement com-
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B
A
Fig. 4.20. (A) Photomicrograph showing basal contact of the upper part of the Upper Chromitite against a gabbro interlayer, Koitelainen intrusion. Parallel nicols, the bar is 2 mm long. (B) Layering in mottled anorthosite a few meters above the Upper Chromitite, Koitelainen intrusion. Match box length 5 cm. Photos: Erkki Halme (A) and Tapani Mutanen (B).
plex and its Paleoproterozoic volcanic cover. The layered sequence of the Koitelainen intrusion resembles that of the Akanvaara intrusion with the exception of the presence of a spatially restricted basal series, grading upwards from dunites through clinopyroxenites to monzodioritic rocks, in the northwestern corner of the intrusion (Hanski et al., 2001c). The chromitite-bearing orthopyroxenites above this series can be regarded as the counterparts of the lowest ultramafic cumulates in the Akanvaara intrusion and were probably generated by a new injection of magma. Chromitite layers 0.2 m to 3 m thick, which are sandwiched within these orthopyroxenites, form the Lower Chromitites. The orthopyroxenites are overlain by a thick, monotonous, mainly noritic gabbro sequence enclosing two thin ultramafic interlayers in its lower part. The lower one contains a 5-cm-thick chromitite layer corresponding to the Upperwort Lower Chromitite of the Akanvaara intrusion. Above the gabbroic unit is the Upper Chromitite (Figure 4.20A). It is 2.2 m thick, is regularly associated with mottled anorthosite (Figure 4.20B) and can be followed over a strike length of more than 60 km. The uppermost part of the 170
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intrusion comprises anorthosites, gabbros, and vanadiferous magnetite gabbros. The intrusion is capped with 400 m of granophyre.
Parental magma According to Mutanen (1997), the weighted average composition of the Akanvaara intrusion is 52% SiO2, 6–8% MgO and 0.7% TiO2, thus corresponding to a low-Ti basalt. Fine-grained gabbros at the bottom of the Akanvaara intrusion also indicate that the first melt that entered the magma chamber was not very primitive. The maximum Mg# of 0.838 measured for orthopyroxene in orthopyroxenites and 0.843 for olivine in the lowest dunitic part of the Koitelainen intrusion also imply this. Compared with other, chromititebearing layered intrusions, such as the Bushveld and Stillwater complexes and the Great Dyke, the parental magmas of the Akanvaara and Koitelainen intrusions were clearly more evolved. Trace element analyses of various cumulates from the Koitelainen intrusion show that the magma had high LREE/HREE and low contents of HFSE compared with other incompatible elements, most probably resulting from assimilation of continental crust. In
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this respect, the magma resembled continental low-Ti flood basalts (Hanski et al., 2001c).
Isotope geology Similar to most of the early Paleoproterozoic layered intrusions of the Fennoscandian Shield (e.g., Amelin and Semenov, 1996), the Akanvaara and Koitelainen intrusions are characterized by low initial Nd values of around –2.0 and thus bear a clear crustal signature in their Nd isotope composition (Hanski et al., 2001c). Furthermore, the initial Nd isotope composition is approximately constant throughout the layered sequence in both intrusions. Mutanen (1997) attributed the genesis of the chromitite layers to extensive postemplacement interaction of the magma with country rocks including Cr-bearing sedimentary rocks. To test this hypothesis, Hanski et al. (2001c) performed an Os isotope study on chromitite samples representing different stratigraphic positions. The initial ␥Os values are shown in Figure 4.19 for the Akanvaara intrusion. These are only slightly suprachondritic and no systematic differences are found relative to the stratigraphic position of the samples. The same is true for the Koitelainen intrusion (Hanski et al., 2001c). On the basis of the combined Nd and Os isotope systematics, Hanski et al. (2001c) suggested that extensive crustal contamination took place deep in the crust and that in situ contamination after the final emplacement of the magma was relatively insignificant. The fact that the Os isotope compositions remained almost immune to crustal contamination, while the Nd isotope compositions were strongly affected, led Hanski et al. (2001c) to conclude that the primary magma of the layered intrusions was a high-Os, low-Nd magma, potentially a komatiite or komatiitic basalt. This magma evolved through ACF processes in a deepcrustal magma chamber to a contaminated, low-Ti basaltic composition, and it was then emplaced into upper crustal Archean gneisses and overlying, broadly coeval volcanic rocks. CHAPTER
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This hypothesis is bolstered by the Nd-Os isotope systematics of the contemporaneous, crustally contaminated komatiitic basalts in the Onkamo Group (Table 4.1) and the Vetreny belt, Russian Karelia (Puchtel et al., 2001).
4.2. ~2220 Ma differentiated sills Mafic differentiated sills intruded into quartzites are abundant in the Sodankylä Group, particularly in the area between Sodankylä and Kittilä (Figures 4.2 and 4.18). Similar intrusions are also found within the Jatulian quartzites and their immediate Archean basement in eastern Finland and the Peräpohja belt (Hanski, 1987; Perttunen, 1991; Salmi, 1986; Vuollo and Piirainen, 1992). In the older literature, these rocks were called “albite diabases” (e.g., Piispanen, 1972). However, as they generally do not form dikes but rather gravity differentiated, concordant sills, have cumulate textures, and contain secondary rather than primary albite, Hanski (1986, 1987) referred to these sills as the gabbro–wehrlite association. Later these rocks have also been called karjalites (Vuollo and Piirainen, 1992). Hanski (1986) pointed out that the presence of these sills within quartzites over large areas is not merely coincidental but has implications for the correlation of Jatulian-type quartzites in eastern and northern Finland. At numerous localities the sills have yielded ages of ~2220 Ma, but this magmatic event has rarely been dated elsewhere in the Fennoscandian Shield (Hanski et al., 2001b; Chapter 5). The type occurrence of the gabbro–wehrlite association in central Lapland is the Haaskalehto intrusion for which Tyrväinen (1983) reported a U-Pb zircon age of 2220 ± 11 Ma. This intrusion is part of a string of intrusive bodies that can be delineated within the Sodankylä Group metasediments close to the southern margin of the Central Lapland greenstone belt (Figure 4.18). Individual sills of the gabbro–wehrlite association may reach several hundred meters
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in thickness and they can often be traced for several kilometers along strike. The maximum length (>100 km) is attained by a sill in the Peräpohja belt. The layered sequence in the sills usually comprises the following cumulus assemblages from the bottom upwards: olivine-clinopyroxene, clinopyroxene, clinopyroxene-magnetite, and plagioclase-clinopyroxene-magnetite. The ultramafic cumulates characteristically contain brown, poikilitic amphibole. Its composition normally corresponds to edenitic hornblende, sometimes kaersutite. Hydrous phases are also represented by primary mica varying from titanian phlogopite to Ti-bearing biotite. Orthopyroxene may be present in the olivine-clinopyroxene cumulates but its abundance is always low. Olivine has a maximum forsterite content of ~82%. It ceased to crystallize when its Fo content was diminished to 70%. Plagioclase in gabbroic cumulates is commonly altered to secondary albite, but some intrusions, like the Haaskalehto sill, contain more calcic plagioclase (up to An47; Hanski, 1987). Chilled margin analyses and petrological features of the cumulates show that the parental magma of the gabbro–wehrlite association was a hydrous, low-Al magnesian basalt. The late appearance of plagioclase as a cumulus phase allowed relatively thick ultramafic zones to be developed in the bottom part of the intrusions. The parental magma was characterized by low Al2O3/TiO2 (5–6). It was also enriched in incompatible elements and had LREE-enriched chondrite-normalized REE patterns with (La/Yb)N of ~ 4–5. The initial Nd isotope composition of this magmatic phase is near chondritic (Huhma et al., 1990, 1996; Figure 4.7). The ~2200 Ma magmatic stage appears to have been short-lived, probably involving only a relatively small number of sills. They intruded over a large area at least 300 km in width and 600 km in length and were later disrupted into discrete blocks by tectonic movements. So far no genetically related volcanic coun172
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terparts have been identified. Also, genetically related dikes in the basement outside the immediate contact zones with the schist belts are rare. The tectonic circumstances during their generation thus seem to have favored emplacement as laterally extensive, concordant sill-like bodies rather than dikes, probably at the time when the hosting sedimentary rocks were still unconsolidated. In contrast to the ~2440 Ma plutonism, the ~2200 Ma intrusions seem to have produced no significant mineral deposits in northern Finland. A small Cu-Au deposit in the Peräpohja belt is the only one that has so far been exploited (Rouhunkoski and Isokangas, 1974).
4.3. ~2050 Ma intrusions Mafic intrusions with an age of ~2050 Ma exist in many places in northern Finland, particularly within the Savukoski Group (Rastas et al., 2001). These are typically conformable or semiconformable dike-like intrusions a few tens of meters in thickness. The Keivitsa–Satovaara complex is the only representative of a sizeable, strongly differentiated mafic–ultramafic body of this age group. An olivine pyroxenite gave a nearly concordant U-Pb zircon age of 2058 ± 4 Ma, which is regarded as the crystallization age of the complex (Mutanen and Huhma, 2001). The Keivitsa–Satovaara complex has recently been described by Mutanen (1997). It comprises two separate intrusions (Keivitsa and Satovaara), which probably originally formed a single intrusive body that was later split by faulting. The complex is located just south of the Koitelainen layered intrusion (see Figure 4.18) and intruded phyllites and black schists of the Savukoski Group. The Keivitsa intrusion, hosting a significant Cu-Ni-PGE sulfide deposit, has been studied in more detail.
Keivitsa intrusion The Keivitsa intrusion is funnel-shaped, and
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Keivitsansarvi
N 2 km S Country rocks Serpentinite
Keivitsa intrusion Granophyre Gabbro
Differentiated mafic sill
Olivine pyroxenite
Black schist
“False ore”
Mica schist
Disseminated Cu-Ni sulfide deposit
Mafic to ultramafic metavolcanic rock Arkose quartzite Fault
Fig. 4.21. Geological map of the Keivitsa intrusion (simplified after Mutanen and Huhma, 2001).
has a surface area of ~17 km2 and estimated maximum thickness of more than 2.5 km (Figure 4.21). The intrusion is divided into the marginal chill zone, ultramafic zone, gabbroic zone, and granophyre. The 0–8 m CHAPTER
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thick marginal chill zone at the lower contact of the intrusion consists of microgabbros, quartz gabbros, and quartz-rich pyroxenites. The lower part of the layered succession is occupied by ultramafic rocks, mainly olivine
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websterites and olivine clinopyroxenites with primary intercumulus brown amphibole and biotite-phlogopite. The overlying gabbro zone comprises pyroxene gabbros, ferrogabbros with pigeonite and fayalite, graphite-bearing gabbros, and V-rich magnetite gabbros. The latter grade into granophyres on top of the intrusion. A large, more than 1-km-long serpentinite-dunite body is found within the gabbroic part of the Keivitsa intrusion (Figure 4.21) and is thought to be a huge xenolith genetically unrelated to its host (Mutanen, 1997). Also centimeter-scale graphitic schist inclusions are common in the intrusion. The geochemical characteristics of the marginal zone, multiple saturation of the magma in olivine and pyroxene from the onset of crystallization, and the relatively low forsterite content (<85%) of olivine attest to a basaltic rather than picritic or komatiitic parental magma. In its upper part, the ultramafic zone contains a low-grade Cu-Ni-PGE sulfide deposit, known as the Keivitsansarvi deposit, attaining several hundred meters in extent (Figure 4.21). As reported by Mutanen (1997) and Hanski et al. (1997), this deposit has peculiar mineralogical, chemical, and isotopic features. The ore is formed by disseminated sulfides with the whole-rock sulfur abundances rarely exceeding 3 wt.%. Despite the low-grade nature of the deposit, the chemical variation between different parts of the ore body is substantial and the spatial relationships of the different ore types are rather complex. Three end-member compositions and several subtypes and transitional varieties can be recognized. Mutanen (1997) designated the end-members as the regular, false, and Ni-PGE ore types, using the calculated Ni content (wt.%) normalized to 100% sulfide Ni(100S) as one criterion. The main ore body is composed mostly of regular ore having Ni(100S) between 4% and 7%. The evolved nature of the parental magma is reflected in a relatively low Ni/Cu ratio of less than 1.0. The Ni-PGE ore type is very low in Cu, has 174
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Ni(100S) higher than 25, and is characterized by high PGE contents. Olivine in this ore type has a higher Fo content (85.0–87.5%) than normally encountered in the intrusion, but even more astonishing is its high NiO content up to 1.7 wt.%. The false ore with a low Ni(100S) of 1–4% is practically devoid of all precious components and has also low NiO and Mg/Fe in primary mafic silicates (Fo in olivine down to 76%). Sulfur isotope analyses have yielded heavier than chondritic ␦34S values: regular ore +2.0 to +4.6‰, false ore +5.5 to +12.4‰, Ni-PGE ore +3.7 to +8.8‰. The country rock sulfides possess still higher ␦34S values in excess of +10‰ (Hanski et al., 1996). These results suggest incorporation of substantial amounts of crustal sulfur into the magma. Crustal contamination is also evident in the radiogenic isotopes. Huhma et al. (1995) reported a Sm-Nd isochron age of 2052 ± 25 Ma and an initial Nd value of –3.5 ± 0.3 for the gabbros and ultramafic rocks related to the regular ore type (Figure 4.7). Paradoxically, the highest REE concentrations and lowest initial Nd values (–6.3 and –6.6) have been obtained for samples of the Ni-PGE ore, which forms the most “primitive” ore type in terms of its metal content and primary silicate mineralogy. The regular and Ni-PGE ore types have similar radiogenic Os isotope compositions with initial ␥Os values around +20, while the initial ␥Os values of the false ores are still more radiogenic (␥Os > +100) (Hanski et al., 1997). Although there are several lines of evidence pointing to a substantial crustal contamination, the exact mechanism of ore formation that would account for all the bizarre geochemical and petrological features of the Keivitsansarvi deposit still waits to be formulated.
5. Lapland granulite belt The Lapland granulite belt transects northernmost Finland as an arcuate, roughly northwest-
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trending, 40- to 90-km-wide belt (Figure 4.1). On the Norwegian side of the border it plunges under the Caledonides and on the Russian side it continues for ~100 km to the east, reaching ~400 km in total length. Together with the Umba and Tersk granulite terranes farther to the southeast in the Kola Peninsula, the Lapland granulite belt forms a core zone of the Lapland–Kola orogen, which developed as a result of collision of two Archean cratons and an intervening juvenile arc complex after the closure of the ’Lapland–Kola’ ocean (Daly et al., 2001). On its southern and western side, the Lapland granulite belt is bordered by the Tanaelv (or Tana) belt, a relatively narrow, strongly deformed, heterogeneous zone comprising high-grade amphibolites and garnet-biotite and quartz-feldspar gneisses (Barbey et al., 1984). In the south, both the Lapland granulite belt and Tanaelv belt were thrust southwards on Archean gneisses (the Belomorian terrane in Russia), whereas in the west they were thrust over Paleoproterozoic supracrustal rocks of the Kittilä–Karasjok greenstone complex. The thrust contact between the Lapland granulite belt and Tanaelv belt dips gently towards northeast, and the corresponding seismic reflectors extend to mantle depths. On the other hand, the shear zones between the northeastern margin of the Lapland granulite belt and the Archean Inari area gneisses are subvertical (Korja et al., 1996). According to Korja et al. (1996), the Lapland granulite belt can be divided into four parallel, northeastdipping major thrust sheets reaching 15 km in thickness. The rock types of the Lapland granulite belt are grouped into three main series: the khondalite series, charnockite series, and migmatites (Barbey and Raith, 1990). The dominant khondalite series comprises sillimanite-garnet gneisses, garnet gneisses, and subordinate calc-silicate rocks and graphitebearing schists. The chemical composition of these rocks suggests that some of their CHAPTER
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protoliths were sedimentary rocks including shales, psammitic graywackes, sandstones or carbonate rocks. Other rocks possess chemical characteristics of andesites and rhyolites and were probably volcanic in origin. The charnockite series consists of metamorphosed plutonic rocks ranging from pyroxenites to norites (enderbites). Migmatites are mostly confined to the middle and upper parts of the granulite belt and are considered as partial melting products of sillimanite-garnet gneisses and garnet gneisses (Barbey and Raith, 1990).
5.1. Metamorphic conditions Peak conditions of granulite facies metamorphism have been estimated at ~850 °C and 8 kbar (Barbey and Raith, 1990). Postkinematic high-grade metamorphism reached P-T conditions of ~830 °C and 7.2 kbar in the southwestern tectonostratigraphically lower part of the belt. There is a decreasing temperature and pressure gradient towards the intensively migmatized northeastern part of the belt, where conditions of 760 °C and 6.2 kbar have been recorded. Tuisku and Huhma (1998a) reported ultramafic and mafic rocks at the southwestern margin of the granulite belt that were metamorphosed under transitional conditions between amphibolite and eclogite facies, suggesting pressures up to 12 kbar.
5.2. Radiogenic isotopes Geological and U-Pb geochronological studies performed in the early 1970’s provided background for understanding the geological evolution of the Lapland granulite belt and adjacent areas (Meriläinen, 1976). Archean ages were confirmed for the gneisses northeast of the granulites (Inari Terrane), whereas U-Pb zircon dating of quartz diorites of the granulite belt and granitoids of the Kuorboaivi schist zone (northwestern part of the Inari area) suggested crystallization ages of ~ 1900–1950 Ga. Important age constraints for metamorphic
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evolution were obtained from monazites from the garnet-cordierite gneisses in the granulite belt (~1910 Ma) and titanites in the Inari terrane (~1900 Ma). Bernard-Griffiths et al. (1984) published a U-Pb zircon age of 1906 ± 5 Ma for the Vaskojoki anorthosite and a similar age for a pyroxene gneiss from the adjacent Tanaelv belt, which could represent the timing of the granulite-facies metamorphism. The Sm-Nd isotope method has been em ployed to evaluate the average crustal residence ages. Results from the Lapland granulite belt suggest a major contribution from Paleoproterozoic sources (Huhma and Meriläinen, 1991). The depleted mantle model ages (DePaolo, 1981) for orthogneisses range from 2300 Ma to 2000 Ma (Bernard-Griffiths et al., 1984, recalculated) and for paragneisses they cluster around 2300 Ma (Huhma and Meriläinen, 1991; Daly et al., 2001). On the northeastern side of the Lapland granulite belt felsic orthogneisses have yielded model ages of 2100–2000 Ma (Barling et al., 1997; Hannu Huhma, unpublished data). This, together with their calc-alkaline geochemical characteristics, suggest the existence of a juvenile, ~1950–1900 Ma arc complex (Barling et al., 1997). The Nd isotope data are compatible with the U-Pb ion probe ages obtained from single zircons from the Lapland granulites (Sorjonen-Ward et al., 1994; Tuisku and Huhma, 1998b). Zircon ages from migmatitic paragneisses range from ~3000 Ma to 1900 Ma, with about a half of the grains having protolith ages between 2100 Ma and 1950 Ma. About one third of the analyzed zircons have ages close to 1900 Ma and the youngest metamorphic grains and overgrowths are ~1880 Ma. This is close to the Sm-Nd age of garnet–whole-rock pairs (Tuisku and Huhma, 1999; Daly et al., 2001). Isotope studies on the Lapland–Kola orogen farther east have yielded results comparable to those quoted above (Bibikova et al., 1993; Daly et al., 2001, and references therein). The obtained geochronologic data 176
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thus indicate a relatively short history of the Lapland–Kola orogen with the deposition, deformation, metamorphism, and significant exhumation having taken place within 100 Ma. Dating of anatectic felsic dikes shows that the major thrusting and peak of granulite facies metamorphism was over by ~1910 Ma (Meriläinen, 1976; Daly et al., 2001; Kaulina et al., 2001; Kislitsyn et al., 2001; Glebovitsky et al., 2001).
6. Summary and discussion 6.1. Mantle plume(s) and cracking of the craton The earliest Paleoproterozoic magmatism in the Fennoscandian Shield is represented by the Salla and Onkamo Group metavolcanic rocks, the ~2440 Ma mafic layered intrusions in northern Finland, and the metavolcanic rocks of the Sumi–Sariolian formations in Russia. Although the Sumi–Sariolian metavolcanic rocks have previously been considered orogenic (Heiskanen et al., 1977) and their geochemical signature has been attributed by some authors to a continental arc environment (Pharaoh et al., 1987), no direct relationship to subduction processes seems to exist. Instead, these rocks were probably deposited in intracratonic basins and their geochemical features are due to crustal contamination or were inherited from the subcontinental lithospheric mantle. In this respect, they are reminiscent of more recent continental flood basalts. Accordingly, the magmatic activity of this stage is related to an incipient mantle plume in the initial stages of continental rifting (Amelin et al., 1995; Puchtel et al., 2001; Hanski et al., 2001c). Indeed, the Salla and Onkamo Groups and the Sumi–Sariolian formations possess many characteristics of rock successions associated with mantle plumes (cf. White and McKenzie, 1995). Firstly, a huge volume of magma was emplaced within a short period
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of time. Although older ages have been reported for some Russian intrusions (Amelin et al., 1995), the ages of both intrusive and extrusive rocks tend to converge to a narrow spread around 2450–2430 Ma (e.g., Amelin et al., 1995; Puchtel et al., 1997; Mutanen and Huhma, 2001; Manninen et al., 2001). Secondly, the associated sedimentary rocks are terrigenous arkoses and polymictic conglomerates (Gaskelberg et al., 1986), probably reflecting a regional domal uplift of the crust due to arrival of a mantle plume. Theoretical calculations have shown that significant plume-induced surface uplift may commence 10–20 Ma prior to the onset of volcanism and reach 1–2 km (White and McKenzie, 1989; Campbell and Griffiths, 1990). Thirdly, the volcanic rocks are found mainly as subaerially erupted lavas and display geochemical and isotopic features similar to those found in rocks of more recent continental flood basalt provinces.
6.2. Cratonic sedimentation and volcanism The short-lived, violent, volcanism-dominated overture was followed by a long depositional period producing mature continental sediments and lesser amounts of intermittent mafic volcanic rocks, all deposited on an ensialic Archean substrate. These Sodankylä Group rocks grade transgressively to deeper-water pelitic and euxinic sediments of the Savukoski Group. The Sodankylä Group contains dolomitic interbeds with stromatolitic structures and elevated carbon isotope compositions and hosts concordant differentiated sills of the ~2200 Ma gabbro–wehrlite association, which all are features traditionally regarded as typical of the Jatulian formations. In the Peräpohja belt, the temporal and lithologic counterpart of the Sodankylä Group is the Kivalo Group (Perttunen et al., 1995). The lowest sedimentary rocks of this group were deposited on tilted, uplifted, and partly eroded CHAPTER
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remnants of the ~2440 Ma intrusions (Perttunen, 1991). The observed unconformity in the succession suggests a long period of denudation in the Peräpohja area before the deposition of the supracrustal rocks of the Kivalo Group. A similar situation is observed in the Pechenga area, Kola Peninsula, where the lowest sedimentary rocks lie unconformably on the partly eroded Mt. Generalskaya layered intrusion (Bayanova et al., 1999). In the Central Lapland greenstone belt, such a major unconformity has not yet been recognized between the Sodankylä Group and the preceding plutonic rocks of the ~2440 Ma magmatism.
6.3. Primitive volcanism and deepening basins Compared with Archean komatiites, the komatiitic and picritic metavolcanic rocks of the Savukoski Group have a different environment as they are associated with graphite- and sulfidebearing pelites. A similar transgressive evolution leading to deposition of pelitic, graphiterich sediments accompanied by primitive, subaqueous volcanic rocks has been documented at least in two other Paleoproterozoic successions in the Fennoscandian Shield, viz., the Pilgujärvi Group of the Pechenga area in the Kola Peninsula and the Suisaarian rocks of the Onega region in Russian Karelia (Hanski, 1992; Puchtel et al., 1995). Despite analogous geological positions, available isotope data suggest that the Pechenga ferropicrites and Onega picritic basalts are more than 50 Ma younger than the ~2050 Ma komatiites and picrites in central Lapland (Hanski et al., 1990; Puchtel et al., 1998). To explain the genesis of komatiites, the currently most favored model invokes decompressional melting of a hot, adiabatically ascending deep-seated mantle plume (e.g., Campbell et al., 1989). This is also an appealing concept for the Central Lapland komatiite–picrite association, particularly in the light
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of the geochemical similarities between the picritic rocks in the Central Lapland greenstone belt and the plume-related picrites in Hawaii (Hanski et al., 2001a). A plume model has also been proposed for the genesis of the Pechenga ferropicrites and Onega basalts (Hanski, 1992; Walker et al., 1997; Puchtel et al., 1998). A single incipient mantle plume is, however, not easy to reconcile with the measured age difference between the Lapland rocks and the other occurrences. This is because volcanic rocks of many flood basalt provinces, thought to be linked to a plume head, were extruded in a very short time interval of only a few million years (e.g., White and McKenzie, 1995). As discussed in more detail by Hanski et al. (2001a), another difficulty in central Lapland, and also in the Pechenga and Onega areas, is that the associated sedimentary rocks provide no evidence for a preceding stage of strong regional surface uplift, which is generally assumed to be associated with the impingement of an upwelling mantle plume upon the base of the rigid lithosphere (e.g., White and McKenzie, 1989). One further complication is that the Pechenga area is located on the opposite side of the Lapland–Kola suture zone relative to the Central Lapland and Onega regions (e.g., Daly et al., 2001), thus causing uncertainty on the relative positions of these areas at the time of the postulated plume event (Puchtel et al., 1998).
6.4. Breakup of a supercontinent? In the Archean basement complex of eastern Finland, outside the Central Lapland greenstone belt, at least five mafic dike swarm generations have been identified, having different orientations and approximate ages of 2500 Ma, 2450 Ma, 2330 Ma, 2100 Ma, and 1970 Ma (Vuollo et al., 1995, 2000; Chapter 5). Isotope and geochemical data show that some of the dikes are genetically related to the ~2450–2440 Ma layered intrusions (Vogel et al., 1998), but the links between the other 178
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dike generations and volcanic or intrusive rocks within the supracrustal belts have not yet been established. Several authors have proposed that the Archean cratons of North America were assembled together with the Archean blocks of the Fennoscandian and Siberian Shields in the late Neoarchean (Roscoe and Card, 1993; Heaman 1997, 1998; Aspler and Chiarenzelli, 1998; Condie, 1998). However, paleomagnetic evidence for this is still somewhat ambiguous (Buchan et al., 2000). If the hypothesis is true, the supercontinent (Kenorland of Williams et al., 1991) experienced more than the five episodes of Paleoproterozoic dike injections listed by Vuollo et al. (1995, 2000). Dike generations with ages of 2230 Ma, 2220 Ma, 2210 Ma, 2040 Ma, and 2020 Ma have been documented from the Canadian Shield (Corfu and Andrews, 1986; Pehrsson et al., 1993; LeCheminant and van Breemen, 1994) with the first three ages approaching the age of the intrusions of the gabbro–wehrlite association in the Fennoscandian Shield. It is generally thought that the supercontinent started fracturing and rifting at ~2450 Ma (e.g., Heaman, 1997). This was related to the upwelling of a mantle plume, which resulted in vast outpourings of crustally contaminated lavas as represented, for example, by the Salla and Onkamo Groups, and emplacement of mafic layered intrusions in northern Finland. Subsequently, several magmatic episodes ensued over hundreds of Ma, producing mafic dikes in the Archean basement and various kinds of lavas in the supracrustal belts. Giant mafic dike swarms in the basement complexes have been regarded as potential feeders of old continental flood basalts and their genesis can hence be related to magmatic activity of mantle plumes (Ernst and Buchan, 1997). The radiating patterns of the dike swarms can even be used to locate plume centers (op. cit.). Besides establishing the potential genetic relationship between the different dike generations and extrusive and
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subvolcanic formations, the challenges of future research include the question of the geodynamic control of these separate magmatic events, particularly the role of mantle plumes. Are all these magmatic phases plume-related? How many mantle plumes were there? These are difficult questions because of inadequate age information and because the consequences of mantle plumes may differ depending on the size of the plume, thermal and chemical properties of the lithospheric mantle, and thickness of the lithosphere. Hawkesworth et al. (1999) calculated that 18 mantle plumes may have emplaced under the Gondwana supercontinent within a time interval of 300 Ma prior to its breakup. It is thus conceivable that several mantle plumes extended to the area of the present-day Fennoscandian Shield during the ~400 Ma period when the Salla, Onkamo, Sodankylä, and Savukoski Groups were deposited. The presence of multiple dike swarms indicates that the Fennoscandian Shield (and the Kenorland as a whole) was under a prolonged (>400 Ma) extensional tectonic regime in the early Proterozoic, during which several abortive events tried to rupture the continent. We do not know which, if any, of the potentially plume-related magmatic events resulted in or were associated with the final fragmentation of the supercontinent. Several lines of evidence suggest that the rifting events at 2500–2200 Ma did not proceed to sea-floor spreading and that the supercontinent remained essentially a coherent block at least until 2200 Ma. These include the worldwide, ~2350 Ma glaciations (Young, 1991), global positive carbon isotope anomalies in ~2200 Ma carbonate rocks (Karhu and Holland, 1996; Melezhik et al., 1999a), general absence of juvenile, subduction-related magmatism at 2400–2200 Ma (Condie, 1998), and lack of 2500–2200 Ma ophiolites. Korsman et al. (1999) suggested that the breakup of the Archean continent and development of a marginal sea in Fennoscandia took place at 2100–2060 Ma. The CHAPTER
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best evidence for the final continental breakup (and subsequent collision) is presented by ophiolite complexes (see Chapter 6). Our data from central Lapland show that the continental fragmentation had advanced to sea-floor spreading by ~2000 Ma. If the suggested timing of supercontinental dispersion is valid, the pre-2200 Ma Sodankylä Group and other correlative (Jatulian) formations were probably deposited in an intracratonic basin. This is also consistent with the hypothesis of Melezhik et al. (1999a) that relates the deposition of stromatolitic, isotopically anomalous “Jatulian” carbonates of Russian Karelia to shallow-water, evaporative, partly restricted peritidal basins. Deposition of the epicontinental sediments of the Sodankylä Group and related formations may reflect an aftermath of the purported plume event at ~2450 Ma and preceding large and rapid regional crustal uplift. Cessation of the plume-induced volcanism and plutonism was followed by gradual thermal subsidence and formation of depositional basins over an extensive area (cf. Campbell and Griffiths, 1990; White, 1997). Thus the origin of the Salla, Onkamo, and Sodankylä Groups may be related to a declining plume activity. The later marine transgression related to the Sodankylä to Savukoski Group transition can be correlated with a clear global maximum in black shale abundance at ~2000 Ma (Condie et al., 2001) or the Shunga event of Melezhik et al. (1999b).
6.5. Ocean floor volcanism The supracrustal rocks discussed above are primarily autochthonous formations deposited on the Archean sialic basement, as is the case for the traditional Karelian formations (see Chapter 7). This is not necessarily true for the Kittilä Group, however. This unit contains volumetrically minor ~2015 Ma acid porphyries. If these porphyries and associated mafic metalavas represent coeval magmatism as
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suggested by geological and petrographical studies and U-Pb isotopic evidence and, if the mafic metavolcanic rocks represent part of an ancient oceanic crust, then the porphyries cannot be products of anatexis of underlying old sialic basement. This model predicts that the porphyries should have a noncrustal isotope signature. Indeed, the initial Nd isotope ratios of the porphyries are remarkably ‘primitive’. They match very closely the initial isotope composition of the associated mantle-derived mafic metavolcanic rocks and deviate drastically from the low initial εNd values of –8 or less expected for 2000 Ma partial melts of Archean sialic crust (Figure 4.6; cf. Huhma, 1986). This result is not only consistent with the prediction of the geotectonic model, but offers a strong proof of a noncontinental environment, provided that such isotopically juvenile felsic rocks can be regarded as diagnostic of oceanic crust. Silica-saturated felsic volcanic rocks found in a continental setting have typically a crustal-like isotopic signature reflecting direct crustal anatexis or open-system fractionation of basaltic parental melts. This applies, for example, to the rhyolitic volcanic rocks associated with continental flood basalts (e.g., Peate, 1997; Nicholsson and Shirey, 1990), including the felsic metavolcanic rocks of the Latvajärvi Formation and the Salla Group (Table 4.1). In contrast to continental rhyolites, felsic igneous rocks in oceanic settings characteristically display a strongly radiogenic Nd isotope composition. These settings include intra-oceanic island arcs (e.g., Brouxel et al., 1987), backarc basins (e.g., Hochstaedter et al., 1990; Gribble et al., 1998), oceanic islands (e.g., Geist et al., 1995), oceanic plateaus (White et al., 1999), and mid-oceanic spreading ridges (Engel and Fisher, 1975). Isotopically primitive felsic rocks have also been discovered in many ophiolite complexes as “plagiogranites” (e.g., Amri et al., 1996; Peltonen et al., 1998) and more rare potassic granites (Nägler and Frei, 1997) and rhyolitic extrusives (Ahmed 180
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and Ernst, 1999; Floyd et al., 1998). We conclude that the ~2015 Ma felsic porphyries and associated mafic metavolcanic and sedimentary rocks of the Kittilä Group represent oceanic bimodal magmatism. Hence we suggest that the Kittilä Group is allochthonous and represents a block of ancient oceanic lithosphere. The scarcity of basalts with typical NMORB chemistry indicates that the bulk of the Kittilä Group metavolcanic rocks do not belong to a normal mid-ocean ridge system, but more likely represent an ancient oceanic plateau or a ridge segment influenced by a mantle plume. Hot and thick, such blocks of oceanic crust are reluctant to subduct and consequently have a moderate chance to be preserved. An analogy with Iceland is further supported by the bimodal magmatism. However, the lithologic and geochemical heterogeneity of the Kittilä Group (OIBs, island arc tholeiites, boninitic dikes, and suprasubduction zone-type ultramafic rocks) indicates that a single geotectonic environment cannot have produced all the different kinds of rock types. Rather, the Kittilä Group is considered a composite formed by separate volcanic terranes that were amalgamated during oceanic convergence and a subsequent collision related to the tectonic emplacement of the greenstone complex. It is not uncommon that ophiolitic complexes contain accreted blocks of volcanic rocks generated in different oceanic domains (e.g., Wirth et al., 1994; Ahmed and Ernst, 1999). One problem in the hypothesis that the Kittilä Group represents a slice of oceanic crust is the nature of one of its volcanic formations, Kautoselkä. The overall chemical composition of the metavolcanic rocks of this unit renders these rocks as hybrids between within plate basalts (WPB) and volcanic arc basalts (VAB) (Figure 4.12B). Chemically comparable highTi basalts have been documented from the Columbia River flood basalt province, for example by Martin (1989). Therefore, a natural explanation would be that the parental magma
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of the Kautoselkä Formation underwent interaction with continental crustal material thus creating low Nb/Th, Nb/La, and initial εNd in the magma (Figure 4.6). However, in contrast to typical continental flood basalts, the Kautoselkä Formation lavas were erupted in a submarine environment, as the associated sedimentary rocks are metagraywackes, phyllites, black schists, siltstones, and fine-grained carbonate rocks, though the amygdaloidal nature of the metavolcanic rocks suggests a relatively shallow water depth. If these volcanic rocks were generated in a continental setting, they may originally have formed a passive margin sequence of a new ocean basin which, due to subsequent collisional tectonics, was amalgamated to the other rocks of the Kittilä Group.
6.6. Acid magmatism related to obduction? The ~1920 Ma felsic dikes and minor plutons in the area of the Kittilä Group provide a minimum age for the suggested tectonic emplacement of the Kittilä Group (Hanski et al., 2001b). Assuming a cause-and-effect relationship, we propose that these rocks actually time this tectonic event. Independent evidence for a compressional tectonic regime at that time comes from two opposing directions. Namely, the closure of the ‘Lapland–Kola’ ocean and subsequent overthrusting of the Lapland granulite belt is thought to have been completed by ~1910 Ma (Daly et al., 2001). On the other hand, the collision of the Svecofennian arc complex from the southwest against the Archean continent probably took place between 1910 to 1885 Ma (Korsman et al., 1999). At the moment, the significance of these two collisional orogenic events for the emplacement of the Kittilä Group is unclear.
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6.7. Foreland basin Terrestrial sedimentation and volcanism The angular unconformity beneath the metasediments of the Lainio and Kumpu Groups is striking. These redbed-type metasediments can be regarded as intracontinental foreland basin deposits accumulated in an arid climate. Clastic material was transported into the basin from various sources including the recently obducted slice of oceanic crust, the adjacent Svecofennian orogen in the southwest, and Archean basement and its Paleoproterozoic cover (the Sodankylä and Savukoski Groups). Conventional U-Pb isotope data on bulk detrital zircons have demonstrated that the Lainio and Kumpu Group rocks contain a significant component of post-Archean detritus (Hanski et al., 2001b). Petrographical, geochemical, and isotope (U-Pb, Sm-Nd) studies on the clasts in the Lainio and Kumpu Group conglomerates have revealed fragments from several well-characterized and dated rock units outcropping in western Lapland, thus effectively constraining the maximum age of deposition. These fragments include ~2050 Ma gabbroic rocks, ~2015 Ma and ~1920 Ma felsic porphyries, and the ~1880 Ma Latvajärvi Formation metavolcanic rocks and Haaparanta suite plutonic rocks. These, together with NORDSIM data on detrital zircons, demonstrate that both the Lainio and Kumpu Group sediments were deposited later than ~1880 Ma. Moreover, geochronologic data on clasts and detrital minerals cannot be used to distinguish the Lainio and Kumpu Groups from each other. On the basis of structural observations, Lehtonen et al. (1998) interpreted that the Lainio Group rocks predate and the Kumpu Group rocks postdate the “main deformation stage” of the Svecofennian orogeny. However, the available geochronologic data do not support this interpretation. We suggest that the observed variation in the degree of deformation of the two groups may
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be a function of location in variably strained regions, rather than their relative age. The age constraints coupled with the lithologic and structural observations from the Kumpu Group strongly argue that the Kumpu Group represents molasse-like sediments deposited after the obduction, folding, metamorphism, and uplift of the underlying ~2000 Ma Kittilä Group. Some earlier investigators have compared the rocks of the Kumpu Group with molasse deposits (Simonen, 1960, 1980), but isotopic evidence for this has become available only recently.
Correlation with Svecofennian sedimentation and volcanism In northern Sweden, a large ~1900 Ma Svecofennian volcanic terrain contains mostly subaerially erupted, felsic to mafic lavas. Perdahl and Frietsch (1993) called this felsic rockdominated lava suite the Kiruna–Arvidsjaur Porphyry Group. The ~1880 Ma Latvajärvi Formation and its time-correlative Tuulijoki Formation in the Kittilä area can be regarded as the sole representatives of the extension of the Kiruna–Arvidsjaur Porphyry Group to northern Finland. According to Lehtonen et al. (1998), the geochemistry of the Latvajärvi and Tuulijoki Formations has a clear calc-alkaline island arc signature and is consistent with an Andino-type continental margin environment suggested for the Kiruna porphyries by Pharaoh and Brewer (1990). On the basis of their geologic position and available isotope data, the metasediments of the Lainio and Kumpu Groups can be correlated with the coarse-clastic, post-1880 Ma metasediments deposited on the Kiruna–Arvidsjaur Porphyry Group metavolcanic rocks in northern Sweden (Hanski et al., 2001b), as exemplified by the metasediments in the Kiruna, Arvidsjaur, and Vargfors areas (Martinsson, 1997; Ödman, 1957; Perdahl and Frietsch, 1994; Offerberg, 1959). Similar deposits can be followed further southeast across the Bothnian Bay to western Finland. These deposits 182
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have been informally assigned to the Upper Svecofennian assemblage (Lundqvist et al., 1996; Kousa et al., 2000). Corresponding coarse-clastic metasediments are also found in northern Norway, including the Çaravarri Formation in the Kautokeino greenstone belt (Siedlecka et al., 1985). Isotope data from the above-mentioned Svecofennian formations and associated granitoids suggest that the uplift, exhumation, and erosion of the synorogenic granitoids and the sedimentation of the Upper Svecofennian assemblage took place rapidly after the crystallization of the synorogenic granitoids, potentially within a time span of less than 10 Ma (Kouvo and Tilton, 1966; Korsman et al., 1988; Nironen, 1989; Persson and Lundqvist, 1997; Billström and Weihed, 1996). Evidently, a similar rapid tectonic evolution led to the deposition of the Lainio and Kumpu Groups in the Finnish Lapland.
Relationship to the exhumation of granulites Because the Kittilä Group rocks were affected by the emplacement of the Lapland granulite belt, the overthrusting of the granulites must have existed later than at ~1920 Ma but not much later than that, if the age of 1912 ± 2 Ma obtained by Kislitsyn et al. (2001) for an undeformed leucosome from the Kolvitsa belt in the southeastern extension of the Tanaelv belt can be applied to the granulite belt proper. On the other hand, Tuisku et al. (2001) used ion probe data to conclude that the Lapland granulite belt was thrust and exhumed at ~1.90–1.87 Ga. The same time interval was also favored by Daly et al. (2001) for the exhumation and cooling of the Lapland and Umba granulite terranes. In any event, the Kumpu and Lainio Group sediments could have received detritus from the uplifting Lapland granulite belt in the northeast. It is interesting to note that Torske (1997) interpreted the molasse-type Çaravarri Formation in the Kautokeino area, northern Norway, as representing postcolli-
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sional foreland basin deposits in front of the Lapland–Kola thrust belt. Also, Evins and Laajoki (2002) suggested that the Kumpu Group conglomerates at Pyhätunturi were deposited in a thrust-sheet-top basin during southward thrusting and nappe formation related to the Lapland–Kola orogeny. Granulite clasts or detrital high-pressure metamorphic minerals have not been documented in the Lainio and Kumpu Group metasediments. If they are absent, this means that either the sediment transport was predominately from other directions and/or the granulite facies rocks were not yet exposed when the Kumpu and Lainio Groups were formed. Available paleocurrent indicators from the Pyhätunturi region are compatible with the former alternative (Räsänen and Mäkelä, 1988), while measurements from the Sirkka area have yielded variable paleocurrent directions (Kortelainen, 1983). A titanite U-Pb age of 1915 ± 15 Ma and a rutile U-Pb age of 1876 ± 22 Ma from the Tanaelv belt (Kaulina et al., 2001) indicate the times of cooling below 700 °C and 450 °C, respectively, and suggest that granulite-facies rocks were not yet eroding when the Kumpu and Lainio Group rocks were deposited.
7. Conclusions The ~600 Ma evolution of the Central Lapland greenstone belt commenced with plume-related, intracratonic rift magmatism producing large layered intrusions and primarily subaerial volcanic rocks ranging from crustally contaminated komatiites to rhyolites. These were covered by epicontinental clastic sediments and later by deeper-water pelitic and carbonaceous sediments. This stage has traditionally been connected with sedimentation taking place during a platform-to-continental margin development. Subsequent submarine eruptions of komatiitic to picritic metavolcanic rocks have been dated at ~2060 Ma. This magmatic event has also been related to mantle CHAPTER
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plume activity, but the interpretation has some problems, for example, the lack of evidence for a preceding strong crustal uplift. The absence of cratonic metasediments, presence of EMORB- and OIB-type mafic metavolcanic rocks accompanied with isotopically juvenile felsic rocks and existence of ultramafic rocks with ophiolitic characteristics suggest that at least part of the ~2000 Ma Kittilä greenstone area represents Paleoproterozoic oceanic crust. This was tectonically emplaced onto the Archean basement at ~1920 Ma and was followed by high-pressure metamorphism and overthrusting of the Lapland granulite belt, intrusion of orogenic felsic plutonic rocks of the ~1880 Ma Haaparanta suite, and extrusion of minor intermediate to felsic volcanic rocks. Deposition of the Lainio and Kumpu Group arenitic and rudaceous sediments in a foreland basin soon after the synorogenic magmatism finished the Paleoproterozoic supracrustal evolution of central Lapland. This took place somewhat after 1880 Ma and completed a prolonged Wilson cycle. Later, postorogenic granites were emplaced at ~1800 Ma. There was an increasing contribution of old sialic crust in the source of felsic plutonic and volcanic rocks with decreasing age between 2000 Ma and 1800 Ma, as indicated by a systematic change of Nd towards more negative values (Figure 4.7).
Acknowledgments We thank the staff of the geochronology laboratory at the Geological Survey of Finland, especially Tuula Hokkanen, Arto Pulkkinen and Irmeli Mänttäri, for assistance in the isotope analyses. Tuomo Manninen is thanked for providing samples from the Salla area for Sm-Nd analysis. We are grateful for many discussions with our colleagues in Rovaniemi, including Tuomo Manninen, Tapani Mutanen, Jorma Räsänen, Pentti Rastas, and Jukka Väänänen. Comments by Tapani Rämö,
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Matti Lehtonen, and an anonymous reviewer are highly appreciated.
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dating of the Palaeoproterozoic Pite conglomerate in northern Sweden. Geol. Surv. Sweden, Res. Papers C830, 41–49. Perttunen, V., 1991. Kemin, Karungin, Simon and Runkauksen kartta-alueiden kallioperä. Summary: Pre-Quaternary rocks of the Kemi, Karunki, Simo and Runkaus mapsheet areas. Geological Map of Finland 1 : 100 000, Explanation to the Map of Rocks, Sheets 2541, 2542+2524, 2543 and 2544. Geol. Surv. Finland, Espoo. Perttunen, V., Vaasjoki, M., 2001. U-Pb geochronology of the Peräpohja schist belt. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Paper 33, 45–84. Perttunen, V., Hanski, E., Väänänen, J., 1995. Stratigraphical map of the Peräpohja Schist Belt. 22nd Nordic Geological Winter Meeting, January 8–11, 1996, Turku, Abstracts, p. 152. Pharaoh, T. C., Brewer, T. S., 1990. Spatial and temporal diversity of early Proterozoic volcanic sequences; comparisons between the Baltic and Laurentian shields. Precambrian Res. 47, 169–189. Pharaoh, T.C., Warren, A., Walsh, N.J., 1987. Early Proterozoic metavolcanic suites of the northernmost part of the Baltic Shield. In: T.C. Pharaoh, R.D. Beckinsale, D.T. Rickard (Eds.), Geochemistry and mineralization of Proterozoic volcanic suites. Geol. Soc. Spec. Publ. 33, 41–58. Piispanen, R., 1972. On the spilitic rocks of the Karelidic belt in western Kuusamo, northeastern Finland. Acta Univ. Ouluensis. Ser. A, Sci. Rer. Natur. 4, Geologica 2, 1–73. Puchtel, I., Bogatikov, O.A., Kulikov, V.S., Kulikova, V.V., Zhuravlev, D.Z., 1995. The role of crustal and mantle sources in the petrogenesis of continental magmatism: Isotope and geochemical evidence from the Early Proterozoic picrites of the Onega Plateau, Baltic Shield. Petrol. 3, 357–378. Puchtel, I.S., Haase, K.M., Hofmann, A.W., Chauvel, C., Kulikov, V.S., Garbe-Schoenberg, C.D., Nemchin, A.A., 1997. Petrology
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and geochemistry of crustally contaminated komatiitic basalts from the Vetreny Belt, southeastern Baltic Shield; evidence for an early Proterozoic mantle plume beneath rifted Archean continental lithosphere. Geochim. Cosmochim. Acta 61, 1205–1222. Puchtel, I.S., Arndt, N.T., Hofmann, A.W., Haase, K.M., Kröner, A., Kulikov, V.S., Kulikova, V.V., Garbe-Schönberg, C.-D., Nemchin, A.A., 1998. Petrology of mafic lavas within the Onega plateau, central Karelia: Evidence for the 2.0 Ga plume-related continental crustal growth in the Baltic Shield. Contrib. Mineral. Petrol. 130, 134–153. Puchtel, I.S., Brügmann, G.E., Hofmann, A.W., Kulikov, V.S., Kulikova, V.V., 2001. Os-isotope systematics of komatiitic basalts from the Vetreny Belt, Baltic Shield: evidence for a chondritic source of the 2.45 Ga plume. Contrib. Mineral. Petrol. 140, 588–599. Radchenko, A., Balagansky, V., Basalaev, A., Belyaev, O., Pozhilenko, V., Radchenko, M. 1994. An explanatory note on Geological Map of the north-eastern Baltic Shield of a scale of 1:500 000. Apatity, Russia. Räsänen, J., 1996. Palaeovolcanology of the Sattasvaara komatiites, northern Finland: evidence for emplacement in a shallow water environment. IGCP Project 336 Symposium in Rovaniemi, Finland, August 21–23, 1996, Program and Abstracts. Turun yliopiston geologian ja mineralogian osaston julkaisuja 38, 69–70. Räsänen, J., Huhma, H., 2001. U-Pb datings in the Sodankylä schist area, the central Finnish Lapland. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Paper 33, 153–188. Räsänen, J., Mäkelä, M., 1988. Early Proterozoic fluvial deposits in the Pyhätunturi area, northern Finland. In: K. Laajoki, J. Paakkola (Eds.), Sedimentology of the Precambrian formations in eastern and northern Finland. Proceedings of the IGCP 160 Symposium at Oulu, January 21–22, 1986, Geol. Surv. Finland, Spec. Paper 5, 239–254.
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Räsänen, J., Hanski, E., Lehtonen, M., 1989. Komatiites, low-Ti basalts and andesites in the Möykkelmä area, Central Finnish Lapland. Report of the Lapland Volcanite Project. Geol. Surv. Finland, Rep. Invest. 88. Räsänen, J., Hanski, E., Juopperi, H., Kortelainen, V., Lanne, E., Lehtonen, M., Manninen, T., Rastas, P., Väänänen, J., 1995. New stratigraphical map of Central Finnish Lapland. 22nd Nordic Geological Winter Meeting, January 8–11, 1996, Turku, Abstracts, p. 182. Rastas, P., Huhma, H., Hanski, E., Lehtonen, M.I., Paakkola, J., Mänttäri, I., Härkönen, I., 2001. U-Pb isotopic studies on the Kittilä greenstone area, Central Lapland, Finland. In: M. Vaasjoki (Ed.), Radiometric Age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Paper 33, 95–141. Roscoe, S.M., Card, K.D., 1993. The reappearance of the Huronian in Wyoming; rifting and drifting of ancient continents. Can. J. Earth Sci. 30, 2475–2480. Rouhunkoski, P., Isokangas, P., 1974. The coppergold vein deposit of Kivimaa at Tervola, N-Finland. Bull. Geol. Soc. Finland 46, 29–35. Salmi, H., 1986. Hyrynsalmen alueen varhaisproterotsooiset mafiset ja ultramafiset juonet ja intruusiot. M.Sc. Thesis, Dept. of Geology, Univ. of Oulu, Finland. (in Finnish) Saverikko, M., 1983. Kummitsoiva komatiite complex and its satellites in northern Finland. Bull. Geol. Soc. Finland 55, 111–139. Saverikko, M., 1985. The pyroclastic komatiite complex at Sattasvaara in northern Finland. Bull. Geol. Soc. Finland 57, 55–87. Sederholm, J.J., 1932. On the geology of Fennoscandia with special reference to the PreCambrian. Explanatory notes to accompany a general geological map of Fennoskandia. Comm. géol. Finlande, Bull. 98, 1–30. Shervais, J.W., 1982. Ti-V plots and the petrogenesis of modern and ophiolitic lavas. Earth Planet. Sci. Lett. 59, 101–118. Siedlecka, A., Iversen, E., Krill, A., Lieungh, B.,
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Morten, O., Sandstad, J., Solli, A., 1985. Lithostratigraphy and correlation of the Archean and early Proterozoic rocks of Finnmarksvidda and the Sørvaranger district. Geol. Surv. Norway, Bull. 403, 7–36. Silvennoinen, A., 1972. On the stratigraphic and structural geology of the Rukatunturi area, northeastern Finland. Geol. Surv. Finland, Bull. 257, 1–48. Silvennoinen, A., 1985. On the Proterozoic stratigraphy of Northern Finland. In: K. Laajoki, J. Paakkola (Eds.), Proterozoic exogenic processes and related metallogeny. Geol. Surv. Finland, Bull. 331, 107–116. Simonen, A., 1960. Pre-Quaternary rocks in Finland. Comm. géol. Finlande, Bull. 191, 1–49. Simonen, A., 1980. The Precambrian in Finland. Geol. Surv. Finland, Bull. 304, 1–58. Simonen, A., 1986. Stratigraphic studies on the Precambrian in Finland. Geol Surv. Finland, Bull. 336, 21–37. Skiöld, T., Öhlander, B., 1989. Chronology and geochemistry of late Svecofennian processes in northern Sweden. Geol. För. Stockholm Förh. 111, 347–354. Sorjonen-Ward, P., Nurmi, P.A., Härkönen, I., Pankka, H.S., 1992. Epigenetic gold mineralization and tectonic evolution of a lower Proterozoic greenstone terrane in the northern Fennoscandian (Baltic) Shield. In: S. C. Sarkar (Ed.), Metallogeny related to tectonics of the Proterozoic mobile belts. IBH Publishing, New Delhi, Oxford, pp. 37–52. Sorjonen-Ward, P., Claoué-Long, J., Huhma, H., 1994. SHRIMP isotope studies of granulite zircons and their relevance to early Proterozoic tectonics in northern Fennoscandia. In: M. A. Lanphere, G. B. Dalrymple, B. D. Turrin (Eds.), Abstracts of the Eighth International Conference on Geochronology, Cosmochronology, and Isotope Geology, Berkeley, California, USA, June 5–11, 1994. U.S. Geol. Surv. Circ. 1107, p. 299. Sorjonen-Ward, P., Nironen, M., Luukkonen, E., 1997. Greenstone associations in Finland. In: M.J. De Wit, L.D. Ashwal (Eds.), Greenstone belts. Oxford Monogr. Geol.
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Geophys. 35, 677–698. Torske, T., 1997. The Caravarri Formation of the Kautokeino greenstone belt, northern Norway; foreland basin sediments in front of the Lapland-Kola thrust belt. COPENA Conference at NGU, August 18–22, 1997, Abstracts and Proceedings, NGU Report 97, 131. Tuisku, P., Huhma, H., 1998a. Eclogite from the SW-marginal zone of the Lapland granulite belt; evidence from the 1.90–1.88 Ga subduction zone. In: E. Hanski, J. Vuollo (Eds.), International ophiolite symposium and field excursion; generation and emplacement of ophiolites through time; Abstracts, Excursion Guide. Geol. Surv. Finland, Spec. Paper 26, p. 61. Tuisku, P., Huhma, H., 1998b. SIMS dating of zircons: metamorphic and igneous events of the Lapland granulite belt are 1.9 Ga old, provenance is Paleoproterozoic and Archaean (2.0–2.9 Ga) and the tectonic juxtaposition about 1.9–1.88 Ga old. In: N. Philippov (Comp.) SVEKALAPKO: EUROPROBE project workshop, Repino, Russia, November 26–29, 1998, Abstracts. St. Petersburg, Ministry of Natural Resources of Russian Federation, State Company “Mineral”, pp. 64–65. Tuisku, P., Huhma, H., 1999. SIMS U-Pb dating of zircons from migmatite khondalites and enderbite from Lapland granulite belt, Finland. J. Conf. Abst. 4, 710–711. Tuisku, P., Huhma, H., Mikkola, P., 2001. Petrology, geochronology and evolution of Lapland granulite belt. Univ. of Oulu, Dept. of Geophysics, Rep. No. 24, Svekalapko, an Europrobe Project, 6th Workshop, Lammi, Finland, November 11 – December 2, 2001, Abstracts, p. 61. Tyrväinen, A., 1983. Sodankylän ja Sattasen kartta-alueiden kallioperä. Summary: PreQuaternary rocks of the Sodankylä and Sattasvaara map sheet areas. Geological Map of Finland 1 : 100 000, Explanation to the Maps of Rocks, Sheets 3713 and 3714. Geol. Surv. Finland, Espoo (in Finnish, English summary) Väänänen, J., 1989. Kolarin alueen vulkaniitit.
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Summary: Volcanic rocks of the Kolari area, western Finnish Lapland. Report of the Lapland Volcanite Project. Geol. Surv. Finland, Rep. Invest. 86, 1–79. Väänänen, J., Lehtonen, M., 2001. Isotopic age determinations from the Kolari-Muonio area, western Finnish Lapland. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Paper 33, 85–93. Vaasjoki, M. (Ed.), 2001. Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Paper 33, 1–279. Vogel, D.C., Vuollo, J.I., Alapieti, T.T., James, R.S., 1998. Tectonic, stratigraphic, and geo-chemical comparisons between ca. 2500–2440 Ma mafic igneous events in the Canadian and Fennoscandian shields. Precambrian Res. 92, 89–116. Vuollo, J., Piirainen, T., 1992. The 2.2 Ga old Koli layered sill: the low-Al tholeiitic (karjalitic) magma type and its differentiation in northern Karelia, eastern Finland. Geol. För. Stockholm Förh. 114, 131–142. Vuollo, J. I., Nykänen, V. M., Liipo, J. P., Piirainen, T. A., 1995. Palaeoproterozoic mafic dyke swarms in the Eastern Fennoscandian Shield, Finland: a review. In: G. Baer, A. Heimann (Eds.), Physics and chemistry of dykes. A. A. Balkema, Rotterdam, pp. 179–192. Vuollo, J., Huhma, H., Pesonen, L., 2000. Mafic dyke swarms – geological evolution of the Palaeoproterozoic in the Fennoscandian Shield. In: L.J. Pesonen, A. Korja, S.-E. Hjelt (Eds.), Lithosphere 2000 – A symposium on the structure, composition and evolution of the lithosphere in Finland. Programme and Extended Abstracts, Espoo, Finland, October 4–5, 2000. Institute of Seismology, Univ. of Helsinki, Report S-41, 107–111. Walker, R.J., Hanski, E.J., Vuollo, J., Liipo, J., 1996. The Os isotopic composition of Proterozoic upper mantle: evidence from the Outokumpu ophiolite, Finland. Earth Planet. Sci. Lett. 141, 161–174. CHAPTER
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PALEOPROTEROZOIC MAFIC DIKES IN NE FINLAND
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Cover page: Gabbronorite dike and older Archean dike from Pääjärvi–Karankaniemi, Russian Karelia. Length of compass is ~13 cm. Photo: Jouni Vuollo
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Vuollo, J., Huhma, H., 2005. Paleoproterozoic mafic dikes in NE Finland. 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. 195–236. © 2005 Elsevier B.V. All rights reserved.
Several mafic dike swarms are found in the eastern and northern parts of the Fennoscandian Shield and can be divided into at least six main groups dated at 2.5 Ga, 2.45 Ga, 2.32 Ga, 2.2 Ga, ~2.1 Ga, and 1.98 Ga. The 2.5 Ga group is found only in the Kola Peninsula, the others occur throughout the Karelian domain. The 2.45 Ga episode includes dikes, layered intrusions, and dike swarms that can be divided to five subgroups: (1) NE-trending boninite–noritic dikes, (2) NW-trending gabbronorite dikes, (3) NW-trending tholeiitic dikes, (4) NW-trending Fe-tholeiitic dikes, and (5) E-trending orthopyroxene- and plagioclase-phyric dikes. Groups 1, 2, and 5 are calc-alkaline, groups 3 and 4 are tholeiitic. The 2.2 Ga low-Al tholeiitic (karjalitic) sills are mostly associated with the Archean–Proterozoic unconformity. These sills are composed of marginal zones, layered series, and granophyre; the layered series contains rocks ranging from wehrlite and clinopyroxenite to magnetite gabbro. The 2.2 Ga magmatic series is referred to as karjalite or low-Al tholeiite because it is poorer in Al and richer in Fe and LREE than ordinary tholeiites and komatiites. The younger dike swarms (2.32–1.98 Ga) are typical continental tholeiitic basalts, and their further subdivision based on chemical composition is virtually impossible. The 2.32 Ga tholeiitic dikes and intrusions are relatively rare, yet wide-spread, and probably represent a significant magmatic episode. The ~2.1 Ga Fe-tholeiitic dike swarm is a dominant feature within the Archean craton and Paleoproterozoic cover in NE Finland, foreboding a breakup event at ~2.0 Ga. The main trend of the dikes is approximately E–W across the Kuhmo block, but a NE trend is dominant in the North Karelia schist belt. This group is geochemically homogeneous, showing pronounced Fe enrichment and high LILE, HFSE, and LREE contents. Thus it is similar to the continental tholeiitic dikes of many Paleoproterozoic swarms. The main trend of the 1.98 Ga dikes is approximately NE in the Kuhmo block and their composition is mainly Fetholeiitic. Comparable dikes have also been found in the North Karelia schist belt and in central Lapland. In North Karelia, they have high LILE, are depleted HFSE, and show flat to slightly LREE-enriched REE patterns, thus corresponding to island arc tholeiites.
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1. Introduction Dike swarms hold one of the keys to the interpretation of plate tectonics, as they provide information on the extensional processes occurring both in the continental crust and in the oceanic lithosphere. Dikes are the primary channels for transporting basaltic magma into the crust from a source area in the mantle and thus they can also be used to assess the nature of parental magma for related lavas and intrusions. Mafic dike swarms are also useful time markers that often precisely register major episodes of crustal rifting. Knowledge of the timing of dike emplacement is essential for understanding the tectonic evolution of rift-related environments and for regional correlation of igneous activity. Dike swarms are extremely important in continental environments, because they are often the only surviving evidence of significant geological events (e.g., rifting, mantle plumes, plate subductions, or crustal “breakup”) and can be used to monitor geological history of the continents over long periods of time (Fahrig, 1987; Halls and Fahrig, 1987; Ernst and Buchan, 2001a,b). Dike swarms are ideal indicators for reconstruction of Precambrian crustal blocks, because they provide information on varying paleostress directions (e.g., Halls and Palmer, 1990; Neuvonen et al., 1997). A subparallel swarm indicates that the stress field was constant over a long distance. Dikes may originate in the mantle or in a crustal source and may intrude vertically, or they may originate from a high-level plutonic complex and be emplaced horizontally or with a strong horizontal component. The swarms may be many hundreds of kilometers long, and individual dikes and sills may extend 100 km in length. Their emplacement is associated with significant extension of the continental crust, and dike densities indicating an extension of 5–10% can be found over large areas (Cadman et al., 1990). Dike swarms are also frequently 198
posttectonic and thus undeformed. They can be arranged along fracture zones, which are closely connected with the disaggregation of continents and crustal blocks. Dike formation has taken place episodically over the last 3000 million years in all continents, and a considerable proportion of the continental mafic dike swarms are either Proterozoic or Late Phanerozoic in age. Paleoproterozoic dike swarms and sills are widespread in all Precambrian shields (Ernst et al., 1996). Several mafic dike swarms comprising a large number of dikes are also found in the eastern and northern parts of the Fennoscandian Shield (Gorbatschev et al., 1987; Vuollo et al., 1995b, 2001). Thus Finland, Russian Karelia, and the Kola Peninsula offer a healthy ground for the study of Paleoproterozoic dikes. The entire Archean craton and Karelian Supergroup up to the Kalevian Group are intersected by voluminous NW-, E-, and NE-trending dike swarms (1–3 dikes/km2), the first general account of which was provided by Aro and Laitakari (1987). A new version of the Fennoscandian dike swarm map now builds on GIS databases (Vuollo et al., 2001). Figure 5.1 shows all the mafic dikes swarms and Figure 5.2 provides an example (aeromagnetic maps and field ground surveys) of how the GIS database was compiled. The first studies of the Paleoproterozoic dikes date back to the late 1960’s (Piirainen, 1969), and later work at the University of Oulu (e.g., Hanski, 1986; Vuollo, 1994) has continued this research. Since 1993, joint studies of Paleoproterozoic dike swarms have been carried out by the University of Oulu, University of Toronto, Royal Ontario Museum (Canada), Geological Survey of Finland, and the Russian Academy of Sciences (Apatity and Petrozavodsk) utilizing geochronology (Vuollo et al., 1995a, 1999, this study), geochemistry (Vuollo et al., 1995b; Vogel et al., 1998), and paleomagnetism (Mertanen, 1995; Mertanen et al., 1999a,b). The aim has been to identify various diking events in the eastern part of the
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Fig. 5.1. Eastern Fennoscandian dike swarm map after Vuollo et al. (2001). C H A P T E R 5 • PA L E O P R O T E R O Z O I C M A F I C D I K E S I N N E F I N L A N D •
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Fig. 5.2. Examples of GIS databases. (A) Low-altitude aeromagnetic map with observation and age determination points. (B) Pseudocolor low-altitude aeromagnetic map showing dikes of the 1.98 Ga dike swarm. (C) Ground survey magnetic map (data from Posiva Co.) with dike observation points, dikes and Sm-Nd age determination sample sites.
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Fennoscandian Shield and their relationship to economically important layered intrusions and ophiolites, and to establish the earliest part of the Proterozoic apparent polar wander path for Fennoscandia (see Chapter 15). These studies have provided valuable information on the Paleoproterozoic geological evolution of the shield and have enabled continental reconstructions, e.g., in the North Atlantic area (Mertanen et al., 1999a). Dike swarms were sampled (Figure 5.1) in several locations in the Karelian province (Taivalkoski, Pudasjärvi, and Kuhmo blocks in eastern Finland; Suoperä and Lake Pääjärvi areas in Russian Karelia), where overprinting by the 1.9 Ga Svecofennian orogeny is thought be minimal (Kontinen et al., 1992). This review paper and the associated dike maps (Figures 5.1 and 5.2) are based on these studies and further details can be found elsewhere (Nykänen et al., 1994; Vuollo, 1994; Mertanen, 1995; Vuollo et al., 1995b, 2000; Vogel et al., 1998; Mertanen et al., 1999a,b; Vuollo and Huhma, 2004). The eastern Fennoscandian mafic dike swarm GIS databases (Salmirinne, 2001; Vuollo et al., 2001) were compiled in 1996– 1999 (in ArcInfo–Arcview) to combine and correlate the existing information (1: 100 000scale geological bedrock maps for dikes in eastern and northern Finland). Databases concerning the Russian part of the dike swarms were compiled by the Russian Academy of Sciences (Apatity and Petrozavodsk) and State Mineral Co. (St. Petersburg). We have also utilized aeromagnetic maps from our study area in the form of TIFF pictures (see Figure 5.2). The observation database involved includes more than 10,000 observations in both Finland and Russia and the petrophysical database extracted from the national database of the Geological Survey of Finland consists of 3700 samples. The geochronological database includes 78 U-Pb and 29 Sm-Nd ages for dikes and sills of different types. The geochemical database consists of more than 1500 whole-rock analyses, and the mineral-
ogical database contains hundreds of silicate and oxide microprobe analyses (see Janhila, 2001). An updated eastern Fennoscandian dike swarm map has been digitized (on the scale of 1:50,000) utilizing all these sources (see Figures 5.1 and 5.2).
2. Geological background The Precambrian rocks of our study area are classified into two main groups – the Archean basement complex (3.1 Ga to 2.6 Ga) and the Paleoproterozoic cover. The latter can be divided into the Karelian and Svecofennian Supergroups (Figure 5.3). The Karelian supracrustal sequence was deposited on the Archean basement, which in central and eastern Finland can be divided into the Kuhmo, Taivalkoski, Pudasjärvi, and Iisalmi blocks (Figure 5.1). The blocks are dominated by granitoids, ‘basements gneisses’, and remnants of greenstone belts. The Karelian Supergroup, which rests unconformably on the Archean basement, is divided into three groups: Sumi–Sariola, Jatuli, and Kaleva (Figure 5.3). The Sumi–Sariola Group consists of immature or moderately mature arkoses, conglomerates, and subaerial metalavas. The first Paleoproterozoic continental rifting phase of the cratonized Archean basement involved emplacement of layered mafic intrusions at ~2.5 Ga (Kola Province) and ~2.45 Ga (Kola and Karelian provinces; Tornio–Näränkävaara belt, central Lapland, and Burakovka) and boninite–noritic–gabbro–noritic–tholeiitic to Fe-tholeiitic dikes in eastern and northern Finland (Figure 5.4). Thick chemical weathering sequences separate the Sumi–Sariola formations in Kainuu and North Karelia. The next rifting phase occurred at ~2.32 Ga and was represented by small intrusions and dikes. Previous studies of the Peräpohja metavolcanic rocks have shown that those in the Runkaus Formation are approximately of the same age (Huhma et al., 1990).
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201
Svecofennian Supergroup
Svecofennian granite suite 1900
Kalevian Ophiolites Group Fe-tholeiitic dikes
2000
Keivitsa & Otanmäki Intrusions Fe-tholeiitic dikes 2100
Jatulian Group 2200
Layered Sills
Karelian Supergroup
2300
Fe-tholeiitic dikes and intrusions
Sariolian Group 2400
Layered Intrusions
Sumi Group
Boninite–norite and tholeiitic dikes
2500 Noritic dikes (Kola Peninsula)
Layered Intrusions (Kola)
Dike swarms
Supracrustal rocks
Dike swarms
Mafic intrusive rocks
Granite rocks Archean basement
Ophiolite suite
Fig. 5.3. Simplified Paleoproterozoic geological evolution scheme for the eastern part of the Fennoscandian Shield together with special emphasis on mafic igneous activity (dike swarms, intrusions, ophiolites). Ages in Ma.
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The Jatulian Group is dominated by mature quartzites and conglomerates and, in the upper part, minor dolomites, pelites, and black schists, traditionally assigned to the ’marine Jatulian’. The sedimentation of the Jatuli formations was accompanied by three mafic igneous events. Low-Al tholeiitic or ‘karjalitic’ magmas intruded through the basement at 2.2 Ga. Most of the karjalites (gabbro–wehrlite association of Hanski, 1984) are found as layered sills. An extensive set of dike swarms that cuts the Archean crust and the Jatulian Group is Fe-tholeiitic, ~2.1 Ga old, and trends E–W in the Kuhmo block. The Keivitsa intrusion in central Lapland (Mutanen and Huhma, 2001) and the Otanmäki layered intrusion in Kainuu (Talvitie and Paarma, 1980) represent the next extensional phase, and there are also signs of ~2.05 Ga dike swarms (see Figure 5.3). At the top of the Karelian sequence, turbiditic rocks of the Kalevian Group are present; these are separated from the Jatulian Group by an unconformity. The 1.95 Ga ophiolites of Outokumpu and Jormua (Kontinen, 1987; Peltonen et al., 1996) are located in the upper Kaleva. Ophiolitic ultramafic rocks have also been identified in central Lapland (Hanski, 1997). A significant sign of the breakup event, which predates the above-mentioned ophiolites, is a 1.98 Ga tholeiitic and Fe-tholeiitic dike swarm intersecting the Jatulian Group supracrustal rocks (North Karelia and central Lapland) and the Archean craton (NW-trend; Kuhmo block). Moreover, there is no indication of comparable magmatic activity in the Kaleva turbidites. The Paleoproterozoic lithological units were deformed and metamorphosed in several stages during the Svecofennian orogeny at ~1.9–1.8 Ga. The Archean basement was involved in these and earlier processes, as shown by numerous basic dike swarms transecting it. Paleomagnetic studies from the Karelian province also show a strong Svecofennian overprint (Mertanen, 1995; Mertanen et al., 1999a). Low-altitude aeromagnetic maps
(Figures 5.2 and 5.23) together with field studies indicate that there are several Archean high-grade terranes (Vuollo et al., 2000). In addition, new seismic reflection surveys have revealed some uplifted areas (granulite blocks?) in the Kuhmo block (depth of reflection Moho is 60 km in the west side of Kuhmo greenstone belt and ~40 km in the east side of it) (Annakaisa Korja, pers. comm., 2003).
3. Mafic dike swarms The mafic dike swarms in Finland have traditionally been divided into three groups (e.g., Aro and Laitakari, 1987): (1) the Jatulian diabase dikes and sills (>1.9 Ga), (2) the Subjotnian diabase dikes (~1.6 Ga), and (3) the Postjotnian diabase dikes (~1.3 Ga). This paper provides a general account of the history and diversity of the Jatulian dike swarms and sills in the Fennoscandian Shield in eastern and northern Finland and Russian Karelia. A map of the areal distribution of mafic dike swarms in the eastern Fennoscandian Shield is shown in Figure 5.1. U-Pb zircon geochronology (summarized by Vuollo, 1994) indicates that there were several dike emplacements between 2.5 Ga and 1.98 Ga. It is suggested that these dike swarms and intrusions can be divided into at least six main groups based on their geochemical composition, age, and mode of occurrence – groups dated at 2.5 Ga, 2.45 Ga, 2.32 Ga, 2.2 Ga, 2.1 Ga, and 1.98 Ga can be recognized (Table 5.1). The 2.5-Ga age group occurs only in the Kola Peninsula, but the others are found throughout the Karelian province (Figures 5.1 and 5.3). These six magmatic events are assigned here to age groups that may contain dikes of slightly different ages and involve rocks of different geochemical compositions. The names chosen for the main dike swarms indicate their principal geochemical characteristics. The U-Pb and Sm-Nd ages for mafic dikes and intrusions in northern and eastern
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Table 5.1. Classification of Paleoproterozoic dike swarms in eastern and northern Finland and Russian Karelia. Eastern Fennoscandian Shield 2.45 Ga dike swarms - connected with layered intrusions; two magma types (1) boninite–norite dikes (high MgO, SiO2, Cr, Ni, and LREE, low TiO2 and Zr), NE-trend (2) gabbronorite dikes (low TiO2, Cr, and Zr), NWtrend (3) low-Ti tholeiitic (NW-trend) and (4) Fe-tholeiitic dikes (E-trend), continental type (5) orthopyroxene-plagioclase-phyric dikes (high SiO2, LREE; calc-alkaline aff.), E-trend
2.32 Ga Fe-tholeiitic dike swarm and intrusions - few dikes and intrusions identified - E-trend?
(5) E-trending orthopyroxene-plagioclasephyric dikes. The younger dike swarms (2.32–1.98 Ga) form a homogeneous group in terms of geochemical composition and, in general, resemble continental tholeiitic basalts. This means that a more elaborate subclassification according to chemistry is virtually impossible. Nevertheless, it has been possible to classify the ~2.1 Ga and 1.98 Ga dikes in the Kuhmo block using their geochemical composition, age, and orientation. The oldest group in the Kuhmo block comprises dikes with a 280˚ trend and the second group dikes with a 320˚ trend (Table 5.1). The 2.2 Ga layered sills are widespread in eastern and northern Finland and are stratigraphically located near the Archean–Paleoproterozoic unconformity.
2.2 Ga low-Al tholeiitic (karjalitic) layered sills - layered intrusions/sills (max. length 150 km) and minor dikes - wehrlite–clinopyroxenite–gabbro–granophyre - widespread in eastern and northern Finland
~2.1 Ga Fe-tholeiitic dike swarm - includes several pulses between 2.13 Ga to 2.05 Ga - continental tholeiitic type - mainly E-trend (Kuhmo block) and minor NW-trend
~1.98 Ga Fe-tholeiitic–tholeiitic dike swarm - predates 1.95 Ga ophiolites - NW-trend (Kuhmo block), continental tholeiitic to IAT type - mainly unaltered (plagioclase and pyroxenes)
Finland are summarized in Figures 5.4, 5.14, 5.15, 5.20, and 5.21 and in Table 5.2. Many of the U-Pb determinations are discordant and heterogeneous, and we have included 38 selected determinations. The discordant analyses indicate ages mainly between 2.2 Ga and 2.0 Ga. The ~2.45 Ga dike swarms (Table 5.1) can be divided into five subgroups based on their field relationships and geochemical and isotopic characteristics: (1) NE-trending boninitenoritic dikes, (2) NW-trending gabbronorite dikes, (3) NW-trending tholeiitic dikes, (4) NW- and E-trending Fe-tholeiitic dikes, and 204
3.1. ~2.45 Ga dike swarms The most conspicuous products of the ~2.45 Ga magmatism (Figure 5.3) are undoubtedly the mafic layered intrusions (Chapter 3), but mafic dikes of this age are also found throughout the Archean areas (Figure 5.4). As noted earlier, this dike swarm is divided into five subgroups (Table 5.1). The current perception of the areal distribution of the ~2.45 Ga dike swarms is shown in Figure 5.4. Further studies will probably show that these dikes are even more voluminous – the map in Figure 5.4 shows only boninite–norite dikes, some gabbronorite dikes, and some tholeiitic and Fe-tholeiitic dikes. One isotope age has been published from the Salla greenstone belt (2383 ± 33 Ma; Manninen and Huhma, 2001), all other existing ages (Vuollo et al., 2000) are from the Archean basement area (Kuhmo, Taivalkoski, and Pudasjärvi blocks, and Russian Karelia; see Table 5.2). One U-Pb baddeleyite result (four fractions) indicates an age of 2446 ± 6 Ma (gabbronorite – NW-trend), while other baddeleyite ages provide only crude estimates: a minimum age of 2395 Ma for a boninite–
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norite dike (NE-trend) and an age of 2378 Ma for a tholeiitic dike (NW-trend).
Boninite–norite dikes The boninite–norite dikes are found in many places in the Kuhmo block (Kilpelä, 1991; Vuollo et al., 1995b), in the vicinity of layered intrusions of the Tornio–Näränkävaara belt (Alapieti et al., 1990; Perttunen, 1991; Iljina et al., 1992; Saini-Eidukat et al., 1997; Chapter 3), and in Russian Karelia (Stepanov, 1994; Vuollo et al., 2002). The dikes have a general northeasterly trend in the Kuhmo block (Figure 5.4). In the Suhanko–Konttijärvi area, they are parallel to the basal contacts of the layered intrusions (Iljina et al., 1992). The dikes are 20–60 m thick and some of them can be traced along strike for >40 km. They trend NE (30–40˚) and are characterized by the presence of coarse plagioclase (35%), orthopyroxene (30%), and clinopyroxene (20%) with minor amounts of olivine (5%), chromite, and Fe-Ti oxides (Figure 5.13A). The dikes are fresh only along the Finnish–Russian border and on the Russian side of the border. Elsewhere pyroxenes are altered to amphibole and olivine to serpentine. Fine-grained chilled margins are occasionally observed. Clinopyroxenes are always Ca-augites (~Wo35En52Fs13) and orthopyroxene is bronzite (~En83Fs13Wo4). Plagioclase grains have a core-to-rim zonation from ~An70 to An30. Small chromite grains (<0.05 mm) occur as inclusions in olivine and pyroxenes. Their Cr2O3 contents varies from 38 to 54 wt.% and the lowest Cr number is 0.56. The boninite–norite dikes always exhibit a cumulus texture and their plagioclase grains are turbid (Figure 5.13B), displaying the so-called “teacolor cloudiness”.
Gabbronorite dikes The dikes of the gabbronorite dike swarm trend NW (310˚) and are fresh on the Russian side and also in some areas on the Finnish
side. Glassy chilled margins have been found in places and, near the contacts, quench textures are seen (Figure 5.13C). These dikes are medium- to coarse-grained, up to 50 m thick, and they consist of clinopyroxene (25%), plagioclase (60%), and orthopyroxene (5–10%) with minor amounts of olivine (1–2%), quartz (<5%), biotite, and Fe-Ti-oxides. As the boninite–norites, they are also altered in many places on the Finnish side. The areal distribution of the gabbronorites is difficult to assess because they have the same trend as the younger dikes. According to geochemical studies (low Cr relative to the boninite–norite dikes; Figures 5.5–5.11), these dikes are probably present on both sides of the border. They can be traced for a distance of some kilometers. The gabbronorite dikes are always cumulus-textured (Figure 5.13D) and plagioclase laths (An50–60) have the same cloudiness as in the boninite–norite dikes. Clinopyroxene is Ca-augite, with an average composition of Wo32En51Fs17. Sometimes the clinopyroxene grains have pigeonite exsolution lamellae with a fishbone texture (Figure 5.13D). Orthopyroxene is bronzite (~En73Fs22Wo5).
Low-Ti tholeiitic dikes A low-Ti tholeiitic dike swarm (trend 340˚) has been dated in the Pudasjärvi block (a minimum age of 2378 Ma) and these dikes are also seen in Kuhmo block both on the Russian and Finnish sides of the border (for geochemistry, see Figures 5.5–5.11). In some places, they are fresh and consist of plagioclase (50%), clinopyroxene (35%), minor amounts of orthopyroxene (2%), quartz, secondary amphibole, and Fe-Ti-V oxides. The dikes are >2 km long and normally 40–70 m wide. Tholeiites are cumulus-textured (Figure 5.13E). Plagioclase is An50–60 and clinopyroxene is Ca-rich augite with an average composition of Wo36En44Fs20. One feature in common with the other ~2.45 dikes is that plagioclase is cloudy.
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Fig. 5.4. Distribution and location of dated samples (green rectangle−U-Pb zircon or baddeleyite age, red rectangle−Sm-Nd age) of the 2.5 Ga dike swarms in the eastern and northern Fennoscandian Shield (Finland and Russian Karelia). Age data: 1−Manninen and Huhma (2001), others−Vuollo et al. (1995a, 2000), and Vuollo and Huhma (2004).
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Table 5.2. U-Pb and Sm-Nd ages (this study) and geographical orientations (trd) of the mafic dike swarms in Kuhmo, Taivalkoski, and Pudasjärvi blocks and Russian Karelia. U-Pb Block Kuhmo Dike swarm 1.1.Boninite–norite 2395(Min)
Trd
Taivalkoski
Trd
Russian Karelia Trd
Pudasjärvi
Trd
40˚
–
60˚
–
40˚
–
280˚
1.2 Gabbronorite 1.3.Tholeiite 1.4. Fe-tholeiite 2. Fe-tholeiite
– – – –
290˚ 320˚ – 270˚
– –
– –
2446 +5/–4
340˚
– 330˚ –
–
280˚
278˚ 340˚ –
–
3. Fe-tholeiite
2332 ± 18 2306 ± 6 –
– 2378 – –
–
–
2118 ± 14
300˚
4. Fe-tholeiite
1981±5
320˚
–
–
–
–
–
350˚?
Taivalkoski εNd
Trd
Sm-Nd Block Dike swarm 1.1.Bon.–norite 1.2 Gabbronorite 1.3.Tholeiite 1.4. Fe-tholeiite 2. Fe-tholeiite 3. Fe-tholeiite 4. Fe-tholeiite 350˚?
Kuhmo
εNd
Trd
2370 ± 70 –1.8 40˚
– – – 2407 ± 35 +1.6
60˚ – – – – 285˚ – 2319 ± 27 +1.8 340˚ 2133 ± 32 +0.6 280˚ – – 2054 ± 40 +0.3 280˚ 2014 ± 33 +0.4 320˚ – – 1992 ± 47 +0.2 320˚
Fe-tholeiitic dikes So far, only one Fe-tholeiitic dike has been dated (Sm-Nd isochron age of 2407 ± 35 Ma and εNd value of +1.6; Vuollo and Huhma, 2004; Table 5.2). Nevertheless, we associate these dikes with the previously recognized 2.45 Ga dike swarms. The dated dike is exposed just near the village of Taivalkoski at a road cut, is 10 m wide, and shows conspicuous chilled margins. The modal composition of the dikes is plagioclase (60%), clinopyroxene (30%), and Fe-Ti-oxides (5–10%) with minor amounts of quartz and biotite (Figure 5.13F).
Orthopyroxene-plagioclasephyric dikes The E-trending orthopyroxene-plagioclasephyric dikes are known only in some places in Russian Karelia and in the Kuhmo block near the border. The dikes are black and their
Russian Karelia εNd
2421 ± 27 2476 ± 30 2349 ± 30
Trd Pudasjärvi
εNd
Trd
40˚ – 280˚ –2.4 340˚ – – +1.7 320˚ 2461 ± 150 +0.5 330˚ – – +1.0 270˚ – –
–
–
–
–
–
–
300˚
width varies from 1 m to 3 m. The presence of relatively large (1–5 mm), well-preserved orthopyroxene phenocrysts and, in some places, large plagioclase laths (Figure 5.13G), is a typical feature of these dikes. The finegrained groundmass consists of plagioclase, orthopyroxene, and clinopyroxene with minor amounts of Fe-Ti-oxide and biotite.
3.2. Geochemical and isotopic characteristics Major and some trace element analyses of different dike swarms (from Kuhmo and Taival koski blocks and Russian Karelia; best-preserved areas) are presented in Figures 5.5–5.11. In these diagrams, it is difficult to separate tholeiitic and Fe-tholeiitic dike swarms of younger ages (2.45–1.98 Ga). On the other hand, groups 1, 2, and 5 of the 2.45 Ga dikes (later called BN dikes) have similar
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3
2500 2250
2.4
2000
Cr (ppm)
TiO2 (wt.%)
1750
1.8 1.2
1500 1250 1000 750
.6
500
0
250 0 0
5
10
15
20
25
0
5
MgO (wt.%) 2.45 Ga dike swarms Boninite–gabbronorite Low-Ti tholeiite Cr (ppm)
2.45 Ga Fe-tholeiite (NW), dated samples 2.32 Ga Fe-tholeiite (E-trending), dated samples
+
15
20
25
500
Orthopyroxene-phyric
E-trending Fe-tholeiitic swarms E -Wdike Fe-tholeiitic (2.45–1.98 Ga)
10
MgO (wt.%)
250
+
NW-trending Fe-tholeiitic dike swarms (2.45– 1.98 Ga)
0 2
2.6 3.2
3.8 4.4
5
5.6
MgO (wt.%)
6.2
6.8
Fig. 5.5. Chemical analyses (320 samples) of the Paleoproterozoic dike swarms (~2.45 Ga, 2.32 Ga, 2.1 Ga, and 1.98 Ga) from the Kuhmo block and Russian Karelia plotted in MgO-TiO2 and MgO-Cr diagrams.
geochemical trends showing a weak calcalkaline affinity, while groups 3 and 4 (later called TH dikes) have a typical tholeiitic trend with a nearly total overlap with younger dike swarms. One distinctive feature of the 2.45 Ga dike swarm compared with the younger dike swarms is the wider range of geochemical compositions. Major element diagrams in Figures 5.5–5.7 demonstrate that the BN dikes have their own trend with a calc-alkaline affinity, whereas the TH dikes and Fe-tholeiites plot in the tholeiitic field. All the younger dike swarms are homogeneous, normal tholeiites. In terms of TiO2 and Cr values (Figure 5.5), 208
the BN dikes fall into two groups: the first has low TiO2 (0.3–0.9%) and high Cr values (150–2300 ppm), the second group has moderate TiO2 (0.9–1.2%) and low Cr (<100 ppm). The 2.45 tholeiitic dikes show the same traits as the BN dikes (moderate TiO2 and low Cr, 50–250 ppm). The 2.45 Ga Fe-tholeiitic dikes have high TiO2 contents (~2.5%) combined with low Cr values (150–300 ppm) and plot in the same fields as all the younger (2.32–1.98 Ga) dike swarms. Jensen cation plots (Figure 5.6) also show that the 2.45 Ga dikes comprise two groups: the BN dikes and the TH dikes, the latter with high-Fe tholeiitic (HTF) affinity. Figure 5.7 shows that the BN dikes
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FeO* + TiO2 TA HFT TD TR CA CD CR
2.45 Ga dike swarms Boninite–gabbronorite Low-Ti tholeiite
BK
Orthopyroxene-phyric PK
Al2O3
MgO FeO* + TiO2
TA TD TR CA CD CR
Al2O3
HFT BK
+
E-trending Fe-tholeiitic dike swarms (2.45–1.98 Ga)
HMT PK
MgO
2.45 Ga Fe-thol. (NW-trending), dated samples 2.32 Ga Fe-thol. (E-trending), dated samples FeO* + TiO2
+
TA TD TR CA CD CR
HFT BK
NW-trending Fe-tholeiitic dike swarms (2.45– 1.98 Ga)
HMT
Al2O3
PK
MgO
Fig. 5.6. Same analyses as in Figure 5.5 plotted on Jensen’s (1976) cation plot.
have much higher SiO2 (50–57%) than the TH dikes. In addition, their FeO*/MgO ratio is, in general, much lower than that of the TH dikes. A calc-alkaline affinity in the BN dikes is also evident. Numerous tectonomagmatic discrimination diagrams have been developed for volcanic rocks, and have also been applied to dikes. A Ti/V diagram (Figure 5.8) clearly shows that the 2.45 Ga dikes (both BN and TH) are distinct from the Fe-tholeiites. They show an island arc (IAT) signature, whereas the Fe-rich tholeiites fall in the within-plate basalt field (WPB). All Fe-tholeiitic dikes fall into the continental flood basalt field in the Al2O3/TiO2 vs. Ti/Zr diagram (Figure 5.9). The BN dikes form a more diverse group plotting on both sides of the boundary between the boninite and volcanic arc basalt field, whereas the TH dikes fall in the mid-ocean ridge basalt (MORB) field. Th/Ta and La/Yb ratios have been used to identify magma sources for mafic dike swarms, particularly in order to distinguish
between the subcontinental lithospheric mantle and mantle plumes sources (Condie, 1997). Figure 5.10 shows selected analyses from the dikes of Kuhmo and Taivalkoski blocks and Russian Karelia. It is evident that the eastern Fennoscandian Paleoproterozoic dikes are compositionally similar to the dike swarms in other shields (Condie, 1997). The BN dikes are mainly in the noritic field and tholeiitic and Fe-tholeiitic dikes (2.45–1.98 Ga) plot in the tholeiitic field regardless of their age (2.45–1.98 Ga). Chondrite-normalized REE patterns for various dike swarms are shown in Figure 5.11. The REE patterns for Fe-tholeiitic dike swarms (2.45–1.98 Ga) are homogenous showing moderate LREE enrichment [(La/Yb)N ~ 2–3)]. The BN dikes usually have relatively high LREE/HREE ratios, with (La/Yb)N of 3–9 for boninite–norites, 4–5 for gabbronorite dikes, and 7–11 for orthopyroxene-plagioclase-phyric dikes. The TH dikes are usually slightly LREE-enriched but may also show flat or LREE-depleted patterns (Fig. 5.11).
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209
80 75
2.45 Ga dike swarms Boninite–gabbronorite Low-Ti tholeiite
SiO2 (wt.%)
70 65
Calc-alkaline
60
Tholeiite
Orthopyroxene-phyric
55
2.45 Ga Fe-tholeiite (NW-trending), dated samples 2.32 Ga Fe-tholeiite (E-trending), dated samples
50 45 0
.5
1
1.5
2
2.5
3
3.5
4
4.5
5
FeO*/MgO 80
80
E-trending Fe-tholeiitic dike swarms (2.45–1.98 Ga)
SiO2 (wt.%)
70
70
65 60
NW-trending Fe-tholeiitic dike swarms (2.45–1.98 Ga)
75
SiO2 (wt.%)
75
Calc-alkaline
Tholeiite
55 50
65 60
Calc-alkaline
Tholeiite
55 50
45 0
.5
1
1.5
2
2.5
3
3.5
4
4.5
5
FeO*/MgO
45 0
.5
1
1.5
2
2.5
3
3.5
4
4.5
5
FeO*/MgO
Fig. 5.7. Same analyses as in Figure 5.5 plotted on the FeO*/MgO vs. SiO2 diagram (Miyashiro, 1974).
Sm-Nd isotope studies carried out on dike groups 1–5 (this study and Vuollo and Huhma, 2004) yield some new information on the 2.45 Ga magmatism and its geochemical character. We obtained reliable ages for groups 1–4 (see Figure 5.12 and Table 5.2). Some preliminary Sm-Nd work has been done on the orthopyroxene-plagioclase-phyric dikes also (group 5 of the ~2.45 Ga dikes). However, the whole-rock plagioclase–orthopyroxene analyses indicate disequilibrium conditions. Our results (Figure 5.12 and Table 5.2) show that groups 1 and 2 represent a (boninitic) magma type with negative initial εNd values (from –2.4 to –1.8), consistent with the results recorded previously for the layered mafic intrusions (Huhma et al., 1990; Turchenko et al., 1991; Amelin et 210
al., 1996; Saini-Eidukat et al., 1997; Hanski et al., 2001; Chapter 3). The data on groups 3 and 4 (tholeiitic–Fe-tholeiitic magma type) show slightly positive initial εNd values (from +0.3 to +1.7). The Tsipringa layered intrusion in Russian Karelia has chondritic εNd values (Figure 5.12). These results suggest that at least the dikes of groups 1, 2, 3, and 4 may be consanguineous with the 2.45 Ga intrusions. Group 1 and 2 dikes could be related to boninite-like parental magmas (Cr-rich magma type and Cr-poor magma type, respectively) and group 3 and 4 to tholeiitic parental magmas; the latter is also the case for the layered intrusions of that age (the Tsipringa intrusion and the Western Koillismaa intrusions; Amelin et al., 1996; Iljina and Hanski, 2003).
• C H A P T E R 5 • PA L E O P R O T E R O Z O I C M A F I C D I K E S I N N E F I N L A N D
10
ARC
20
OFB WPS
100
0
5
10
15
V (ppm)
Ti/1000 (ppm) 650 600 550 500 450 400 350 300 250 200 150 100 50 0
2.45 Ga dike swarms Boninite–gabbronorite Low-Ti tholeiite
50
10
ARC
20
20
Orthopyroxene-phyric
25
E-trending Fe-tholeiitic dike swarms (2.45–1.98 Ga)
OFB WPS
50
100
0
5
10
2.45 Ga Fe-tholeiite (NW-trending), dated samples 2.32 Ga Fe-tholeiite (E-trending), dated samples NW-trending Fe-tholeiitic dike swarms (2.45–1.98 Ga) 650 10 ARC 20 OFB 600
V (ppm)
V (ppm)
650 600 550 500 450 400 350 300 250 200 150 100 50 0
15
20
25
Ti/1000 (ppm)
550 500 450 400 350 300 250 200 150 100 50 0
WPS
50
100
0
5
10
15
20
25
Ti/1000 (ppm)
Fig. 5.8. Same analyses as in Figure 5.5 plotted on the Ti/1000 vs.V (Shervais, 1982).
3.3. ~2.32 Ga dike swarm and intrusions Until recently, there have been few indications of ~2.3 Ga magmatic events in the eastern part of the Fennoscandian Shield – the first zircon age of 2331 ± 33 Ma was published by Paavola (1988). Geological evidence shows that the Runkaus Formation metavolcanic rocks in the Peräpohja area are older than the 2.2 Ga layered sills but younger than the 2.45 Ga layered intrusions (Perttunen and Vaasjoki, 2001). Huhma et al. (1990) obtained a Sm-Nd isochron age of 2330 ± 180 Ma for these metalavas. New age determinations (e.g., 2349 ± 30 Ma and 2332 ± 18 Ma, Vuollo et al., 2000, and this study; 2306 ± 6 Ma and 2319 ± 27 Ma, Vesa Nykänen, pers. comm., 2003; 2295
± 5 Ma Jorma Paavola, pers. comm., 2003) together with previous results clearly indicate that the ~2.32 Ga magmatism represents a significant magmatic event in the eastern part of the shield. Figure 5.14 shows the location of samples dated so far. Four Sm-Nd results are just over 2.3 Ga within errors (Figure 5.14) and U-Pb ages are approximately 2300 Ma (intrusions) and 2320 Ma (dikes). The areal distribution of the 2.32 Ga dikes and intrusions is difficult to estimate, because of few available ages and the fact that, geochemically, these igneous rocks are inconspicuous (see Figures 5.5–5.11). The dikes consist of plagioclase (50%), clinopyroxene (30–40%), quartz (5%), Fe-Ti-V oxides (5%), and minor amounts of olivine (< 2%),
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211
1000
2.45 Ga dike swarms Boninite–gabbronorite Low-Ti tholeiite
Ti/Zr
Continental flood basalt MORB Volcanic arc basalt
100 Ocean-island basalt
Orthopyroxene-phyric 2.45 Ga Fe-tholeiite (NW-trending), dated samples 2.32 Ga Fe-tholeiite (E-trending), dated samples
Boninite 10 1
10
100
Al2O3/TiO2 1000
1000
E-trending Fe-tholeiitic dike swarms (2.45–1.98 Ga) MORB
100
Continental flood basalt
Ti/Zr
Ti/Zr
Continental flood basalt
Volcanic arc basalt
MORB
100
Ocean-island basalt
NW-trending Fe-tholeiitic dike swarms (2.45–1.98 Ga)
Ocean-island basalt
Volcanic arc basalt
Boninite
Boninite 10
10 1
10
100
Al2O3/TiO2
1
10
100
Al2O3/TiO2
Fig. 5.9. Same analyses as in Figure 5.5 plotted on the Al2O3/TiO2 vs. Ti/Zr (Wilson and Versfeld, 1994) diagram.
biotite, apatite, and epidote. The composition of cloudy plagioclase (Figure 5.13H) is ~An50 and clinopyroxene is Ca-augite to subalkaline augite, averaging Wo30–35En40Fs30. The average composition of small olivine grains is ~Fo50. Aeromagnetic maps and field observations show that the trend of the dike swarm is roughly E (95˚) and that the dikes are broken and can be traced only for a few kilometers.
3.4. ~2.2 Ga layered sills and dikes The karjalitic layered sills and intrusions form a conspicuous mafic–ultramafic magmatic suite that is widely spread in eastern and northern Finland, including all the Karelian 212
schist belts (Figure 5.15). These differentiated sills and intrusions and/or parts of them have been called by a variety of names: karjalite (Väyrynen 1938; Vuollo and Piirainen, 1992), hypabyssal spilite (Piirainen 1969), and the gabbro–wehrlite association (Hanski, 1986). Various alternatives have also been proposed for the parental magma, including olivine basalt (Meriläinen, 1961), tholeiite (Piirainen, 1969), Fe-picrite (Hanski, 1986), and low-Al tholeiite (karjalite) (Vuollo and Piirainen, 1992). Ages have been obtained for karjalitic sills from Peräpohja (2210–2220 Ma, Perttunen, 1991; Perttunen and Vaasjoki, 2001), central Lapland (2220 ± 11 Ma, Tyrväinen,
• C H A P T E R 5 • PA L E O P R O T E R O Z O I C M A F I C D I K E S I N N E F I N L A N D
20 Norite swarms
Th/Ta
10
2.45 Ga dike swarms Boninite–gabbronorite Low-Ti tholeiite
Tholeiite swarms
Orthopyroxene-phyric 2.45 Ga Fe-tholeiite (NW-trending), dated samples 2.32 Ga Fe-tholeiite (E-trending), dated samples
1 .5
1
10
30
La/Yb 20
20 10
Norite swarms
Th/Ta
Th/Ta
10
Tholeiite swarms
E-trending Fe-tholeiitic dike swarms (2.45–1.98 Ga)
1 .5
1
10
Norite swarms
Tholeiite swarms
NW-trending Fe-tholeiitic dike swarms (2.45–1.98 Ga)
1
30
La/Yb
.5
1
10
30
La/Yb
Fig. 5.10. Selected chemical analyses (158 samples) of the Paleoproterozoic dike swarms (2.45 Ga, 2.32 Ga, 2.1 Ga, and 1.98 Ga) from the Kuhmo block and Russian Karelia plotted on the La/Yb vs. Th/Ta diagram (Condie, 1997).
1983; 2210 ± 4 Ma, Rastas et al., 2001; 2222 ± 6 Ma, Räsänen and Huhma, 2001; 2209 ± 10 Ma, Lauerma, 1995), Kuusamo (2206 ± 9 Ma, Silvennoinen, 1991; 2216 ± 4 Ma, Evins and Laajoki, 2001), and Kuhmo (2172 Ma, Hyppönen, 1983). The data on the Koli sill in North Karelia give a minimun age of 2170 Ma (Vuollo and Huhma, 2004). New U-Pb ion microprobe data also show ages slightly older than 2.2 Ga. These figures allow us to conclude that the emplacement of the karjalitic sills and dikes took place at ~2220 Ma (Figures 5.3 and 5.15). Stratigraphically, the 2.2 Ga sills and intrusions are restricted to the vicinity of the
unconformity between the Archean basement and Paleoproterozoic metasediments, and they are found within both of them. Most often they lie parallel to the overlying Karelian metasediment series (Karelian belt) and are thus referred to as sills. Most typically they are found as sills varying from a few kilometers up to 150 km in length and are quite thin (200–400 m) relative to their length. Later tectonic movements have fragmented the originally continuous intrusions. As exemplified by the Koli layered sill (see Figures 5.16 and 5.17), the karjalitic intrusions contain only one magmatic cycle and were crystallized from a single magma pulse
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213
Sample/Chondrite
100
2.45 Ga dike swarms Boninite–gabbronorite Low-Ti tholeiite
10
Orthopyroxene-phyric 2.45 Ga Fe-tholeiite (NW-trending), dated samples 2.32 Ga Fe-tholeiite (E-trending), dated samples
1 La Ce Pr Nd SmEuGd Tb Dy Ho Er Tm Yb Lu
10
100
Sample/Chondrite
Sample/Chondrite
100
E-trending Fe-tholeiitic dike swarms (2.45–1.98 Ga)
NW-trending Fetholeiitic dike swarms (2.45–1.98 Ga)
10
1
1 La Ce Pr Nd SmEuGd Tb Dy Ho Er Tm Yb Lu
La Ce Pr Nd SmEuGd Tb Dy Ho Er Tm Yb Lu
Fig. 5.11. Examples of chondrite-normalized REE patterns for 2.45 Ga, 2.32 Ga, 2.1 Ga, and 1.98 Ga dikes from the Kuhmo block and Russian Karelia.
that crystallized into a highly differentiated structure with a simple internal stratigraphy (e.g., the modified differentiation index, MDI, varies from 20 to 80; Figure 5.17). The layered structure of the sill is similar to that of many mafic–ultramafic layered intrusions, comprising a marginal series, a layered series with occasional rhythmic layering (Figure 5.18I), and a granophyre. The upper marginal zone consists of a fine-grained chilled margin (Figure 5.18J) and underlying coarse-grained clinopyroxenites (Figures 5.18G–H). Fractionation of the layered series proceeded normally, i.e., from the bottom upwards, but the order is reversed in the upper marginal zone (Figures 5.16 and 5.17). The layered series contains wehrlites and clinopyroxenites in the lower part (Figures 5.18A–D) and, in the upper part, gabbroic rocks (Figures 5.18D–E) grade via 214
coarse-grained gabbro to granophyre (Figure 5.18F). The granophyre crystallized from a residual magma beneath the upper marginal zone. The rocks of the layered series are cumulus-textured and include cumulus minerals olivine (Fo81–74), chromite, diopsidic clinopyroxene, titanian magnetite, plagioclase, and apatite. Intercumulus phases include edenitic hornblende, ilmenite, clinopyroxene, phlogopite, plagioclase, and titanian magnetite. Primary orthopyroxene occurs only in some 2.2 Ga intrusions in central Lapland (Hanski, 1987). Primary magmatic Ti-rich amphibole (edenitic hornblende) is found as an intercumulus phase in the karjalites and indicates high pH2O conditions during crystallization (Figure 5.16). Geochemically, the karjalitic sills form a peculiar magmatic series (cf. Figure 5.19). In
• C H A P T E R 5 • PA L E O P R O T E R O Z O I C M A F I C D I K E S I N N E F I N L A N D
Fe-tholeiite (NWtrending) Low-Ti-tholeiite Boninite–gabbronorite
3 2 1
εNd 0 -1 -2 -3 -4 0
1
(Nd/Sm)N
2
3
Kivakka (2445 ± 1 Ma) Tsipringa (2441 ± 1 Ma) Lukkulaisvaara (2442 ± 2 Ma) Penikat (2428 ± 35 Ma) Koitelainen (2439 ± 3 Ma) Akanvaara (2435 ± 7 Ma)
Fig. 5.12. Sm-Nd data for the 2.45 Ga Fennoscandian layered intrusions (Huhma et al., 1990; Amelin et al., 1996; Hanski et al., 2001, and Mutanen and Huhma, 2001) and dike swarms (Vuollo and Huhma, 2004).
the CMA diagram (Figure 5.19A), they show an almost continuous, curved trend from the MgO apex towards the Al2O3 apex. The curvature towards to the CaO corner is due to the dominant role of clinopyroxene as a cumulus phase. The Jensen cation plot (Figure 5.19B) displays a trend lying above the komatiitic and normal tholeiitic trends. The composition of the cumulates, granophyre, and the chilled margin is consistent with an overall LREE-enriched character of the parental magma (Figure 5.19C). This was a low-Al tholeiite (~10 wt.% Al2O3), i.e., karjalite, rich in iron (~13 wt.% FeOtot) and LREE [(La/Yb)N = 5.8]. A low Al2O3/TiO2 ratio of 5 to 6 is also a characteristic feature of the 2.2 Ga layered sills.
3.5. ~2.1 Ga dike swarms Emplacement of the ~2.1 Ga Fe-tholeiitic dikes was a widespread magmatic event that affected all parts of central and northern Finland. It
gave rise to a dense, predominantly NW- and E-trending dike network (Figure 5.20) that now intersects the Karelian formations and the Archean basement. New ages (see below) from the Kuhmo block and central Lapland show that there were many stages of dike intrusion between 2.1 Ga and 1.98 Ga. Regional differences can nevertheless be observed; e.g., the nortwesterly orientation is dominant in the North Karelia and Kainuu schist belts (Figure 5.20), whereas in the Archean basement of the Kuhmo block, the main orientation is to the east (Anttila et al., 1991; Kilpelä, 1991; this study). Unfortunately, the scarcity of outcrops and resultant lack of field data mean that the distribution of the dikes in Lapland must be presented on analogical grounds, based on their stratigraphic position and geochemistry. Available material (Lehtonen et al., 1989), however, suggests that the Fe-tholeiitic dikes are also the main group in Lapland. In the Archean basement (Kuhmo block), the Fe-tholeiitic dikes form a highly regular swarm in which individual dikes vary from a few centimeters up to 200 m in width with a typical range between 10 and 100 m. They can be followed for a few hundred meters to several tens of kilometers along the strike. On a large scale, the dike swarms have an ‘en echelon’ structure (cf. Rickwood, 1990). The dikes are homogeneous in their modal composition, contain principally hornblende and plagioclase, and have been referred to as metadiabases (e.g., Piirainen, 1969; Pekkarinen, 1979; Perttunen, 1991; Lehtonen et al., 1992). In the northern part of the Kuhmo block (on the eastern side of the Kuhmo greenstone belt), these ~2.1 Ga Fe-tholeiitic dikes have preserved their primary mineral composition. These dikes are ophitic and comprise 30–40% plagioclase (~An60), 30–50% clinopyroxene (Ca- to sub-alkaline-augite), and minor amounts of Fe-Ti-V oxide (5–10 %), quartz, biotite, and apatite. Plagioclase is typically stained (Figure 5.13I). Several U-Pb ages are available for the Fe-
C H A P T E R 5 • PA L E O P R O T E R O Z O I C M A F I C D I K E S I N N E F I N L A N D •
215
OLIV
OLIV OPX PLAG PLAG 500 m
500 m
B
A
CPX
CPX
500 m
1000 m
D
C
CPX
PLAG 1000 m
E
216
1000 m
F
• C H A P T E R 5 • PA L E O P R O T E R O Z O I C M A F I C D I K E S I N N E F I N L A N D
PLAG
CPX
OPX
1000 m
1000 m
H
G
PLAG PLAG
CPX 1000 m
I
1000 m
J
Fig. 5.13. Photomicrographs of rock types from Paleoproterozoic dike swarms. 2.45 Ga: (A) Boninite−norite dike, cumulus texture (sample VE3), crossed nicols; (B) Boninite–norite dike, cumulus texture with plagioclase grains covered by “tea-color cloudiness” (sample VE3); (C) Gabbronorite dike, sample near the contact, quench-texture (sample XD20); (D) Gabbronorite dike, clinopyroxene grain with pigeonite exsolution lamellae, fish-bone texture (sample XD6), crossed nicols; (E) Tholeiitic dike with a cumulus texture (sample 1-UD-93); (F) Fine-grained Fe-tholeiitic dike (sample WD9); (G) Orthopyroxene-phyric dike (sample 42-VEN-94). 2.32 Ga: (H) Fe-tholeiitic dike with cloudy feldspar (sample XD17). ~2.1 Ga: (I) Fe-tholeiitic dike with faintly cloudy feldspar (sample VEPO28–12.55). 1.98 Ga: (J) Fe-tholeiitic dike with faintly cloudy feldspar (sample KD12). Abbreviations: PLAG−plagioclase, OLIV−olivine, CPX−clinopyroxene, OPX−orthopyroxene. If not mentioned, the scale bar is 500 µm in length. Photos: Jouni Vuollo.
tholeiitic dikes in North Karelia (2115 ± 6 Ma, Pekkarinen, 1979; 2113 ± 4 Ma, Pekkarinen and Lukkarinen, 1992; 2105 Ma, Huhma, 1986), Peräpohja (2118 ± 14 Ma, Perttunen, 1987; 2117 ± 6 Ma, Perttunen and Vaasjoki,
2001), Varpaisjärvi (2106 ± 6 Ma, Jorma Paavola, pers. comm., 2003), and Kuusamo (2078 ± 8 Ma, Silvennoinen, 1991). These point to an overall age of ~2.1 Ga for the emplacement of the Fe-tholeiitic dikes (Figures 5.3 and
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360000
740000
740000
340000
6 2349±30 Ma 5 2332±18 Ma Pudasjärvi block
Suoperä– Paanajärvi area
Taivalkoski block
4 2319±27 Ma
720000
720000
2306±6 Ma
4
FINLAND
RUSSIA Kuhmo block
3 2350±50 Ma
2 Iisalmi block
2295±5 Ma
2.32 Ga dike swarms 2331±33 Ma
700000
700000
1
Unclassified (>1980 Ma) 2.32 Ga dike swarms
3
Sm-Nd age determination 2270±40 Ma
U-Pb age determination Archean Paleoproterozoic 340000
360000
Fig. 5.14. Distribution and location of dated samples of the ~2.32 Ga dike swarm in the
eastern and northern parts of the Fennoscandian Shield (Finland and Russian Karelia). Age data: 1−Paavola (1988); 2−Hölttä et al. (2000); 3−Jorma Paavola, pers. comm. (2003); 4−Osmo Nykänen, pers. comm. (2003); 5−Vuollo et al. (2000); 6−Vuollo and Huhma (2004). 218
• C H A P T E R 5 • PA L E O P R O T E R O Z O I C M A F I C D I K E S I N N E F I N L A N D
340000
10
360000
2185±35 Ma
7
2210±5 Ma
4 2210±4 Ma
10
2220±11 Ma
2231±30 Ma
1
7 2223±43 Ma
2213±5 Ma
3 2209±10 Ma
2 2217±10 Ma 2 5 2205±220 Ma
2
9 2216±18 Ma
6
2215±15 Ma
2206±9 Ma
2 2216±10 Ma
Pudasjärvi block
Taivalkoski block
Kuhmo block
720000
720000
2
740000
740000
2222±6 Ma 5
10 2187±44 Ma
7
2231±19 Ma
FINLAND
~2200 Ma
RUSSIA
8
8 Min 2186 Ma
2.2 Ga layered sills
Iisalmi block
7
7 Min 2172 Ma
2170±40 Ma
U-Pb age determination Sm-Nd age determination (ion microprobe) U-Pb age determination (conventional) Archean Paleoproterozoic 340000
Koli intrusion ~2200 Ma
700000
700000
2.2 Ga layered sills
3
7
2212±30 Ma 7
7
2203±49 Ma
360000
Fig. 5.15. Distribution and location of dated samples of the ~2.2 Ga layered intrusions/sills
in eastern and northern Finland. Age data: 1−Tyrväinen (1983); 2−Perttunen and Vaasjoki (2001); 3−Lauerma (1995); 4−Rastas et al. (2001); 5−Räsänen and Huhma (2001); 6−Silvennoinen (1991); 7−Vuollo and Huhma (2004); 8−Hyppönen (1983); 9−Evins and Laajoki (2001); and 10−Eero Hanski, pers. comm. (2003). C H A P T E R 5 • PA L E O P R O T E R O Z O I C M A F I C D I K E S I N N E F I N L A N D •
219
Straticgrafic height (m) 330
100% cumulus %intercumulus Oliv Chrom Flog Hbl Cpx Plag 0 50 100 50 100 (Fo%) (%Mg) (Ca:Mg:Fe) Magn aCp
Plag
+ + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + + +
Qz
300
pamC
250
Plag. amCp
200
150
Magn.
amCp Cpx
Plag
100
aCp
50 Olivine oaCh
Edhb Cpx+ Flog
0
o(c) Crom Cah
Wehrlite Wehrlite o(c)Cah oaCh Clinopyroxenite/ magnetite clinopyroxenite
80.4 80.5 81.1 81.0 80.4 77.0 74.4
72.6
63.5
68.4 65.1 60.0 69.0 44 42 14
63.8 44 41 15 68.0 43 43 14 69.1 43 43 13
70.1 43 45 12
43 46 11
71.6 43 47 10
71.3 43 48 9 71.7 43 49 8 73.4 45 48 7 79.7 79.6 81.5 82.5 79.2 81.1 79.2
Granophyre Magnetite gabbro
Zirc Apat Allan Quarz
Upper marginal zone (chilled margin and clinopyroxenite) Archean basement complex
Fig. 5.16. Stratigraphy of the ~2.2 Ga Koli sill, North Karelia. Solid line = cumulus mineral, dashed line = intercumulus mineral, plus sign = subsolidus mineral. Mineral abbreviations for cumulate (C) names after Irvine (1982): olivine (o), chromite (c), augite (a), edenitic hornblende (h), plagioclase (p), and titanian magnetite (m).
220
• C H A P T E R 5 • PA L E O P R O T E R O Z O I C M A F I C D I K E S I N N E F I N L A N D
Height (meters) 350
MDI
MgO (%)
Cr (%)
aCp 300
pamC 250
amCp
200
150
aCpm
100
50 oaCh 0(c)Cah 0 0
50
100
Wehrlite o(c)C Wehrlite oaC
0
20
40
0
0.2
0.4
Magnetite gabbro pamC Granophyre
Clinopyroxenite aC
Upper clinopyroxenite aC
Magnetite clinopyroxenite amC
Chilled margin
Fig. 5.17. Variation of MDI (modified differentiation index by von Gruenewaldt, 1973), MgO, and Cr as a function of stratigraphic height in the Koli layered sill.
5.20). New U-Pb ages from central Lapland (2046 ± 9 Ma, 2060 ± 8 Ma, and 2052 ± 7 Ma; Rastas et al., 2001; and 2054 ± 14 Ma and 2046 ± 18 Ma; Räsänen and Huhma, 2001) show that a significant magmatic event occurred also at ~2.05 Ga. This event is also registered by the Keivitsa and Otanmäki intrusions. For the dikes within the Archean basement, we have two Sm-Nd ages from the Kuhmo block: 2133 ± 33 Ma and 2054 ± 40 Ma (Vuollo et al. 2000, and this study).
Geochemically, the Fe-tholeiitic dikes form a relatively homogeneous group from North Karelia through Kainuu to Lapland (Vuollo et al., 1992; Vuollo, 1994, and this study). They are quartz-normative, sub-alkaline tholeiitic basalts (see Figures 5.5–5.11) and form a set of continental dike swarms of the type frequently found in shield areas (e.g., Tarney and Weaver, 1987).
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222
• C H A P T E R 5 • PA L E O P R O T E R O Z O I C M A F I C D I K E S I N N E F I N L A N D
Fig. 5.18. Photomicrographs of rock types in the Koli layered sill. (A) Olivine-(chromite) cumulate, sample 26-JIV-85; (B) Contact between olivine-(chromite) cumulate and olivine-clinopyroxene cumulate, sample 33B-JIV-85; (C) Clinopyroxene cumulate, sample 35-JIV-85; (D) Sharp contact between clinopyroxene-magnetite and clinopyroxene-plagioclase-magnetite cumulates, sample 43-JIV-85; (E) Laminated plagioclase-clinopyroxene-magnetite cumulate, sample 47-JIV-85; (F) Granophyre, sample 51-JIV-85; (G) Upper clinopyroxenite, sample 53-JIV-85; (H) Chilled margin, sample 54-JIV-85. Outcrops: (I) Rhythmic layering near the contact of wehrlite and clinopyroxenite (sample 31-JIV-85); (J) Upper chilled margin against Archean granite (sample 54-JIV-85). Abbreviations after Irvine (1982). Photos: Jouni Vuollo.
3.6. ~1.98 Ga dike swarm Previous studies (Vuollo, 1994) have shown that this dike swarm is not as voluminous as the ~2.1 Ga dike swarm. However, recent geochronological (Vuollo and Huhma, 2004) and field studies, combined with aeromagnetic data (Figure 5.2), show that the ~1.98 Ga dikes are found throughout the Archean basement and the Karelian formations (Figure 5.21). The swarm consists of up to 70-m-wide dikes that form prominent NW-trending (320–340˚), linear features >120 km in length (Kuhmo block). Data from the Veitsivaara nuclear waste study area (material from Posiva Co.) give valuable information on the trends and mutual age relationships of these ~2.1 Ga and ~1.98 Ga dike swarms (see Figures 5.2C and 5.22C). In this relatively small study area, the ~2.1 Ga dikes have an easterly trend (280˚) and the ~1.98 Ga dikes trend nortwest (330˚). We have used these data to make a generalization over the entire Kuhmo block. The 1.98 Ga dike swarm is the youngest
Paleoproterozoic mafic dike swarm observed within the Fennoscandian Archean craton. It is also the least deformed (Figure 5.3J) and is closely connected to the 1.95 Ga ophiolites (Kontinen, 1987; Vuollo et al., 1992; Hanski, 1997). These coarse- to medium-grained, ophitic dikes (Figure 5.13) consist of subhedral, coarse plagioclase (30–50%) in a matrix of fine-grained anhedral–subhedral clinopyroxene (20–30%), Fe-Ti-V oxide (5–20%) with quartz (3%), biotite (4%), uralitic amphibole (10%), and occasional orthopyroxene (1–3%). Plagioclase is turbid labradorite (An60), clinopyroxene is Ca-augite, and orthopyroxene varies from bronzite to hypersthene. In some places, fine-grained chilled margins of the dikes can be observed. There is one U-Pb baddeleyite age of 1981 ± 4 Ma from the Archean basement area (Figure 5.21; Vuollo et al., 1995a). Sm-Nd isotope data (a dike with the same NW-trend) give the same age value within errors (1992 ± 47 Ma; see Figures 5.2C and 5.22C). Other U-Pb ages from Lapland (2003 ± 4 Ma and 1995 ±
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A
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Fig. 5.19. Analyses of the low-Al tholeiites (Koli sill; see Fig. 5.15) plotted on (A) CMA, (B)
Jensen’s (1976) diagram, and (C) REE patterns for cumulates and chilled margin. 9 Ma; Rastas et. al., 2001) and North Karelia (1972 ± 5 Ma; Vuollo et al., 1992) are rather scattered but show that the ~1.98 Ga swarm is present throughout the eastern part of the Fennoscandian Shield. According to the earlier geochemical studies from North Karelia and the Kuhmo block (Vuollo et al., 1992) the 1.98 Ga dikes are Fetholeiitic to tholeiitic in composition (Figures 5.5–5.11). It is thus impossible to separate them geochemically from the older tholeiitic 224
dike swarms. Analyses from the Kuhmo block show that all the ~1.98 Ga dikes are typical continental Fe-rich tholeiites; tholeiites from the North Karelia schist belt, however, have a weak island arc tholeiitic (IAT) affinity (Vuollo et al., 1992).
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Fig. 5.20. Distribution and location of dated samples of the ~2.1 Ga dike swarms in eastern and northern Finland. Age data: 1−Perttunen (1987); 2−Perttunen and Vaasjoki (2001); 3−Tyrväinen (1983); 4−Rastas et al. (2001); 5−Räsänen and Huhma (2001); 6−Silvennoinen (1991); 7−Pekkarinen (1979); 8−Pekkarinen and Lukkarinen (1992); 9−Huhma (1986); 10−Jorma Paavola, pers. comm. (2003); and 11−Vuollo et al. (2000) and Vuollo and Huhma (2004). C H A P T E R 5 • PA L E O P R O T E R O Z O I C M A F I C D I K E S I N N E F I N L A N D •
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4. Tectonic significance of the dike swarms 4.1. Paleoproterozoic rifting events in the Archean Kuhmo block The earliest (~2.45 Ga) dikes were emplaced when a late Archean supercontinent (see Heaman, 1997) began to breakup. The large mafic layered intrusions were formed at this stage. These intrusions, which are known for their Cr and PGE ores, include the Tornio–Näränkävaara belt (Chapter 3) and the Koitelainen and Akanvaara intrusions (Chapter 4) in Finland and also the Oulanka complex (Lukkulaisvaara, Tsipringa, and Kivakka intrusions) near the Finnish border in Russia, and the Burakovka intrusion east of Lake Onega (Alapieti et al., 1990; Chapter 3). The boninite–noritic– gabbronoritic–tholeiitic–Fe-tholeiitic dikes, which cut the Archean basement, were formed slightly before or almost contemporaneously with the layered intrusions. The dikes are distributed over a broad area extending from Pudasjärvi via the Taivalkoski and Kuhmo blocks to Russian Karelia. Overall, the mafic layered intrusions and associated dike swarms appear to herald the breakup of the Archean crust and formation of the continental rift system at the beginning of the Paleoproterozoic Era. These 2.45 Ga intrusions and dikes are intimately associated with the volcano-sedimentary sequences of the Sumi group (Lehtonen et al., 1992; Manninen, 1991; Strand, 1993; Hanski and Huhma, Chapter 4). The metavolcanic rocks of this age are found in central Lapland, Salla–Kuusamo, and Kainuu (Räsänen et al., 1989; Manninen, 1991; Huhma et al., 1996). It is significant that the Sumi–Sariola felsic metavolcanic rocks near the layered Oulanka complex have yielded a U-Pb age of 2434 ± 24 Ma (Turchenko et al., 1991) and that potassic granites are associated with the 2.45 Ga mafic magmatism (Luukkonen, 1988; Lauri and Mänttäri, 2002; Russia: Stepanov, 1994). 226
The new data (Figures 5.4 and 5.22A) from the Kuhmo block and Russian Karelia indicate that paleostress directions varied within a short period of time around ~2.45 Ga. As shown in Figure 5.22A, the paleostress trends are orthogonal and point to small changes in the tension field. The reason for this variation may relate to processes in the mantle, e.g., rising mantle plumes, tapping of different levels of magma sources, etc. Sm-Nd isotope studies indicate that the earliest boninite-like parental magma type had negative initial εNd values, probably arising from an Archean subcontinental lithospheric mantle (Puchtel et al., 1997; Hanski et al., 2001). Tholeiitic and Fe-tholeiitic dikes (~2450 Ma) have roughly zero to positive εNd values and low Th/Ta and La/Yb ratios, indicating a depleted or primitive mantle source or a modest degree of crustal contamination. Young large igneous provinces (LIP) have been regarded as evidence for the presence of mantle plumes and/or active hotspots (e.g., White and McKenzie, 1989; Campbell and Griffiths, 1990; Ernst and Buchan, 1997; Eldholm and Coffin, 2000), whereas radiating giant mafic dike swarms have been used to identify mantle plumes in older cratonic regions such as the Canadian Shield (e.g., Ernst and Buchan, 2001b). However, such huge coherent regions (continuous Archean crustal blocks) are not present in eastern Fennoscandia (see Figures 5.2 and 5.23), and potential reactivation of the same tensional directions (Fig. 5.22) makes it difficult to identify mantle plume centers. Figure 5.22B shows the main paleostress trends for the ~2.1 Ga and ~1.98 Ga Fe-tholeiitic dike swarms, indicating a difference of ~40–50º in their extensional fields. The 1.98–1.95 Ga mafic magmatic events were extremely significant for the ore-forming processes of the Fennoscandian Shield, as indicated by the ore-bearing ophiolite formations (a breakup event at ~2.0 Ga) in Finland. The latest Paleoproterozoic Fe-tholeiitic–tho-
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Fig. 5.21. Distribution and location of dated samples of the ~1.98 Ga dike swarms in eastern
and northern Finland. Age data: 1−Vuollo et al. (1992); 2−Rastas et al. (2001); 3−Vuollo et al. (2000); and 4−Vuollo and Huhma (2004).
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leiitic dike swarms (1.98–1.97 Ga) formed throughout the Archean continental crust almost contemporaneously. However, it must be remembered that some other dike swarms, such as the 2.45 Ga Fe-tholeiites and tholeiites and 2.32 Ga Fe-tholeiites could also show the same paleostress directions as the younger swarms. The ~2.13–2.05 Ga Fe-tholeiitic dike swarms running in an E–W direction throughout the Archean crust are indicative of pronounced crustal extension and incipient rifting. A clear thinning of the continental crust towards the west, together with marine sedimentary environments, is detectable in the marine Jatuli formations of the Tohmajärvi area in the southern part of the North Karelia schist belt, whereas the Ilomantsi area (Archean craton) farther east represents thicker continental crust (see Nykänen et al., 1994).
4.2. Uplifted Archean high-grade terranes The eastern part of the Fennoscandian Shield contains several Archean cratonic blocks surrounded by Proterozoic cover rocks. Archean cratons and especially their margins have been demonstrated to be reworked, e.g., show reset isotope ages (Kontinen et. al., 1992; Kontinen, 2002). Hence, dike swarms can be used to monitor the postcratonization history of Archean shields (Halls and Zhang, 1998). For example, petrographic (cloudy feldspar) and paleomagnetic (magnetic polarity) data on dikes indicate that Paleoproterozoic (~2.0 Ga) uplift has occurred in the interior of the Superior Province. A few speculations can be made about the uplifted Archean high-grade terranes in the Kuhmo, Taivalkoski, and Pudasjärvi blocks and Russian Karelia in view of our dike swarm studies and knowledge of high-grade terranes (Varpaisjärvi: Hölttä et al., 2000; Russian Karelia: Korsakova et. al., 1987, Ylo Systra, pers. comm., 2003; Pudasjärvi block: Huhma 228
and Mutanen, 2002; Taivalkoski block: Kontinen et al., 1992, and Jorma Räsänen, pers. comm., 2003). Dike swarm investigations together with previous field surveys have clearly defined areas of high-grade metamorphism (uplifted Archean granulites?) in different parts of the Fennoscandian Shield (Figure 5.23). One indication of these “dry Archean areas” is the preservation of primary magmatic minerals (cloudy feldspar, see Figure 5.13, and fresh pyroxenes) in the examined dike swarms. However, paleomagnetic studies (Mertanen, 1995; Chapter 15) have shown that all Paleoproterozoic dike swarms exhibit strong Svecofennian overprinting. Figure 5.23 shows the proposed uplifted Archean highgrade terranes based on our dike swarm studies and existing field data. The most extensive areas (the Kuusamo–Pääjärvi block and the Vodlozero–Viianki block) are located near the Finnish-Russian border and can be clearly delineated on aeromagnetic maps. Kontinen (2002) showed that the granulites of the Varpaisjärvi block were uplifted in the Archean time. On the other hand, in some parts of the Kuhmo block, all the examined dikes have cloudy feldspars, suggesting the same deep erosional level. This means that the uplift of the Archean granulites may have occurred after emplacement of the Paleoproterozoic dike swarms. Dike swarms are ideal tools for reconstructing Precambrian crustal blocks, because they provide information on both long-term and variable paleostress directions (e.g., Halls and Palmer, 1990; Neuvonen et al., 1997). Paleomagnetic studies from the Varpaisjärvi block (Neuvonen et al., 1997) show clear evidence for rotation of the granulitic crustal blocks by ~16˚. According to Neuvonen et al. (1997), the rotation of the blocks took place before intrusion of the dikes. Further integrated studies (including geochemistry, age determination, paleomagnetism, seismic reflection surveys, and field mapping) will probably clarify how the various crustal blocks
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Fig. 5.22. The 2.45 Ga, 2.1 Ga, and 1.98 Ga mafic dikes and related paleotress directions (heavy arrows) in the Kuhmo block, east-central Finland.
were reassembled and if any rotation or tilting was, in fact, involved.
Acknowledgments The work presented here was carried out at the Department of Geology, University of Oulu,
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Fresh to partly fresh primary pyroxenes Fig. 5.23. Uplifted Archean granulites drawn according to aeromagnetic maps and mineralogical (cloudy feldspar and primary mineral composition) studies of dike rocks in the Taivalkoski, Pudasjärvi, and Kuhmo blocks.Varpaisjärvi granulites according to Korsman et al. (1997).
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Finland and, Geological Survey of Finland, Rovaniemi, in the framework of the research project “Early Proterozoic mafic magmatism and related ore deposits in eastern and northern Finland”. The project was financed by the Ministry of Trade and Industry, Geological Survey of Finland, and the University of Oulu. The Academy of Finland and the Finnish International Geological Correlation Programme (IGCP) committee are also thanked for financial support. We wish to express our sincere gratitude to Professor Emeritus Tauno Piirainen for numerous discussions during the project and Professor Eero Hanski for reading the manuscript and making many valuable suggestions. Eero Hanski, Asko Kontinen, Vesa Nykänen, and Jorma Paavola are acknowledged for putting their unpublished age data at our disposal. We also thank Viena Arvola for drawing the diagrams.
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with English abstracts and figure and table captions) Cadman, A., Tarney, J., Park, R.G., 1990. Intrusion and crystallisation features in Proterozoic dyke swarms. In: A.J. Parker, P.C. Rickwood, D.H. Tucker (Eds.), Mafic Dykes and Emplacement Mechanisms. Balkema, Rotterdam, 13–24. Campbell, I.H., Griffiths, R.W., 1990. Implications of mantle plume structure for the evolution of flood basalts. Earth Planet. Sci. Lett. 99, 79–93. Condie, K.C., 1997. Sources of Proterozoic mafic dyke swarms: constraints from Th/Ta and La/Yb ratios. Precambrian Res. 81, 3–14. Eldholm, O., Coffin, M.F., 2000. Large igneous provinces and plate tectonics. In: M.A. Richards, G. Gordon, R.D. van der Hilst (Eds.), The history and dynamics of global plate motions. Am. Geophys. Union, Geophys. Monogr. 121, 309–326. Ernst, R.E., Buchan, K.L., 1997. Giant radiating dyke swarms; their use in identifying pre–Mesozoic large igneous provinces and mantle plumes. In: J.J. Mahoney, M.F. Coffin (Eds.), Large Igneous Provinces; Continental, Oceanic, and Planetary Flood Volcanism. Geophys. Monogr. 100, 297–333. Ernst, R.E., Buchan, K.L., 2001a. Large mafic magmatic events through time and links to mantle–plume heads. In: R.E. Ernst, K.L. Buchan (Eds.), Mantle Plumes: Their Identification Through Time. GSA Spec. Pap. 352, 483–575. Ernst, R.E., Buchan, K.L., 2001b. The use of mafic dike swarms in identifying and locating mantle plumes. In: R.E. Ernst, K.L. Buchan (Eds.), Mantle Plumes: Their Identification Through Time. GSA Spec. Pap. 352, 247–265. Ernst, R.E., Buchan, K.L., West, T.D., Palmer, H.C., 1996. Diabase (Dolerite) dyke swarms of the World, first edition Geol. Surv. Canada Open–File 3241, includes map, scale 1:35 000 000 at equator, 1–104. Evins, P.M., Laajoki., 2001. Age of the Tokkalehto metagabbro and its significance to the lithostratigraphy of the early Proterozoic Kuusamo supracrustal belt, Northern Finland. Bull. Geol. Soc. Finland 73, 5—16. Fahrig, W.F., 1987. The tectonic settings of continental mafic dyke swarms: Failed arm and early passive margin. In: H.C. Halls, W.F.
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Fahrig (Eds.), Mafic dyke swarms. Geol. Assoc. Can., Spec. Pap. 34, 331–348. Gorbatschev, R., Lindh, A., Solyom, Z., Laitakari, I., Aro, K., Lobach–Zhuchenko, S.B., Markov , M.S., Ivliev, A.I., Bryhni, I., 1987. Mafic Dyke Swarms of the Baltic Shield. In: H.C. Halls, W.F. Fahrig (Eds.), Mafic dyke swarms. Geol. Assoc. Can., Spec. Paper 34, 361–372. Gruenewaldt, G. von, 1973. The modified differentiation index and the modified crystallization index as parameters of differentiation in layered intrusions. Trans. Geol. Soc. S. Africa 76, 53–61. Halls, H.C., Fahrig, W.F. (Eds.), 1987. Mafic Dyke Swarms. Geol. Assoc. Can., Spec. Pap. 34, 1–503. Halls, H.C., Palmer, H.C., 1990. The tectonic relationship of two Early Proterozoic dyke swarms to the Kapuskasing Structural Zone: a paleomagnetic and petrographic study. Canadian J. Earth Sci. 27, 87–103. Halls, H.C., Zhang, B., 1998. Uplift structure of the southern Kapuskasing zone from 2.45 Ga dike swarm displacement. Geology 26, 67–70. Hanski, E., 1984. Geology of the gabbro–wehrlite association in the eastern part of the Baltic Shield. University of Oulu, Arkeeisten alueiden malmiprojekti, Oulun yliopisto. Report 20, 1–78. Hanski, E., 1986. Gabbro–Wehrlite association in the eastern part of the Baltic shield. In: G.H. Friedrich, A.D. Genkin, A.J. Naldrett, J.D. Ridge, R.H. Sillitoe, F.M. Vokes (Eds.), Geology and Metallogeny of Copper Deposits, Society for Geology Applied to Ore Deposits Berlin-Heidelberg, Springer Verlag, Spec. Publ. 4, 151–170. Hanski, E., 1987. Differentioituneet albiittidia baasit – gabro-wehrliittiassosiaatio. Summary: Differentiated albite diabases – gabbro wehrlite association. In: K. Aro, I. Laitakari (Eds.), Suomen diabaasit ja muut mafiset juonikivilajit – Diabases and other mafic dyke rocks in Finland. Geol. Surv. Finland, Rep. Invest. 76, 35–44. Hanski, E., 1997. The Nuttio serpentinite belt, Central Lapland: An example of Paleoproterozoic ophiolitic mantle rocks in Finland. Ofioliti 22, 35–46. Hanski, E., Walker, R.J., Huhma, H., Suominen, I., 2001. The Os and Nd isotopic systematics of the c. 2.44 Ga Akanvaara and Koitelainen
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mafic layered intrusions in northern Finland. Precambrian Res. 109, 73–102. Heaman, L.M., 1997. Global mafic magmatism at 2.45 Ga: Remnants of an ancient large igneous province? Geology 25, 299–302. Hölttä, P., Huhma, H., Mänttäri, I., Paavola, J., 2000. P–T–t development of Archaean granulites in Varpaisjärvi, central Finland II. Dating of high-grade metamorphism with the U–Pb and Sm–Nd methods. Lithos 50, 121–136. Huhma, H., 1986. Sm-Nd, U-Pb and Pb-Pb isotopic evidence for the origin of the Early Proterozoic Svecokarelian crust in Finland. Geol. Survey Finland, Bull. 337, 1–48. Huhma, H., Cliff, R.A., Perttunen, V., Sakko, M., 1990. Sm–Nd and Pb isotopic study of mafic rocks associated with early Proterozoic continental rifting: the Peräpohja schist belt in northern Finland. Contrib. Mineral. Petrol. 104, 369–379. Huhma, H., Mutanen, T., 2002. Oldest rocks of the Fennoscandian Shield in the Pudasjärvi granulite belt. In: K. Korkka-Niemi (Ed.), Geologian tutkijapäivät 13.-14.3.2002 Helsinki: ohjelma, tiivistelmät, osallistujat. Helsinki: Geologian valtakunnallinen tutkijakoulu, 22–23. Huhma, H., Mutanen, T., Hanski, E., Räsänen, J., Manninen, T., Lehtonen, M., Rastas, P., Juopperi, H., 1996. Sm–Nd isotopic evidence for contrasting sources of the prolonged Palaeoproterozoic mafic–ultramafic magmatism in northern Finland. In: IGCP project 336 symposium in Rovaniemi, Finland, August 21–23, 1996: program and abstracts. Turun yliopiston geologian ja mineralogian osaston julkaisuja 38, 17. Hyppönen, V., 1983. Ontojoen, Hiisijärven ja Kuhmon kartta-alueiden kallioperä. Summary: Pre-Quaternary rocks of the Onto joki, Hiisijärvi and Kuhmo map-sheet areas. Suomen geologinen kartta 1:100 000. Kallioperäkarttojen selitykset. 4411 Ontojoki, 4412 Hiisijärvi and 4413 Kuhmo. Geol. Surv. Finland. 1–60. Iljina, M.J., Alapieti, T.T., McElduff B.M., 1992. Platinum-group element mineralization in the Suhanko–Konttijärvi intrusion, Finland. Austr. J. Earth Sci. 39, 303–313. Iljina, M., Hanski, E., 2003. Multimillion-ounce PGE deposits of the Portimo layered igneous complex, Finland. 9th International Platinum Symposium, July 21—25, 2002,
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Billings, Montana, USA. Abstracts pdf. 4 p. Irvine, T.N., 1982. Terminology for layered intrusions. J. Petrol. 23, 127–162. Janhila, M., 2001. Fennoskandian kilven itäosan emäksisten juoniparvien (2440–1980 Ma) mineralogia. M.Sc. Thesis, University of Oulu, Finland. 1–122. (in Finnish) Jensen, L.S., 1976. A new cation plot for classifying subalkalic volcanic rocks. Ontorio Dep. Min., Misc. Pap. 66, 1–22. Kilpelä, M., 1991. Itä-Kainuun varhaisproterotsooiset diabaasit. M.Sc. Thesis. University of Turku, Finland. 1–113. (in Finnish) Kontinen, A, 1987. An early Proterozoic ophiolite – the Jormua mafic–ultramafic complex, northeastern Finland. In: G.Gaál, R. Gorbatschev (Eds.), Precambrian geology and evolution of the Central Baltic Shield. Special Issue, Precambrian Res. 35, 313–341. Kontinen, A., 2002. Proterozoic tectonothermal over print in the eastern Finland Archaean complex and some thoughts of its tectonic setting. In: K. Korsman, P. Lestinen, (Eds.) Raahe–Laatokka –symposio. Kuopio 20.–21.3.2001. Laajat abstraktit. 12. Geologian tutkimuskeskus, arkistoraportti, K 21.42/2002/1, 42–62. (in Finnish) Kontinen, A., Paavola, J., Lukkarinen, H., 1992. K-Ar ages of hornblende and biotite from Late Archaean rocks of eastern Finland— interpretation and discussion of tectonic implications. Geol. Surv. Finland, Bull. 365, 1–31. Korsakova, M.A., Ivanov, N.M., Vakar, Ye.V., Muradymov, G.S., Lebedeva, N.I., 1987. Schematic tectono–geologic map of the Baltic (Fennoscandian) Shield. A supplement map to M. Korsakova, V. Proskuryakov, N. Ivanov, J. Nuutilainen (1988). The Archaean greenstone belts of the Baltic Shield. In: E. Marttila (Ed.), Archaen geology of the Svennoscandian Shield. Proceedings of a Finnish-Soviet Symposium in Finland on July 28–August 7, 1986. Geol. Surv. Finland, Spec. Pap. 4, 21–27. Korsman, K., Koistinen, T., Kohonen, J., Wennerström, M., Ekdahl, E., Honkamo, M., Idman, H., Pekkala, Y., 1997. Bedrock map of Finland 1:1000 000. Geological Survey of Finland, Espoo. Lauerma, R., 1995. Kursun ja Sallan kartta-alueiden kallioperä. Summary: Pre-Quaternary rocks of the Kursu and Salla map-sheet ar-
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evolution of the Fennoscandian Shield – a palaeomagnetic study with emphasis on the Svecofennian orogeny. Ph. D. Thesis with original articles (I-IV). Geol. Surv. Finland, Espoo. 1–187. Mertanen, S., Halls, H. C., Vuollo, J. I., Pesonen, L. J., Stepanov, V. S., 1999a. Paleomagnetism of 2.44 Ga mafic dykes in Russian Karelia, eastern Fennoscandian Shield – implications for continental reconstructions. Precambrian Res. 98, 197–221. Mertanen, S., Vuollo, J. I., Halls, H. C., Pesonen, L. J., Huhma, H., Kamo, S., Hölttä, P., Paavola, J., Stepanov, V. S., 1999b. Palaeomagnetism of 2.0–2.4 Ga mafic dykes in the eastern Fennoscandian Shield. In: European Geophysical Society 24th General assembly: society symposium, solid earth geophysics & geodesy. Geophys. Res. Abstracts 1 (1), 144. Miyashiro, A., 1974. Volcanic rock series in island arcs and active continental margins. American J. Sci. 274, 321–355. Mutanen, T., Huhma, H., 2001. U-Pb geochronology of the Koitelainen, Akanvaara and Keivitsa layered intrusions and related rocks. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Pap. 33, 229–246. Neuvonen, K.J., Pesonen, L.J., Pietarinen, H., 1997. Remanent magnetization in the Archaean basement and cutting diabase dykes in Finland, Fennoscandian Shield. In: L.J. Pesonen (Ed.), The Lithosphere in Finland – A Geophysical Perspective. Geophysica 33, 111–146. Nykänen, V.M., Vuollo, J.I., Liipo, J.P., Piirainen, T.A., 1994. Transitional (2.1 Ga) Fe–tholeiitic – tholeiitic magmatism in the Fennoscandian Shield signifying lithospheric thinning during Palaeoproterozoic extensional tectonics. Precambrian Res. 70, 45–65. Paavola, J., 1988. Lapinlahden kartta-alueen kallioperä. Summary: Pre-Quaternary rocks of the Lapinlahti map-sheet area. Suomen geologinen kartta 1:100 000, kallioperäkarttojen selitykset, lehti 3332 Lapinlahti. Geol. Surv. Finland. 1–60. Pekkarinen, L.J., 1979. The Karelian formations and their depositional basement in the Kiihtelysvaara–Värtsilä area, East Finland. Geol. Surv. Finland, Bull. 301, 1–141.
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Pekkarinen, L.J., Lukkarinen, H., 1992. Paleoproterozoic volcanism in the Kiihtelysvaara – Tohmajärvi district, eastern Finland. Geol. Surv. Finland, Bull. 357, 1–30. Peltonen, P., Kontinen, A., Huhma, H., 1996. Petrology and geochemistry of metabasalts from the 1.95 Ga Jormua Ophiolite, northeastern Finland. J. Petrol. 37, 1359–1383. Perttunen, V., 1987. Pudasjärven graniittigneissikompleksin luoteisosan mafiset juonet. Abstract: Mafic dykes in the northwestern part of the Pudasjärvi Granite Gneiss Complex. In: K. Aro, I. Laitakari (Eds.), Suomen diabaasit ja muut mafiset juonikivilajit – Diabases and other mafic dyke rocks in Finland. Geol. Surv. Finland, Rep. Invest. 76, 29–34. Perttunen, V., 1991. Kemin, Karungin, Simon ja Runkauksen kartta-alueiden kallioperä. Summary: Pre-Quaternary rocks of the Kemi, Karunki, Simo and Runkaus mapsheet areas. Geological map of Finland, 1:100 000. Explanation to the maps of Pre-Quaternary rocks, sheets 2541 Kemi, 2542+2524 Karunki, 2543 Simo and 2544 Runkaus. Geol. Surv. Finland. 1–80. Perttunen, V., Vaasjoki, M., 2001. U-Pb geochronology of the Peräpohja Schist Belt, northwestern Finland. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Pap. 33, 45–84. Piirainen, T., 1969. Initialer Magmatismus und seine Erzbildung in der Beleuchtung des Koli–Kaltimogebiets. Bull. Geol. Soc. Finland 41, 21–45. Puchtel, I S., Haase, K.M., Hofmann, A.W., Chauvel, C., Kulikov, V.S., Garbe–Schönberg, C.-D., Nemchin, A.A. 1997. Petrology and geochemistry of crustally contaminated komatiitic basalts from the Vetreny Belt, southeastern Baltic Shield: Evidence for an early Proterozoic mantle plume beneath rifted Archean continental lithosphere. Geochim. Cosmochim. Acta 61, 1205–1222. Räsänen, J., Hanski, E., Lehtonen, M.I., 1989. Komatiites, low-Ti basalts and andesites in the Möykkelmä area, Central Finnish Lapland. Report of the Lapland Volcanite Project. Geol. Surv. Finland, Rep. Invest. 88, 1–41. Räsänen, J., Huhma, H., 2001. U-Pb datings in the
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Sodankylä schist area, central Finnish Lapland. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Pap. 33, 153–188. Rastas, P., Huhma, H., Hanski, E., Lehtonen, M.I., Härkönen, I., Kortelainen, V., Mänttäri, I., Paakkola, J., 2001. U-Pb isotopic studies on the Kittilä greenstone area, central Lapland, Finland. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Pap. 33, 95–141. Rickwood, P.C., 1990. The anatomy of a dyke and the determination of propagation and magma flow directions. In: A.J. Parker, P.C. Rickwood, D.H. Tucker (Eds.), Mafic dykes and emplacement mechanisms; proceedings of the Second international dyke conference. A.A. Balkema, Rotterdam–Brookfield. Vol. 2, 81–100. Saini-Eidukat, B., Alapieti, T.T., Thalhammer, O.A.R., Iljina, M.J., 1997. Siliceous, highmagnesian parental magma compositions of the PGE-rich Early Paleoproterozoic layered intrusion belt of northern Finland. In: P. Rongfu (Ed.), Proceedings of the 30th International Geological Congress, Beijing, China, 4–14 August 1996. Vol. 9: Energy and Mineral Resources for the 21st Century, Geology of mineral deposits, mineral economics. VSP, Utrecht, 177–197. Salmirinne, H., 2001. Paikkatietojärjestelmän hyödyntäminen geofysikaalisissa ja geologisissa tutkimuksissa. Suomen mafisten juo ni kivien paikkatietokanta. Phil.Lic. Thesis, University of Oulu, Finland, 1–62. (in Finnish) Shervais, J.W., 1982. T–V plots and the petrogenesis of modern and ophiolitic lavas. Earth Planet. Sci. Lett. 59, 101–118. Silvennoinen, A., 1991. Kuusamon ja Rukatunturin kartta-alueiden kallioperä. Summary: Pre-Quaternary rocks of the Kuusamo and Rukatunturi map-sheet areas. Geological map of Finland 1:100 000. Explanation of the Maps of Pre-Quaternary rocks, 4524 + 4542 Kuusamo and 4613 Rukatunturi. Geol. Surv. Finland. 1–62. Stepanov, V.S., 1994. Magmatiti Pjaozerskovo blocka (petrohimisekie osobennosti I poste-
dovatelnost obrazovania kompleks). (Magmatic rocks of the Pjaozerskij block (petrochemical characteristics and formation of the complex)). In: V.S. Stepanov (Ed.), Dokembrii Severnoi Karelii (Precambrian of the Northern Karelia). Petrozavodsk, 118–170. (in Russian). Strand, K., 1993. Sedimentation in a Palaeoproterozoic (Karelian) rift to cratonic margin transition, Finland. Acta Univ. Ouluensis A 242, 1–35. Tarney, J., Weaver, B.L., 1987. Geochemistry and petrogenesis of Early Proterozoic dyke swarms. In: H.C. Halls, W.F. Fahrig (Eds.), Mafic Dyke Swarms. Geol. Ass. Canada, Spec. Pap. 34, 81–94. Talvitie, J., Paarma, H., 1980. Precambrian basic magmatism and the Ti-Fe ore formation in central and northern Finland. In: J. Siivola (Ed.), Metallogeny of Baltic Shield. Proceedings of the symposium held in Helsinki, Finland, June 12–21, 1978. Geol. Surv. Finland, Bull. 307, 98–107. Turchenko, S.I., Semenov, V.S., Amelin, Ju.V., Levchenkov, O.A., Neymark, L.A., Buiko, A.K., Koptev–Dvornikov, Je.V., 1991. The Early Proterozoic riftogenic belt of Norhern Karelia and associated Cu-Ni, PGE and Cu-Au mineralizations. Geol. För. Stockholm Förhandl. 113, 70–72. Tyrväinen, A., 1983. Sodankylän ja Sattasen kartta-alueiden kallioperä. Summary: PreQuaternary rocks of the map-sheet areas of Sodankylä and Sattanen. Suomen geologinen kartta 1:100 000. Kallioperäkarttojen selitykset, 3713 Sodankylä, 3714 Sattanen. Geol. Surv. Finland. 1–59. Väyrynen, H., 1938. Notes on the geology of Karelia and the Onega region in the summer of 1937. Bull. Comm. géol. Finlande 123, 66–80. Vogel, D.C., Vuollo, J.I., Alapieti, T.T., James, R.S., 1998. Tectonic, stratigraphic, and geochemical comparisons between ca. 2500–2440 Ma mafic igneous events in the Canadian and Fennoscandian Shields. Precambrian Res. 92, 89–116. Vuollo, J., 1994. Palaeoproterozoic basic igneous events in Eastern Fennoscandian Shield between 2.45 Ga and 1.97 Ga, studied by means of mafic dyke swarms and ophiolites in Finland. Acta Univ. Ouluensis Ser. A 250, 1–47. Vuollo, J., Huhma, H., 2004. Radiometric age de-
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terminations from mafic intrusions, dykes and their bearing on the timing of Precambrian mafic magmatic events. Geol. Surv. Finland, Spec. Pap. (in prep.) Vuollo, J., Piirainen, T., 1992. The 2.2 Ga old Koli layered sill: The low-Al tholeiitic (karjalitic) magma type and its differentiation in northern Karelia, eastern Finland. Geol. Fören. Stockholm Förh. 114, 131–142. Vuollo, J., Piirainen, T., Huhma, H., 1992. Two Early Proterozoic tholeiitic diabase dyke swarms in the Koli–Kaltimo area, Eastern Finland – their geological significance. Geol. Surv. Finland, Bull. 363, 1–32. Vuollo, J., Kamo, S., Mertanen, S., Halls, H., Stepanov, V., Nykänen, V., Pesonen, L., 1995a. U–Pb geochronology and paleomagnetism of mafic dyke swarms in the Fennoscandian Shield: a Canadian Shield connection. In: A. Agnon, G. Baer (Eds.), Third International Dyke Conference, September 4–8, 1995, Jerusalem, Israel. Program & Abstracts, 79. Vuollo, J.I., Nykänen, V.M., Liipo, J.P., Piirainen, T. A., 1995b. Palaeoproterozoic mafic dyke swarms in the Eastern Fennoscandian Shield, Finland: a review. In: G. Baer, A. Heimann (Eds.), Physics and Chemistry of Dykes. A. A. Balkema, Rotterdam, 179–192. Vuollo, J.I., Salmirinne, H., Pesonen, L., Stepanov, V. Fedotov, G., Philippov, N., 1999. The eastern Fennoscandian mafic dyke swarms GIS–database – a tool for integrated geoscientific studies. In: Sudbury 1999. GAC/MAC Joint Annual Meeting, Sudbury, Ontario, 26th to 28th May, 1999: Abstract Volume. GAC/MAC Annual Meeting 24, 133.
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Vuollo, J., Huhma, H., Pesonen, L.J., 2000. Mafic dyke swarms – geological evolution of the Palaeoproterozoic in the Fennoscandian Shield. In: L.J. Pesonen, A. Korja, S.-E. Hjelt (Eds.), Lithosphere 2000 - A symposium on the structure, composition and evolution of the lithosphere in Finland. Programme and extended abstracts, Espoo, Finland, October 4–5, 2000. Institute of Seismology, University of Helsinki. Report S–41, 107–111. Vuollo, J.I., Salmirinne, H., Pesonen, L., Stepanov, V. Fedotov, G., Frank–Kamenetsky, D 2001. The Eastern Fennoscandian Shield Mafic Dyke Swarms GIS Databases. In: M. Watkeys (Ed.), 4th International Dyke Conference, 26–29 June, 2001, Ithala Game Reserve KwaZulu–Natal, South Africa, Programme and Abstracts, 2. Vuollo, J.I., Huhma, H., Stepanov, V., Fedotov, G., 2002. Geochemistry and Sm-Nd isotope studies of a 2.45 Ga dyke swarm: hints at parental magma composition and PGE potential to Fennoscandian layered intrusions. In: A. Boudreau (Ed.), 9th International Platinum Symposium, 21–25 July, 2002, Billings, Montana, USA: Extended Abstracts. Billings, MT: Duke University, 469–470. White, R., McKenzie, D., 1989. Magmatism at rift zones; the generation of volcanic continental margins and flood basalts. J. Geophys. Res. 94, 7685–7729. Wilson, A.H., Versfeld, J.A., 1994. The early Archaean Nondweni greenstone belt, southern Kapvaal Craton, South Africa, Part II. Characteristics of the volcanic rocks and constraints on magma genesis. Precambrian Res. 67, 227–320.
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Chapter 6
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P. Peltonen
Cover page: Sheeted dike complex consisting of subparallel EMORB dikes. Plagioclasephyric dikes (with a drill hole) are cut by sligthly younger aphyric dikes with chilled margins. The Jormua ophiolite. Photo: Asko Kontinen.
Peltonen, P., 2005. Ophiolites. 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. 237–278. © 2005 Elsevier B.V. All rights reserved.
The Precambrian of Finland includes three mafic–ultramafic rock complexes – Jormua, Outokumpu, and Nuttio – that are interpreted to represent ancient ophiolites. Of these, the Jormua ophiolite is most complete containing all the salient units of an ophiolite: mantle tectonites, sheeted dikes, gabbros and plagiogranites, and lavas and hyaloclastites. The mantle unit of Jormua ophiolite is, however, atypical for ophiolites as it mainly consists of subcontinental lithospheric mantle. Pyroxenitic and hornblenditic dikes intrusive into the mantle peridotites contain Archean zircon xenocrysts, whereas the gabbros, plagiogranites, and volcanic rocks crystallized at ~1.95 Ga. These features imply that the Jormua ophiolite was formed within a passive ocean–continent transition zone where old subcontinental lithospheric mantle became unroofed along low-angle detachment faults and was injected by basaltic melts from an asthenospheric source. The Outokumpu “ophiolite” consists of several tens of allochthonous mantle tectonite massifs that are not associated with sheeted dikes complexes or abundant basalts. Compared to Jormua, Outokumpu peridotites are more depleted and are also devoid of pyroxenitic/hornblenditic veins and associated mantle metasomatism. These massifs are, however, intruded by gabbros that yield similar crystallization ages as Jormua gabbros, suggesting that Outokumpu and Jormua share a common origin. The Outokumpu-type massifs probably represent mantle from the same passive margin environment as the Jormua ophiolite but from a more oceanic setting. The more residual character of the Outokumpu peridotites and the presence of Cu-rich sulfide mineralization are indicative of high heat flow, and the Outokumpu massifs are interpreted to represent oceanic mantle from ridge-axis discontinuity, which is a favorable site for both peridotite exposure and hydrothermal activity. The third ophiolite, Nuttio, was formed in a different geodynamic setting than Jormua and Outokumpu. Highly depleted peridotite and chromite compositions, together with presence of boninitic, tholeiitic, and calc-alkaline dikes suggest that Nuttio metaperidotite massifs represent fragments of oceanic mantle from a fore-arc basin or intra-oceanic island arc. Whether the origin of Nuttio ophiolite is somehow related to that of the Jormua and Outokumpu, as the similar isotope age constraints would suggest, remains unclear.
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1. Introduction The Precambrian of Finland is unique in exposing several allochthonous rock assemblages that closely resemble ophiolites, i.e., fragments of ancient oceanic lithosphere. This chapter attempts to provide a state-of-the-art review of what is currently known about these rocks. The three early Proterozoic ophiolites in Finland, Jormua, Outokumpu, and Nuttio, display distinct internal characteristics. They differ from each other with respect to the ophiolite units present/preserved, chemical composition of mantle tectonites, and nature of their intrusive and extrusive rock types (Table 6.1). The Jormua ophiolite represents a complete Ligurian-type ophiolite with moderately depleted peridotites being intruded by sheeted dikes and capped by a thin basaltic lid composed of enriched mid-ocean ridge type lavas and hyaloclastites, isotropic gabbros, and plagiogranites (Kontinen, 1987). In contrast, the Outokumpu and Nuttio rocks do not contain much else than strongly depleted and serpentinized mantle peridotites. Nevertheless, their ophiolitic character is not determined solely by the petrology of the ultramafic and mafic lithologic units but also by their regional setting. Some important features of the Finnish ophiolites become evident from their areal distribution depicted in Figure 6.1. They all occur in eastern and northern Finland within allochthonous <2.0 Ga old sedimentary or volcano-sedimentary sequences obducted on the Archean (3.5–2.5 Ga) basement and its autochthonous (2.5–2.0 Ga) sedimentary cover deposits. The Jormua ophiolite and Outokumpu-type serpentinite massifs are found close to the western margin of the Karelian craton within the Kainuu schist belt and North Karelia schist belt, respectively. Some ser pentinite massifs are also located within the sedimentary-deficient “suture” between the Iisalmi complex (IC) and Eastern Finland complex (EFC). Both schist belts consists of basically similar metasedimentary strata, 240
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which can be subdivided into the autochthonous 2.5–2.0 Ga Jatuli strata and younger, 2.1– <1.92 Ga, Kaleva strata (e.g., Chapter 7). The Kaleva assemblage consists of two main tectonostratigraphic units: the autochthonous lower Kaleva and allochthonous upper Kaleva tectofacies (terminology of Kaleva after Kontinen, 1987). The upper Kaleva consists of marine turbiditic metagraywackes with thick intercalations of black schists and, in the Outokumpu region, metaperidotite massifs close to its basal contact. In the Kainuu schist belt the association of the Jormua ophiolite with allochthonous upper Kaleva is less evident (Figure 6.2) as some of the ophiolite slices are intimately associated with imbricated slices of both Archean basement and riftogenic and marine sediments of lower Kaleva. The ophiolitic Nuttio serpentinites in northern Finland are also located within the western margin of the Karelian continent (Figure 6.1), but in a truly distinct lithological setting compared to Jormua and Outokumpu. They are sporadically distributed along the eastern margin of the Kittilä volcanic complex. The whole Kittilä complex has been interpreted to represent an allochthonous sheet of early Proterozoic oceanic lithosphere (Hanski et al., 1995). The chemical heterogeneity of the associated volcanic rocks (EMORB, WPB, OIB, VAB, BON) suggests that this allochthon consists of several smaller amalgamated volcanic terrains. Interestingly, the obduction of the Kittilä allochthon occurred roughly at the same time (~ 1.92 Ga) as that of the Jormua ophiolite and the Outokumpu serpentinites (Chapter 4). In the following, description of the regional stratigraphic context is kept to a minimum, as this is extensively discussed elsewhere in this volume (Chapter 7). Furthermore, although the chemical composition of the rocks will be scrutinized in some detail, no analytical data are tabulated. An interested reader can find most of the whole-rock analyses that were used in the construction of the diagrams from the original contributions by Kontinen
Table 6.1. Characteristics of the Jormua, Outokumpu, and Nuttio mafic–ultramafic complexes. Rock type
Jormua western block central block
Crustal units Extrusive unit Sheeted dike complex Gabbros/plagiogranites Ultramafic cumulates Sulfide mineralization
Outokumpu
Nuttio
eastern block
– – – – –
– ++ – – –
++ ++ ++ ? –
? – + – +
? – – – –
Intrusive to mantle tectonites EMORB-dikes IAB-dikes Gabbroic mantle dikes Chromitite pods OIB-type dikes Cpx + amph ± grt mantle dikes Boninitic dikes
? – – – ? ++ –
++ – + – + – –
++ – ++ + – – –
+ – + + – – ?
– + – + – – +
Mantle tectonites Lherzolites (>3 wt.% Al2O3) Depleted lherzolites (1< Al2O3 < 3 wt.%) Harzburgites and dunites (<1 wt.% Al2O3)
+ ++ –
– ++ +
+ ++ +
– + ++
– – ++
– absent, + minor component, + + abundant
(1987), Hanski (1997), and Peltonen et al. (1996b, 1998).
2. Significance of ancient ophiolites Although ophiolites are rather common in younger, particularly Mesozoic orogenic belts (Tethyan), they are notably rare in the Archean and early Proterozoic rock record. An Archean, 2.505 Ma ophiolite with most of the major components required by the Penrose Conference ophiolite definition (Anonymous, 1972) has recently been reported from the North China Craton (Kusky et al., 2001). The second oldest ophiolites reported are the early Proterozoic ophiolites in Finland and Canada (Kontinen, 1987; Peltonen et al., 1996b; 1998; Scott et al., 1991). An apparent implication of the recognition of ancient ophiolites is that they attest to the operation of moderntype plate tectonic processes at that time.
Evidence for plate tectonics is also seen in the composition of even older basaltic rocks (e.g., Condie, 1990), and therefore, it is likely that oceanic crust during the Proterozoic and Archean was formed in a similar manner as at the present-day ocean floor. It is important to keep in mind, however, that because the Paleoproterozoic ophiolites are still few in number and seem to have formed in island arc, mature ocean/oceanic island, and passive margin settings (Kontinen, 1987; Peltonen et al., 1998; Dann, 1991; Scott et al., 1991), they bear little information on the properties of the oceanic lithosphere in major Paleoproterozoic oceanic basins. On a more local scale, however, the presence of ophiolites and their characteristic features have great potential to yield information on the processes that lead to continental breakup, evolution of continental margins, and subsequent tectonic evolution of orogenic belts. As will be discussed later, this is particularly true also for the suture zone CHAPTER
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te uli an Gr
“SUTURE”
lt be
Kittilä allochthon Nuttio serpentinite belt Proterozoic cover
Archean
PC EFC
Jormua ophiolite complex IC
Outokumpu-type ultramafic massifs
Svecofennian mobile belt
“SU
TU
RE
”
100 km
Fig. 6.1. Distribution of Paleoproterozoic ophiolites in Finland. Modified from Kontinen (1987), Korsman et al. (1997), and Hanski and Huhma (Chapter 4).
between the Archean Karelian craton and the early Paleoproterozoic Svecofennian mobile belt (Figure 6.1) in eastern Finland. In addition to providing proof for the operation of modern-type plate tectonics and recognition of ancient plate boundaries, Archean and Paleoproterozoic ophiolites are sources of other fundamental information. The ophiolite basalts, for instance, provide an uncontaminated window to the processes that took place in the ancient oceanic mantle. For example, the Jormua ophiolite (Finland) and 242
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Purtuniq ophiolite (Cape-Smith thrust–foldbelt, Canada) contain both depleted (MORB) and enriched (OIB) type basaltic rocks. This is convincing evidence that such mantle sources were already isolated during the earliest Proterozoic and their interaction was similar to that in the modern mantle (Scott et al., 1991; Peltonen et al., 1998). The mantle sections of ophiolites also permit direct study of the petrology, geochemistry, and physical processes of the upper mantle at the scale of several kilometers. Importantly, the lowest
ophiolitic units may expose not only oceanic lithospheric mantle, but some ophiolites (Ligurian-type) that record the early stages of continental breakup may also expose exhumed old subcontinental lithospheric mantle, SCLM (e.g., Rampone and Piccardo, 2000). Ophiolites that have lower units composed of continental mantle bear some similarities to orogenic lherzolite massifs, both of which provide more representative SCLM samples than small mantle xenoliths.
3. Age constrains for Finnish ophiolites The first isotope age determinations of the Finnish ophiolites were published by Huhma (1986) who reported an U-Pb zircon age of 1972±18 Ma for the Horsmanaho gabbro pegmatite enclosed in one of the Outokumpu-type serpentinite massifs (Table 6.2). This age was based on four highly concordant fractions of euhedral, elongated magmatic zircon grains, which are believed to reliably date the emplacement of the gabbro. As lavas and sheeted dikes are practically absent in the vicinity of the Outokumpu-type massifs, it is unclear if this age corresponds to seafloor magmatism. The possibility remains that the gabbros were emplaced into the uppermost part of the mantle diapir before the final breakup of the continental crust and the formation of a new oceanic basin. In such a case, the Horsmanaho gabbro would provide the maximum age for the continental breakup and “ophiolite formation.” Samples from the crustal unit of the Jormua ophiolite, one high-level gabbro and one plagiogranite, yield ages of 1960 ± 12 Ma and 1954 ± 11 Ma, respectively (Kontinen, 1987). These determinations were made from clear and euhedral tetragonal zircon prisms and constrain the age of the seafloor magmatism reasonably well. To define the age relationship between the crustal unit of the Jormua ophio-
lite and gabbroic dikes that intrude the mantle tectonites deeper in the ophiolite stratigraphy, Peltonen et al. (1998) dated a pegmatitic variety of one gabbroic feeder dike. This yielded a high-precision U-Pb zircon age of 1953 ± 2 Ma, which is consided to most reliably date the formation of oceanic crust at Jormua. Western block peridotites particularly, but locally also central block tectonites, have been intruded by “dry” clinopyroxenite and hydrous hornblendite dikes. Recently, zircons from two such clinopyroxenite dikes were dated by the NORDSIM ion microprobe (Peltonen et al., 2003). Two distinct types of zircon grains were recovered from clinopyroxenitic dike 24 F (Table 6.2). The first subgroup consists of euhedral, brownish grains with clear magmatic growth zonation. These zircons yielded concordant Archean ages with an average of 2747 ± 8 Ma (n = 7). One grain records even an older age of 2858 ± 14 Ma. The second subgroup of zircons consists of colorless rounded grains which lack magmatic growth textures. They record concordant ages between 2020 and 1940 Ma (see below for details). Nd isotope data clearly imply that the clinopyroxenite dikes are Proterozoic. Therefore, Archean grains in these dikes must represent xenocrysts inherited from older metasomatic veins deeper in the mantle, suggesting that the host serpentinites represent Archean subcontinental lithospheric mantle. The younger age group consists of ~2.1 Ga igneous grains that were partially recrystallized in the 1.95 Ga event. Hornblenditic, OIB-type, and carbonatitic mantle dikes also record crystallization ages of ~2.1 Ga, thus predating the 1.95 Ga ocean-floor magmatism. Altogether, zircon chronology indicates that the Jormua ophiolite includes both Archean and Proterozoic units – consistent with its postulated origin within the ocean–continent transition zone (Peltonen et al., 1998). The exact age of the Nuttio serpentinite massifs is not known. The only age constraint comes from Sm-Nd isotope analyses of two CHAPTER
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Table 6.2. Isotope age data for the Precambrian ophiolites in Finland. Lithology
Sample
mineral/method
Age
Ref.
Gabbro Plagiogranite Gabbroic feeder dike Hornblenditic mantle dike Clinopyroxenitic mantle dike
A729 A196 A1402 A1403 24F
Clinopyroxenitic mantle dike OIB mantle dike Carbonatitic mantle vein
23B A1529 60-L
zr/conventional zr/conventional zr/microcapsule zr/microcapsule xenocrystic zr/sims zr/sims xenocrystic zr/sims zr/sims purple, high-U zr/sims bright, low-U zr/sims
1960 ± 12 Ma 1954 ± 11 Ma 1953 ± 2 Ma > 1.94 Ga ~2860–2730 Ma ~2040–1960 Ma ~3110–2800 Ma ~2020–1960 Ma ~2.1 Ga 1948 ± 30 Ma
1 1 2 2 3 3 3 3 3 3
zr/conventional
1972 ± 18 Ma
4
Sm/Nd, TDM
< 2.1 Ga
5
Jormua
Outokumpu Gabbro pegmatite
A235
Nuttio Calc-alkaline dike in serpentinite
zr = zircon References: (1) Kontinen, 1987; (2) Peltonen et al., 1998; (3) Peltonen et al., 2003, (4) Huhma, 1986; (5) Hanski and Huhma, Chapter 4.
cross-cutting calc-alkaline dikes crosscutting serpentinites. They yielded a TDM of ~2.1 Ga, which can be regarded as the maximum crystallization age for the dikes (Hanski and Huhma, Chapter 4). However, the timing of the separation of the Nuttio ultramafic rocks from the convecting mantle has not been constrained.
4. The Jormua ophiolite The Jormua ophiolite in eastern Finland is the prime example of a Precambrian ophiolite (Kontinen 1987). In fact, it is the oldest mafic–ultramafic rock complex where the transitional contacts between all the main ophiolite units can be demonstrated, providing strong evidence for their consanguinity. All the major components of a typical ophiolite are present (Figures 6.2 and 6.3): (i) a unit of massive and pillow lavas devoid of terrigenic 244
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sediment intercalations, (ii) a sheeted dike complex with dike-in-dike intrusive features implying extensional setting, (iii) cumulate gabbros, and (iv) a mantle tectonite unit. Importantly, most of the gabbroic cumulates and sheeted dikes occur as intrusions within the upper mantle tectonites. Thick units of layered ultramafic and gabbroic cumulate rocks, which characterize many younger ophiolites and modern oceanic crust, are poorly developed at Jormua. Overall, the pseudostratigraphy of the Jormua ophiolite closely resembles that of lherzolite-type ophiolites (LOT) believed to have formed at slow spreading ridges or continental rift zones (Nicolas, 1989). The younger analogues for the Jormua ophiolite include, e.g., the western Alps ophiolites (Lanzo, Liguria, Apennines, Corsica) and the Trinity ophiolite (USA). The internal structure and postulated paleogeographic setting of the Jormua ophiolite are thus different from those of the other two early Proterozoic
3545
Eastern block
P P
Central block 7140
cpx OIB OIB
C 7140
? cpx hbl cpx hbl cpx hbl cpx
hbl
Western block
3545 4 km
Allochthonous rocks Upper Kaleva tectofacies (deep marine metasediments) Pillow lava and pillow breccia Sheeted dikes
~1.95 Ga
Gabbro P
Plagiogranite
OIB
OIB-type “early dikes”
hbl
Hornblendite dikes
cpx
Clinopyroxenite dikes
~2.1 Ga
C
Serpentinite (mostly Archean subcontinental mantle) Chromitite pods
Autochthonous rocks Lower Kaleva tectofacies (riftogenic marine sediments, 2.1–1.95 Ga) Jatuli tectofacies (shallow marine, epicratonic sediments, 2.3–2.1 Ga) Archean basement (>2.8–2.5 Ga)
Fig. 6.2. Geology of the Jormua ophiolite (modified after Kontinen, 1998b). CHAPTER
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ophiolites: the 1.99 Ga Purtuniq ophiolite in northern Quebec, considered to represent a mature open ocean or back-arc basin (Scott et al., 1991), and the 1.83 Ga Payson ophiolite in Arizona believed to be of intra-arc origin. The feature that makes Jormua unique among ancient ophiolites is that it exposes extensive areas of upper mantle rocks, which enable the direct study of mantle processes at the scale of several kilometers. Since the original description of the Jormua ophiolite (Kontinen, 1987) it has become evident that the complex is much more variable than originally assumed. Detailed mapping of the mantle section has revealed major differences between individual tectonic blocks (Peltonen et al., 1996b; 1998). These are summarized in Table 6.1. The eastern block is similar to “normal” Ligurian-type ophiolites as it consists of depleted mantle harzburgites that have been intruded by MORB dikes, gabbroic pods, and small podiform chromitites (Figure 6.2). Dike complex or, locally, mantle tectonites, are directly overlain by a thin, originally 100–400-m-thick, extrusive unit, which mainly consists of pillow and massive lavas (Figure 6.4A). The lavas are overlain by basic tuffs interbedded with carbonate rocks which in turn are structurally overlain by upper Kaleva graywacke and black shales. The central block is in many respects similar to the eastern block, but some important differences are evident. This block mainly consists of mantle peridotites and sheeted dikes, while gabbros and lavas are almost absent. This suggests that the central block represents a somewhat deeper section of the original ophiolite stratigraphy. In fact, the central block might represent a nearly vertical section through the lower oceanic crust and uppermost mantle as the volume of dikes gradually decreases from the spectacular 100% in dike-in-dike sets (Figure 6.4B) in the southeast to sporadic “deep dikes” intruding the peridotites (Figure 6.4C) in the northwestern part (Figure 6.2). The central block peridotites have been intruded also by a 246
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distinct generation of mafic–ultramafic dikes that are not found within the eastern block. These “early dikes” are oriented parallel to the mantle tectonite foliation and are crosscut by MORB dikes. The chemical character of the “early dikes” is similar to that of ultramafic lamprophyres and OIB (Peltonen et al., 1996b). The western block is truly distinct from the eastern and central blocks and bear more similarities to the orogenic lherzolite massifs than to oceanic mantle. Major differences between the western block and the remaining ophiolite include: (a) the western block peridotites are less depleted compared to those of the other blocks, (b) the western block is not associated with any of the rock types of the crustal unit (MORB dikes or lavas), (c) the western block is spatially associated with imbricated slices of the Archean basement and riftogenic autochthonous lower Kaleva sediments instead of allochthonous upper Kaleva sediments, and (d) the western block mantle tectonites are extensively veined by clinopyroxenitic and hornblenditic high-pressure cumulate dikes (Figure 6.3, 6.4G, 6.4H). These differences between the individual blocks bear a major role in reconstructing and presenting the paleogeographic setting for the Jormua ophiolite (Section 8 of this chapter).
4.1. The crustal unit Petrology of the basalts The basaltic–gabbroic crustal part of the Jormua ophiolite is rather thin, only 100–400 m on average. Locally, field observations indicate that basaltic flows deposited directly onto mantle peridotites. The extrusive rocks include pillow lava and massive lava flows with minor hyaloclastites. Kontinen (1987) argued that the presence of hyalo clastite inter pillow matrix and the vesicle-rich pillow flows imply eruption in a shallow-water environment. Most of the basaltic rocks in the Jormua ophiolite, however, are found as dikes. They occur as an extensive dike complex > 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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A
B
C
D
E
F
G
H
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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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249
Lavas and dikes
Chondrite normalized
1000
A
100 Jormua
10
Outokumpu 1 La
Ce Pr
Nd
Sm Eu
Gd Tb
Dy Ho Er Tm Yb
B
500
Gabbros Chondrite normalized
Jormua plagiogranites 100 Jormua gabbros
10
Outokumpu gabbros
1
La
Ce Pr
Nd
Sm Eu
Gd Tb
Dy Ho Er Tm Yb
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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251
FeOtot
P
0 100
Jormua upper-level gabbros
10
90
20
Jormua gabbroic feeder dikes
80
Outokumpu gabbro stocks
30
70 P
40
60
THOLEIITIC 50 60
P
50 40
P CALC-ALKALINE
770
Jormua plagiogranites
30
P P 80
20
P 90
10
100 0
10
20
30
40
Na2O+K2O
50
60
70
80
90
0 100
MgO
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
A
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 (<1 m wide) dikes have been completely altered to chlorite and amphibole. Thicker dikes contain metagabbroic portions in their cores. The larger stock-form bodies have chlorite schists along their margins and less strained and alCHAPTER
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A
B
C
D
E
F
Fig. 6.11. (A) Medium-grade metamorphic mineral paragenesis antigorite (gray) + olivine (granular) + tremolite (bladed) in mantle tectonite, crossed polarizers, width of the image is ~5 mm. (B) Stable mineral paragenesis orthopyroxene + carbonate + olivine in mantle tectonite, crossed polarisers, width of the image is ~5 mm. (C) High-grade mineral paragenesis olivine (black) + orthopyroxene (white, retrograded by talc) in ultramafic rock. (D) Eskolaite (Cr2O3) crystal in Cu-ore from the Outokumpu sulfide mine. (E) Chromian diopside crystals in carbonate-rich skarn. (F) Polished slab of deformed gabbro stock intruded into Outokumpu-type peridotite massif. Photos: (A), (B), (C), and (F) by Jaakko Säntti; (D) and (E) by Jari Väätäinen.
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Table 6.3. Metamorphic mineral parageneses of the Outokumpu nappe (Figure 6.8) from east to west (after Säntti et al., in preparation). Zone
Characteristic mineral assemblage
Equilibration temperature
Antigorite zone Talc zone Anthophyllite zone Enstatite zone
antigorite ± olivine ± tremolite olivine+talc olivine+anthophyll. ± cummington. ± talc olivine+enstatite ± anthophyll. ± Mg-Al spinel
500–550 °C 550–680 °C 660–700 °C 700–770 °C
5.3. Cu-Co-Zn-Ni±Au sulfide deposits
Fig. 6.12. About 42 million tons of ore averaging 3.1 wt.% Cu, 1 wt.% Zn, 0.2 wt.% Co, 0.1 wt.% Ni, and 0.6 ppm Au was mined between 1913 and 1988 from sulfide deposits intimately associated with the ultramafic massifs. Photo: Archives of the Outokumpu Mining Ltd.
tered metagabbros in their core parts. Uralite pseudomorphs after pyroxene are preserved in many samples, whereas plagioclase is usually recrystallized into granoblastic mass, which may still preserve the outlines of the original coarse-grained plagioclase. Gabbro stocks are frequently rodingitized and consist of the mineral assemblage clinopyroxene+hastingsitic amphibole+grossular garnet (primarily hydrogarnet)+epidote–zoisite. Rodingitization of the gabbros implies that ultramafic massifs were at least partly serpentinized before the regional metamorphism.
The Outokumpu-type ultramafic massifs are intimately associated with polymetallic CuCo-Zn-Ni±Au sulfide deposits (Figure 6.12). The mining history of the district extends from 1913 to 1988 involving exploitation of three major deposits with a total production of ~42 Mt of ore. The detailed description of the ore deposits and historical development of the deposit modeling falls outside this review but some salient features of the ores are outlined below. Most of the Outokumpu-type sulfide deposits are thin (<1–10 m), narrow (<50–400 m) and up to >6.5-km-long lenses and sheets of sulfides with quartz, diopside, and tremolite as principal gangue (Vähätalo, 1953; Koistinen, 1981). Contacts of the ore lenses are frequently sharp and in some cases the sulfide material intrudes and brecciates the wall rocks. These features imply that the final emplacement of the sulfide mass was structurally controlled. The sulfide sheets are frequently semimassive and the amount of quartz exceeds that of sulfides. The main sulfide minerals are pyrite, pyrrhotite, chalcopyrite, cubanite, and sphalerite. Accessory sulfides include Co-pentlandite, stannite, and cobaltite. Galena is extremely rare. Gold is present in metallic form as Au+Ag±Hg grains. Pyrite is abundant in the Outokumpu and Kylylahti deposits, whereas all the other deposits are pyrrhotite-dominated. Ore structures and textures reflect metamorphic recrystallization CHAPTER
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and subsequent annealing. Rare colloform textures in the pyritic parts of Outokumpu and especially the Kylylahti deposit are probably the only relicts of primary depositional features (Loukola-Ruskeeniemi, 1999). Geochemically, Outokumpu-type sulfide occurrences can be divided into two distinct subtypes, (1) those with Cu-dominated sulfide compositions, i.e., Cu-Co-Zn-Ni±Au deposits, and (2) those with Ni-rich sulfide phase notably poor in copper, i.e., Ni-disseminations (Huhma, 1970). All the major deposits exploited belong to the Cu-ore subtype. The Ni-disseminations refer to low grade (0.2–0.5 wt.% Ni) sulfides in the margins of the ultramafic massifs (Parkkinen and Reino, 1985). The present composition of the CuCo-Zn-Ni±Au deposits resulted from mixing between two end member sulfides having a distinct origin and age. These end members are (a) the Cu-rich proto-ore that deposited from hydrothermal solutions in an ultramafic sea-floor environment, and (b) the younger Ni-sulfide disseminations that formed within altered marginal zones of the ultramafic massifs during obduction. Field evidence suggests that mixing and homogenization of these end members took place during remobilization of the sulfides and produced the polymetallic CuCo-Zn-Ni±Au ores. The Outokumpu-type ores represent an uncommon ore type which seems to have no clear ancient or modern analogue elsewhere (Papunen, 1987). The distinctive features include the lack of associated volcanic rocks and hydrothermal sediments, close association with mantle tectonite massifs, high Ni and Co content of Cu-Zn sulfide ore, low contents of trace elements such as Bi and Se, low Se/S, and extremely low abundances of Pb in the ore and its mantle-like initial isotope composition (Vaasjoki, 1981).
6. The Nuttio serpentinite belt The third Paleoproterozoic ophiolite – the 262
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75-km-long Nuttio serpentinite belt – has been described from central Lapland, northern Finland (Figure 6.1). Kontinen (1981) came to the conclusion that the thin slivers of serpentinized ultramafic rocks along fault zones on the eastern flank of the Kittilä greenstone belt (KGB) share many features common with residual mantle tectonites (Figure 6.13). Hanski et al. (1995) suggested that the protoliths for the Nuttio serpentinites were fragments of oceanic lithospheric mantle. These ultramafic massifs have gone through a complex metamorphic history that has largely destroyed their primary minerals. Typically, they consist of olivine and lizardite cores surrounded by successive shells of antigorite serpentinites, carbonate-serpentine rocks, talc-carbonate rocks, and finally a reaction zone (blackwall) consisting of monomineralic talc, actinolite, and chlorite seams against country rocks (Kontinen, 1981). Most of the olivine in Nuttio-type serpentinites is strain-free and is believed to be a metamorphic mineral produced through serpentine dehydration reaction or, in some cases, recrystallized primary olivine. However, in the core parts of the larger massifs, olivine is present as large strained grains with smaller neoblasts in between, and the rock texture thus closely resembles that of porphyroclastic mantle tectonites (Harte, 1976). Cumulus textures are absent. The Nuttio “ophiolite” is, however, far from being a complete lithological sequence required by the definition of an ophiolite (Anonymous, 1972). Still, several lines of evidence suggest an affinity with oceanic lithosphere. Some of the strongest evidence for an ophiolitic origin comes from the geotectonic setting of the serpentinites. They are located within a major lithologic boundary between the eastern autochthonous cratonic domain and western allochthonous marine domain (Figure 6.13). The western domain, i.e., the Kittilä greenstone belt, consists of NMORB-, EMORB-, IAT-, and MORB-type volcanic rocks, which according to their
Younger 1.9–1.8 Ga granites b b b
b
b b b
Kittilä allochthon (~2.0 Ga) Serpentinite massifs
b
Mg-tholeiitic and minor felsic metavolcanic rocks BIF, Fe-sulfide, and Fe-carbonate schist Fe-tholeiitic metavolcanic rocks
b b b
b b
b b
Proterozoic autochthonous cover b
Archean basement
10 km
Fig. 6.13. Simplified geology of the Kittilä allochthon emphasizing the distribution of ophiolitic serpentinite massifs (Nuttio serpentinite belt) along its eastern margin. Modified from Lehtonen et al. (1998).
depleted mantle-like initial Nd isotope ratio (Hanski and Huhma, Chapter 4) lack any interaction with the Archean basement. Finegrained leucocratic igneous rocks associated with the Kittilä greenstone belt lavas yield equally depleted Nd isotope compositions. Hanski (1997) considered this, together with the absence of cratonic sediments, to indicate that much of the KGB could represent an allochthonous nappe consisting of ancient oceanic lithosphere and its volcanic-sedimentary cover sequence overthrusted onto the craton (Salla, Onkamo, Sodankylä, and Savukoski Groups). The detailed stratigraphy of the region is beyond the scope of this paper and is described in Chapter 3 of this volume. The serpentinite bodies of the Nuttio “ophiolite” range in thickness from 20 to 400 m and in length from 100 to 1000 m (Hanski, 1997). Alteration of the bodies has been so pervasive that cores of large chromite grains probably represent the only relicts of primary minerals. The relict chromite grains have very high Cr/(Cr+Al) and low TiO2. The interiors
of some larger ultramafic massifs contain highly strained olivine porphyroclasts which may also be primary. The serpentinites contain <0.5 wt.% Al2O3 and <0.05 wt.% TiO2 and thus their protoliths were highly depleted dunites and harzburgites. Most of the serpentinite samples have depleted REE patters typical of residual peridotites but some are clearly LREE enriched with La up to seven times chondritic. Similarly, LREE-enriched mantle tectonites were also described from the Jormua ophiolite where their origin is related to percolation and chromatographic fractionation of tholeiitic and alkaline melts in the refractory peridotite (Peltonen et al., 1998). With respect to the intrusive rocks, the Nuttio ophiolite is remarkably different from the Jormua and Outokumpu ophiolites. While the intrusive rocks in the Jormua and Outokumpu mantle tectonites are dikes and cumulates derived from EMORB- and OIB-type melts, the Nuttio dikes include both ultramafic boninitic and mafic tholeiitic and calc-alkaline dikes with island arc geochemical affinities (Hanski CHAPTER
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and Huhma, Chapter 4). The ultramafic dikes are very low in incompatible elements (REE 0.7–2 times chondrite) but high in Cr and Ni. Hanski (1997) considered them to be similar to highly depleted boninites. The mafic dikes include both tholeiitic dikes with flat REE patterns and a suite of rather primitive calcalkaline basaltic dikes. The dike compositions, together with the highly depleted composition of both serpentinites and chromites led Hanski (1997) to suggest that the Nuttio metaperidotites represent fragments of oceanic mantle from a fore-arc basin of an intra-oceanic island arc.
7. Comparative geochemistry of the Finnish ophiolites 7.1. Metaperidotites Major element compositions of these ultramafic massifs are not particularly useful because their pervasive serpentinization. However, the SiO2/MgO ratios of the Jormua serpentinites are still very close to those at fresh mantle tectonites elsewhere and it is probable that the serpentinization conserved both MgO and SiO2. In this case, however, the volume of the peridotites must have increased (O’Hanley, 1996). Aluminium is a largely immobile element and thus whole-rock Al2O3 contents can be used as an approximate measure of the basaltic component in the peridotites (e.g., McDonough and Frey, 1989). Primitive mantle is estimated to contain 4.44 wt.% Al2O3 (McDonough and Sun, 1995) and, in the simplest case, any values less than that are indicative of extraction of basaltic melt from the peridotite. The Al2O3 abundances of the Jormua, Outokumpu, and Nuttio peridotites imply that they are all depleted relative to the primitive mantle (Figure 6.14). The least depleted peridotites are those of the Jormua ophiolite, particularly those of the western block, but even they have lost a considerable 264
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amount of basaltic component. Outokumputype massifs generally contain only < 1 wt.% Al2O3, their compositions correspond to highly residual harzburgites and dunites. Nuttio samples are even more depleted. Several other moderately incompatible elements such as Sc, V, and Ti show consistent behavior relative to aluminum. Although these mantle tectonites are strongly altered and metamorphosed, several immobile elements can still be applied to study their pre-alteration igneous history. For example, Peltonen et al. (1998) discussed the mobility of REE and came to the conclusion that they had behaved as immobile elements, with the exception of La that was slightly leached from some samples. Chondrite-normalized REE patterns (Figure 6.15) imply that several distinct types of peridotites are present in Jormua, Outokumpu, and Nuttio. Although Al2O3 (and Sc, V, and Ti) abundances are considerably depleted, LREE contents and (La/Yb)N of many peridotites are relatively high and, in fact, enriched relative to primitive mantle. This implies complex post-melting history for the peridotites, involving mantle metasomatism by infiltrating fluids or melts. Importantly, the Al2O3 abundances show a good correlation with HREE implying that the HREE abundances have not been disturbed at the post-melting stage. Such correlations become, however, progressively weaker through MREE to LREE suggesting interaction of peridotites with fluid or melt that had a high LREE/HREE. The serpentinites from the eastern and western blocks of the Jormua ophiolite (Figure 6.15) have distinct characteristics. Most of the eastern block samples have low REE abundances, being indicative of their highly residual character. Some samples exhibit U-shaped patterns, which suggests that peridotites were first depleted by an extensive melt extraction and afterwards, probably much later, enriched in LREE (e.g., McDonough and Frey, 1989). The western block peridotites show distinct, steep fractionated or sinusoi-
5 primitive mantle: Al2O3 = 4.4 wt.%
Al2O3(wt.%)
4 3 2 1
0
Eastern Central Western
Jormua
Lu
Mi
Ou
Outokumpu
Nuttio
Fig. 6.14. Box-and-whisker diagram emphasizing the Al2O3 contents of the mantle peridotites (now metaserpentinites) samples from the Jormua, Outokumpu, and Nuttio ophiolites. The vertical lines refer to median and 90th, 75th, 25th, and 10th percentiles. Outliers outside the 10th and 90th percentiles are indicated as black dots. As Al2O3 contents can be related to the degree of partial melting of the mantle peridotite (McDonough and Frey, 1989), this diagram implies that, compared to primitive mantle, most Jormua peridotites are moderately depleted, whereas samples from Outokumpu and especially from Nuttio are strongly depleted in basaltic constituents. Lu–Luikonlahti, Mi–Miihkali, and Ou–Outokumpu (see Figure 6.8).
dal patterns compared to the eastern block peridotites. It is not a coincidence that the western block peridotites became extensively veined by hornblenditic mantle dikes, which show similar mantle-normalized pattern as the peridotites (Figure 6.7). Hornblenditic dike material, probably as percolating melt, is the apparent candidate for the cause of the enriched patterns of the western block peridotites. Outokumpu peridotites have REE abundances close to or below the detection limit of the used analytical procedure (ICP-MS). This is consistent with their Al2O3 abundances, which are lower compared to those of Jormua samples. The sensitivity of the analytical method does not permit the detection of possible U-shaped patterns for Outokumpu metaserpentinites with confidence. They are, however, similar to the eastern block peri-
dotites of the Jormua ophiolite. Outokumpu peridotites seem to completely lack signs of mantle metasomatism that are ubiquitous in the Jormua western block peridotites. Equally, as noted above, the Outokumpu massifs are devoid of veining by alkali melts. Nuttio serpentinites yield variable mantlenormalized REE patterns: they may be either LREE depleted, flat, or LREE enriched (Hanski, 1997). Again, due to analytical limitations, the REE distribution of the most depleted samples remains obscure. Enriched patterns show a gradual rise towards La and do not resemble the sinusoidal patterns of the western block of the Jormua ophiolite. In fact, the CAdikes (see above) of Nuttio are strongly LREE enriched and infiltration of 3–4 wt.% of such CA basalt filtrated into the residual peridotite would explain their patterns. In summary, the chemical composition of the peridotites indiCHAPTER
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Primitive mantle-normalized
1
0.1
Primitive mantle-normalized
Jormua/eastern block serpentinites
10
Jormua/western block serpentinites
1
0.01 La Ce Pr Nd
Sm Eu Gd Tb Dy Ho Er Tm Yb
La Ce Pr Nd
Sm Eu Gd Tb Dy Ho Er Tm Yb
Outokumpu serpentinites 1
0.1
0.01
Primitive mantle-normalized
Primitive mantle-normalized
10 Nuttio serpentinites (Hanski, 1997) 1
0.1
0.01 La Ce Pr Nd
Sm Eu Gd Tb Dy Ho Er Tm Yb
La Ce Pr Nd
Sm Eu Gd Tb Dy Ho Er Tm Yb
Fig. 6.15. Primitive mantle-normalized rare earth element patterns for metaserpentinites (altered mantle tectonites) from the Precambrian ophiolites of Finland. Normalization values from McDonough and Sun (1995).
cates that distinct processes have taken place in the mantle sections of Jormua, Outokumpu, and Nuttio.
7.2. Lavas and dikes The lavas and dike rocks associated with the Jormua, Outokumpu, and Nuttio ophiolites span an extensive compositional spectrum. Chondrite-normalized REE patterns for the Jormua and Outokumpu lavas and dikes are presented in Figure 6.5A, whereas corresponding diagrams for Nuttio samples can be found elsewhere in this volume (Chapter 3). All Jormua main suite dikes and lavas have broadly flat REE patterns 10–20 times chondrite, being similar to transitional MORB in this respect (e.g., Sun and McDonough, 1989). The “early” 266
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OIB-type dikes have distinct, steeply fractionated patterns indicative of derivation from an enriched mantle source within the stability field of garnet. Their chemical composition is modified due to alteration, but Peltonen et al. (1996b) considered them to be similar to ultramafic lamprophyres. The Outokumpu lavas are distinguished from the Jormua basalts by their lower REE abundances (4–10 times chondrite) and LREE-depleted patterns. This is consistent either with derivation from a more depleted mantle source, or that they represent higher degree mantle melts than the Jormua main suite basalts. Lava and dike analyses (excluding Jormua OIB-type dikes) have been plotted on the TiZr diagram of Pearce (1982) in Figure 6.16. These two elements have equal bulk partition-
ing coefficients during partial melting and differences in the Ti/Zr can thus be related to source heterogeneity. All main suite samples from Jormua plot within the elongated field of MORB with the Ti/Zr between NMORB and EMORB (Figure 6.16). Lava samples (Jaakko Säntti and Asko Kontinen, unpublished data) from Outokumpu (Losomäki) have lower abundances of Ti and Zr and plot within the field of island arc tholeiites. However, the low LILE abundances, low LREE/HREE and high compatible element abundances of Losomäki samples (see below) are incompatible with the indicated island arc setting. Rather, the similar Ti/Zr and the fact that they plot on the linear extension of the Jormua samples imply that these lavas represent high-degree melts from similar mantle source as the Jormua “main suite” EMORB. The dikes that crosscut the Nuttio serpentinites seem to have multiple origins: three samples plot within the field of calc-alkali basalts, two resemble MORB, and three boninitic dikes of Hanski (1997) are extremely depleted in incompatible elements. Rare ultramafic dikes from Outokumpu ultramafic massifs are closely similar to the boninitic Nuttio dikes. It is also noteworthy that the “boninitic” and ultramafic samples plot within the extension of the narrow compositional field defined by the less primitive Jormua and Outokumpu basalts. This suggests that also the Nuttio dikes could have been derived from a similar, yet more depleted source. High abundances of Cr, Ni, and Sc imply that practically all lavas and dikes from Jormua, Outokumpu, and Nuttio represent rather primitive mantle melts. On the Cr vs. Zr plot (Figure 6.17), for example, most samples follow the subhorizontal melt trend. The Jormua data illustrate the apparent complexity of the magmatism at oceanic ridges. The Jormua lavas have high Cr content at given Zr value and plot close to the trend defined solely by partial melting. On the contrary, the sheeted dikes have experienced some chromite + olivine
fractionation and – in spite of their position beneath the lavas in the ophiolite stratigraphy – most of them cannot represent feeders for the lavas. Furthermore, “deep dikes”, which are located below the sheeted dike complex, are the most fractionated samples and cannot be considered feeders for any of the other basalts. Compared to Jormua, the Outokumpu (Losomäki) lavas are even more primitive and have compositions indicating extensive (30–40%) mantle melting (Figure 6.17). Some of these lavas, however, underwent extensive chromite + olivine ± plagioclase fractionation and resemble the evolved MORB dikes of the Nuttio complex. The ultramafic dikes from Outokumpu, together with their compositionally equal counterparts from Nuttio, have unreasonably high Cr abundances to be primary melts from undepleted mantle. Instead, they must represent melting products of a mantle source that had been already depleted by earlier melt extraction. The Zr-Nb relationship is particularly informative in studying the relative contributions of depleted and enriched mantle sources to the chemical composition of oceanic basalts (Figure 6.18). As discussed above, two distinct suites of basalts are present in Jormua. The “main suite” basalts have Nb/Zr equal to or somewhat higher than that of primitive mantle and are thus similar to EMORB. The “early dikes” (ultramafic lamprophyres) have Nb/Zr similar to OIB. Peltonen et al. (1996b) concluded that the chemical (and Nd isotope) composition of the main suite basalts would be compatible with mixing of a depleted mantle source with a relatively uniform proportion of OIB-type source (mixing lines in Figure 6.18). Outokumpu lavas record distinct characteristics compared to those of Jormua. They have Nb/Zr equal to that of primitive mantle with absolute concentrations far lower than observed in Jormua lavas. This is compatible with the previous conclusions (above) that the Outokumpu lavas were derived from a roughly similar mantle source as the Jormua basalts CHAPTER
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11000 10000 9000 8000
N
Ti (ppm)
7000 6000 ‘ 5000
E
4000 3000 2000 1000
Zr (ppm) 0 0
+
50
100
150
200
Jormua “main suite” lavas and dikes
Island arc tholeiite
Outokumpu basalts
MORB, calc-alkali basalt, island arc basalt Calc-alkali basalt
Nuttio dikes
MORB Fig. 6.16. Ti vs. Zr diagram for the lavas and dikes from Jormua, Outokumpu, and Nuttio ophiolites (OIB-type Jormua dikes are not shown). Note that the Jormua lavas are characterized by higher absolute abundances than Outokumpu (Losomäki) lavas, but their Ti/Zr ratios are equal. This suggests similar source characteristics and that the Jormua and Outokumpu lavas can be related by varying degrees of partial melting. The Nuttio boninites are extremely depleted in these trace elements. N and E refer to average NMORB and EMORB compositions of Sun and McDonough (1989), respectively. Compositional fields after Pearce (1982).
but by a significantly higher degree of partial melting. The former are completely devoid of the OIB-type component.
8. Environments of ophiolite formation The origin and tectonic evolution of the Finnish Precambrian ophiolites, especially those of Jormua and Outokumpu, are intimately related 268
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to the evolution of the Karelian continental margin (e.g., Koistinen, 1981; Kon tinen, 1987). Therefore, detailed study of these mafic–ultramafic complexes can yield a wealth of information concerning the timing and mode of the breakup of the Karelian Archean craton, subsequent formation of the (passive) continental margin and its later geodynamic evolution during the Svecofennian orogeny. Tectonic evolution of the craton margin environment has been reviewed extensively
Primitive mantle 80% 60%
1000
40%
30%
20%
10%
5%
Cr (ppm)
plagioclase olivine
orthopyroxene
100
chromite, clinopyroxene
Y (ppm) 10 0
10
20
Jormua sheeted dikes Jormua lavas Jormua deep dikes
30
40
Nuttio dikes
+ +
Outokumpu (Losomäki) lavas Outokumpu ultramafic dikes
Fig. 6.17. Cr vs.Y diagram for the Jormua, Outokumpu, and Nuttio lavas and dikes. The subhorizontal line is the partial melting line (with melt percentages indicated) from Pearce (1982), primitive mantle composition is according to (McDonough and Sun, 1995), and mineral vector calculations after Peltonen et al. (1996b). Note that most of the samples plot close to the partial melting line implying that the sample suite can be related by varying the degree of mantle melting, whereas fractional crystallization has had only minor effect.
elsewhere (e.g., Koistinen, 1981; Kontinen, 1987; Gaál and Gorbatchev, 1987; Kohonen, 1995; Peltonen et al., 1996b, 1998; Korsman et al., 1999; Laajoki, Chapter 7; Lahtinen et al., Chapter 11) and are not be repeated here. Recent studies have emphasized the great complexity of the mantle section of the Jormua ophiolite, implying that it is rather atypical ophiolite as it includes fragments of both the Archean subcontinental lithospheric mantle (SCLM) and younger oceanic lithosphere.
Table 6.1 summarizes the characteristic features of the three blocks of the Jormua ophiolite. Based on the composition of the mantle peridotites and the presence of high-pressure clinopyroxenite, hornblendite, and garnetite dikes similar to those found in fragments of SCLM elsewhere (mantle xenoliths, orogenic lherzolite massifs), Peltonen et al. (1998) suggested that the western block represents a piece of the ancient lithospheric mantle, exposed beneath the Archean crust by detachCHAPTER
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1000
.7
15
b=
r/N
:Z
le nt
a
ve
m
OIB
iti
im
Zr (ppm)
Pr
100
N
E fractional crystallization source enrichment
partial melting
10 1
10
100
Nb (ppm) Jormua “early dikes”
+
Outokumpu (Losomäki) lavas
Jormua ”main suite” lavas and dikes Fig. 6.18. Zr vs. Nb diagram for Jormua and Outokumpu basalts. Primitive mantle ratio and NMORB, EMORB, and OIB compositions according to Sun and McDonough (1989). Calculated mixing lines between NMORB (N), EMORB (E), and OIB are indicated.
ment faulting. Zircons dated from these dikes by ion microprobe (Peltonen et al., 2003; Table 6.2) and the Re-Os study of the peridotites and chromitites by Tsuru et al. (2000) have unequivocally shown that the western block is Archean in age. The oldest dikes yielded up to 3.1 Ga zircons suggesting that the peridotites are still older. Furthermore, the Re-Os study indicates that not only the western and central blocks but probably all Jormua peridotites represent Archean SCLM (Tsuru et al., 2000). Thus the Jormua ophiolite may consist of two main components of distinct origin and age: (a) strongly streched Archean subcontinental lithospheric mantle, and (b) younger 270
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~1.95 Ga gabbros and volcanic rocks derived from an unexposed asthenospheric diapir that intruded the shallow remnants of the SCLM. These features imply that the Jormua ophiolite formed within the transition zone where the continental lithosphere graded into an oceanic regime (Figure 6.19). In younger terrains, such lithological successions are seldomly exposed, but have been found for example in the Zabargad Island of the Red Sea area where continental mantle became exhumed due to extreme crustal thinning and detachment faulting during the final stages of continental breakup (e.g., Bonatti et al., 1981). Also the passive margins of modern
oceans expose Jor mua-like lithological sequences. One such location is the West Iberia Margin where ocean–continent transition zone between rifted and thinned continental crust and true oceanic crust has been studied in detail. There, seismic studies have identified a ~100 km wide zone of partially serpentinized peridotites exposed at the seafloor (Chian et al., 1999). In this zone, scarce basaltic rocks, locally pillowed, were deposited directly on subcontinental lithospheric peridotites – a likely scenario for Jormua, too (Figure 6.19). These peridotites enclose strongly sheared and metamorphosed gabbro intrusions and alkaline pyroxenites. Not all these intrusions are believed to be comagmatic with the volcanic rocks but are interpreted to represent magmas that under plated the continental crust already before the final rifting (Cornen et al., 1999). Similar lithologic associations as identified at Jormua are also present in the ophiolites of the External Ligurides, northern Apennines, Italy. These consist mainly of subcontinental (Proterozoic) mantle unroofed at a non-volcanic continental margin along low-angle detachment faults. Mantle tectonites became injected by minor basaltic magmas that crystallized as basaltic flows or gabbro pods within the peridotite (Rampone and Piccardo, 2000). In terms of ophiolite stratigraphy, the Outokumpu-type ultramafic massifs represent a far less complete ophiolite than Jormua. The absence of layered cumulate units and sheeted dike complexes, and the scarcity of volcanic rocks in Outokumpu are in apparent contrast with the ophiolite model. Reinterpretation of the carbonate rocks and quartz rocks as strongly altered peridotites instead of seafloor sediments (Kontinen, 1998a), further blurs the ophiolite connection. Thus, one is left with ultramafic mantle peridotite massifs with some gabbroic dikes and chromitite bodies, enclosed by graywacke–black schist metasediments thrusted onto the craton margin. Some constraints on their origin is provided by the chemical composition of the serpentinites. As
noted above, Outokumpu serpentinites have more depleted whole-rock compositions than Jormua peridotites and bear no evidence of mantle metasomatism or veining by alkaline pyroxenitic or hornblenditic dikes. These compositional features are believed to exclude subcontinental lithospheric mantle origin for the Outokumpu-type ultramafic massifs, and is more consistent with them having derived from an oceanic asthenospheric mantle diapir. In spite of the affinity with oceanic lithosphere, these massifs are not associated with an oceanic crustal unit (cumulates, sheeted dikes complex, lavas). The almost complete absence of the oceanic crustal unit may result from selective tectonic preservation. However, as the crustal unit is practically absent from all Outokumpu-type massifs, it is more likely that these massifs were derived from an oceanic basin where abyssal peridotites were directly exposed at the seafloor (Figure 6.19). Modern analogues for such a peridotitic seafloor have been located in a number of regions (Bonatti et al., 1990; Brun and Beslier, 1996; Cannat, 1993; Cannat and Casey, 1995; Nicholls et al., 1981; Pickup et al., 1996), and are sometimes also associated with hydrothermal venting and deposition of sulfides onto the ultramafic seafloor (Bogdanov et al., 1997; Murphy and Meyer, 1998). The third ophiolite, Nuttio, was formed in a different geodynamic setting than Jormua and Outokumpu. Highly depleted peridotite and chromite compositions, together with the presence of boninitic, tholeiitic, and calc-alkaline dikes suggest that the Nuttio metaperidotite massifs represent fragments of oceanic mantle from a fore-arc basin or intraoceanic island arc. The composition of peridotites or intrusive dikes are distinct from the other two ophiolites, but the available isotope age data indicate that the ocean floor volcanic rocks of the Kittilä allochthon formed roughly at the same time as the gabbros of Outokumpu and Jormua. Furthermore, as these volcanic rocks have been intruded by a 1.92 Ga granoCHAPTER
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Lithosperic detachment faulting ≥ 1.95 Ga
Onset of seafloor spreading ≤ 1.95 Ga
J
O
Upper/lower crust of the Archean craton Asthenospheric 1.95 Ga mantle diapir “Old” subcontinental (Archean) lithosperic mantle, SCLM
J
Protolith for the Jormua ophiolite
O
Protolith for the Outokumpu ophiolite
Mid-ocean ridge lavas and gabbros “Young” oceanic lithosperic mantle Faults active during lithospheric boudinage Faults active during lithospheric detachment Faults active during onset of sea-floor spreading
Fig. 6.19. A schematic lithosphere-scale model illustrating a possible tectonic setting for the Jormua and Outokumpu ophiolites within a magma-poor passive margin at ~1.95 Ga (modified from Whitmarsh et al., 2001). See text for details.
diorite (Hanski and Huhma, Chapter 4), the obduction of the Kittilä greenstone belt and associated ophiolitic serpentinites probably took place between 2.0 and 1.92 Ga, which approximates the timing of the emplacement of the Jormua and Outokumpu ophiolites as well (~1.92–1.90 Ga). Whether the emplace272
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ment of the Nuttio ophiolite is related to the same collisional event as Jormua and Outokumpu remains speculative. One possible plate configuration that relates Nuttio, Jormua, and Outokumpu to the same orogenic event is presented in Chapter 11 of this volume.
9. Concluding remarks In summary, the petrology and internal structure of Jormua and Outokumpu indicate that these ophiolites are related to continental breakup and subsequent initial stages of oceanization. The volcanic rocks of both ophiolites are devoid of subduction-zone geochemical components (Peltonen et al., 1996), which, together with the absence of any other subduction related igneous rocks west of Jormua, implies that the breakup was not within an ensialic back-arc/fore-arc basin (e.g., Park, 1988; Kohonen, 1995). Detachment faulting that resulted in exposures of continental mantle at the seafloor at ~1.95 Ga and the formation of the thin basaltic lid of the ophiolites was more likely related to passive rifting and formation of a non-volcanic continental margin. The Jormua and Outokumpu ophiolites are believed to have been derived from distinct positions within such a margin. Jormua represents fragments of seafloor consisting of continental mantle, which was exposed relatively close to the continental crust. The affinity of Jormua to the continental mantle is evident not only in the petrology of the mantle section but also in the chemical composition of the basaltic rocks. The lavas and sheeted dikes of Jormua contain a small but uniform proportion of the OIB-type end member which is believed to result from the thermal erosion of the continental mantle by an ascending asthenospheric mantle diapir. In contrast, the Outokumpu lavas are devoid of this enriched component, which is indicative of more advanced stages of oceanization, when the mantle diapir did not have significant interaction with the ancient subcontinental lithospheric mantle. This is consistent with the lack of SCLM characteristics in the Outokumpu-type ultramafic massifs. The higher degree of mantle melting and formation of massive peridotite-associated copper sulfide deposits at Outokumpu, however, require a different thermal subregime
compared to that of Jormua. These features are consistent with the Outokumpu protoliths representing oceanic mantle from a ridge-axis discontinuity, which was a favourable site for both peridotite exposure and hydrothermal activity. Soon after their formation, these passive margin ophiolites became covered by upper Kaleva slope–rise turbidites. The youngest detrital zircons in these metasediments have been dated at 1.92 Ga (Claesson et al., 1993), which constrains the maximum age for the ophiolite obduction. The minimum for the obduction is provided by a 1.87 Ga granite intruding the upper Kaleva schists. Within the Outokumpu nappe, some ultramafic massifs are found immediately east of the Svecofennian–Archean suture zone. This, together with the nappe tectonics (Koistinen, 1981), suggest that the Outokumpu-type massifs formed west of the suture zone between the Svecofennian mobile belt and the Karelian craton. This is less evident in the case of Jormua and a major uncertainty remains whether this ophiolite formed in a continental rift zone close to its present location (e.g., Kontinen, 1987; Kohonen, 1995) or whether the Jormua rocks were tectonically transported to their present location from the passive margin west of the suture zone across the Archean Pudasjärvi and Iisalmi complexes (PC and IC in Figure 6.1). Both of these settings are equally capable of producing passive margintype ophiolites.
Acknowledgments Several colleagues are thanked for stimulating discussions and putting unpublished data at the author’s disposal, especially Asko Kontinen, Jaakko Säntti, Hannu Huhma, Eero Hanski, Irmeli Mänttäri, Jarmo Kohonen, Raimo Lahtinen, and Kalevi Korsman. I am particularly indebted to Asko Kontinen who also carefully reviewed the manuscript. Over the past ten years this author’s work on the Finnish CHAPTER
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ophiolites has been made possible through the generous support of the Geological Survey of Finland (GTK), Outokumpu Mining Oy, and the Academy of Finland.
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of magmatism in serpentinized peridotites from the 15°N region. In: R.L.M. Vissers, A. Nicolas (Eds.), Mantle and lower crust exposed in oceanic ridges and in ophiolites. Dordrecht, Kluwer Academic Publisher, 5–34. Chian, D., Louden, K.E., Minshull, T.A., Whitmarsh, R.B., 1999. Deep structure of the ocean–continent transition in the southern Iberia Abyssal Plain from seismic refraction profiles. Ocean Drilling Program (Legs 149 and 173) transect. J. Geophys. Res. 104, 7443–7462. Claesson, S., Huhma, H., Kinny, P. D., Williams, I. S., 1993. Svecofennian detrital zircon ages – implications for the Precambrian evolution of the Baltic Shield. Precambrian Res. 64, 109–130. Condie, K. C., 1990. Geochemical characteristics of Precambrian basaltic greenstones. In: R. P. Hall, D. J. Hughes (Eds.), Early Precambrian basic magmatism. London, Blackie, 40–55. Conquéré, F., 1971. Les pyroxénolites à amphibole et les amphibolites associées aux lherzolites du gisement de Lherz (Ariège, France): un example du rôle de l’eau au cours de la cristallisation fractionnée des liquides issus de la fusion partielle de lherzolites. Contrib. Mineral. Petrol. 33, 32–61. Cornen, G., Girardeau, J., Monnier, C., 1999. Basalts, underplated gabbros and pyroxenites record the rifting process of the West Iberian margin. Mineral. Petrol. 67, 111–142. Dann, J.C., 1991. Early Proterozoic ophiolite, central Arizona. Geology 19, 590–593. Eskola, P.E., 1933. On the chrome minerals of Outokumpu. Bull. Comm. géol. Finlande 103, 26–44. Fabriès, J., Lorand, J-P., Guiraud, M., 2001. Petrogenesis of the amphibole-rich veins from the Lherz orogenic lherzolite massif (Eastern Pyrenees, France): a case study for the origin of orthopyroxene-bearing amphibole pyroxenites in the lithospheric mantle. Contrib. Mineral. Petrol. 140, 383–403. Geol. 69, 451–467. Gaál, G., Gorbatschev, R., 1987. An outline of the Precambrian development of the Baltic Shield. Precambrian Res. 35, 15–52.
Gaál, G., Parkkinen, J., 1993. Early Proterozoic ophiolite-hosted copper-zinc-cobalt deposits of the Outokumpu type. In: R.V. Kirkham et al. (Eds.), Mineral deposit modeling. Geol. Assoc. Canada, Spec. Pap. 40, 335–341. Gaál, G., Koistinen, T., Mattila, E., 1975. Tectonics and stratigraphy of the vicinity of Outokumpu, North Karelia, Finland: including a structural analysis of the Outokumpu ore deposit. Geol. Surv. Finland Bull. 271, 1–67. Green, D.H., 1969. The origin of basaltic and nephelinitic magmas in the earth’s mantle. Tectonophysics 7, 409–422. Haapala, P., 1936. On serpentine rocks in Northern Karelia. Bull. Comm. géol. Finlande 114, 1–83. Hanski, E., 1997. The Nuttio serpentinite belt, central Lapland: an example of Paleoproterozoic ophiolitic mantle rocks in Finland. Ofioliti 22, 35–46. Hanski, E., Pankka, H., Walker, R.J., 1995. The Nuttio serpentinite belt, Central Lapland: a new example of Palaeoproterozoic ophiolitic rocks mantle rocks in Finland. Nordic Winter Meeting, January 8-11, 1996, Turku. Abstracts, p. 62. Harte, B., 1976. Rock nomenclature with particular relation to deformation and recrystallisation textures in olivine-bearing xenoliths. J. Geol. 85, 279–288. Huhma, M., 1970. Nickel, cobalt and copper in some rocks of the Outokumpu region. Bull. Geol. Soc. Finland 42, 67–88. 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, 1–40. Huhma, A., Huhma, M., 1970. Contribution to the geology and geochemistry of the Outokumpu region. Bull. Geol. Soc. Finland 42, 57–88. Irvine, T.N., Baragar, W.R.A., 1971. A guide to the chemical classification of the common volcanic rocks. Can. J. Earth Sci. 8, 523–548. Kohonen, J., 1995. From continental rifting to collisional crustal shortening – Paleoproterozoic
Kaleva metasediments of the Höytiäinen area in North Karelia, Finland. Geol. Surv. Finland, Bull. 380, 1–79. Koistinen, T. J., 1981. Structural evolution of an early Proterozoic stratabound Cu-CoZn deposit, Outokumpu, Finland. Trans. Roy. Soc. Edinburgh: Earth Sciences 72, 115–158. Kontinen, A., 1981. Kittilän vihreäkivialueen itäreunalla Nolppiossa sijaitsevan pienehkön serpentiniittipahkun petrografia ja petrologia. M.Sc. Thesis, University of Helsinki, Finland. (in Finnish) Kontinen, A., 1987. An early Proterozoic ophiolite – the Jormua mafic–ultramafic complex, northern Finland. Precambrian Res. 35, 313–341. Kontinen, A., 1998a. The nature of the serpentinites, associated dolomite-skarn-quartz rocks and massive Co-Cu-Zn sulphide ores in the Outokumpu area, eastern Finland. In: E. Hanski, J. Vuollo (Eds.), International ophiolite symposium and field excursion: generation and emplacement of ophiolites through time, August 10-15, 1998, Oulu, Finland. Abstracts, Excursion Guide. Geol. Surv. Finland, Spec. Pap. 26, 33. Kontinen, A., 1998b. Geological map of the Jormua area 1:50 000. Geol. Surv. Finland, Spec. Pap. 26, Appendix. Kouvo, O., Vuorelainen, Y., 1958. Eskolaite, a new chromium mineral. Am. Mineral. 43, 1098–1106. Korsman, K., Koistinen, T., Kohonen, J., Wennerström, M., Ekdahl, E., Honkamo, M., Idman, H., Pekkala, Y, (Eds.), 1997. Suomen kallioperäkartta, Bedrock map of Finland 1:1 000 000. Geol. Surv. of Finland, Espoo. Korsman, K., Korja, T., Pajunen, M., Virransalo, P., 1999. The GGT/SVEKA transect: structure and evolution of the continental crust in the Paleoproterozoic Svecofennian orogen in Finland. Int. Geol. Rev. 41, 287–333. Kusky, T.M., Li, J.-H., Tucker, R.D., 2001. The Archean Dongwanzi ophiolite complex, North China Craton: 2.505-billion-year-old oceanic crust and mantle. Science 292, No. 5519, 1142–1145.
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Lehtonen, M., Airo, M.-L., Eilu, P., Hanski, E., Kortelainen, V., Lanne, E., Manninen, T., Rastas, P., Räsänen, J., Virransalo, P., 1998. Kittilän vihreäkivialueen geologia: Lapin vulkaniittiprojektin raportti. Summary: The stratigraphy, petrology and geochemistry of the Kittilä greenstone area, northern Finland: A report of the Lapland Volcanite Project. Geol. Surv. Finland, – Rep. Invest. 140, 1–144. Loukola-Ruskeeniemi, K., 1999. Origin of black shales and serpentinite-associated Cu-ZnCo ores at Outokumpu, Finland. Econ. Geol. 94, 1007–1028. McDonough, W.F., Frey, F.A., 1989. Rare earth elements in upper mantle rocks. In: B.R. Lipin, G.A. McKay (Eds.), Geochemistry and mineralogy of rare earth elements, Reviews in Mineralogy 21, Min. Soc. Am. 99–145. McDonough, W.F., Sun, S.-s., 1995. The composition of the Earth. Chem. Geol. 120, 223–253. Murphy, P.J., Meyer, G., 1998. A gold-copper association in ultramafic-hosted hydrothermal sulphides from the Mid-Atlantic Ridge. Econ. Geol. 93, 1076–1083. Nicolas, A. 1989. Structures of Ophiolites and Dynamics of Oceanic Lithosphere. Dordrecht, Kluwer. Nicholls, I.A., Ferguson, J., Jones, H., Marks, G.P., Mutter, J.C., 1981. Ultramafic blocks from the ocean floor southwest of Australia. Earth Planet. Sci. Lett. 56, 362–374. O’Hanley, D.S., 1996. Serpentinites. Records of tectonic and petrological history. Oxford University Press. Papunen, H. 1987. Outokumpu-type ores. In: T. A. Häkli (Ed.), Otto Trüstedt symposium in Finland on June 3–5, 1985. Geol. Surv. Finland, Spec. Pap. 1, 41–50. Park, A.F. 1988. Nature of the Early Proterozoic Outokumpu assemblage, eastern Finland. Precambrian Res. 38, 131–146. Park, A. F., Bowes, D.R., 1981. Metamorphosed and deformed pillows from Losomäki: evidence of sub-aqueous volcanism in the Outokumpu association, eastern Finland. Bull. Geol. Soc. Finland 53, 135–145.
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Parkkinen, J., Reino, J., 1985. Nickel occurrences of the Outokumpu type at Vuonos and Keretti. In: H. Papunen, G. I. Gorbunov (Eds.), Nickel-copper deposits of the Baltic Shield and Scandinavian Caledonides. Geol. Surv. Finland, Bull. 333, 178–188. Pearce, J. A., 1982. Trace element characteristics of lavas from destructive plate boundaries. In: R. S. Thorpe (Ed.), Andesites; orogenic andesites and related rocks, John Wiley & Sons, 525–548. Peltonen, P., Kontinen, A., Johanson, B., 1996a. Eskolaiitin koostumuksesta. In: P. Peltonen, K. Korsman, R. Salminen (Eds.), Tutkimuksia geologian alalta II. Ann. Univ. Turkuensis C (126), 109–116. (in Finnish) Peltonen, P., Kontinen, A., Huhma, H., 1996b. Petrology and geochemistry of metabasalts from the 1.95 Ga Jormua Ophiolite, Northeastern Finland. J. Petrol. 37, 1359–1383. Peltonen, P., Kontinen, A., Huhma, H., 1998. Petrogenesis of the mantle sequence of the Jormua Ophiolite (Finland): Melt migration in the upper mantle during Palaeoproterozoic continental break-up. J. Petrol. 39, 297–329. Peltonen, P., Mänttäri, I., Huhma, M., Kontinen, A., 2003. Archean zircons from the mantle – Jormua Ophiolite revisited. Geology 31, 645–648. Pickup, S.L.B., Whitmarsh, R.B., Fowler, C.M.R., Reston, T.J., 1996. Insight into the nature of the ocean–continent transition off West Iberia from a deep multichannel seismic reflection profile. Geology 24, 1079–1082. Rampone, E., Piccardo, G.B., 2000. The ophiolite – oceanic lithosphere analogue: new insights from the Northern Apennines (Italy). In: Y. Dilek, E.M. Moores, D. Elthon, A. Nicholas (Eds.), Ophiolites and oceanic crust: new insights from field studies and the ocean drilling program. Geol. Soc. Am. Spec. Pap. 349, 21–34. Säntti, J., 1996. Cr-Fe-spinellin, oliviinin ja enstatiitin koostumus ja alkuperä Outokummun ja eräissä muissa karelidien metaultramafiiteissa. M.Sc. Thesis, University of Helsinki, Finland. (in Finnish) Scott, D. J., St-Onge, M. R., Lucas, S. B., Helms-
taedt, H., 1991. Geology and chemistry of the early Proterozoic Purtuniq ophiolite, Cape Smith Belt, northern Quebec, Canada. In: Tj. Peters (Ed.), Ophiolite genesis and evolution of the oceanic lithosphere. Dordrecht: Kluwer, 825–857. Sun, S-s., McDonough, W. F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: A. D. Saunders, M. J. Norry (Eds.), Magmatism in the ocean basins. Geol. Soc. Spec. Publ. 42, 313–345. Treloar, P. J., 1987. The Cr-minerals of Outokumpu – their chemistry and significance. J. Petrol. 28, 867–886. Tsuru, A., Walker, R.J., Kontinen, A., Peltonen, P., Hanski, E., 2000. Re-Os isotopic systematics of the 1.95 Ga Jormua ophiolite complex, northeastern Finland. Chem. Geol. 164, 123–141. Vaasjoki, M., 1981. The lead isotopic composition of some Finnish galenas. Geol. Surv. Finland, Bull. 316, 1–30. Vähätalo, V.O., 1953. On the geology of the Outokumpu ore deposit in Finland. Bull. Comm. géol. Finlande 164, 1–98.
Vétil, J.-Y., Lorand, J.-P., Fabriés, J., 1988. Conditions de mise en place des filons de pyroxénites à amphibole du massif ultramafique de Lherz (Ariége, France). C.R. Acad. Sci. Paris 307(II), 587–593. Vuollo, J., Liipo, J., Nykänen, V., Piirainen, T., Pekkarinen, L., Tuokko, I., Ekdahl, E., 1995. An early Proterozoic podiform chromitite in the Outokumpu ophiolite complex, Finland. Econ. Geol. 90, 445–452. Walker, R. J., Hanski, E., Vuollo, J., Liipo, J., 1996. The Os isotopic composition of Proterozoic upper mantle: evidence for chondritic upper mantle from the Outokumpu ophiolite, Finland. Earth Planet. Sci. Lett. 141, 161–173. Will, T.M., Powell, R., Holland, T.J.B., 1990. A calculated petrogenetic grid for ultramafic rocks in the system CaO-FeO-MgO-Al2O3SiO2-CO2-H2O at low pressures. Contrib. Mineral. Petrol. 105, 347–358. Whitmarsh, R.B., Manatschail, G., Minshull, T.A., 2001. Evolution of magma-poor continental margins from rifting to seafloor spreading. Nature 413, 150–154.
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Chapter 7
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K. Laajoki
Cover page: Cross-bedded arkositic hematite-bearing quartzite, Finnish Lapland. Photo: Jari Väätäinen. •
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Laajoki, K., 2005. Karelian supracrustal 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. 279–342. © 2005 Elsevier B.V. All rights reserved.
The Karelian formations comprise Paleoproterozoic supracrustal rocks that flank late Archean basement blocks in eastern and central Finland. These rocks were folded and metamorphosed at lower-greenschist facies to upper amphibolite facies during the Svecofennian orogeny at ~1.9 Ga. They consist of several fault- and thrust-bounded fragments that can be divided into six tectofacies separated by unconformities. The four lowermost (Sumi, Sariola, Kainuu, Jatuli) and the fifth (Lower Kaleva) represent continental–epicontinental and rift–marginal basin deposits, respectively, whereas the sixth (Upper Kaleva) probably represents an allochthonous marine basin. The predominantly volcanic Sumi tectofacies marks the initial rifting of the Archean basement. The Sariola tectofacies was formed during renewed rifting of the basement and is separated from it by a ~2350 Ma weathering crust. It consists of immature fluvial rudites and arenites and locally preserved basinal turbiditic–tempestitic sequences. Basic–intermediate volcanic rocks abound in some basins and glaciofluvial–glaciomarine deposits are also present. The Sariola stage was followed by a period of intense chemical weathering with resultant kaolinitic crust (now kyanite-bearing quartzites). The Kainuu tectofacies consists of fluvial and deltaic sandstones overlain by transgressive marine sandstones, heteroliths, and carbonate rocks deposited during continued extension of the continental crust. The Jatuli tectofacies is separated from the Kainuu tectofacies by a major erosional unconformity and mainly consists of fluvial and shallow marine feldspathic arenite, quartz arenite, and shelf dolomite with well-preserved stromatolites. The Lower Kaleva comprises heterogeneous autochthonous–parautochthonous sequences, deposited unconformably on Jatuli or the Archean basement, and is characterized by black shales, banded iron-formations, and turbiditic arenites and graywackes. Available basin models include a rift or rifted continental margin. The probably allochthonous Upper Kaleva consists of monotonous basinal turbidites that were deposited after the formation of the ~1.95 Ga Jormua and Outokumpu ophiolites. In Kainuu and Savo, these rocks were thrust to the east atop older tectofacies and Archean basement rocks.
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1. Introduction Precambrian supracrustal formations in eastern Finland, the ages of which vary from ~2500 Ma to ~1850 Ma, occur in fragmentary erosional basins within the east-central Fennoscandian Shield (Figure 7.1). They are poorly exposed, metamorphosed in lower greenschist to upper amphibolite facies, complexly deformed, and separated from each other by major tectonic zones. Thus it is difficult to apply conventional basin analysis methods to them. The most hampering thing is the lack of continuous vertical stratigraphic sections and the small size of the outcrops. The result is that unconformities and their correlative surfaces, which are essential in sequence stratigraphy, can be observed in a few places only. Consequently, interbasinal correlations can be done only in a broad way (see for instance, discussion in Kohonen and Marmo 1992, pp. 56–59). It is, however, possible to apply conventional facies analysis methods to the outcrops with primary sedimentary features. In fact, this method has been applied quite successfully to the Finnish sedimentary formations – see several articles in Laajoki and Paakkola (1988). Isotopic and other geochemical methods (dating of detrital zircons, Sm-Nd provenance studies, C isotopes of carbonate rocks, and REE) offer additional important tools, but the data available are rather scanty (Huhma, 1986, 1987; Claesson et al., 1993; Karhu, 1993; Kortelainen, 1998; Vaasjoki, 2001). This review concerns the Paleoproterozoic sedimentary basins in eastern (North Karelia and Savo) and northern (Kainuu, Kiiminki, Peräpohja, Posio–Kuusamo) Finland (Figure 7.1). Their sequences and bounding unconformities are described in stratigraphic order with the main emphasis on the stratigraphy and sedimentology of the best-studied occurrences. Volcanism, metamorphism, and deformation of the belts are discussed only briefly. 282
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2. Geological setting and basin classification 2.1. Regional distribution of the supracrustal belts All the basins in question are found east or northeast of the Raahe–Ladoga zone and south of the Central Lapland granitoid and Kemijärvi complexes within the area known as the Karelian domain (Figure 7.1; Gaál and Gorbatschev, 1987). This domain consists of a late Archean basement subdivided into the Kuhmo, Iisalmi, Manamansalo, and Pudasjärvi complexes (or Kianta, Iisalmi, and Ranua terrains; Chapter 2) and Paleoproterozoic sedimentary rocks with minor volcanic units. The latter, known as the Karelian formations or the Karelian supergroup, are found as separate supracrustal belts within, around or in between the basement blocks. The major supracrustal belts are as follows (Figures 7.1 and 7.2): The Kainuu belt (1) has been squeezed between the Archean Kuhmo, Iisalmi, Manamansalo, and Pudasjärvi complexes (Figure 7.3). Its eastern and southwestern margins with the Kuhmo and Iisalmi complexes are mainly autochthonous. The central part is intensely deformed and contains thrust/shear slices of supracrustal rocks and ophiolite and basement complexes. The Central Puolanka Group, the depositional age of which is a little problematic, occupies the western part of the belt. Its lower part passes gradually into the paragneisses of the Oulujärvi shear zone. The Hirvaskoski shear zone forms the tectonic contact in the northwest. The North Karelia belt (2) is composed of a thin autochthonous–parautochthonous quartzite-dominated occurrence rimming the Kuhmo complex in the east (Figure 7.4). Its bulk is underlain by the Höytiäinen basin, which is bordered by the Outokumpu nappe complex (3) in the west. The contact of the latter with the Paleoproterozoic Svecofen-
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Central Lapland complex belt
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Fig. 7.1. A simplified geological map of the central part of Finland (mainly from and appropriately amended from Lundqvist et al., 1996, and Korsman et al., 1997). Notice that faults and overthrust surfaces (thick lines) divide the bedrock into blocks and hamper correlation between supracrustal formations between different areas. The hatched blue line marks the Raahe–Ladoga zone, a deep-seated fault zone that defines the boundary between the Karelian domain (includes the Archean basement complexes and Karelian formations) in the northeast and the Svecofennides in the southwest. Maps of the Kainuu, North Karelia, Kuusamo, and Peräpohja belts are given in Figures 7.3, 7.4, 7.5, and 7.6, respectively. Numbers 1–16 refer to the columns in Figure 7.2. CHAPTER
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1. Kuusijärvi
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nian crust (Svecofennides) is tectonic (e.g., Lundqvist et al., 1996). The Kuopio belt (4) is a small belt southwest of the Iisalmi block deposited on small basement domes (Figure 7.4). The other boundaries are arbitrary, as the better-preserved rocks pass into gneisses and migmatites closer to the Raahe–Ladoga zone. The Salahmi belt (5) is a tiny, but well preserved belt rimming the western margin of the Iisalmi complex (Figure 7.1). The northeastern part of the rather large Kiiminki belt (6) lies autochthonously on the Pudasjärvi complex, whereas its central parts are folded and likely thrust to the northeast (Figure 7.1). Svecofennian granitoids and the Mesoproterozoic Muhos Formation occupy the southwestern margin. The Kuusamo belt (7) rims the northern part of the Kuhmo complex (Figure 7.5). It continues on the Russian side as the Paanajärvi belt (Figure 7.1). The Kemijärvi complex and the Salla belt border it in the west and north, respectively. These contacts are tectonic–metamorphic in the west and delineated by a fault zone in the north. The Peräpohja belt (8) rims the Pudasjärvi complex in the northwest (Figure 7.6).
Its border with the Central Lapland granitoid complex is tectonic–metamorphic, and it is separated from the Kalix belt in Sweden and the Kuusamo belt by the Bothnia and Hirvaskoski shear zones, respectively. In addition to these main belts there are several smaller ones of which the most important are the Saari–Kiekki belt (9) within the Kuusamo block (Figure 7.1) and relict supracrustal sequences closely associated with the Kuusijärvi (Figure 7.5) and Suhanko (Figure 7.6) layered mafic intrusions within the northern parts of the Kuhmo and Pudasjärvi complexes, respectively.
2.2. Metamorphism The supracrustal rocks of all the belts have been metamorphosed and thus the original volcanic and sedimentary rocks are now diverse metamorphic rocks, which, especially close to the borders of the Svecofennides and Central Lapland and Kemijärvi complexes and within the Oulujärvi shear zone, are migmatized and intruded by Svecofennian plutonic rocks. In eastern Finland, the grade of metamorphism increases westwards, so that the
Fig. 7.2. (facing page) Simplified stratigraphic columns of the basins reviewed (Figure 7.1). Successions 1 and 2 include sequences that were deposited in Sumi–Sariola rifts within and on the margins the Kuhmo complex, respectively; sequences of successions 3 and 4 were deposited on the present northeastern and southwestern margins of the Iisalmi–Pudasjärvi complex, respectively. Sources for the column data (keyed to the numbers before locality names): 1−Karinen (1998); 2−Luukkonen (1989); 3−Pekkarinen (1979); 4−Kohonen and Marmo (1992); 5−Gehör and Havola (1988); 6−Kontinen (1986); 7−Laajoki (1991); 8−Silvennoinen (1972) and Pekkala (1985); 9−Laajoki (2000); 10−Peltonen et al. (1996); 11−Laajoki (1991); 12−Paavola (1984); 13−Aumo (1983); 14−Korkiakoski and Laajoki (1988); 15−Honkamo (1985); 16−Perttunen (1991). Legend: 1−Sumi layered intrusions; 2−Sariola conglomerates and metasandstones; 3−Arenitic turbidites; 4−Tempestitic semipelites etc.; 5−Quartzite with basal conglomerate; 6−Basic metavolcanic rock; 7−Tuffite; 8−Dolomite; 9−Utajärvi conglomerate; 10−Lower Kaleva turbidites and mica schists etc.; 11−Lower Kaleva iron-formations; 12−Upper Kaleva turbidites and mica schists; 13−Kolmiloukkonen conglomerate and metasandstone; 14−Sariola glacigenic deposits; 15−Jormua ophiolite complex; 16−Tectonic contact. III and IV dolomite sections refer to Karhu’s (1993) carbon isotope stages III and IV, respectively. The red line marks the nonconformity between the Archean basement and its Paleoproterozoic cover. Formations in the West Puolanka and West Kuusamo/ East Posio columns area: Pj−Puolankajärvi; Av−Akanvaara; Pk−Pärekangas; Kv−Karkuvaara; A−Ahola; Nv−Nilovaara; Kk−Kirintökangas. CHAPTER
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27°
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Intrusive rocks Otanmäki alkaline gneiss Paleoproterozoic granitoids
Nenäkangas Jokijyrkkä Hetehongikko
Western part Vihajärvi Group Somerjärvi Group
Kolkonkangas
65°
65°
Central Puolanka Group
Fig. 7.10 Hepoköngäs
SB
Eastern margin and central part Pyssykulju Formation
on
e
Pu
Paragneisses
ea
rz
Upper Kaleva
sh
Jormua ophiolite complex Lower Kaleva
O
ulu
jär
vi
Pitukansuo
R
SB
Haapala quarry Haapalanmäki
P J
NB
East Puolanka & correlative groups Korvuanjoki & correlative groups Kurkikylä & correlative groups Late Archean basement complex
Kajaani
NB V
64°
64°
IISALMI COMPLEX
50 km 27°
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rocks of the eastern part of the Kainuu and North Karelia belts belong to the greenschist facies and the western parts of the Kainuu, Savo, Kuopio, Salahmi, and Kiiminki belts are upper amphibolite facies (e.g., Campbell et al., 1979; Tuisku and Laajoki, 1990). The primary features of the rocks are relatively well preserved in the east, whereas they may be completely destroyed in the west where the rocks are banded gneisses or granite-veined migmatites. Within the Peräpohja belt, the grade of metamorphism increased towards the Central Lapland granitoid complex in such a way that its southern margin was metamorphosed in lower-greenschist facies. In the north, the grade is upper amphibolite facies and the supracrustal rocks pass into paragneisses and migmatites thermometamorphosed during the emplacement of the Central Lapland granitoid complex (Lappalainen, 1994; Perttunen et al., 1996). In Kuusamo, the eastern and southern parts represent greenschist facies, whereas upper-amphibolite facies gneisses occur in the western part close to the tectonic-metamorphic contact with the Kemijärvi complex. As this review emphasizes the sedimentation of the formations, the effects of metamorphism are not considered and the metasedimentary rocks are usually named according to their protolith.
2.3. Tectonic features The Paleoproterozoic tectonic evolution of the supracrustal belts in eastern and central Finland involves four main Svecofennian phases. The first two represent folding and overthrust stages, the latter two are more brittle
and produced N-, SW-, and NW-striking shear zone patterns (e.g., Koistinen, 1981; Ward, 1987; Laajoki and Tuisku, 1990; Kärki et al., 1993; Kärki and Laajoki, 1995). Tectonism of the Peräpohja and Kuusamo belts differs from the more southern belts in that their early tectonic evolution may be more closely related to the Central Lapland and granulite belts than to the Svecofennides. The Peräpohja belt is characterized by E-trending D2 fold and fault structures and is crosscut by D3 shear zones (subparallel to the contact zone between the Central Lapland granitoid complex and the Peräpohja belt) and by younger NW-trending brittle faults (Perttunen et al., 1996). The early E–W grain of the Kuusamo belt was affected by a shear zone located between it and the Salla belt, and the Hirvaskoski shear zone, resulting in a complex fold interference structure.
2.4. Basin classification The basins and their lithologic units are all considered Paleoproterozoic, although there is some doubt as to the depositional age of the Central Puolanka Group. The basins are filled by the supracrustal rocks for which the informal collective names “Karelian formations” and “Karelian supergroup” have been used. There is no strict understanding of how the Karelian units should be classified (Laajoki, 1986a). A rather useful way is to group them informally into the Sumi, Sariola, Kainuu (or Lower Jatuli), Jatuli (or Middle and Upper Jatuli), Lower Kaleva, and Upper Kaleva subunits; these are treated by some authors as lithostratigraphic units (e.g., Meriläinen, 1980; Ojakangas et al., 2001). Melezhik et al.
Fig. 7.3. (facing page) Geological map of the Kainuu schist belt combined from the maps by Kontinen (1989, 1993), Havola (1981), Laajoki (1991), and Lundqvist et al. (1996). The thick lines depict faults, of which only the most obvious ones are marked. Within the Kainuu schist belt there are probably overthrust surfaces. The area of Figure 7.10 is framed. NB and SB Nuasjärvi and Salmijärvi basins, respectively. Locations: J−Jormua; P−Paltamo; Pu−Puolanka; R−Ristijärvi;V−Vuokatti. CHAPTER
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S U P R AC RU S TA L RO C K S
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28° 30´
Keyritty
IISALMI COMPLEX
KUHMO COMPLEX
Nilsiä
Juuanvaarat
Pisa
Koivusaari Fm.
Suhmura thrust
63°00' Herajärvi Group
Kuopio Höytiäinen
Kyykkä Group
ahe
Ra
Sotkuma
ado
–L
Hyypiä & Raatevaara Groups
ga
N
Joensuu
OUTOKUMPU NAPPE COMPLEX
e
zon Oravisalo
SVECOFENNIDES
Tohmajärvi
50 km 28° 30´ Svecofennian plutonic rocks
Lower Kaleva volcanic rocks
Upper Kaleva pelites, psammites & gneisses Serpentinites of the Outokumpu ophiolite
Jatuli volcanic rocks
Lower Kaleva pelites & psammites
Archean basement
Kainuu & Jatuli (quartzite)
Fig. 7.4. Simplified geological map of North Karelia and eastern Savo (mainly according to Lundqvist et al., 1996; southernmost part from Korsman et al., 1997).
(1994) used them as large-scale informal units and correlated them to formally established lithostratigraphic groups. In this study, each of them refers to a group of formations or sequences that can be considered as coeval, but that were not necessarily deposited in a 288
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single basin. Within the North Karelia and Kainuu belts, the major units are bounded by unconformities and can thus be considered synthems (Salvador, 1994). Ojakangas et al. (2001) interpreted them as unconformity-bounded second-order
K A R E L I A N S U P R AC RU S TA L
ROCKS
28° 30´ SA
Lithostratigraphy Limestone–dolomite Fm. Amphibole schist Fm. Dolomite Fm.
LL A BE
Rukatunturi Fm.
LT
Greenstone Fm. III Siltstone Fm. Greenstone Fm. II
R U S S I A
Sericite quartzite/Nilovaara Fms. Ahola Fm.
KL
Greenstone I/Karkuvaara Fms. Archean basement
F Ki
66°10´
66°10´
Posio
HL Kuusamo
+
Kuusijärvi
+ + + + +
Intrusive rocks + + ~1.8 Ga granitoids + 2.2 Ga metadiabases 2.4 Ga layered mafic intrusions
+
0
+ +
50 km
28° 30´
Fig. 7.5. Simplified geological map of the Kuusamo belt (mainly from Silvennoinen et al., 1992, and Korsman et al., 1997). Abbreviations: HL−Himmerkinlahti; KL−Kolmiloukkonen; KiF−Kitka fault.
depositional sequences. In this article they are treated as tectofacies. A tectofacies is defined to include all the formations formed during a specific tectonic phase of the depositionalvolcanic history of a basin or nearby basins (Laajoki, 1990, 1991). This is a broader usage than that of Krumbein and Sloss (1951) who defined tectofacies as “the laterally varying tectonic aspects of a stratigraphic unit.” In terms of sequence stratigraphy, they are CHAPTER
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KARELIAN
related to Hubbard’s (1988) megasequences. The Karelian tectofacies are summarized in Figure 7.2 and Table 7.1. In the following, the tectofacies are often referred to simply by their names (Sumi, Sariola, etc).
S U P R AC RU S TA L RO C K S
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289
CENTRAL LAPLAND GRANITOID COMPLEX
OV
Rovaniemi
KV
66°28´
66°28´
Narkaus
Keinokangas
Suhanko
Southern part 11
S W E D E N
10
8
3 Northern part 14 2
7
1
9 Penikat
30 km
Figure 7.8 Kemi
Younger 5 intrusive rocks 4 15
6
13 12
25° 30´
Fig. 7.6. Geological map of the Peräpohja belt (simplified from Perttunen et al., 1995). Legend: 1−Archean basement; 2−2.42-Ga layered mafic intrusions. Paleoproterozoic formations of the southern part (3–11): 3−Sompuvaara and Runkaus Formations with 2.2 Ga metadiabase; 4−(Palo)Kivalo; 5−Jouttiaapa; 6−Kvartsimaa; 7−Tikanmaa; 8−Poikkimaa and Hirsimaa; 9−Rantamaa; 10−Väystäjä; 11−Martimo. Northern part (12–14): 12−Ounasvaara (OV) and other quartzite formations; 13−Korkiavaara (KV) and other arkosite formations; 14−Pöyliöjärvi Formation. 15−Haaparanta and other younger plutonic rocks. Area of Figure 7.8 and location of the Keinokangas porphyry are indicated.
3. Sumi tectofacies
3.1. Supracrustal rocks
The oldest Karelian supracrustal rocks, mainly acid and intermediate volcanic rocks, are most abundant in Russian Karelia and Kola Peninsula, where they are known as the Sumi formations and are closely associated with 2440-Ma layered mafic intrusions. The latter are also common in Finland, whereas only relics of the Sumi supracrustal rocks are found.
The only occurrence of the Sumi supracrustal rocks in the area examined in this chapter is a thin unit of acid volcanic rocks and granophyre on the Kuusijärvi layered mafic intrusion (Figure 7.5; Karinen, 1998; Karinen and Salmirinne, 2001; Lauri et al. 2003), but they may also occur in the basal parts of the Kuusamo belt (Räsänen, 1999). Felsic volcanic rocks of this group are, however, abundant at the eastern extension of the Kuusamo belt
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Kaleva
Table 7.1. Karelian tectofacies. Tectofacies/unconformity
Age frame (Ma) Type occurrence/type unconformity
Upper Kaleva*) (allochthonous)
<1870–<1920
?Lower Kaleva
1970–?2060
Lower Kaleva
1970–?2060
2060–>2200
Sub-Jatuli unconformity >2200–2350 (intruded by 2200 Ma sills)
Karelia sensu stricto
Kainuu
2350–<2440
Unconformity upon Sumi. Nonconformity upon basement Sumi
Nuasjärvi, Höytiäinen Martimo
Turbidites with P- and Mn-bearing BIFs and black shales
Shallow-water turbidite basin
Jatuli/Upper Kaleva boundary in Kainuu and North Karelia
Marked by polymictic conglomerates with abundant Jatulian clasts
?Continental breakup
North Karelia East Puolanka
Fluvial–shallow marine arenites ± dolomites and associated shales and tuffs
Epi-/pericontinental basins
Nenäkangas
Marked by polymictic conglomerates with abundant Kainuan clasts
Block faulting and tilting and local deep erosion
(a) Pärekangas
(a) Shallow-marine hetero- (a) Highstand deposits liths and volcanic rocks (b) Fluvial ?half-graben (b) Quartz-pebble conglomerate–sericite basins schist–quartzite sequence
(b) Korvuanjoki, Akanvaara
Sub-Kainuu unconformity Sariola
Interpretation
Monotonous turbidites as- Marine basin Jormua and Outokumpu sociated with ophiolites nappe complexes tectonic contact? Fluvial–muddy shelf Conglomerate–arenite– Vihajärvi tempestite sequence
Sub-Lower Kaleva unconformity Jatuli
Distinctive features
≥2440
Hokkalampi
Lateritic weathering profile Chemical weathering period
(a) Puolankajärvi (basal parts pass to paragneisses) (b) Kurkikylä, Saari–Kiekki
(a) Medium–thin bedded turbidites
(a) ?lowstand fan”Basinal Sariola”
(b) Mixed immature sedimentary-volcanic sequence ± glacigenic units on weathered basement
(b) Narrow rift basins developed on late Archean continent. ”Continental Sariola”
Sääperi, Särkilampi, Hetehongikko Kemi–Penikat
(a) Sharp unconformity (a) Inversion of Sumi rifts with the layered intrusions (b) Gradual change from (b) Mainly physical fresh basement to regolith weathering under arid–semiarid conditions
Kemi–Penikat–Koillismaa
Acid metalavas and layered Initial rifting of the Archean intrusions basement
Late Archean basement of the Kuhmo, Iisalmi, and Pudasjärvi complexes. *)The depositional basement and basal units of the Upper Kaleva are unknown, but it may have deposited upon the the Jormua and Outokumpu ophiolite complexes. Contact with the Lower Kaleva is presumably tectonic. The Upper Kaleva is generally considered allochthonous. This is supported by the fact that the detrital zircon populations dated so far do not allow a distinction between this tectofacies and the Svecofennian metasediments (Claesson et al., 1993).
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in Paanajärvi, Russia (Systra, 1996). Small occurrences are also reported from central Lapland (Manninen et al., 2001; Räsänen and Huhma, 2001).
3.2. 2440-Ma layered intrusions In Chapter 3 Iljina and Hanski examine the geology and petrogenesis of the layered mafic intrusions, and Alapieti and Lahtinen (2002) described platinum-group element mineralization. In the area reviewed (Figure 7.1), mafic layered intrusions are found in three main localities: within the Archean basement in the northern parts of the Kuhmo (Kuusijärvi in Figure 7.5) and Pudasjärvi (Suhanko in Figure 7.6) complexes, and between the Karelian formations and the Archean basement along the northwestern margin of the Pudasjärvi complex (Kemi–Penikat and Narkaus in Figure 7.6). It is important to notice that they reflect the sub-Sariola erosion related to rift inversion (next Section), that relics of Sariola cover rocks have been preserved upon two of them (Section 5.5), and that their ~2440 Ma age has been well established by several diverse methods (Hanski et al., 2001; Perttunen and Huhma, 2001, and references therein).
4. Sub-Sariola unconformity There are two types of unconformities below the Sariola rocks: an in situ breccia–satrolite zone upon the late Archean basement rocks and the one developed upon the 2440-Ma layered mafic intrusions. The Archean basement/Sariola contact is marked by a weathering crust or regolith. Probably the best-exposed example is found in the Pasvik–Pechenga area, northern Norway and Kola Peninsula, where its age is estimated to be between 2450 and 2350 Ma (Sturt et al., 1994). In eastern Finland, it has been described from Sääperi (Pekkarinen, 1979,
292
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Fig. 29), Särkilampi (op cit., Figs. 12 and 32) and Ilvesvaara (Kohonen and Marmo, 1992) in North Karelia, Lippumäki (Aumo, 1983) in the Kuopio area, Hetehongikko (Strand, 1988) in Kainuu, Kuntijärvi (Silvennoinen, 1972, Fig. 4) in Kuusamo, and Runkausvaara (Perttunen, 1991) in Peräpohja. In each case, a few meter thick zone of weathered/fractured basement granitoid is overlain by in situ granitoid breccia (Figure 7.7A), which passes into immature Sariola basal conglomerates (Figure 7.7B). Pekkarinen (1979) and Sturt et al. (1994) favored arid or semiarid climatic conditions, whereas Kohonen and Marmo (1992) attributed the origin of the Ilvesvaara regolith to mechanical disintegration by continental ice sheet. The Hetehongikko regolith is attributed to physical in situ weathering (Laajoki et al., 1989). The fluvial deposits overlying the Pasvik regolith preserve unweathered pyrite-magnetite clasts and terrigenous pyrite grains, which support an arid or semi-arid paleoenvironment, possibly with some deficiency of oxygen in the atmosphere–hydrosphere system during the beginning of the deposition of the Paleoproterozoic formations on the Fennoscandian Shield (Sturt et al., 1994). The unconformity below the Hepoköngäs Formation is exceptional as it does not contain regolith, but the polymictic Sariola conglomerate lies directly on the Archean basement (Figure 7.7C). It may represent a washed paleosurface. The unconformity between the layered intrusions and the Sariola rocks is exposed above the Kemi intrusion (Perttunen, 1991) and has been penetrated by several drill holes. It differs from the one developed on Archean granitoids, being sharp with only a thin, if any, weathering crust (Figure 7.7D). The petrologic studies of the layered mafic intrusions prove that all the hanging wall rocks and significant parts of the layered mafic intrusions themselves were eroded to varying depths before deposition of the basal Sariola conglomerates. This has been most convincingly reported
K A R E L I A N S U P R AC RU S TA L
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A
B
C
D
Fig. 7.7. Photographs of primary features of the sub-Sariola unconformity. (A) In situ weathering breccia zone upon Late Archean granitoid of the Kuhmo complex. Laanhongikko, Kurkikylä, Kainuu belt. Slide 2-145. (B) Basal conglomerate above the in situ breccia. All the clasts are basement orthogneiss in scanty muscovite-rich matrix. Kainuu belt. Slide 2-49. (C) Unconformity between the Kuhmo complex (lower part) and polymictic matrix-supported boulder conglomerate. Hepoköngäs, Kainuu belt. Slide 3-18. (D) Unconformity between the Kemi layered mafic intrusion (lower part) and the Sariola conglomerate. Clasts are mostly basement granitoids. Near Kemi airport, Peräpohja belt. Slide 31-108. Photos: Kari Strand (A, B) and Kauko Laajoki (C, D).
from the Penikat intrusion (Figure 7.8) and the Narkaus intrusion (Huhtelin et al., 1989). As the crystallization depths of the intrusions are not known, no exact numbers can be given to the thickness of the hanging-wall Sumi cover and the total rock column eroded. It can be estimated, however, that it was at least a few kilometer thick. The amount of erosion varies significantly from one structural block to another (Figure 7.8). Sumian volcanic rocks are preserved above the Koillismaa intrusion (Lauri et al., 2003) and they are separated from the Sariola basal conglomerate by a knife-sharp contact (cf. Karinen, 1998). The CHAPTER
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pre-Sariola fragmentation of most of the intrusions into several blocks indicates that the erosion was related to post-Sumi block movements. Melezhik and Sturt (1994) attributed the likely coeval erosion in Russia to rift inversion. This is supported by the fact that the block movements in northern Finland seem to have occurred only in the areas of the layered mafic intrusions. The sharpness and lack of regolith on this unconformity type poses an open question. One possible explanation is that the eroded parts of the layered mafic intrusions represented topographically higher levels than the S U P R AC RU S TA L RO C K S
•
293
550
4
N
Suomujärvi block
5
4 1–3
1–3 4
Sub-Sariola unconformity
1–3
4 1–3
4
7370
7370
1–3 4
AlaPenikat block
5
1–3
Penikat Layered Intrusion
KeskiPenikat block
4 1–3
Paleoproterozoic rocks of the Kemi belt Sub-Sariola unconformity 5 Megacyclic unit 5 4
Megacyclic unit 4 AP reef
1–3 Megacyclic units 1–3 & Marginal series
5 km
Late Archean basement
SSW
NNE Eroded part 5 4
3 km
5
Sub-Sariola unconformity
1–3 5 km
Fig. 7.8. Generalized geological map and reconstructed cross-section of the Penikat layered mafic intrusion between the late Archean basement rocks of the Pudasjärvi complex and the Paleoproterozoic supracrustal rocks of the Peräpohja belt (simplified from Halkoaho, 1994, Figs. 2 and 4). For location see Figure 7.6. Note that if the Keski-Penikat block is excluded, the sub-Sariola unconformity erodes progressively deeper levels of the layered intrusions from the Ala-Penikat block to the Sompujärvi block and that the intrusion was faulted into several blocks before or during this erosion period. The Keski-Penikat block has been lifted onto higher paleotopographic levels than the nearby blocks and has been eroded down to the AP reef of the Megacyclic unit 4.
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enclosing Archean basement, and were thus more apt to be worn away.
5. Sariola tectofacies The type locality of Sariola rocks is in Russian Karelia where Eskola (1919) described tillite-like conglomerates and arkosites on the gneiss granite basement. The most important Sariola occurrences in eastern Finland are those in North Karelia and Kainuu where these rocks were deposited nonconformably on the Archean basement of the Kuhmo block. They comprise mainly immature conglomerates, with abundant basement granitoid and gneiss clasts and arkosites, but basic–intermediate lavas are also common in Kainuu. Glaciogenic rocks give a special importance to these occurrences. The tectofacies classification used in this review has not been generally applied to the Karelian formations in the Kuusamo and Peräpohja areas. However, the Sariola tectofacies is well presented above the layered mafic intrusions and as the lowermost units of the Kuusamo and Peräpohja belts. Also the Karkuvaara and Ahola Formations as well as the latter’s correlative, the Puolankajärvi Formation, are included in this tectofacies. If the opinion of Sturt et al. (1994) – that the physical weathering crust encountered in the lower part of the Finnish Karelian formations represents the same crust as the Pasvik regolith in Norway – is correct, the maximum age limit of Sariola can be set at 2440 to 2330 Ma. This is also indicated by the occurrence of 2340 Ma (Mustalampi type) detrital quartz in the lowest Sariola arkosites at Kiihtelysvaara (Pekkarinen, 1979).
5.1. North Karelia The Sariola formations rim the southwestern margin of the Kuhmo block and the northern margin of the Sotkuma dome (Figure 7.4). They are characterized by mineralogically and CHAPTER
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KARELIAN
texturally immature fluvial conglomerates and arkosic sandstones, which have received their detritus from the late Archean basement. Good sedimentological descriptions are available in Pekkarinen (1979), Marmo et al. (1988), and Kohonen and Marmo (1992). In the following, only the glaciogenic Sariola rocks will be discussed in more detail.
Glaciogenic rocks of the Urkkavaara Formation Marmo and Ojakangas (1983, 1984) were the first to convincingly describe glaciogenic rocks within the Sariola tectofacies. The best evidence (see discussion by Ojakangas, 1985) reported was the presence of dropstones in thinly laminated units associated with diamictites in the Urkkavaara Formation, which occupies the middle part of the Kyykkä Group on the northern margin of the Sotkuma dome (Figure 7.4). The Urkkavaara Formation overlies the gravelly Ilvesvaara Formation, which was deposited nonconformably on the basement, whereas the Hokkalampi paleosol forms the erosional upper contact. The Urkkavaara Formation (maximum thickness 265 m) is subdivided into seven members (Kohonen and Marmo, 1992) whose deposition was associated with a grounded or floating glacier (Marmo and Ojakangas, 1983, 1984). The first member, lower siltstone-argillite (>15 m) with dropstones/lonestones in laminated and graded siltstone, is interpreted to have been deposited in front of the glacier as a silt-clay rythmite sequence. The lower graded sandstone member (15 m) with oversized clasts/lonestones (Figure 7.9A) was deposited closest to the glacier by turbidity currents during glacial advances. Glacial retreat revived siltstoneargillite deposition (member 3; 2–40 m) and this was followed by diamictites (member 4; 0–10 m, Figure 7.9B) in a more distal aqueous environment. The second advance of the glacier deposited the upper graded sandstone member (50 to 70 m) on top of the diamictites. The overlying but transitional parallel-bedS U P R AC RU S TA L RO C K S
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A
B
C
D
E
F
G
H
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ded conglomerates (member 6; 50 m) are interpreted as products of sedimentation from subaqueous melt-water tunnels during the same stage. These are overlain by proglacial sandur and/or esker conglomerates and pebbly sandstones (member 7; >50 m). Kohonen and Marmo (1992) correlated the Pölkkylampi and Särkilampi Formations in the Kiihtelysvaara area (Pekkarinen and Lukkarinen, 1991) with the Urkkavaara Formation and showed that the basement here is overlain by diamictites and siltstones containing dropstones. The diamictites seem to correlate with the basal “pre-Jatuli” breccias of Pekkarinen (1979).
5.2. Eastern part of the Kainuu belt The Sariola rocks are found as larger units in Kurkikylä, where they lie upon the Kuhmo complex, and as small erosional relics on the small allochthonous Väyrylänkylä basement wedge within the belt. These have been mapped as the Kurkikylä and Honkajärvi Groups, respectively (Laajoki, 1991). Both groups contain significant volcanic units and resemble in this respect more the Sariola rocks in the Kuusamo–Paanajärvi belt than those in North Karelia. Glaciogenic rocks have been described from the Honkajärvi Group. There is no unequivocal evidence for Sariola in the southeastern margin of the Kainuu belt, although the Ohravaara Group
(Gehör and Havola, 1988) as well as the feldspar-bearing quartzites and arkosites of the Syväjoki Formation of the Hyrynsalmi Group (Kontinen, 1986) may be part of it. The latter group is underlain by a ~50-m-thick regolith of in situ weathered Archean basement in several places (op. cit.), but it is not known whether it represents a physically (sub-Sariola) or chemically (sub-Kainuu) weathered zone.
Glaciogenic rocks of the Honkajärvi Group Small patches of the Sariola sedimentary and volcanic rocks nonconformably cover the western margin of the tectonic wedge of Väyrylänkylä nappe (Figure 7.10). The sequence (at least 190 m thick) begins either with lavas and pyroclastic rocks or arkosic conglomerates and arkosites. The latter are overlain by a deformed, epidote-amphibolite facies association of granitoid-clast diamictites and laminated schists with dropstones (Figure 7.9C). Thin, massive, faintly graded sandy turbidite beds are also found (Figure 7.9D). The association is interpreted to have been deposited in a glaciomarine rift basin extending towards the interior of the Archean basement (Strand and Laajoki, 1993).
Kurkikylä Group The Kurkikylä Group forms a narrow, almost EW-striking belt in the southeastern corner of the region (Strand, 1988). There is a several
Fig. 7.9. (facing page) Photographs of Sariola rock types. (A) Thinly-bedded psammitic metaturbidites of the Urkkavaara Formation with a lonestone overlain by a thicker massive turbidite bed. Urkkavaara, North Karelia belt. Slide 18-177. (B) Diamictite of the Urkkavaara Formation. Urkkavaara, North Karelia belt. Slide 18-110. (C) Diamictite association of the Honkala Formation. Dmm–massive diamictite, Fld–laminated siltstone with dropstones (arrows). Nurmela, Kainuu belt. Slide 2-424. (D) Massive sandstone (Sm) in the Honkala Formation. Nurmela, Kainuu belt. Slide 2-422. (E) Medium-bedded massive metaturbidites of the Puolankajärvi Formation. Hakasuo, Paltamo, Kainuu belt. Slide 12-517. (F) Thin and graded-bedded staurolite-mica schist (metaturbidites) of the Puolankajärvi Formation. Kainuu belt. Slide 7-322. (G) Hummocky cross-stratified metapsammite of the upper part of the Puolankajärvi Formation. Kainuu belt. Slide 7-184. (H) Polyphase-folded turbiditic paragneiss within the Karkuvaara Formation. Posio, Kuusamo belt. Pegmatite veins on the left. Slide 30-92. Photos: Kauko Laajoki. CHAPTER
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S U P R AC RU S TA L RO C K S
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297
9
3
540
5 km
8
18 12
N
17 4
8
4
1
2
5
14
9
8
13 12
7
8
16
6
13
3
1
8
17 Jalka-aho syncline
4 16 3
3
9
1
4
12
17 3
2
2
13
15 16
3
9
17 14
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Fig. 7.10. Detailed map of the central part of the Puolanka area showing tight folding and faulting (simplified from Laajoki, 1991). Legend: 1−Paragneisses and granitoids of the Oulujärvi shear zone; 2−Puolankajärvi Formation; 3−Akanvaara Formation; 4−Pärekangas Formation; 5−lapilli tuff of the Central Puolanka Group; 6−Archean basement of the Kuhmo complex; 7−Sariola metasediments and metavolcanic rocks; 8−East Puolanka Group; 9−Somerjärvi Group; 10−Jatuli tuffites; 11−Jatuli dolomites; 12−Black schists and iron-formations of the Lower Kaleva; 13−Nonmagnetic mica schists of the Lower Kaleva; 14−Pyssykulju Formation; 15−Mäntykangas Formation (Vihajärvi Group); 16−Jalka-aho Formation (Vihajärvi Group); 17−Ultramafic rocks; 18−Metadiabase and metagabbro; 19−Fault (mostly inferred); 20−Top direction.VBW−Väyrylänkylä tectonic basement wedge.
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meters thick physical weathering crust (Figure 7.7A) between it and the Archean basement, characterized by a cracked zone in the Archean granitoid and an in situ breccia zone beginning with angular granitoid fragments that are hardly distinguishable from their surroundings. Upwards in the sequence, the fragments become smaller and the amount of sericite-rich interstitial material increases. The breccia gradually turns into a breccia conglomerate (Figure 7.7B), the material of which has slightly moved from its original site. The plutonic rocks of the Kuhmo block served as the provenance for this coarseclastic lowermost part, which is overlain by a fining-upwards sequence of fluvial arkosic conglomerates, sands, and clays. Volcanic activity associated with this phase produced 300- to 500-m-thick subaerial basaltic lavas (the Matinvaara Formation) that were locally deposited directly on the Archean basement. The lavas eroded and their clasts, together with basement granitoid clasts, were deposited as mixed volcanic-granitoid boulder conglomerates in an alluvial fan. The present erosional surface forms the top of this group, which seems to fill a shallow basin developed on the Archean basement.
5.3. Western part of the Kainuu belt The western margin of the Kainuu belt north of Oulujärvi is occupied by the Central Puolanka Group, which passes into the paragneisses of the Oulujärvi shear zone to the west (Figure 7.10). The depositional age of the Central Puolanka Group is still a matter of dispute. Originally, the group was considered Proterozoic and was included in the Kainuu belt (Laajoki, 1986). However, recent isotope studies indicate that the group could be late Archean (~2700 Ma) in age (Kontinen et al., 1996; Huhma et al., 2000). If this is correct, there should be a major tectonic-metamorphic nonconformity between the group and the Kainuu belt, representing a time gap of at least CHAPTER
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200 Ma. The newest field observations indicate, however, that the contact is more likely an intra-Karelian unconformity (Section 5.6; Laajoki and Wanke, 2002) and, accordingly, the Central Puolanka Group is considered Proterozoic. In the terms of tectofacies, the Puolankajärvi Formation and the two overlying formations characterized by immature arenites and pelites are included in the Sariola (Puolankajärvi) and Kainuu (the other two) tectofacies.
Puolankajärvi Formation The Puolankajärvi Formation consists of a nearly vertical, at least 500-m-thick and 70km-long sequence of sandy turbidites and overlying tempestites west of the paragneisses of the Oulujärvi shear zone (Figure 7.10). Its lower part (500 to 1000 m) consists of alternating members of massive or graded-bedded arkosites (Ta, Tab, Tabc turbidites, Figure 7.9E) and graded-bedded pelite–muddy sandstones (thin-bedded Tae turbidites, Figure 7.9F), which were deposited by high-concentration and low-concentration turbidite currents, respectively (Laajoki and Korkiakoski, 1988). The upper part (100 to 200 m), interpreted to have been deposited by storm waves and other shelf-processes, consists of micaceous sandstones showing hummocky cross-stratification (Figure 7.9G) and combineflow-origin ripple cross-lamination. On the basis of its turbiditic nature and solely Archean provenance (Kontinen et al., 1996; Huhma et al., 2000), the Puolankajärvi Formation is considered a more metamorphosed and distal equivalent of the Sariola turbidites of the Honkajärvi Group (Section 5.2).
5.4. Saari–Kiekki belt A small (3 km by 15 km) fault-bounded volcanic–sedimentary Saari–Kiekki belt is found within the Kuhmo block close to the Russian border (Figure 7.1; Luukkonen, 1989). The sequence begins with a ~500-m-thick unit of S U P R AC RU S TA L RO C K S
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breccia conglomerate interpreted as a talus deposit whose material was disaggregated by mechanical erosion caused by faulting of the Archean basement. This is overlain by a ~1000-m-thick volcanic unit of basalts, basaltic andesites and andesites, which compositionally correspond to high-Mg tholeiites, except for the intermediate volcanic rocks that are calc-alkaline. The uppermost unit consists of immature arenites and polymictic conglomerates with both basement granitoid and volcanic clasts derived from the belt. As a whole, this sedimentary–volcanic belt bears many similarities with the Kurkikylä Group.
The sedimentary–volcanic Sariola sequence above the Kuusijärvi intrusion is ~2000 m thick (Karinen, 1998; Karinen and Salmirinne, 2001). It begins with the 500-m-thick Unikumpu Formation, which mainly comprises clast- and matrix-supported polymictic boulder conglomerates. Most of the clasts represent basement ortho- and paragneisses, vein quartz and volcanic rocks, but there are also granophyre clasts from the hanging wall of the underlying intrusion and acid volcanic rocks likely derived from the eroded Sumi cover. The conglomerate is overlain by a ~1500-m-thick volcanic sequence that consists of a lower 150-m-thick unit of pyroclastic rocks and lava breccias and a principal 1400-m-thick unit of basaltic andesites–andesites; the latter are geochemically similar to the Sariola lavas in the Kuusamo belt. A small occurrence of polymictic conglomerate similar to those of the Unikumpu Formation overlies the Suhanko layered mafic intrusion (Isohanni, 1971).
chean granitoids at the northern margin of the Kuhmo complex (Silvennoinen, 1972). The lower part contains only basement granitoid clasts, but it passes into matrix-supported conglomerate in which also volcanic clasts are found in a tuffitic matrix. This type was referred to as a volcanic breccia by Silvennoinen (op. cit.) and Piispanen (1972), but the author considers it a matrix-supported conglomerate or diamictite that may include glaciogenic detritus, as the observed tillite-like features imply. The conglomerate contains also quartz porphyry clasts. A composite sample of three clasts gives a U-Pb zircon age of 2405 ± 6 Ma (Silvennoinen, 1991) or a slightly higher age (Hanski et al., 2001), indicating that the porphyry clasts represent Sumi acid volcanic rocks exposed in the nearby Paanajärvi area. The main part of Sariola consists of the possibly up to 500-m-thick Greenstone I Formation. It begins with a volcanic breccia–conglomerate, but is mainly composed of amygdaloidal or massive, non-magnetic basaltic andesite–andesite metalavas (Veki, 1991). A tuffitic schist considered a weathering product of the lavas overlies it; in this study, it is considered the upper boundary of Sariola in Kuusamo. On the basis of similar geochemistry and lithology, Räsänen and Vaasjoki (2001) indicated that the Karkuvaara Formation (the Posio greenstones, Figure 7.5) can be correlated with the Greenstone I, and it is included here in the Sariola tectofacies. It contains sandy metaturbidite interunits (Figure 7.9H) and thin sillimanite mica schist units between lava beds and is overlain by the poorly exposed turbiditic Ahola Formation (at least 100 m thick). The latter correlates with the Puolankajärvi Formation (columns 9 and 11 in Figure 7.2).
5.6. Kuusamo belt
5.7. Peräpohja belt
The Karelian sequence in Kuusamo begins with a thin (0 to 20 m) basal conglomerate, which lies directly on the paleoweathered Ar-
The volcanic–sedimentary lower part of the Karelian sequence in Peräpohja was previously subdivided into the Lower, Middle, and
5.5. Sariola cover of the layered intrusions within the basement complexes
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Upper Jatuli (Perttunen, 1989). Recently it has been considered a single group, Kivalo (Perttunen et al., 1995). However, the tectofacies classification is also readily applicable in this area. Of the nine formations of the Kivalo Group, the two lowermost, the Sompujärvi and Runkaus Formations, are here included in the Sariola tectofacies (Figure 7.6). The Sompujärvi Formation is known for its typical Sariola conglomerates lying either on the Archean basement or on the layered mafic intrusions (Figure 7.7D). In addition to conglomerate, this ~50-m-thick unit also contains arkosites in its upper part. Importantly, the conglomerate overlying the Narkaus layered mafic intrusion contains PGE-bearing clasts from the intrusion (Alapieti et al., 1989). The Runkaus Formation (40 to 100 m) consists mainly of subalkaline, tholeiitic basalts (Perttunen, 1989). The ~2250 Ma Pb-Pb age of secondary titanite-leucoxene provides a minimum age for the Runkaus Formation (Huhma et al., 1990). A ~2200 Ma layered metadiabase sill intrudes the contact with the overlying quartzite. These results and the age of the layered mafic intrusions bracket the age of the Sariola tectofacies between 2420 Ma and 2250 Ma in this area.
5.8. Other Sariola occurrences Typical Sariola basal conglomerates with well-rounded basement and quartz cobbles in arkosic matrix, gradually overlain by arkosites and sericite quartzites, are reported from the Nilsiä region. They are best preserved in NW-trending Reittiö belt within the Iisalmi complex (Paavola, 1984). A similar sequence is found in Lippumäki, in the Kuopio area (Aumo, 1983). Likely Sariola rocks are also present in Paltamo (Kärki, 1988).
6. Sub-Kainuu unconformity A period of intense chemical weathering followed the deposition of the Sariola rocks. CHAPTER
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This is best exemplified by the greenschist facies Hokkalampi paleosol in North Karelia. Marmo’s (1992) thorough study of this occurrence is based on extensive mapping of outcrops, almost 3000 m of drill core and over 500 chemical analyses. The original soil developed either on late Archean granitoids or glaciogenic Sariola rocks. The paleosol consists of quartz-sericite schists up to 80 m thick with an increasing proportion of kyanite and andalusite toward the top and a gradual increase in the values of the chemical index of alteration (CIA). It is subdivided into three zones. (1) An upper zone composed of kyanite/ andalusite-quartz schist (Figure 7.11A) corresponding to a zone of intense kaolinization above the water table, characterized by removal of all soluble minerals. The soil weathering residue consisted essentially of kaolinite and quartz. The present alumina content and CIA value reach 30 and 96%, respectively. (2) An intermediate zone of quartz-sericite schist, corresponding to a mixture of kaolinite and quartz. (3) A lower zone of quartz-feldsparsericite rock corresponds to a zone of partially disintegrated parent rock with low CIA values (60 to 65) and showing a gradual contact with the underlying parent rock. The Hokkalampi occurrence contains almost 30 million tons of rock with 3 to 4 million tons of kyanite and andalusite. The chemical maturity and great thickness of the Hokkalampi paleosol and associated metasediments are interpreted as recording intense chemical weathering in a warm and humid climate, comparable to a modern tropical climate. Assuming a modern annual rate of kaolinization of 0.1 to 0.01 mm, the formation of the Hokkalampi paleosol would have lasted several million years. Pekkarinen (1979) described a zone above basal arkosites in the Kiihtelysvaara area where sericite content increases upward. Pekkarinen and Lukkarinen (1991) attributed this unconformity to chemical weathering. Kohonen and Marmo (1992) did not consider S U P R AC RU S TA L RO C K S
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A
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Fig. 7.11. Photographs of primary features of the sub-Kainuu weathered crust. (A) Kyanite quartzite with quartz and tourmaline (black) clasts. Hokkalampi, North Karelia belt. Slide 13-123. (B) Foliated kyanite quartzite in Hallakulma, Kainuu belt. Slide 3-152. (C) Sheared, ~60-cm-thick sericite schist (paleoweathered crust) between the Late Archean granitoids of the Iisalmi complex (on the left) and the Kainuan quartzite of the Salahmi belt. Lähdemäki, Salahmi belt. Slide 14-447. (D) Basal breccia between the Late Archean granitoids of the Iisalmi complex and the Kainuan quartzite of the Salahmi belt. Clasts are mostly basement granitoids and vein quartz. Lähdemäki. Slide 14-337. Photos: Kauko Laajoki.
this unconformity as an equivalent of the Hokkalampi zone, although the distance between these two places is only ~50 km, but proposed that the unconformity described by Pekkarinen (1979) should lie at a lithostratigraphically much higher level. A paleosol similar to Hokkalampi is found in Hallavaara, in the very northeastern tip of the Kainuu belt (Figure 7.3). Here kyanite quartzites (Figure 7.11B), andalusite-kyanite quartzites, chloritoid schists, and sericite schists with CIA values of 98, 92, 86 and 72–77, respectively, are found (Strand, 1988). 302
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The Hallavaara paleosol was developed on the Sariola sedimentary and volcanic rocks and probably also on late Archean granitoids. In comparison to Hokkalampi, it is, however, poorly exposed and less extensive. These two penecontemporaneous occurrences ~300 km apart indicate that deep chemical weathering covered large areas in the eastern Fennoscandian Shield. The lateral extension of the unconformity should underlie the Akanvaara Formation of the Central Puolanka Group, but evidence for it may have been destroyed by metamorphism.
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In Nilsiä, Paavola (1984, p. 46) described sericite quartzites with quartz pebbles between the typical Sariola basal formations and the overlying orthoquartzite. Both the stratigraphic position and rock sequence indicate existence of a sub-Kainuu unconformity in the area. The regolith under the Salahmi belt on the western margin of the Iisalmi block is also considered sub-Kainuu, because it consists of chloritoid schist developed upon Archean amphibolites. Locally, a thin sericite-schist overlies the granitoids of the basement (Figure 7.11C). Where the schist is missing, the basement is overlain by a basal vein quartz and orthogneiss-clast conglomerate, which passes into a rather mature quartzite (Figure 7.11D; Korkiakoski and Laajoki, 1988). In Kuusamo, Silvennoinen (1972) considered the tuffite schist above Greenstone I as the weathering product of the latter and stated that the contact with the overlying sericite quartzite is gradual. The author studied the same contact at Vitikkovaara where the contact is completely exposed at the bottom of a ditch. Here a well preserved Greenstone I lava is overlain by an epiclastic schist with volcanic plagioclase clasts. This passes into a sericite schist with granules and pebbles of quartz, basement feldspar, and orthogneiss. The author considers this a typical sub-Kainuu weathering crust developed on a basic substratum (= Sariola metalavas). Juopperi (1976) also considered the chlorite schist overlying the Kuusijärvi Formation a sub-Kainuu weathering zone. In the Peräpohja area, the lower contact of the Kainuu tectofacies is placed above the Runkaus Formation. This is, however, problematic as the Runkaus Formation/Kivalo (recently renamed Palokivalo) Formation contact is unexposed or is intruded by a metadiabase sill (Perttunen, 1991). The fact that the Kivalo quartzite overlies both the Archean basement and the Runkaus Formation indicates that deposition was preceded by a significant eroCHAPTER
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sion period. Within the deeper-water western parts of the Kainuu and Kuusamo belts, the unexposed Puolankajärvi/Akanvaara and Ahola/Nilovaara contacts are supposed to represent the sub-Kainuu unconformity or its correlative conformity.
7. Kainuu tectofacies The chemical weathering period described in Section 6 was followed by erosion and sedimentation, which apparently eroded much of the paleosol and deposited sequences characterized by aluminous alluvial sediments. This tectofacies has its type locality in Kainuu, where it was first described by Väyrynen (1928, 1933). It should be noted, however, that a large part of Väyrynen’s (op. cit., 1954) Kainuu facies is nowadays included into the Jatuli tectofacies and, vice versa, a large part of Väyrynen’s Jatuli facies is included in the Kainuu tectofacies (for historical remarks, see Laajoki, 1991, p. 37).
7.1. Kainuu belt Korvuanjoki Group in Kainuu The Korvuanjoki Group (800 to 1200 m) is the type group of the Kainuu tectofacies. It consists almost solely of alluvial fan and braided river deposits (Strand, 1988) and begins with a typical Kainuan basal pebble conglomerate with vein quartz clasts (Figure 7.12A). This grades into an almost pure orthoquartzite (>95% quartz), which forms the bulk of the group. The material of the Korvuanjoki Group is mineralogically clearly more mature than that of the Kurkikylä Group, which indicates that its source was the sub-Kainuu chemical weathering crust. In its uppermost part, this thick quartz-arenitic succession contains detrital potassium feldspar grains indicating a partial reworking of the pre-existing weathering crust and a change in weathering conditions. S U P R AC RU S TA L RO C K S
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Middle and Upper part of the Central Puolanka Group The turbiditic–tempestitic Puolankajärvi Formation of the Sariola tectofacies is overlain by the Akanvaara Formation, which is a ~800m-thick, upwards maturing quartzite unit. It begins with trough cross-bedded feldspathic quartzites (Figure 7.12B) overlain by a rather monotonous parallel laminated quartzite, which passes to low-angle cross-bedded micaceous quartzite in its upper part (Figure 7.12C). The lower-middle part of the formation is interpreted fluvial, whereas the upper part is shallow marine. The Akanvaara Formation is overlain rather sharply, but conformably, by the Pärekangas Formation (at least 500 to 800 m thick). It is poorly exposed, but the outcrops available display well preserved primary structures (Laajoki, 1994). The lower part comprises thin-bedded heterolith units with combine-flow ripple cross-bedding and graded bedding interbedded with thicker hummocky cross-stratified (Figure 7.12D) and low-angle cross-bedded quartzite units. The middle part of the formation is more pelitic and quartzite interbeds are common again in the uppermost part. The upper part of the formation is volcanic–sedimentary, consisting of alternating beds of pelite, quartzite, tuffs (Figure 7.12E), and lavas. The Akanvaara–Pärekangas couple is interpreted to be a Type 1 sequence. The
lower-middle and upper parts of the Akanvaara Formation represent lowstand and transgressive systems tracts, respectively, whereas the Pärekangas Formation records the highstand systems tract.
7.2. North Karelia In North Karelia, on top of the Hokkalampi paleoweathering zone or upon its correlative zones, there are sericite quartzites with frequent matrix-supported quartz-pebble conglomerate interbeds, which pass gradually upwards to purer orthoquartzite. In the Koli area, the lower conglomeratic part is interpreted as consisting of mass flow deposits with an abundance of detritus from the Hokkalampi weathering profiles and flash flood deposits overlain by proximal braided river or alluvial fan deposits (the ~60-m-thick Vesivaara Formation; Kohonen and Marmo, 1992). These are overlain by braidplain bar and channel conglomerates and diverse fluvial gravels and sands (the ~250-m-thick Koli Formation). If the unconformity in Kiihtelysvaara is considered sub-Kainuu, then the Haukilampi Formation (Pekkarinen and Lukkarinen, 1991) belongs to the Kainuu tectofacies. Väyrynen (1933) and Nykänen (1971) considered this “lower Jatulian quartzite formation” as Kainuu, whereas Ojakangas et al. (2001) included it in their Lower Jatuli Group.
Fig. 7.12. (facing page) Photographs of Kainuu rock types. (A) Typical quartz-pebble conglomerate portions in Kainuan basal quartzite. Korvuanjoki Group, Kurkikylä, Kainuu belt. Slide 2-246. (B) Overturned large-scale cross-bedding in the basal feldspathic part of the Akanvaara Formation. Huosiuslampi, Kainuu belt. Slide 8-95. (C) Low-angle cross-bedded sericite quartzite. Upper part of the Akanvaara Formation, Pärekangas, Kainuu belt. Photo 10860. (D) Heterolith of the Pärekangas Formation with lens and wavy bedding and hummocky cross-stratification (under the compass). Lehtomäki, Paltamo, Kainuu belt. Slide 12-588. (E) Lapilli tuff of the uppermost part of the Pärekangas Formation. Haapala quarry, Kainuu belt. Photo 10552. (F) Folded cross-bedded quartzite sericite quartzite.Välivaara, Kuusamo belt. Photo 10924. (G) Tidal heterolith of the Erivaaransuo Formation with deformed mud cracks. Erivaaransuo, eastern Kuusamo belt. Photo 10621. (H) Combine-flow ripples in scapolite spotted heterolith, the Kirintökangas Formation. Kirintökangas, western Kuusamo belt. Slide 30-166. Photos: Kari Strand (A) and Kauko Laajoki (B–H). CHAPTER
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7.3. Kuusamo and Kuusijärvi The base of the Kainuu tectofacies is convincingly defined above the Greenstone Formation I, more problematic is the position of the boundary between Kainuu and Jatuli. In this paper, the top of Greenstone II is considered to represent this boundary. The Kainuu rocks in the eastern main part of the Kuusamo belt begin with the Sericite Quartzite Formation (100 to 200 m), which consists of typical Kainuan fluvial sericitefeldspar quartzite with occasional quartzpebble layers (“arkose conglomerate”) and dominant sericite quartzite (Figure 7.12F; Piispanen, 1972; Silvennoinen, 1972, 1991). The quartz clasts in the lower part of the formation are distinctively blue. They are similar to the quartz xenocrysts/vesicle fills in the Sumi porphyries in the Paanajärvi area, indicating that they were derived either directly from this source or from Sariola sediments rich in porphyry detritus. The sericite quartzite is overlain by the heterolithic Erivaaransuo Formation (>50 m thick) that consists of several tidal parasequences (Figure 7.12G) and records a period of transgression (Laajoki, 2000). Pekkala (1985) described stromatolites from the 30- to 35-m-thick dolomite member in the upper part of this formation. In the west, in Posio, the Kainuu tectofacies begins with quartzites of the Nilovaara Formation, which is intruded by a ~2200 Ma metagabbro sill (Evins and Laajoki, 2001). The dominant rock type is a cross-bedded fluvial sericite quartzite that is feldspathic in the lower part and becomes more mature upwards. It commonly contains fuchsite and can be considered a typical representative of the traditional Lapponian quartzite (see Chapter 4). The formation is a likely correlative of the Sericite Quartzite Formation in the east and the Akanvaara Formation in Kainuu (but is much thicker, ~2000 m). It is overlain by the shallow-water Kirintökangas Formation (>500 m) of alternating tidal sandstone and 306
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tempestitic heterolith units (Figure 7.12H; Laajoki, 1997) and a yet unnamed poorly exposed silty quartzite unit. It seems possible that the Kirintökangas Formation represents the more distal equivalent of the Erivaaransuo Formation. In terms of sequence stratigraphy and in agreement with the Central Puolanka Group, the Nilovaara and the Kirintökangas–Erivaaransuo Formations seem to represent lowstand-transgressive and highstand systems tracts, respectively. The highstand systems tract is terminated locally in the east by the agglomerates and pillow lavas of Silvennoinen’s (1972) Greenstone II. In the Kuusijärvi area, the Kainuu tectofacies is represented by the 500-m-thick Hautavaara Formation of sericite quartzite with minor mudstones in its upper part (Karinen, 1998).
7.4. Peräpohja Typical basal Kainuan rock types (quartzpebble conglomerate and sericite schist) seem to be missing in the Peräpohja area. However, the rather sericite-rich and feldspathic quartzites of the 1000- to 2000-m-thick Palokivalo (or Kivalo) Formation (Ojakangas, 1965) and local quartz-pebble conglomerates in the lower part of the Kallinkangas section (Perttunen, 1991) indicate that this formation can be included in the Kainuu tectofacies. This is confirmed by correlative quartzites in the Rovaniemi (Perttunen et al., 1996) and Vanttauskoski (Salonsaari, 1990) areas, which are even richer in micas and may contain fuchsite. The upper 50 m of the Kallinkangas section contains mud-cracked mudstone beds (Mikkola, 1960; Ojakangas, 1965).
7.5. Other occurrences Outside the main areas of the Karelian formations in Kainuu and North Karelia there are several small quartzite occurrences.
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Due to their size and incomplete sequences as indicated by several tectonic contacts (Lundqvist et al., 1996), it is hard to identify the tectofacies. One occurrence, that might belong to Kainuu is the Lähdemäki Formation (40 to 80 m thick) in the Salahmi belt (Korkiakoski and Laajoki, 1988). It consists of fluvial feldspathic quartzites and arkosites and is separated from the Archean basement by a thin chemical weathering crust/breccia (Figures 7.11C and D). It is also likely that the lower quartzite sequence of the Nilsiä belt (Figure 7.1) belongs to the Kainuu tectofacies as it lies either on the Archean basement or the Sariola formations (Paavola, 1984) and is overlain by a conglomerate with quartzite clasts (Section 5.8). This is supported by the relatively old U-Pb zircon age of the metadiabase sill (1967 ± 24 Ma, op. cit.) that intrudes this lower part. At the northern end of the Höytiäinen basin, the quartzites of eastern Juuanvaarat (Figure 7.4) correspond to the Koli belt, but are separated from it by faults (Kohonen et al., 1991).
8. Sub-Jatuli unconformity Evidence for the unconformity under the Jatuli tectofacies is sparse. The most convincing case is the alluvial Nenäkangas conglomerate in Kainuu (Laajoki, 1988a). At its type locality, this conglomerate lies upon the Kainuu tectofacies quartzite and contains, in addition to basement orthogneiss and vein quartz clasts, also quartzite clasts, which seem to have been derived from the underlying Kainuu tectofacies (Figure 7.13A). Thus, the sediments of the Korvuanjoki Group must have been lithified to the extent that they could supply rock fragments. The Nenäkangas unconformity itself is not exposed. Some 25 km south of Nenäkangas, a similar conglomerate is found close to the Archean basement, from which it is separated by a 20-m-thick member of basal breccia and conglomerate as well as CHAPTER
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laminated biotite- and chlorite-bearing phyllites. The basal breccia consists of basement fragments in a chloritic matrix (Figure 7.13B). The chloritic matrix, which cannot be an in situ weathering product of the basement granitoids, and the sharp upper contact of the breccia indicating post-breccia erosion, support the idea that this contact might not correlate with the sub-Sariola nonconformity in Hetehongikko. It may be possible that the argillitic Katajalampi Member belongs to the Sariola tectofacies, but if so, the Nenäkangas-type conglomerate of the Jatuli tectofacies was deposited on Sariola rocks. In any case, there must be a big erosional unconformity somewhere under the Nenäkangas-type conglomerate also at this locality. The sub-Jatuli unconformity is placed here at the base of the basal breccia. Other convincing evidence for a major period of sub-Jatulian erosion in Kainuu is the polymictic conglomerate of the Pitukansuo Member (up to 65 m thick), which begins the Somerjärvi Group (Kangas, 1985; Laajoki 1991). This group is the western equivalent of the East Puolanka Group, from which it is separated by the folded and faulted Salmijärvi basin (Figure 7.10). The well-rounded cobbles and boulders in the clast-supported conglomerates are mostly orthoquartzite, sericite quartzite, arkosite, and vein quartz with minor siltstone and schist (Kangas, 1985). The depositional basement of the Somerjärvi Group is not exposed in this area and the group most likely has tectonic contacts with the Central Puolanka and Vihajärvi Groups. However, the likely southern extension of the lower contact of the Somerjärvi Group is exposed at the Haapala quarry, ~30 km south of Pitukansuo (Figure 7.3). Here a metalava of the Pärekangas Formation is separated from the Jatuli orthoquartzite by a 10- to 15-cmthick biotite-chlorite schist with microcline porphyroblasts and well-rounded quartz and polycrystalline quartz clasts (Figure 7.13C). The schist is interpreted as a mixture of local S U P R AC RU S TA L RO C K S
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Pärekangas lava detritus and quartz-rich sand derived from a more distant source. As there is no structural or metamorphic evidence for a nonconformity between the Central Puolanka Group and the Jatuli orthoquartzite (cf. Section 5.3), this contact is considered an intra-Karelian unconformity, which marks a stratigraphic gap between the Kainuu and Jatuli tectofacies. Kohonen and Marmo (1992) correlated the lower contact of the Jero Formation, North Karelia, with the unconformity at Nenäkangas and considered it to be the base of a major depositional cycle. The base of the Jero Formation is typically gradational with the underlying Koli Formation, but an erosional contact was also observed. In places, the underlying Koli quartzite has been fragmented in situ and the breccia fragments are matrixed by coarse subarkose. The basal conglomerates of the Jero Formation are polymictic with basement and orthoquartzite clasts. Piirainen (1968) considered the corresponding boundary to be “a bigger hiatus”, and Piirainen et al. (1974) attributed it to block movements and erosion of uplifted blocks. Sorjonen-Ward (1997) attributed the Jero Formation to a transient episode of basement rejuvenation and rifting associated with intrusion of 2200-Ma mafic sills. Ojakangas et al. (2001) correlated the unconformity between their Lower and Upper Jatuli groups with the Nenäkangas unconformity.
Paavola (1984, pp. 29, 30) attributed the conglomerate beds with quartz and orthoquartzite pebbles in the middle sericite quartzite part of the Nilsiä belt to a short erosional period. The author is apt to consider this erosion surface as sub-Jatulian. It might be questionable to use unconformity observations far apart from each other for correlation purposes. However, the sequences under and above the unconformities are so similar in Kainuu and North Karelia that the author tends to correlate them with each other and informally calls their basal surface the “sub-Jatuli unconformity.” It is supposed to mark a significant period of deep erosion, as indicated by the presence of ubiquitous quartzite clasts from the underlying Karelian sequences in the overlying sequence. Satrolites and basal conglomerates are found between the Archean basement and a thin sequence of epicontinental quartzites around the basement domes and nappe cores in the Sotkuma (Gaál et al., 1975) and Heinävesi areas (Koistinen, 1993). A little higher in the stratigraphy, there are local quartzite conglomerates and polymictic conglomerates, but these areas are so deformed that no explicit correlation with the Sariola and Jatuli rocks in North Karelia has been done. In the Kuusamo and Peräpohja areas, the Greenstone III and Jouttiaapa Formation, respectively, mark the Kainuu/Jatuli bound-
Fig. 7.13. (facing page) Photographs of the sub-Jatuli unconformity and Jatuli rock types. (A) Matrixsupported Nenäkangas conglomerate of the basal part of the East Puolanka Group. Nenäkangas, Kainuu belt. Slide 18-75. (B) In situ breccia above the granitoid of the Kuhmo complex (on the left) overlain by chloritic schist. Kolkonkangas, Kainuu belt. Slide 3-2. (C) Contact between the foliated metalava of the Pärekangas Formation (on the left) and Jatuli quartzite. Top to the right. Mixed detritus zone above the metalava is indicated by the white line. Photo 10456. (D) Cross-bedded quartzite. Sl = low-angle cross-bedded quartzite. Paleocurrents by Kari Strand. Siikavaara, Kainuu belt. Slide 4’-184. (E) Alternating tidal mudstone and quartzite beds. Lower part of the Rukatunturi Formation. Ronkonriutta, western Kuusamo belt. Photo 10637. (F) Stromatolites in the Kvartsimaa Formation. Kvartsimaa, Peräpohja belt. Slide 31-69. (G) Laminated turbiditic tuffite with one thick bed of the Tikanmaa Formation. Ossaus, Peräpohja belt. Slide 31-231. (H) Stromatolitic and laminated dolomite of the Rantamaa Formation. Rantamaa quarry, Peräpohja belt. Photo 11229. Photos: Kauko Laajoki and Pekka Härmä (D). CHAPTER
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ary. It is possible that the base of the Himmerkinlahti Member (Laajoki, 2000) and the Kolmiloukkonen Formation in Posio marks this unconformity, but as depositional ages of these occurrences are not known, they are problematic (Section 14.4).
9. Jatuli tectofacies The Jatuli tectofacies consists of quartzite-dominated sequences which, especially in the north, contain volcanic units in the middle parts and carbonate rocks, tuffites, and schist units in the upper parts. Type areas are North Karelia and Kainuu, where various lithostratigraphic classifications have been established. As systematical sequence stratigraphy has so far only been applied to the East Puolanka Group, interbasinal correlations are arbitrary. Carbon isotope studies by Karhu (1993; Chapter 16) indicate that the Jatuli may be subdivided into two stages, of which the older one is distinguished by the carbon isotope stage III (CIS III) carbonates enriched in 13C (δ13C values from about 8 to12.5‰). These rocks were deposited at ~2200 to 2100 Ma. The carbonates of the second Jatuli stage were deposited between ~2100 to 2060 Ma during the carbon isotope stage IV (CIS IV), which records an approximately 10% drop in the δ13C values of sedimentary carbonates. In the following, all references to CIS III and CIS IV carbonates are taken from Karhu (1993).
9.1. Koli and Kiihtelysvaara areas in North Karelia After deposition of the quartz sandstone of the Koli Formation, a new sedimentary cycle commenced with deposition of the immature Jero Formation (~1000 m). This formation begins with a braidplain and alluvial plain conglomerate member whose material was derived from the underlying formations and also probably from the Archean basement 310
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(Kohonen and Marmo, 1992). The overlying homogeneous fluvial arkosite member seems to have a gradational contact with the overlying subarkosic to quartz arenitic Puso Formation (<1200 m) of nearshore/shelf deposits with minor fluvial input. If the correlation between the Koli and Kiihtelysvaara areas suggested by Kohonen and Marmo (1992) is used, the Jatuli tectofacies in Kiihtelysvaara would consist of two quartzite formations separated by a thin (30 to 80 m) metalava unit, the Koljola Formation. The lower quartzite formation, Haukilampi (<550 m; Pekkarinen, 1979; Pekkarinen and Lukkarinen, 1991), consists of a gradual sequence from sericite quartzite with quartzpebble conglomerates to orthoquartzite. Based on conventional U-Pb zircon dating of the associated metadiabases, the age of the Koljola Formation is ~2120 to 2100 Ma (op. cit.). The extrusion of this lava unit was followed by a period of erosion as verified by volcanic clasts in the overlying upper quartzite formation (Pekkarinen, 1979). This surface may be the counterpart of the ~2080-Ma sub-Ludicovian unconformity in Russian Karelia (Ojakangas et al., 2001). The Kalkunmäki Formation (Pekkarinen and Lukkarinen, 1991) is overlain by the Viistola CIS III dolomites, minor volcanic lavas and tuffites, hematite rock, and the Petäikkö CIS IV dolomites and black schists, which Väyrynen (1933) included in his Marine-Jatuli facies. It is significant that the Koljola Formation and the overlying Kalkunmäki Formation (<250 m) and “Marine Jatuli” sequence (? >400 m) are missing in the Koli area; this may be due to erosion or tectonic removal of these units. As the 2200-Ma low-Al tholeiite sills (“karjalites”) have intruded up to the middle part of the Jero Formation (Vuollo, 1994; Chapter 5), the age of the formation is at least 2200 Ma.
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9.2. East Puolanka Group and corresponding groups in Kainuu The East Puolanka Group comprises the northeastern margin of the Kainuu belt. It was deposited unconformably on top of the Korvuanjoki Group or the Archean basement. As discussed above, this is demonstrated by the basal Nenäkangas conglomerates. The at least 2300- to 2600-m-thick East Puolanka Group consists mainly of mature quartzites (Figure 7.13D) with minor mudstones and heteroliths and one thin metalava unit. Their depositional systems (fluvial, high wave-energy shoreline, tidally influenced shoreline, and inner shelf) define four transgressive systems followed by highstand periods (Figure 7.14; Strand, 1993; Strand and Laajoki, 1999). The upper contact of the group is tectonic, which has resulted in the absence of “Marine Jatuli” rocks. Although poorly exposed, the Somerjärvi Group (>1000 m) offers a good section of the Jatuli formations on the western margin of the Kainuu belt (Figure 7.10). It begins with the polymictic quartzite clast-dominated conglomerates of the Pitukansuo Member (Section 7.8) overlain by mostly fluvial quartzites of the Eskosenvaara Formation (Kangas, 1986; Laajoki, 1991). In its upper part, this formation contains dark homogeneous hematite-matrixed orthoquartzite indicating a gradation to the “Marine Jatuli”. The latter is represented by a volcanic unit of several discontinuous tuffite bodies and the overlying thick (<200 m) CIS IV dolomite-marl formation. The southern extension of this group in Paltamo is known as the Melalahti Group, the upper part of which contains CIS IV dolomites, mica schists, and black schists (Kärki, 1988). The quartzites of the East Puolanka Group can be traced to Paltamo, where they turn towards the northeast to Ristijärvi as the thick (up to 3500 m) Hyrynsalmi Group (Kontinen, 1986) and continue to Sotkamo as the Vuokatti Group (Gehör and Havola, 1988). The Hyrynsalmi Group can be divided into two CHAPTER
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transgressive series, which are separated from each other by a thin (~50 m) paleoweathered metalava unit. The uppermost units consist of low-angle, cross-bedded shoreface orthoquartzite, dolomite, and phyllite with thin tuffite interbeds, thus demonstrating that this sedimentation phase ended also here in a continental shelf environment (“Marine Jatuli”) with CIS III dolomites accompanied by some volcanic activity.
9.3. Kuusamo In the west, Greenstone II is overlain by the Siltstone Formation (200 m), Greenstone Formation III (200 m), the Rukatunturi Quartzite Formation (600 to 800 m), the Dolomite Formation (50 to 100 m), and the Amphibole Schist Formation (>250 m, Silvennoinen, 1972). Pekkala (1985) added the Limestonedolomite Formation (100 to 200 m) on top of the Amphibole Schist Formation, but it has a tectonic contact with the main Kuusamo belt. Greenstone II and part of the Rukatunturi Formation are cut by 2209 ± 9 Ma and 2078 ± 4 Ma metadiabases. Hanski et al. (2001) were doubtful of the latter age and suspect that this diabase could also belong to the 2200 Ma group. Strand and Laajoki (1999) correlated the lower tidal part of the Rukatunturi Formation (Figure 7.13E) with the transgressive system that begins the second sequence in the East Puolanka Group (Figure 7.14). In Karhu’s (1993) classification, the Dolomite Formation and the Limestone-dolomite Formation belong to carbon isotope stages III and IV, respectively. Most of the Jatuli seems to be missing in Posio, where only thin layers of conglomerates and overlying siltstones are found. The poorly exposed Himmerkinlahti Member (>8 m) is interpreted as a minor coarse-grained alluvial braid/braidplain delta. It is considered to indicate a Type 1 sequence boundary, which is tentatively correlated with the one represented by the Greenstone II in the west (Laajoki, 2000). S U P R AC RU S TA L RO C K S
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Dominant depositional environment
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Relative water depth
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Tidal channels Sand shoals of inner self
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Influence of shoreline processes Tidal channels Foreshore & upper shoreface tidal flat
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Backbeach lagoon stacked foreshore & upper shoreface
Vuorivaara Fm.
Incised valley
Kiskonkoski & Naulaperä Fms.
HST Condensed section
Alluvial plain
1 TST (LST)
Fig. 7.14. Sequence stratigraphy of the East Puolanka Group (modified from Strand and Laajoki, 1999). LST, TST, and HST – lowstand, transgressive, and highstand system tracts, respectively.
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This is, however, tentative, as the conglomerate contains a lot of albitite clasts, whose age is not known and which were most likely derived from the albitite dikes cross-cutting nearby quartzite formations (see Section 14.4.) Another problematic unit is the polymictic Päästäispuro conglomerate, which seems to overlie the Kirintökangas Formation and contains diverse quartzite, metabasite, schist, and granitoid clasts in epidotized gneissic quartz-feldspar matrix (Section 14.4).
9.4. Peräpohja The Jatuli tectofacies is best represented in the southern part of the belt, north of the city of Kemi. It consists of the metavolcanic 2090 ± 70 Ma (Sm-Nd age; Huhma et al., 1990) Jouttiaapa Formation (300 to 1000 m), the Kvartsimaa Formation (50 to 200 m) of tidal orthoquartzites, CIS III stromatolitic dolomites (Figure 7.13F), and one polymictic conglomerate interbed (Ukonköngäs), the turbiditic tuffites (Figure 7.13G) of the Tikanmaa Formation (200 to 300 m), and the tidal Rantamaa Formation (100 to 300 m) of CIS IV stromatolitic dolomites (Figure 7.13H) with thin orthoquartzite interbeds (Perttunen, 1989, 1991). The Jouttiaapa Formation, which consists of tens of subareal lava flows (Perttunen, 1989), may mark the change from the Kainuu tectofacies to the Jatuli tectofacies. The stromatolites have been described by Härme and Perttunen (1963), Krylov and Perttunen (1978), and Kortelainen (1998). According to Kortelainen (op. cit.), who carried out a detailed sedimentological and isotope geochemical study of the lower greenschist facies CIS IV Rantamaa Formation, the dominant stromatolites are domes or columnar and indicate a supra intertidal depositional environment for the formation. She also described subtidal carbonate turbidites. As a whole, the basin into which the Rantamaa carbonates deposited was a moderately low energy epeiric platform and/or ramp. On the CHAPTER
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basis of Perttunen’s (1991) lithostratigraphic and Karhu’s (1993) and Kortelainen’s (1998) isotope studies, the Jatuli in Peräpohja can be considered as the type area of CIS III and CIS IV in the Fennoscandian Shield.
9.5. Other occurrences There are several smaller quartzite and dolomite occurrences, which might be included in the Jatuli (see Kousa et al., 2000). In the southern end of the Höytiäinen basin is the Tohmajärvi volcanic complex (Nykänen et al., 1994), which contains both pillow lavas and pyroclastic ashes and with which the Hammaslahti copper ore is closely associated. The complex is ~2100 Ma old (Huhma, 1986) and thus belongs to Jatuli not to Kaleva, within which it is exposed in an anticline. In the Kuopio area, a thin (<200 m) quartzite formation is overlain by a CIS IV dolomiteskarn unit (<200 m) (Aumo, 1983). Also the Nilsiä, Ala-Siikajärvi, and Juuanvaarat quartzite belts contain in their upper parts dolomites (both CIS III and CIS IV are present in Ala-Siikajärvi) and other rocks typical of the “Marine Jatuli” (Paavola, 1984; Kousa et al., 2000). Ekdahl (1993) included part of the volcanic rocks, mica schists, and migmatites of the Pielavesi area in Jatuli.
10. Sub-Lower Kaleva unconformity The traditional Kaleva includes all the Karelian metagraywackes, mica schists, mica gneisses, and associated rocks that overlie the Jatuli tectofacies or, in places, the Archean basement. They fill the cores of the Nuasjärvi and Höytiäinen basins in Kainuu and North Karelia, form the bulk of the Outokumpu nappe complex, and form a faulted belt from Kuopio via Salahmi to the Kiiminki belt (Figure 7.1). These formations are vaguely subdivided into Lower and Upper Kaleva. S U P R AC RU S TA L RO C K S
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No unconformity between them has been described, and the subdivision is based on lithological and geochemical correlations and the the 1953 ± 2 Ma Jormua ophiolite (Kontinen, 1986; Peltonen et al., 1996; Chapter 6). Conventional detrital zircon isotope studies demonstrate that the Lower Kaleva had mainly an Archean provenance, whereas in the Upper Kaleva a Proterozoic provenance is dominant (Huhma, 1990). The marked unconformity between the Jatuli and lowermost Kaleva formations was shown first by Väyrynen (1933), who described several basal Kaleva formations in North Karelia and Kainuu. These are conglomerates–sedimentary breccias with Jatuli quartzite and other clasts in micaceous or arenitic matrix. Kohonen (1995) questioned this interpretation stating that the traditional basal Kaleva conglomerates in North Karelia are found at several different stratigraphic levels. As it may be, their occurrence implies a significant erosional unconformity between Jatuli and Kaleva. Conglomerates (Figure 7.15A) indicating the presence of a sub-Lower Kaleva unconformity have been described in North Karelia by Piirainen et al. (1974), Pekkarinen (1979), and Kohonen (1995). In Kainuu, the classical basal Kaleva unit within the Nuasjärvi basin is the Rieskavaara breccia (Väyrynen, 1933; Gehör and Havola, 1988). At Juurikka, Hyrynsalmi, a thin basal breccia that lies directly on top of the Jatulian quartzite is overlain by a polymictic conglomerate with poorly to well-rounded clasts of quartzite, granite and mica schist (Kontinen, 1986). According to Kontinen (1986) the lowermost Kaleva unit, the Kotila Formation, rests unconformably on the various units of the underlying Jatuli rocks and, in places, on the Archean basement, indicating strong block movements and deep local erosion before the deposition of the Kotila Formation. This unconformity has also been located in Puolanka (Laajoki, 1991). In the Kiiminki belt, the Lower Kaleva 314
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rocks lie nonconformably upon the Archean basement (Figure 7.15B; Honkamo, 1985). A similar relationship is obviously also present in the Salahmi belt (Korkiakoski and Laajoki, 1988). It is noteworthy that no coarse-clastic rocks indicative of sub-Lower Kaleva unconformity have been described from the Kuopio and Peräpohja area. In the latter, however, the Kainuu/Jatuli rocks pass into the schists/turbidites of the (Lower) Kaleva (Ojakangas, 1965; Salonsaari, 1990; Perttunen et al., 1996). In these cases, the Lower Kaleva contact is placed above the carbonate-bearing units that belong to Karhu’s (1993) CIS IV.
11. Lower Kaleva tectofacies 11.1. Kainuu belt The Lower Kaleva consists of a rather heterogeneous autochthonous–parautochthonous sequence of basal turbiditic conglomerates and breccias, quartz wackes, graywackes and shales as well as banded iron-formations. Kontinen (1986, 1998) attributed these rocks to a 2100-1950 Ma rift phase. In the Vuokatti area, this tectofacies begins with the heterogenic Torikylä Formation of diverse sedimentary breccias, e.g., the Rieskavaara breccia, and conglomerates, of which a part may be attributed to submarine canyons/channels and turbiditic sandstones (Gehör and Havola, 1988). These are overlain by the mixed chemical–clastic Tuomivaara Formation (5 to 50 m) containing quartz (chert)-banded iron-formation units with abundant mica schist and banded amphibolite units and a few turbidite sandstone interbeds. The amphibolite units most likely represent volcanoclastic rocks. Two uppermost units are the 50- to 100-m-thick Ruokonen Formation of turbiditic conglomerate and sandstone, and the Naapurinvaara Formation of monotonous massive quartz arenite and subarkosite.
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In the Hyrynsalmi area, the Lower Kaleva consists of local basal breccias and conglomerates with minor quartzites overlain by sandstones with black schist and iron-formation interbeds (Kontinen, 1986). Their origin is associated with rifting of the continental crust after the Jatuli, which formed steep fault ridges and escarpments in submarine basins of the half graben type. The breccia conglomerates represent submarine rockslides and the metasandstones are turbidite deposits. Similar rocks are also found in Melalahti (Kärki, 1988). In Puolanka, Kaleva fills the Salmijärvi basin (Figure 7.10) and begins with sedimentary breccias and turbidites deposited unconformably on the CIS IV dolomites. They are followed by quartz-rich turbidites (Figure 7.15C), black schists, mica schists, and ironformations. Quartz (chert) banded iron-formations (BIFs, Figure 7.15D) are a peculiar feature of the Lower Kaleva in Kainuu. They are exceptionally rich in C and P and include up to 2-cm-thick phosphorite bands (Figure 7.15E). These have negative Ce anomalies and are commonly rich in Mn (Laajoki, 1975; Laajoki and Saikkonen, 1977; Gehör and Laajoki, 1987; Gehör, 1994). They were previously considered Lake Superior-type and were included in the “Marine Jatuli.” As they are now known to belong to Kaleva (Kontinen, 1986) they seem to represent a unique phosphorous BIF type. The BIFs are interbedded with and overlain by sulfide-bearing black schists. A Ni-CuZn deposit hosted by the Lower Kaleva black schists is found in Talvivaara. The precursors of the black schists (sapropelitic muds) with average Cgraf and S contents of 7% to 8% and 8 to 9%, respectively, were deposited in anoxic conditions (Loukola-Ruskeeniemi and Heino, 1996).
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11.2. Höytiäinen basin, North Karelia The Höytiäinen basin is a narrow structure formed by faulting and mainly contains mica schists and turbiditic, mica-bearing metasandstones (Figure 7.4). The contact between the basin and the eastern quartzite belt is unconformable so that on top of the Kiihtelysvaara CIS III dolomitic series is a ~50-m-thick turbiditic metaconglomerate with clasts of Jatulian quartzites and lavas (Pekkarinen, 1979). The metaconglomerate is overlain by a graded turbiditic quartzite and the metapelites and graywackes of the Höytiäinen basin (Ward 1987, 1988; Kohonen, 1995). This sharp lithologic contact is apparently tectonic, at least locally. Kohonen (1995) interpreted it as a step-like, faulted rim of an ancient sedimentary basin, where the metaconglomerates represent several different stratigraphic levels with the oldest conglomerates lying farther towards the center of the basin. Ward (1987, 1988) interpreted the Höytiäinen basin as an en echelon transtensional, intracratonic rift basin. The volcanic activity at Tohmajärvi was apparently associated with the initial phase of rifting. Kohonen (1995) subdivided the Höytiäinen basin fill into five lithologic assemblages, which were interpreted to represent various shelf and basinal sediments ascribed to the development of the basin – a Wilson cycle between 2110 and 1880 Ma. A Sm-Nd isotope study indicates that the Lower Kaleva basins may have been dominated by a Jatuli/Archean source (Huhma, 1987).
11.3. Kuopio area The CIS IV Petonen Formation is overlain by a significant volcanogenic unit, the Vaivanen Formation (at its thickest ~200 m), which mainly consists of mafic, pillow metalavas (Aumo, 1983). The correlative Koivusaari Formation contains a 2062 ± 2 Ma felsic tuff (Pekkarinen and Lukkarinen, 1991). These formations have been included in Jatuli (e.g., Kousa et al., S U P R AC RU S TA L RO C K S
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2000). Here they are included in Lower Kaleva because the age of the Koivusaari Formation is close to that of the Keinokangas porphyry, which is found within the Lower Kaleva in Peräpohja (Section 11.6). In addition, the CIS IV dolomites in other areas are not known to be overlain by volcanic rocks but their upper contact is considered to represent the subLower Kaleva unconformity. The Koivukangas and Keinokangas volcanic assemblages seem to represent bimodal volcanism associated with the initial Kaleva rifting. The thick mica schist–mica gneiss unit containing intercalated beds of carbonate and black schist overlies the Vaivanen metalava and is generally included in the Lower Kaleva (Aumo, 1983). Ekdahl (1993) described from Pielavesi, northwest of Kuopio, dolomites, black schist, and associated rocks and included them into the “Marine Jatuli”. According to the classification used in this paper, however, at least part of them may belong to Upper Kaleva.
11.4. Salahmi belt
Late Archean Iisalmi basement (Figure 7.15F). This passes into the basinal Rotimojoki Formation (~1000 m) which has exceptionally well-preserved primary structures including diverse Bouma sequences (Figure 7.15G) and inverse- to normally graded conglomerates (Korkiakoski and Laajoki, 1988). The formation consists of three members: (1) the lower member of arkosic Ta(b)e turbidites of smooth mid/lower fan, (2) the middle member mainly composed of massive Ta(ab)e turbidites of channeled mid-fan, (3) the Ilkonaho member mainly composed of arenite-clasted turbiditic conglomerates–graded arkosites of upper/mid fan channels and laminated–thinly bedded and graded pelites–siltstones of levees and interchannel deposits of upper/mid fan. The small Itämäki belt, consisting of Kalevian staurolitemica schists and metapsammites, lies north of Salahmi. A similar small occurrence of turbiditic rocks is found in Otanmäki (Puumalainen, 1986). Nd isotope data on the schists from the Itämäki and Otanmäki belts imply Proterozoic provenance (Finnilä, 2000).
11.5. Kiiminki belt
The part of the Salahmi belt (Laajoki and Luukas, 1988) included in the Lower Kaleva begins with the coarse-clastic Haajainen Formation composed of basal debris flow turbiditic conglomerates with diverse granitoid and gneiss clasts, most likely derived from the
The relatively large Kiiminki belt is exceptional among the Lower Kaleva belts as it contains thick volcanic units between two thick turbidite (graywacke) units. In the southeast, the lowermost formation is the quartz- and
Fig. 7.15. (facing page) Photographs of primary features of the sub-Lower Kaleva unconformity and Lower Kaleva rock types. (A) Polymictic Lower Kaleva conglomerate with Jatuli quartzite clasts. Kortevaara, North Karelia belt. Slide 18-20. (B) Lower Kaleva metagraywacke deposited nonconformably on the granitoid of the Pudasjärvi complex. Mäkipalo, Kiiminki belt. (C) Medium-bedded quartz arenite metaturbidite. Roninkangas, Kainuu belt. Slide 4-318. (D) Quartz-banded mixed silicate-oxide facies iron-formation. Tuomivaara, Kainuu belt. Slide 12-278. (E) Three phosphorite bands (2–3 mm, black) in the Tuomivaara chert and iron-mineral banded iron-formation. Kainuu belt. (F) Basal conglomerate of the Haajainen Formation with mostly basement granitoid and gneiss clasts. Haajainen, Salahmi belt. Slide 14-362. (G) Upwards thinning Tae turbidite in the Rotimojoki Formation. Top to the left. Rotimojoki, Salahmi belt. Slide 14-152. (H) Clast-supported quartzite-clast conglomerate beds separated by a thin laminated sandstone cap. Top to the left. Taivalkoski, Peräpohja belt. Slide 31-132. Photos: Mikko Honkamo (B), Seppo Gehör (E), and Kauko Laajoki. CHAPTER
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granitoid-clasted Kalliomaa conglomerate deposited on the Archean Pudasjärvi complex (Kähkönen et al., 1986). The overlying Utajärvi Formation (Figure 7.1; probably 2 km to 3 km thick), consisting of arkosites and matrix-supported conglomerates similar to those in the Haajainen Formation, passes both laterally and vertically into the lower turbidite pile of the Vuotto Formation, which locally has been deposited directly on top of the Pudasjärvi Archean block (Figure 7.15B; Honkamo, 1985). This formation is separated from the upper turbidite unit of the Ylikiiminki Formation by several discontinuous volcanic units of basaltic tholeiites (Honkamo, 1987) with associated schists, minor carbonate rocks with rather low δ13CTot values (0.16‰ to 1.97‰; Karhu, 1993), BIF, chert, and local conglomerate and quartzite. Stratigraphic relations within the belt are not reliably known. Lundqvist et al. (1996) considered the upper turbidite sequence to be allochthonous, but Kousa et al. (2000) described the entire Utajärvi–Ylikiiminki sequence continuous. The 2093 ± 35 Ma zircon U-Pb age of a felsic porphyry clast from the polymictic Koiteli conglomerate overlying (or interfingering) the metalava unit gives a maximum age for the deposition of the conglomerate (Honkamo, 1988).
11.6. Peräpohja The Martimo Formation in the southern part of the belt overlies the Rantamaa Formation. This is a monotonous sequence of phyllites and mica schists with minor black schists traditionally included in Lower Kaleva. The formation is folded and its total thickness is unknown. A significant quartzite-pebbled conglomerate is found in Taivalkoski (Härme, 1949; Enkovaara et al., 1953; Perttunen, 1991). Unfortunately, this occurrence was not studied in detail before it was buried under a dam reservoir. On leftover outcrops, sedimentary facies includes pebbly cross-bedded quartzite, 318
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normally (Figure 7.15H) or inversely graded, clast-supported conglomerate, and rippled quartzite. The quartzite clasts may represent intra-Kaleva material, and the sequence as a whole resembles the Ilkonahonkallio member in Salahmi (see above). It may represent a channel fill of a submarine fan. In the north, in Mustamaa, brecciated/thinly banded uranium-bearing phosphorites have been described from the Rantamaa/Martimo contact zone. Their stratigraphic position and origin are not clear and they occur in dolomite, tuffite, chlorite schist, and marl (Yrjölä, 1982). This sequence is overlain by mica schists and metagraywackes (Korkalo, 1971) that, farther to the north, are overlain by the mafic volcanic rocks of the Väystäjä Formation (Figure 7.6). The Väystäjä Formation contains the 2050 ± 8 Ma Keinokangas porphyry (Perttunen and Vaasjoki, 2001), the only felsic effusive rock in the Peräpohja area. This porphyry is associated with felsic agglomerates/conglomerates that contain phosphorite fragments (Eeronheimo, 1979). It is not clear whether these fragments are from the Mustamaa occurrence or from the phosphorite bands in the Lower Kaleva BIFs in Kainuu. Their REE distribution and low U content point to the latter possibility. A turbiditic mica schist of the Martimo Formation has a Pb-Pb age of about 2575 Ma, indicating a dominant Archean provenance (Perttunen and Vaasjoki, 2001).
12. Sub-Upper Kaleva unconformity Upper Kaleva turbidites are supposed to have been deposited upon the Jormua and Outokumpu ophiolite complexes (Kontinen, 1987; Peltonen et al., 1996). Ward (1987) and Sorjonen-Ward (1997) considered that, although the eastern margin of his Savo province lithofacies (Upper Kaleva in this study) is allochthonous, evidence for local deposition upon Archean basement has been preserved,
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including an unconformity within the inlier at Oravisalo. However, Korsman et al. (1997) correlated the basal supracrustal rocks rimming the Oravisalo basement dome (Figure 7.4) with the Höytiäinen basin (Lower Kaleva in this study) and considered the contact with the surrounding mica schists (Upper Kaleva) tectonic (Figure 7.16A). As no direct observations of the sub-Upper Kaleva unconformity have been reported, the depositional basement of Upper Kaleva is unknown. If this tectofacies was deposited upon the ophiolite complexes, there should be at least a 30- to 50-Ma time gap between it and the Jormua and Outokumpu ophiolites, as the depositional age of Upper Kaleva is less than 1920 Ma (Claesson et al., 1993) and the ophiolites have been dated at 1953 ± 23 Ma (Peltonen et al., 1996) and 1972 ± 18 Ma (Huhma, 1986), respectively.
13. Upper Kaleva tectofacies 13.1. Upper Kaleva in Kainuu The Upper Kaleva is rather monotonous, consisting mainly of medium- to thick-bedded, mica-bearing massive or graded metasandstones or graywackes and thin-bedded mica schists. These rocks occupy the main core of the Nuasjärvi basin (Figure 7.3; Lundqvist et al., 1996; Kontinen and Peltonen, 1998). They were deposited by turbidity currents in deep water, or at least below the level of storm activity. Kontinen (1986) considered the transition from Lower to Upper Kaleva sedimentary. The latter was subsequently associated with the Jormua ophiolite complex in the allochthon that was probably emplaced as a thin thrust sheet across the foreland to Kainuu during an early stage of the Svecofennian orogeny, and later imprecated and folded by thick-skinned deformation (Peltonen et al., 1996; Kontinen and Peltonen, 1998). Detrital zircons from a sample near the Jormua ophiolite give a CHAPTER
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maximum depositional age of ~1939 Ma for Upper Kaleva in the Nuasjärvi basin (Claesson et al., 1993).
13.2. Upper Kaleva within the Outokumpu nappe complex and the Kuopio–Pielavesi area The Outokumpu nappe complex comprises allochthonous Proterozoic rocks between the Höytiäinen basin and the Raahe–Ladoga zone (Figure 7.1). It is structurally complex, consisting of several nappes of both Upper Kaleva rocks and basement slices. The lowermost nappes may also contain Lower Kaleva rocks. The rocks in the eastern, less metamorphosed part of the complex are monotonous metasandstones, mainly mass-flow deposits and turbidites (Figure 7.16B; Ward, 1988). Farther to the west, the Outokumpu association consists of several thrust sheets and contains a characteristic serpentinite-ophiolite band with associated volcanic rocks and chemical sediments as well as Outokumpu-type copper ores (Koistinen, 1981; Loukola-Ruskeeniemi, 1999). Closer to the Svecofennides, the rocks pass into paragneisses (Figure 7.16C) and migmatites. A Nd isotope study indicates that the Upper Kaleva rocks consist of sediments derived from orogenic sources to the west with a significant potassium feldspar-poor Archean component (Huhma, 1987). Ward (1987, 1988) favored passive margin sedimentation prograding either onto ocean floor or backarc, which derived its material from both the Archean basement and a primitive Proterozoic substrate. Detrital zircons of the sample collected near the Outokumpu ophiolite indicate a maximum depositional age of ~1924 Ma for Upper Kaleva in this area (Claesson et al., 1993). Based on extensive geochemical data, Kontinen and Sorjonen-Ward (1991) concluded that the Upper Kaleva sediments of S U P R AC RU S TA L RO C K S
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the Nuasjärvi basin and the Outokumpu nappe complex are so strikingly similar that they shared a common provenance and accumulated within a single large-scale depositional system.
14. Problematic younger Karelian formations In Puolanka, there are one larger and three small sedimentary formations that are considered younger than the Jatuli, but whose sequences deviate so much from typical Kaleva within the Kainuu belt that their correlation with this tectofacies is questionable. Due to unresolved fault tectonics, the stratigraphy of the northern part of the Peräpohja belt is also problematic. In addition, two unique conglomerate-quartzite units are found in Kuusamo. Although fragmentary and small, these occurrences give significant information on the younger Karelian sedimentation.
14.1.Vihajärvi Group and Haapalanmäki and Jokijyrkkä conglomerates The Vihajärvi Group forms a narrow (0.5 km to 1 km by 18 km) and tight syncline between the Central Puolanka and Somerjärvi Groups on the western margin of the Kainuu belt (Figures 7.3 and 7.10). It was previously considered to lie lithostratigraphically between these groups (Laajoki, 1991). New Nd isotope results of metapelites show εNd values from –4.9 to –0.6 and TDM from 2680 to 2226 Ma, which clearly indicate a mixed Archean–Proterozoic provenance and correlation with the Upper Kaleva, rather than Jatuli tectofacies (Kontinen et al., 1996). On the basis of its quite exceptional lithostratigraphy and problematic present position, this group is not included in the Upper Kaleva in this study, but its relation to the Kaleva rocks in the Nuasjärvi basin remains unsolved. The primary contacts of the group are not exposed, 320
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but are probably tectonic. The Vihajärvi Group (>700 m) forms a unique fluvial–storm-dominated muddy shelf sequence, which begins with a basal polymictic clast-supported conglomerate with quartzite, schist, and volcanic clasts, of which the major part may have been derived from the Central Puolanka Group (Figure 7.17A; Kangas, 1985). This is overlain by braided alluvial plain arenites, which gradually grade into a pelitic unit at least 420 m thick. The sandstones of the transitional part show diverse storm-generated facies of shoreline–inner shelf, including hummocky cross-stratification (Figure 7.17B), graded bedding, finingupwards cycles of parallel laminated muddy arenites (Figure 7.17C), and slump-structures (Figure 7.17D). The uppermost pelites–siltstones are thinly-laminated with solitary massive or graded sandstone interbeds, indicating outer shelf conditions. A strongly deformed conglomerate is exposed in a road cut along the Oulu–Kajaani highway at Haapalanmäki, south of Kivesvaara in Paltamo (Heino, 1983). This strongly deformed conglomerate lies with an angular unconformity on the Central Puolanka Group (Figure 7.17E). It must also lie unconformably on Jatuli tectofacies, as the nearby contact between this group and the Jatuli quartzite is vertical (Figure 7.13C). This indicates that even the Jatuli was tilted or folded before the deposition of the Haapalanmäki conglomerate. The quartzite cobble–boulder conglomerates are matrix- or clast-supported and may show reverse grading, indicating that they represent alluvial deposits derived from the nearby Jatuli quartzites. Similar quartzite-clast conglomerates seem to lie unconformably on the Central Puolanka Group at Jokijyrkkä (Enkovaara et al., 1953; Laajoki, 1991). Kontinen (1998) included the Vihajärvi Group in Upper Kaleva, but, as discussed above, the Upper Kaleva of the Nuasjärvi basin likely represents an allochthonous deep water sequence, whereas the Vihajärvi Group rather
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B
A
C Fig. 7.16. Photographs of primary features of Upper Kaleva rock types. (A) Deformed contact/unconformity (see the text) between the basement gneisses and Upper Kaleva metaturbidites. Oravisalo, Outokumpu nappe complex. Slide 18-188. (B) Thick-bedded metaturbidites. Outokumpu nappe complex. Slide 18-187. (C) Gneissic metaturbidites. Humaljärvi, Outokumpu nappe complex. Slide 18-160. Photos: Kauko Laajoki.
is autochthonous and represents a fluvial-shelf system. The Haapalanmäki conglomerate may represent a debris flow deposits of an alluvial fan, which derived its clasts from the nearby quartzites or quartzite-clast conglomerates.
14.2. Pyssykulju Formation Pyssykulju at Puolanka is a small hill surrounded by poorly exposed Lower Kalevian rocks at the northern end of the Salmijärvi basin (Figure 7.3). In the lower part of the hill, thin-bedded, massive hematite pigmented turbiditic metasandstone and mica schist are found and are overlain by a thin, cross-bedded conglomerate with quartzite clasts in orthoquartzitic matrix. The conglomerate CHAPTER
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is, in turn, overlain by pure orthoquartzite (Laajoki, 1988b). According to Nd isotope data, at least the material of the mica schist in the lower part of the formation seems to have been derived from Archean bedrock (TDM 3153 Ma; Kontinen et al., 1996). Strong fractionation of rare earth elements in the analyzed samples indicates, however, that the presence of younger material cannot be excluded (Finnilä, 2000). The clasts of the overlying conglomerate were most likely derived from the nearby Jatuli tectofacies (Laajoki, 1988b). It might be that the lower, turbiditic part belongs to Lower Kaleva, upon which the overlying conglomerate–quartzite sequence was unconformably deposited. The contact is not, however, exposed. S U P R AC RU S TA L RO C K S
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14.3. Northern margin of the Peräpohja belt The northern part of the Peräpohja belt is cut by an E-trending fault zone, which separates the pelitic Pöyliöjärvi and associated formations from the main southern part. This tectonic zone is revealed by geological (Figure 7.6) and aeromagnetic maps and by the Korkiavaara Formation north of the fault zone. Bulk detrital zircons give an apparent age of 1985 Ma for the provenance of an arkosite in the Korkiavaara Formation, indicating that the Upper Kaleva units may be present in this northern area (Perttunen and Vaasjoki, 2001). The Korkiavaara Formation is said to overlie, with gradational contacts, the Kalliovaara quartzite and underlie the Pöyliöjärvi Formation. The former seems to be of the Kainuan type and belongs to the southern part of the Peräpohja belt, whereas the latter comprises mica schists and gneisses. The fact that the area consists of D2 antiforms and synforms and that the rocks are gneisses without clear top indicators, the stratigraphic relationships given should be considered only tentative. On the basis of available data (Salonsaari, 1990; Lappalainen, 1994; Väänänen et al., 1997; Perttunen et al., 1996) it is likely that the Kalliovaara and Pöyliövaara Formations may be
allochthonous in respect to the main southern part of the Peräpohja belt. Although the apparent provenance age of the Korkiavaara Formation is relatively young, a direct correlation of these formations with the allochthonous Upper Kaleva in Kainuu and North Karelia is not necessarily warranted.
14.4. Himmerkinlahti Member and Kolmiloukkonen Formation in Posio The Himmerkinlahti Member (HL in Figure 7.5; Laajoki, 2000) is found as small solitary outcrops and, consequently, its relations with other parts of the Kuusamo belt are not reliably known. The main significance of the Himmerkinlahti Member is that it contains conglomerate beds with abundant Karelian metabasite, albitite (Figure 7.17G), and other lithic clasts, which show evidence of postdepositional alteration. The member was interpreted to represent a minor coarse-grained alluvial braid/braidplain delta and thus yield evidence for a period of significant subaerial erosion during deposition of the early Proterozoic Karelian formations in northern Finland. Its rocks are probably part of a lowstand prograding wedge of the Type 1 sequence of the Kuusamo belt. However, as this occurrence lies close to the tectonic boundary (the Kitka
Fig. 7.17. (facing page) Photographs of the problematic Karelian rock types. (A) Deformed basal conglomerate of the Vihajärvi Group with quartzite and schist clasts. Brownish feldspar quartzite clasts are derived from the Akanvaara Formation. Mustavaara, Kainuu belt. Slide 4-4. (B) Parallel-laminated sandstone (below the 16-cm-long scale) overlain by hummocky cross-stratified sandstone, the Jalkaaho Formation. Jalka-aho, Kainuu belt. Slide 4-208. (C) Thinning upwards tempestite sequences, the Jalka-aho Formation. Jalka-aho, Kainuu belt. Slide 4-210. (D) Synsedimentation deformation structures in parallel-laminated sandstone, the Jalka-aho Formation. Jalka-aho, Kainuu belt. Slide 4-215. (E) Deformed angular unconformity between the siltstone of the vertical Pärekangas Formation (on the left = east) and Haapalanmäki quartzite-cobble conglomerate. Haapalanmäki road cut, Kainuu belt. Photo 10812. (F) Matrix-supported quartzite-boulder conglomerate. Haapalanmäki road cut, Kainuu belt. Photo 10807. (G) Pebbly and hematite-laminated sandstone of the Himmerkinlahti Member. Note the angular pink albitite clasts right of the scale. Himmerkinlahti, western Kuusamo belt. Slide 30-559. (H) Polymictic epidotized and gneissic conglomerate with diverse quartzite, metabasite, and granitoid clasts. Päästäispuro, western Kuusamo belt. Slide 30-24. Photos: Kauko Laajoki. CHAPTER
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fault) between the eastern and western parts of the Kuusamo belt, its stratigraphic position needs to be verified by isotope dating. Another exceptional conglomeratic unit is found ~30 km northeast of Himmerkinlahti and forms a tight syncline within the highly deformed Kainuan rocks of the western Kuusamo belt (Figure 7.5). This Kolmiloukkonen Formation seems to lie unconformably upon the Kirintökangas Formation. Its base is an epidotized polymictic conglomerate with diverse quartzite, metabasite, and granitoid clasts (Figure 7.17H). The rock is highly metamorphosed and its matrix is gneissic. The bulk of the formation is made up of micaceous quartzite and quartzitic mica schist. It is possible that these units may be younger than the bulk of the Kuusamo belt and could be correlated with the molassic Kumpu Group in Central Lapland (see Chapter 4).
15. Karelian metadiabases The late Archean basement and its Sariola– Kainuu–Jatuli cover are cut by several generations of metadiabases, which offer a valuable tool for relative dating of the deposition of the Karelian supracrustal rocks. Vuollo (1994) (see also Chapter 5) classified these dikes in eastern Finland into four groups, which are, from the oldest to the youngest, as follows: (1) ~2.45-Ga mostly boninitic (noritic) dikes associated with layered intrusions of the same age group. (2) 2.2-Ga low-Al tholeiites (karjalites), intruded into Kainuu and the lower part of Jatuli, typically found as differentiated (layered) sills. (3) ~2.1-Ga Fe-tholeiites; these are terminated by the Koljola lava (Pekkarinen and Lukkarinen, 1991), and seem to a give minimum age for the overlying CIS III dolomites. (4) 1.98-Ga Fe-tholeiitic–tholeiitic dike swarm dated only in North Karelia where they 324
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cut the uppermost Jatuli quartzites. They have not been found in Kaleva, indicating that the deposition of the latter took place after 1.98 Ga. The age pattern seems to be a little different in the Peräpohja and Kuusamo areas, where only metadiabases of the 2.2-Ga age group have so far been dated with confidence (Hanski et al., 2001). Felsic dikes are common within the Peräpohja belt, but due to their Archean xenocrystic zircons, their emplacement ages are unknown.
16. Previously proposed basin models Since Hietanen (1975) published her tectonic model for the Svecofennides, depositional settings of the Karelian basins have been modeled in the terms of plate tectonics. The various ideas are summarized in Table 7.2.
16.1. Continental and pericontinental Karelia (sensu stricto) basins As can be seen from Table 7.2, most authors classify the Sumi, Kainuu, and Jatuli tectofacies, which may be called collectively Karelian basin deposits sensu stricto, as continental, cratonic or epicontinental sequences. However, closer basin analyses and reconstructions are hard to carry out as only poorly exposed and arbitrary cross-sections of deformed basins are visible and paleocurrents are difficult to measure. The available data (Ojakangas, 1965; Marmo et al., 1988; Strand, 1993; Kohonen and Marmo, 1993) indicate that most of the paleocurrents trend towards the northwest. Data from Russian Karelia (Sokolov and Heiskanen, 1985; Ojakangas et al., 2001) support this general trend, which indicates that the major source of sediments was from the Kuhmo complex.
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16.2. Kaleva basins Flysch-type Kaleva sequences are closely connected to the development of the Svecofennian island arc(s) in the west and both marginal basin/foredeep and deep oceanic basin models have been proposed (Table 7.2). Gaál and Gorbatschev (1987) suggested that the sequences should be included in the Svecofennides. The Kaleva depositional models depend very much on how the origin of Jormua and Outokumpu ophiolites is interpreted. Kontinen (1987) originally favored a passive margin or intracontinental rift environment for the Jormua ophiolite. Park (1988) argued that the Jormua and other Karelian ophiolites represent a back-arc basin. Lahtinen (1994) discussed models in which the Outokumpu ophiolite had been generated in a supra subduction zone environment either near a distant pre-existing island arc or intraoceanic island arc. There is little evidence for rift sedimentation and magmatism around 1950 Ma within the Kainuu belt. Because the Jormua ophiolite has been emplaced together with <1920 Ma monotonous deep-water turbidites, Peltonen et al. (1996, Figure 15) considered the Jormua ophiolite an allochthonous unit formed during continental breakup somewhere along the western margin of the Iisalmi–Pudasjärvi complex, and upon which the Upper Kaleva slope-rise turbidites with unknown distant (non-Archean) provenance were deposited immediately after post-rift thermal subsidence. Nironen (1997, Figure 3) discussed two models: (1) the pre-Svecofennian sea had opened around 2100 to 2000 Ma and at 1950 Ma, and a marginal basin developed at the attenuated continental margin; and (2) the continental breakup occurred at 1950 Ma, supporting the latter hypothesis. For more details, see also the article by Peltonen (Chapter 6) in this volume. The fact that the Karelian ophiolites and Upper Kaleva were transported from a distant unknown “root area” means that the latter’s CHAPTER
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relation to the Lower Kaleva is open. Unfortunately, the sedimentation age of the Lower Kaleva in its type areas in North Karelia or Kainuu is not known. It is usually bracketed between 1970 Ma and 1920 Ma on the basis that it is not cut by the 1970 Ma metadiabases that cut the underlying Jatuli rocks (Vuollo, 1994), and by the maximum sedimentation age of the Upper Kaleva turbidites based on detrital zircons (Claesson et al., 1993). However, the fact that the sediment material in Lower Kaleva seems to be older than in Upper Kaleva (Huhma, 1990) cannot be used as evidence for the relative sedimentation ages of these two units - it only indicates provenance differences.
17. Paleogeographic reconstructions The evolution of the Karelian belts is not yet thoroughly understood and rather different opinions on the matter exist. This is because the Karelian formations are widely distributed in Finland, northeastern Russia, Norway, and Sweden where they have been studied separately and utilizing different approaches. The sequence stratigraphic approach has only been applied in rare cases and dating of sedimentation of different units has been rather arbitrary and sporadic. Most of the significant fault and thrust zones are unexposed and, consequently, their bearing on stratigraphy is in most cases only arbitrary. This holds especially for the Kajaani tectonic zone running from North Karelia via Kainuu to Kuusamo (Figure 7.1), the allochthonous Upper Kaleva units, and the northern part of the Peräpohja belt. However, the continental Karelian sequences, which are mostly autochthonous–parautochthonous, are so well known that rather reliable basin models can be established. In contrast, the Kaleva sequences are, for the most part, allochthonous and the tectonic significance of their main tectonic boundaries is unresolved. S U P R AC RU S TA L RO C K S
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17.1. Continental–marginal Karelian sequences The brief paleogeographic model outlined for the continental–marginal Sariola–Jatuli sequences in Figure 7.18. was reconstructed on the basis of the following facts and hypotheses: 1. The Karkuvaara Formation (Figure 7.5) represents a transposed part of the greenstone that possibly once connected the correlative greenstones of the Kuusamo and Peräpohja belts (Greenstone I and Runkaus Formations). 2. Structural observations in the western Kuusamo belt indicate that its tectonic transportation along the Hirvaskoski shear zone was to the northeast. 3. The Karkuvaara Formation and the southeastern margin of the Peräpohja basin are located at the southwestern extension of the present southwestern corner of the Kuusamo belt. 4. The estimated crustal shortening across the Kajaani tectonic zone is at least 50 km (~1.5 times the length of the Karkuvaara Formation). 5. The northeastern boundary of the Pudasjärvi–Iisalmi complex was approximately parallel to the present-day southwestern boundary of the Kuhmo complex before this crustal shortening. 6. Sumi–Sariola rift valleys controlled the deposition of fluvial–marine deposition of the Kainuu tectofacies. 7. Distribution of Jatuli stages I and II is based on Karhu’s (1993) CIS III and CIS IV observations. The preliminary sequence stratigraphic reconstruction for the northern part of the Kainuu belt and the Kuusamo belt in Figure 7.19 is based on the correlation of the Central Puolanka Group with the Ahola, Nilovaara, and Kirintökangas Formations in Posio and correlation the northeastern part of the Kainuu 326
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belt with the southern part of the Kuusamo belt (Strand and Laajoki, 2000). Figure 7.19 also shows correlation within the upper part of the Jatuli (“Marine Jatuli”). This is based on both lithostratigraphy and carbon isotope data (Karhu, 1993).
17.2. Kaleva sequences As the tectonic history of the Lower and Upper Kaleva tectofacies is known only fragmentarily, no paleogeographic reconstruction is presented for them. Most sophisticated models of the Kaleva development in North Karelia and Savo are found in Ward (1987, 1988) and Kohonen (1995).
18. Synopsis The following summary is based on the author’s own concepts built on studies carried out together with several students in Kainuu and Kuusamo, and the extensive literature available (see also Table 7.1).
18.1. Karelia (sensu stricto) basin development (1) Sumi: the first Karelian basin phase. The Archean crust started to extend and rift ~2500 Ma ago (Melezhik and Sturt, 1994). At this stage, mainly NW-trending relatively narrow rift basins developed into which the Sumi sediments were deposited and bimodal lavas erupted at ~2450 Ma. This phase is very well represented on the Russian side in the Paanajärvi area (Systra, 1996) and also in the Finnish Lapland (Hanski et al., 2001). In the area reviewed it is represented mainly by subvolcanic layered mafic intrusions and minor relics of hanging wall acid volcanic rocks above the Kuusijärvi intrusion (Lauri et al., 2003). The rift basins of Suhanko and Kuusijärvi at least were formed and filled by Sumi sediments and lavas (Figure 7.18A).
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Table 7.2. Väyrynen’s concepts and various plate tectonic related tectonosedimentary models proposed for the Karelian tectofacies in eastern Finland.The nomenclature used in this chapter is in bold. Traditional Karelian formations Author
Jatuli sensu stricto Jatuli (+ Marine Kainuu Jatuli (Lower Jatuli, (Upper Jatuli Ojakangas et al., + Ludikovian, 2001) Ojakangas et al., 2001) “Jatuli continent”
Sariola (¨pre-Jatuli¨, Pekkarinen, 1979) Väyrynen (1933, 1954)
“Marine” Chemical weathering Cratonic shallow-water, partly marine
“Dry”
Pekkarinen (1979) Strand (1988)
Cratonic grabenhalf-graben Intracratonic rift
Luukkonen (1989)
Continental rifting and faulting
Laajoki (1988a)
Incipient rifting
Strand and Laajoki (1999) Gehör and Havola (1988) Kontinen (1986)
Rift
Upper Kaleva (Western Kaleva) (allochthonous & ?autochthonous)
Flysch-type
Open sea Shelf
Continental platform
Basinal
Stable cratonic setting
Half-graben rift basins Rift-phase related
Cratonic to epicratonic Anorogenigic craton cover (preorogenic) (Karelian sensu stricto) Cratonic versus marginal sequencies Continental-epicontinental Shelf Shelf
Ward (1987, 1988) Sorjonen-Ward (1997)
Lower Kaleva (Eastern Kaleva) (autochthonous)
Divergent continental margin
Narrow sea–inland basins Half-graben
Kontinen and Peltonen (1988) Gaál and Gorbatchev (1987) Karhu (1993) Ekdahl (1993) Park et al. (1984) Park (1986, 1991)
Kaleva
Jatuli sensu Väyrynen (1933, 1954) or Karelian sensu Gaál and Gorbatschev (1987)
Stable platform
Kohonen (1996) Korkiakoski and Laajoki (1988) Laajoki (1988b)
Deep marine basin Deep allochthonous marine basin
Developed on rifted passive continental margin. Orogenic (Svecofennian) Back arc Flysch from the arc in the SW Back arc basin (shallow water Outokumpu association) and Kaleva flysch Marginal sea (alIntracratonic rift lochthonous) Rifting and subsidence leading to passive margin formation From syn-rift to post-rift passive margin/foredeep Foredeep
Transition from divergent to convergent tectonics
Rift-phase
Basinal
(Foredeep)
Rift-bound turbidite basin on continental margin Back arc basin
Gaál (1990) Gaál (1982) Peltonen et al. (1996)
Slope-rise of passive margin W of Iisalmi block Marginal basin or intracontinental rift Passive margin sedimentation
Honkamo (1987) Korsman et al. (1999) CHAPTER
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KARELIAN
S U P R AC RU S TA L RO C K S
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327
28° 00´
28° 00´
PKTS Kuusamo
Puolankajärvi fan ??? KR
lex
lmi
nita
?
Kuopio
100 km
A–Sumi–Sariola
B–Kainuu 28° 00´
28° 00´
PKTS
Puolankajärvi fan
Kuusamo
KS
JS
lmi plex
com
nita
•
lex
Iisa
rvi-
nita
lex
cog
mp
a in
co mi
r Ter
isal
cog 7
RS
Kuopio
D–Jatuli stage 2 (=CIS IV)
K A R E L I A N S U P R AC RU S TA L
R U S S I A
omp
asjä
vi-I a in
CHAPTER
Kainuu fluvial– marine systems 100 km
C–Jatuli stage 1 (=CIS III) •
RS
AS
Kemi
Legend
Kuopio
100 km
328
65° 45´
lex
sjär r Ter Kainuu fluvial– marine systems
JS
KS
SuS
omp
a Pud
Legend
CIS IV
Pud
mo c
AS
R U S S I A
Kuusamo
KuS
CIS III
c amo Kuus
a Kuus
SuS Kemi
PKTS
Puolankajärvi fan
KuS
CIS III
65° 45´
RS
nita
cog
plex
Kainuu fluvial– marine systems
?
JS
com
a in
cog
plex
a in
com
r Ter
lmi
Iisa r Ter
?
RR Kuopio
omp
Iisa
rvi-
rvi-
Legend
100 km
AS
asjä
Kemi
mo c
65° 45´
lex
asjä
Kemi
SuS
omp
Pud
mo c
SuR
KS
Pud
65° 45´
a Kuus
KuR
R U S S I A
a Kuus
R U S S SR I A
Kuusamo
KuS
Puolankajärvi fan
ROCKS
CIS IV
Most likely, the structural development of the Kurkikylä, Saari–Kiekki, and Reittiö rifts also started at this stage, but they were filled later during the Sariola stage. How the Kuhmo and Iisalmi–Pudasjärvi basement complexes were related at this stage is not known. (2) The phase of pre-Sariola physical weathering and erosion. The Sumi sedimentation was followed by a significant period of physical weathering in arid–semiarid conditions at ~2350 Ma (Sturt et al., 1994). The deeply penetrating erosion, caused by rift inversion, removed almost completely all the Sumi supracrustal rocks and upper parts of the faulted layered mafic intrusions associated with them. (3) Sariola: the second Karelian basin phase. Strand and Laajoki (1993, Figure 2) suggested, that a fault-bounded rift basin was developed between the Kuhmo and Pudasjärvi–Iisalmi complexes before the Sariola glaciation; in this paper the basin is considered to be a narrow seaway or depression. The Sariola fluvial polymictic conglomerates and arkoses were deposited on top of the ~2350 Ma weathering crust and reactivated Sumi rift basins. Basic–intermediate volcanic activity was significant in the northeastern part of the Kainuu belt (e.g., the Kurkikylä Group), in the Saari–Kiekki rift, and in Posio–Kuusamo and Peräpohja. The location of sedimentation and volcanism was controlled by rift structures that started to develop at the Sumi stage (Figure 7.18A). Locally, the basins subsided under sea level or were transgressed by a sea into which
glaciomarine sediments were deposited. This stage may represent the global Huronian glaciation. It is possible that during the lowstand caused by the glaciation, immature detritus was carried by gravity flows into deeper water along the valleys incised into the Sumi–Sariola rift basins. The turbiditic lower part of the Puolankajärvi Formation could represent these sediments, but this is rather speculative, as no bounding surface of this hypothetical lowstand fan has been found (Figure 7.18A). (4) Period of chemical weathering. After the Sariola phase, Karelian areas lay in a subtropical–tropical climate, probably close to the equator. The previously formed rocks were subjected to strong chemical weathering (Marmo, 1992). The intensity of the weathering was accentuated by the relatively low oxygen and correspondingly high carbon dioxide contents of the atmosphere (Karhu, 1993). The weathering produced kaolin-bearing deposits, later metamorphosed into kyanite-bearing quartzites (e.g., Hokkalampi) during the Svecofennian orogeny. (5) Kainuu: the third Karelian basin phase. The exposure of the Sumi rocks to erosion indicates fall of relative sea level. After the period of chemical weathering and coeval lowstand, an ocean started to spread from the northwest along the Sumi–Sariola rift basins, which also acted as the deposition areas of fluvial sediments derived from the Archean continent. As a result, localized transgressive fluvial, delta and paralic sediments were formed at the margins of the Kuhmo block and
Fig. 7.18. (facing page) Schematic paleogeographic reconstruction of the Sariola–Jatuli sequences. Volcanism is excluded. Fixed point−the Kuhmo complex (see Figure 7.1). (A) Sumi−Sariola tectofacies: Rift basins (KR−Kuusijärvi; KuR−Kurkikylä; RR−Reittiö; SR−Saari–Kiekki; SuR−Suhanko). The existence of the Puolankajärvi fan is questionable at this stage (see the text). (B) Kainuu tectofacies: Kainuu fluvial–marine systems (KuS−Kuusijärvi; KS−Korvuanjoki; AS−Akanvaara; JS−Jero; RS−Reittiö; SuS−Suhanko), and the Pärekangas–Kirintökangas tempestitic system (PKTS) are supposed to form one Type 1 sequence. (C) First stage of the Jatuli tectofacies, only the approximate known distribution area of the CIS III carbonate platform (green color) is shown. (D) Second stage of the Jatuli tectofacies, only the approximate known distribution area of the CIS IV carbonate platform (yellow color) is shown. CIS III and CIS IV refer to Karhu’s (1993) carbon isotope stages III and IV, respectively. CHAPTER
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KARELIAN
S U P R AC RU S TA L RO C K S
•
329
Peräpohja RM TM KM
CIS IV
Kuusamo
Posio
CIS III
SF TU
East Puolanka
TST EPG
West Puolanka CPG PK AV PJ
HST LST-TST LST-HST
KK NV AH
LD AF DF
4
HST 3 li TST Jatu HST 2 TST HST 1 u inu TST (LST) Ka iola KG Sar KuG
Archean basement
RF GIII EV SQ G1
Sumi
Fig. 7.19. Schematic sequence stratigraphic diagram of the Karelian sequences in northern part of the Kainuu belt and their approximate correlation with some of the units in Kuusamo, Posio, and Peräpohja (not to scale). The sequence subdivision and relative water depth curve of the East Puolanka Group (EPG) are simplified from Figure 7.14. The relative water-depth curve (blue) of the Central Puolanka Group (CPG) is tentative only. Abbreviations of the groups: KG−Korvuanjoki; KuG−Kurkikylä. Abbreviations for the formations: AF−Amphibole schist; AV−Akanvaara; DF−Dolomite Formation; G I and G III−Greenstone I and III, respectively; KK−Kirintökangas; KM−Kvartsimaa; LD−Limestone–dolomite; NV−Nilovaara; PJ−Puolankajärvi; PK−Pärekangas; RF−Rukatunturi; RM−Rantamaa; SF−Salmijärvi; SQ−Sericite quartzite; TM−Tikanmaa. TU−Tuffite units in Puolanka. Other: HST, LST, and TST−highstand, lowstand, and transgressive system tracs, respectively; CIS III and CIS IV−carbon isotope stages III and IV, respectively.
also partly on the Pudasjärvi–Iisalmi block. Marine sands, clays, and carbonates were deposited farther away from the continent, perhaps somewhere within and north of the area of the present Gulf of Bothnia. During the highstand, the tempestitic heteroliths of the Kirintökangas and Pärekangas Formations and probably also the tidalites of the Erivaaransuo Formation were deposited. Some volcanic activity occurred towards the end of the highstand phase. As an entity, the Kainuu series seems to represent a Type 1 sequence. (6) Period of pre-Jatuli erosion and faulting. During this period, at least some block movements took place and exposed deeper 330
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•
lithified parts of the Kainuu (–Lapponi) and older tectofacies (possibly even the Archean basement) for erosion. (7) Jatuli stages I and II: basin phase. The Jatuli marks transition from rift topography controlled fluvial-dominated system to open sea conditions. Mainly fluvial sediments were deposited in the beginning, but later the sea advanced onto the continent and a continental shelf was formed, with characteristic changes in relative sea level, deposition of carbonates and minor sands, and volcanic activity. This series may be considered a continental prism, which formed on the top of the Kainuu basins. It is probable that, at this stage, the sea was
K A R E L I A N S U P R AC RU S TA L
ROCKS
large and opened towards the northwest, but shallow bays extended onto the continent. The transgressions during this phase were probably directed from the northwest towards the southeast. The rifting of the continental crust also continued during the Jatulian. This is demonstrated by the volcanic rocks within the Jatulian and especially by the abundant volcanic activity towards the end of the period, the products of which are best observed in the Kemi–Kalix area. The Jatuli deposits in East Puolanka can be subdivided into four Type 1 sequences (Figures 7.14 and 7.19). Their correlation with other areas, the western margin of the Kainuu belt included, is not possible at this stage. With the aid of carbon isotope studies and carbonate lithostratigraphy, the Jatuli can be, however, subdivided into two major stages. The first stage comprises the bulk of siliciclastic sedimentation ending with the formation of the CIS III carbonate platform (Figure 7.18C). The second stage started with widespread volcanism and associated epiclastic sedimentation and was followed by the CIS IV carbonate platform; the topmost part of the Karelian formations (sensu stricto) preserved in the area reviewed (Figure 7.18D).
18.2. Lower Kaleva development After Jatuli, in the course of the development of the Svecofennian sea and island arc systems and 1.97 to 1.95 Ga juvenile oceanic crust southwest of the Karelian continent, the Karelian continent was split into the North Karelia–Kainuu and Kuopio–Kiiminki basins (cf. Ward, 1987; Kohonen,1995). Their configurations were not controlled by the previous Sumi–Sariola rifts and the basins developed independently. The North Karelia basin may represent a rather open basin on rifted continental margin, whereas the Kainuu basin was probably a narrow restricted intracratonic basin characterized by deposition of chemiCHAPTER
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KARELIAN
cal (BIF) and euxinic carbon-rich sediments. The lower part of the Kuopio–Kiiminki basin was deposited in a marginal basin or in a riseslope environment of the newly formed Kaleva continental margin. The upper part including volcanic rocks may represent an allochthonous basin from farther southwest.
18.3. Upper Kaleva development As the Upper Kaleva is likely allochthonous, its relation with Lower Kaleva and the Svecofennides is unknown. It may represent a major submarine fan system prograding over the subsiding passive margin, just prior to collision, with sediments being derived from the craton or more distant sources (Kontinen and Sorjonen-Ward, 1991). The Kaleva and the underlying rocks were folded by the Svecofennian orogeny. It is apparent that sedimentation occurred also during this orogeny, but their potential to survive the subsequent erosion was small, as they were located in the uppermost parts of the sedimentary sequences. It is also difficult to separate basinal rift-stage turbiditic sediments from orogenic “flysch” because of strong metamorphism.
18.4. Closing comments This chapter treats only a small part of the Paleoproterozoic supracrustal rocks deposited on or marginal to the late Archean basement blocks of the Fennoscandian Shield and the depositional models should be considered only tentative. As to the evolution of the coeval and in many respects similar rocks in Finnish Lapland, the reader is referred to Chapter 4.
Acknowledgments I thank the editors, Richard Ojakangas, and Jarmo Kohonen for useful comments on the contents and linguistic formulation of the S U P R AC RU S TA L RO C K S
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manuscript and Kristiina Karjalainen for the line drawings. Seppo Gehör, Paavo Härmä, Mikko Honkamo, and Kari Strand are thanked for slides. Special thanks go to the former students of the University of Oulu with whom the author has worked and whose studies are referred to in this paper. They are: Paul Evins, Jarmo Finnilä, Seppo Gehör, Pekka Härmä, Tuomo Karinen, Aulis Kärki, Esko Korkiakoski, Jouni Luukas, Kari Strand, and Pekka Tuisku.
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(in Finnish) Karinen, T., Salmirinne, H., 2001. Koillismaan kerrosintruusiokompleksin läntisen osan geologinen evoluutiomalli. Geol. Surv. Finland, Internal report M 19/3543/2001/2, 1–23. (in Finnish) Kärki, A., 1988. Stratigraphy of the Kainuu Schist Belt and palaeosedimentology of its Kalevian metasediments at Melalahti, northern Finland. In: K. Laajoki, J. Paakkola (Eds.), Sedimentology of the Precambrian formations in eastern and northern Finland. Proceedings of ICGP 160 symposium at Oulu, January 21–22, 1986. Geol. Surv. Finland, Spec. Pap. 5, 149–164. Kärki, A., Laajoki, K., 1996. An interlinked system of folds and ductile shear zones - late stage Svecokarelian deformation in the central Fennoscandian Shield, Finland. J. Struct. Geol. 17, 1233–1248. Kärki, A., Laajoki, K., Luukas, J., 1993. Major Palaeoproterozoic shear zones of the central Fennoscandian Shield. In: R. Gorbatschev (Ed.), The Baltic Shield. Special Issue. Precambrian Res. 64, 207–223. Kohonen, J., 1996. From continental rifting to collisional crustal shortening - Paleoproterozoic Kaleva metasediments of the Höytiäinen area in North Karelia, Finland. Geol. Surv. Finland, Bull. 380, 1–79. Kohonen, J., Marmo, J., 1992. Proterozoic lithostratigraphy and sedimentation of Sariola and Jatuli-type rocks in the Nunnanlahti–Koli–Kaltimo area, eastern Finland; implications for regional basin evolution models. Geol. Surv. Finland, Bull. 364, 1–67. Kohonen, J., Luukkonen, E., Sorjonen-Ward, P., 1991. Nunnanlahti and Holinmäki shear zones in North Karelia: evidence for major early Proterozoic ductile deformation of Archean basement and further discussion of regional kinematic evolution. In: S. Autio (Ed.), Geol. Surv. Finland, Current Research 1989–1990. Geol. Surv. Finland, Spec. Pap. 12, 11–16. Koistinen, T.J., 1981. Structural evolution of an
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Chapter 8
SVECOFENNIAN SUPRACRUSTAL ROCKS
Y. Kähkönen
Cover page: Mafic volcanogenic graywacke. Pulesjärvi–Kolunkylä complex, eastern shore of Lake Näsijärvi. The vertical dimension corresponds to ~0.5 m in nature. Photo:Yrjö Kähkönen.
Kähkönen, Y., 2005. Svecofennian supracrustal 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. 343–406. © 2005 Elsevier B.V. All rights reserved.
A substantial part of the Paleoproterozoic crust of southern and central Finland is characterized by ~1.92 Ga turbiditic sedimentary and ~1.90–1.88 Ga arc-type volcanic rocks. Fluvial to shallow-water sedimentary rocks, felsic schists, sedimentary carbonates, black shales, and MORB- and WPL-like basalts are locally also common. The rocks were metamorphosed predominantly at high-T, low-P amphibolite facies and they record two major orogenic periods: ~1.89–1.86 Ga and ~1.83–1.81 Ga; a poorly constrained ~1.91–1.90 Ga event is also inferred. Three major terranes are identified: Savo belt, central Svecofennia, and southern Svecofennia. The Savo belt, adjacent to the Archean craton, is characterized by ~1.93 –1.92 Ga bimodal volcanic rocks (εNd ~ +3) emplaced in a relatively immature arc. In central Svecofennia, arc-type ~1.905 to 1.89 Ga volcanic and related fluvial to turbiditic sedimentary rocks of the Tampere belt (εNd ~0) belong to an evolved arc system presumably associated with a ~2.1–2.0 Ga microcontinent. The Pirkanmaa belt to the south represents the associated subduction zone complex. The same arc system is observed in Pohjanmaa, western Finland. The subduction zone complex and stratigraphically low parts of the Tampere belt are dominated by turbidites, but there are also black shales, NMORB- to WPL-like pillow basalts (εNd ~ 0 to +3), and minor cherts. The turbidites include an Archean component as well abundant 2.0–1.92 Ga detritus presumably derived from the Savo belt and the inferred microcontinent. In southern Svecofennia, metamorphism and migmatization at 1.83–1.81 Ga largely obliterated primary features but the sedimentary rocks tend to be more pelitic than in central Svecofennia. The Häme belt is characterized by ~1.89–1.88 Ga arc basalts to rhyolites and overlying mafic lavas with rift affinity. The Häme belt represents a less mature setting than the bulk of the Tampere and Uusimaa belts. The Saimaa area is dominated by sedimentary rocks with diverse sources and deposition ages. The supracrustal rocks of the Uusimaa belt include mudrocks, graywackes, as well as EMORB-, WPL-like, and arc-type (1.90–1.88 Ga) volcanic rocks; felsic volcanic and sedimentary rocks and sedimentary carbonates are abundant in the western part of the belt. The volcanic and related plutonic rocks have εNd values of –1 to +3. Most of the supracrustal rocks of the Uusimaa belt (and related Bergslagen field in Sweden) were presumably deposited close to or within a ~2.1–2.0 Ga microcontinent. In southern Svecofennia, ~1.86 Ga detrital zircons in sporadic quartz arenites indicate deposition after or during an orogenic peak at 1.87–1.86 Ga.
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1. Introduction The Precambrian of Finland saw the advent of plate tectonics when Hietanen (1975) suggested that the Svecofennian fold belt (Figure 8.1) formed in an island arc environment. This was followed by an idea of two diachronous subduction zones, ~1.92 Ga and ~1.88 Ga (Gaál and Gorbatschev, 1987). At about the same time, Park (1985) explained the bulk of the Svecofennian domain as a collage of exotic volcanic arcs, with arcs in the south being 100–150 Ma younger than those in the north. Windley (1992) considered the domain a type example of accretionary orogens characterized by growth and amalgamation of juvenile island arcs, slices of oceanic crust, oceanic plateaus, and microcontinental blocks, and their accretion to a continent. This idea is still plausible, but the age zonation of Svecofennian arcs (Park, 1985) is not supported by the age distribution of the metavolcanic rocks (Table 8.1; Billström and Weihed, 1996; Lundström et al., 1998). Anyhow, the Svecofennian domain registers significant growth of continental crust ~2.0–1.8 Ga ago (Patchett and Kouvo, 1986; Huhma, 1986; Lahtinen and Huhma, 1997; Rämö et al., 2001). The growth does not, however, seem to have been as rapid as thought two decades ago, because increasing isotope evidence indicates that the Svecofennian domain may include extensive yet concealed ~2.1–2.0 Ga components (see
also Chapter 11). The ideas of crustal growth build essentially on isotope studies, and the concept of presumed subduction-related settings has been derived from studies of supracrustal rocks. This article presents a review of Svecofennian supracrustal rocks in Finland and discusses their role in crustal evolution. Major emphasis is placed on the primary character, tectonic setting, and geochemical (including isotope) composition of the metavolcanic rocks. Several important issues, however, arise from the metasedimentary rocks as well.
2. Geologic setting 2.1. General aspects The Svecofennian domain covers an 800 km by 800 km area of Finland and Sweden. In the east and northeast, the domain is bounded by the Archean craton and its supracrustal cover (Karelian deposits), in the west by the Caledonides, and in the southwest by the 1.75–1.50 Ga Southwest Scandinavian domain. The southern boundary of the Svecofennian domain has conventionally been positioned under the Baltic Sea, but Gorbatschev and Bogdanova (1993) included the Precambrian bedrock covered by Paleozoic sedimentary rocks in the Baltic States, Poland, and western Russia in the Svecofennian realm.
Fig. 8.1. (facing page) Generalized geological map of southern and central Finland based on Koistinen (1994), Lundqvist et al. (1996), Korsman et al. (1997), Kähkönen (1999), and data therein. The continuous and hatched lines show the boundaries of the supracrustal belts, areas and fields mainly according to Nironen et al. (2002). The continuous line delineating the eastern margin of the Savo belt is the northeastern boundary of the Svecofennian domain. In the Pohjanmaa belt, the hatched line separates the Evijärvi and Ylivieska fields. The southern hatched line separates the Tampere and Pirkanmaa belts. The continuous line between the Pirkanmaa belt and the Häme belt marks the suture between central Svecofennia and southern Svecofennia. The inferred ~2.0 Ga Keitele microcontinent roughly equates to the Central Finland granitoid complex by area, and the Uusimaa belt represents part of the presumed ~2.1–2.0 Ga Bergslagen microcontinent (Chapter 11). Abbreviated localities: H-linna– Hämeenlinna, Ha–Haukivuori, Ik–Ikaalinen, Ka–Kankaanpää, Ki–Kiikoinen, La–Lavia, Su–Suodenniemi, Va–Vammala.
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Oulu
of Bo th nia
Raahe
Pohjanmaa belt
ulf
Vihanti
G
Jormua Ylivieska
Svecofennian domain
Kuusaa Pyhäsalmi
Savo belt
Evijärvi Perho
Alajärvi
Seinäjoki
Saarijärvi Haapamäki Parkano
Siipyy Ka
Pielavesi
Pihtipudas
Vimpeli
Rautalampi
Ik
Hirsilä
La
Outokumpu
Central Finland granitoid complex
Kuru
Tampere belt
Karelian domain
Luhanka
Joroinen HaVirtasalmi RantaJuva salmi Ristiina
Su
Saimaa area
Pirkanmaa belt
Ki
Va
Urjala
PunkaharjuParikkala
H-linna
61°N
Lahti
Häme belt
Forssa
30°E
Hyvinkää
Turku Nauvo–Korppoo Kemiö
Järvenpää
100 km
Kisko
Uusimaa belt Helsinki
Pellinki
Archean rocks
Gabbros and diorites
Karelian and Svecofennian quartz arenites Svecofennian and Karelian mica schists, mica gneisses, and migmatites Svecofennian and Karelian mafic to intermediate metavolcanic rocks
1890–1870 Ma granitoids
Svecofennian felsic to intermediate schists
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1840–1820 Ma granitoids 1650–1540 Ma rapakivi granites 1400–1200 Ma Jotnian sedimentary rocks and diabases
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Table 8.1. Ages of Svecofennian volcanic and related rocks in Finland. Belt/Locality Savo belt Pyhäsalmi region Kettuperä Riitavuori Pyhäsalmi Pielavesi region Kotajärvi Vihanti Virtasalmi region Joroinen Pohjanmaa belt Seinäjoki
Rock type
Age ± 2σ (Ma)
Reference
Comments
gneissic tonalite rhyolite plagioclase porphyry
1930 ± 15 Ma 1921 ± 2 Ma 1875 Ma
Lahtinen and Huhma (1997) Kousa et al. (1994) Helovuori (1979)
subvolcanic
diorite/lava plagioclase porphyry
1882 ± 2 Ma 1978 ± 17 Ma
Salli (1983) Vaasjoki and Sakko (1988)
rhyolite gneissic tonalite
1906 ± 4 Ma 1903 ± 10 Ma
Vaasjoki and Sakko (1988) Huhma (1986)
plagioclase porphyry
1886 ± 3 Ma
Mäkitie and Lahti (1991)
minimum age from a total zircon fraction
hypabyssal
Belts within the Central Finland granitoid complex Pihtipudas granites, porphyries, and a dacite 1883 ± 20 Ma Saarijärvi quartz-feldspar porphyry 1887 ± 2 Ma Parkano rhyolite 1907 ± 13 Ma rhyolitic dyke 1872 ± 12 Ma Kuru felsic tuff 1897 ± 2 Ma
Aho (1979) pooled age Nironen (2003) hypabyssal Vaasjoki and Lahti (1991) Vaasjoki and Lahti (1991) Tiainen and Kähkönen (1994)
Tampere belt Koskuenjärvi Sammatinjärvi Pirttiniemi Pukala Tervakivi Tesoma Kalkku Takamaa Lempiänniemi Varissaari Häme belt Valijärvi Aulanko Hyvinkää
rhyolite dacite felsic tuff subvolcanic porphyry high-K rhyolite felsic tuff feldspar porphyry dacite plagioclase porphyry gabbro
1904 ± 4 Ma 1898 ± 4 Ma 1898 ± 8 Ma 1896 ± 4 Ma 1892 ± 3 Ma 1892 ± 2 Ma 1889 ± 19 Ma 1889 ± 5 Ma 1880 ± 7 Ma 1885 ± 5 Ma
Kähkönen et al. (1989) Kähkönen et al. (1989) Kähkönen et al. (2004) Talikka (2003) Kähkönen et al. (2004) Kähkönen et al. (2004) Kähkönen et al. (1989) Kähkönen et al. (1989) Kähkönen et al. (1989) Patchett and Kouvo (1986)
lava-like andesite granodiorite plagioclase porphyry gabbro
1888 ± 11 Ma 1886 ± 14 Ma 1880 ± 3 Ma 1880 ± 5 Ma
Vaasjoki (1994) Patchett and Kouvo (1986) Suominen (1988) Patchett and Kouvo (1986)
1888 ± 11 Ma 1891 ± 4 Ma
Reinikainen (2001) Käpyaho (2001)
1895 ± 2 Ma 1878 ± 3 Ma 1870 ± 10 Ma 1891 ± 13 Ma 1898 ± 9 Ma 1887 ± 14 Ma
Väisänen and Mänttäri (2002) Väisänen and Mänttäri (2002) U. Mäkelä (1989) Huhma (1986) pooled age Väisänen and Mänttäri (2002) analysis by SIMS Patchett and Kouvo (1986) pooled age
Uusimaa belt Kemiö–Järvenpää field Norrlammala felsic tuffite Kuovila felsic tuff or tuffite Orijärvi area Orijärvi Fm. rhyolitic flow Kisko Fm. dacite Iilijärvi altered subvolcanic dyke Orijärvi granodiorite Orijärvi granodiorite Pellinki field andesite
The Svecofennian domain is mainly composed of granitoids but also contains a significant proportion of schists and gneisses; approximately one third of the Svecofennian bedrock is of supracrustal origin (Simonen, 348
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analysis by SIMS
1980; Korsman et al., 1997; Koistinen et al., 2001). The supracrustal rocks are typically turbiditic mica gneisses that include metamorphosed black shales and mafic metavolcanic rocks of MORB to WPL affinity. However,
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more typically Svecofennian volcanic rocks1) are ~1.91–1.88 Ga arc-type basalts to rhyolites, having fluvial to turbiditic sedimentary rocks with volcanic provenance as essential related components. In addition, quartz-feldspar schists and gneisses, mainly volcanic and volcanogenic sedimentary rocks, as well as sedimentary carbonates are abundant in certain areas. Quartz arenites are rare. The Svecofennian supracrustal rocks were metamorphosed in greenschist/amphibolite facies to granulite facies conditions with high-T amphibolite-facies rocks being most common. In general, the regional metamorphism was of high-T, low-P type and culminated at two major stages. Indications of 1.89–1.86 Ga metamorphism can be found in the entire Svecofennian domain in Finland and it produced trondhjemite migmatites at peak conditions (Korja et al., 1994; Korsman et al., 1999; Väisänen, 2002). A subsequent 1.83–1.81 Ga stage only affected the late Svecofennian granite–migmatite zone of southern Finland (Ehlers et al., 1993) and was characterized by potassic granite leucosomes. This stage also resulted in granulite-facies rocks surrounding thermal domes. It is possible that metamorphism started during a collision in the northeast at 1.91–1.90 Ga (see Nironen, 1997; Chapter 11) although this event is not particularly well constrained. Different parts of the Svecofennian domain show broadly similar structural sequences with the two earliest deformation phases characterized by overthrusts and recumbent folds with north to northeast vergence (Korsman et al., 1999). S1 schistosity has only been found as inclusion trails in porphyroblasts and S2 schistosity is the dominant tectonometamorphic feature. Early structures and schistosities were generally turned subvertical during D3 at 1.89–1.86 Ga, and subsequent deforma1)
In this article, primary volcanic and sedimentary terminology is preferred over metamorphic names, and rock types are mostly referred to without the prefix “meta”. CHAPTER
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tion mainly took place within shear zones. Nevertheless, large-scale metamorphism and deformation still characterized the 1.83–1.81 Ga late Svecofennian granite-migmatite zone of southern Finland. According to Nironen (1997), the tectonometamorphic evolution of the Finnish Svecofennian domain can be ascribed to accretional or collisional events at 1.91–1.90 Ga (in the northeast) and 1.89–1.87 Ga (effective over the entire area). Due to continued convergence, an intracratonic transpressional zone developed in the south and was followed by extensional collapse along the tranpressional zone and formation of the late Svecofennian granite-migmatite zone of southern Finland at 1.83–1.81 Ga.
2.2. Proterozoic cover deposits of the Archean craton The Archean Karelian craton underwent several phases of Proterozoic rifting, resulting in 2.45 Ga layered mafic intrusions in north-central Finland (Alapieti, 1982; Chapter 3) and 2.5–2.4 Ga and 2.2–1.97 Ga rift-related volcanic and dike rocks in Lapland and eastern Finland (Vuollo, 1994; Chapters 4 and 5). These events also led to the 2.0 Ga and 1.97–1.95 Ga ophiolites of northern and eastern Finland (Koistinen, 1981; Kontinen, 1987; Peltonen et al., 1996; Chapters 4 and 6). Proterozoic sedimentary rocks, known as Karelian formations, were deposited in continental to epicontinental and rift to marginal basin settings. These include, from the bottom, (1) Sariolian deposits such as regolithic basal conglomerates and breccias as well as glaciogenic diamictites and siltstone–argillites; (2) a paleosol; (3) Kainuan and Jatulian quartz arenites, conglomerates, tuffs, sedimentary carbonates, and black shales; and (4) Kalevian graywackes and mudrocks with minor conglomerates and banded iron-formations (Marmo and Ojakangas, 1984; Kohonen and Marmo, 1992; Kohonen, 1995; Chapter 7).
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The Kalevian deposits are divided into two groups (cf. Chapter 7). The Lower Kaleva is mainly autochthonous and consists of turbiditic conglomerates, breccias, quartz wackes, graywackes, mudrocks, and banded iron-formations, whereas the younger group, the Upper Kaleva (known also as Western Kaleva), is allochthonous and dominated by monotonous graywackes. The provenance of the Upper Kaleva graywackes was predominantly relatively immature 2.0–1.92 Ga volcanic arc(s) and Archean crust (Lahtinen, 2000). The youngest, 1.93–1.92 Ga detrital zircons in the Upper Kaleva graywackes (Claesson et al., 1993) indicate deposition more than 500 Ma after the earliest (2.45 Ga) rift events. Probably, the Upper Kaleva graywackes were deposited in a major submarine fan system largely on 1.95 Ga oceanic crust but locally on Archean crust (Ward, 1987; Chapter 7). This was evidently associated with early phases of collision of a collage of 2.10–1.92 Ga arc systems and microcontinents with the Archean continent (Lahtinen, 2000). Carbon isotope studies (Karhu, 1993; Chapter 16) indicate that the Jatulian sedimentary carbonates show an excursion to high δ13C (~10 at ~2.2–2.1 Ga). This is not shown by the older or younger Karelian and Svecofennian sedimentary carbonates and provides constraints for discussion regarding the age of deposition of some sedimentary units in the Svecofennian domain.
2.3. Division of the Svecofennian domain Division of the Svecofennian domain includes a plethora of terms. Hietanen (1975) suggested that the area consists of a single, Svionian island arc and that an inter-arc (Bothnian) basin existed between the arc and the Archean continent. Gaál and Gorbatschev (1987) divided the Svecofennian domain into three parts: the northern and southern Svecofennian provinces rich in subduction-related interme350
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diate to felsic volcanic rocks, and the central Svecofennian province (the Bothnian basin of Hietanen, 1975) that has supracrustal rocks dominated by graywackes and mudrocks. The presence of MORB-like volcanic rocks close to the boundary between the central and southern Svecofennian provinces in southern Finland led Lahtinen (1994, 1996) to suggest a suture between these provinces. Korsman et al. (1997) divided the Svecofennian domain into (1) the Primitive arc complex of central Finland (1.93–1.87 Ga), (2) the Accretionary arc complex of central and western Finland (1.90–1.87 Ga), and (3) the Accretionary arc complex of southern Finland (1.90–1.82 Ga). The boundary between the latter two approximately coincides with the boundary between the Central and Southern Svecofennian provinces, whereas the boundary between the Primitive arc complex of central Finland and the Accretionary arc complex of central and western Finland deviates from the division of Gaál and Gorbatschev (1987). In this article, I will mainly apply the division of Korsman et al. (1997) but the term Savo belt (see below) is used instead of Primitive arc complex of central Finland, whereas the other two terranes are called central Svecofennia and southern Svecofennia. Supracrustal rocks of the Svecofennian domain can be divided into belts, areas, regions, and fields, which are in many cases separated by faults and intrusions. The supracrustal rocks of southern Svecofennia are divided into the Uusimaa belt in the south, the Häme belt in the north, and the Saimaa area in the northeast (Figure 8.1). The supracrustal rocks of central Svecofennia are divided into the Pirkanmaa and Tampere belts in the south and the Pohjanmaa belt in the west and north. The Pohjanmaa belt is further divided into the Evijärvi and Ylivieska fields. In addition to these, the Central Finland granitoid complex includes numerous minor supracrustal belts. The Pirkanmaa and Tampere belts are interpreted to be coupled and to comprise rem-
SVECOFENNIAN
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nants of an arc system. Accordingly, the bulk of the Tampere belt represents a volcanic arc, whereas the gneiss- and migmatite-dominated Pirkanmaa belt mainly belongs to the associated subduction zone complex (accretionary prism). Analogously, the Ylivieska field is a part of the volcanic arc in the Pohjanmaa belt, and the Evijärvi field represents the subduction zone complex. Geophysically, the Pirkanmaa belt and Evijärvi field delineate a zone of high crustal conductivity that curves from southern Finland via western Finland to northern Sweden (Korja, 1993; Korja and Hjelt, 1993; Chapter 11). As discussed by Lahtinen and Huhma (1997) and Rämö et al. (2001), relatively low εNd values in the Central Finland granitoid complex, which forms the bulk of central Svecofennia, may indicate the presence of ~2.1–2.0 Ga protocrust. Lahtinen et al. (Chapter 11) coin this pre-1.92 Ga component the Keitele microcontinent and roughly associate it with the Central Finland granitoid complex. A relatively old (~2.1–2.0 Ga) Proterozoic block or nucleus may also be present in the southernmost part of southern Svecofennia. It extends to Bergslagen in south-central Sweden and is coined by Lahtinen et al. (Chapter 11) the Bergslagen microcontinent. These authors also suggest a ~1.95 Ga arc crust to have formed in the Savo belt and a pre-1.92 Ga crust in the Häme belt.
2.4. U-Pb zircon ages and Nd isotopes Practically all U-Pb zircon ages from the Svecofennian volcanic rocks in Finland come from arc-type sequences. The volcanic rocks of the Savo belt are typically 1.93–1.92 Ga, older than the volcanic rocks of the other Finnish Svecofennian belts; the latter range from ~1.905 Ga to 1.88 Ga (Table 8.1). The Savo belt also includes 1.90–1.88 Ga volcanic rocks (Korsman et al., 1997). Several Svecofennian plagioclase and quartz porphyries, most of CHAPTER
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these dikes, have yielded ages in the 1.88–1.87 Ga range. The Svecofennian plutonic rocks in Finland show a largely similar age distribution. The oldest of these are gneissose 1.93–1.92 Ga granitoids in the Savo belt, and this belt also includes 1.90–1.88 Ga granitoid rocks (Helovuori, 1979; Lahtinen and Huhma, 1997; Chapter 10). In central Svecofennia, 1.89–1.88 Ga synkinematic granitoids and associated mafic rocks represent the most abundant type and are only slightly older than 1.885–1.87 Ga undeformed high-K granites of within-plate affinity (Rämö et al., 2001). In southern Svecofennia, the ages of orogenic plutonic rocks range from 1.90–1.89 Ga to ~1.81 Ga (Huhma, 1986; Väisänen et al., 2002; Chapter 10). The initial εNd values of the Svecofennian plutonic rocks in Finland show some variation from area to area but rocks derived from Archean crust are absent. The plutonic rocks of the Savo belt have mostly εNd (at 1.875 Ga) values of +2 to +3 (Lahtinen and Huhma, 1997; see also Rämö et al., 2001). In central Svecofennia, the bulk of the plutonic rocks show εNd values of –1 to +1, whereas those in the northwest are more juvenile with εNd of ~ +3. Most of the plutonic rocks of southern Svecofennia have εNd values of +1 to +3; those in the far southwest have, however, somewhat lower values (Rämö et al., 2001; see also Chapters 9 and 10). Nd isotope studies on Svecofennian volcanic rocks are relatively scarce but in general they have yielded results compatible with those of the the plutonic rocks (Figure 8.2). The bulk of the volcanic rocks of the Savo belt have high εNd (at 1.93–1.92 Ga) values of ~ +3. In central Svecofennia, the εNd (at 1.90 Ga) values range from ~ –1 to +4, and a similar variation is shown by those of southern Svecofennia. In the former terrane, the highest values come from the northwest (Evijärvi), whereas in the latter the lowest and highest values are found in the southwest (near Orijärvi).
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351
+3
+1
εNd –1
–3
2.0
1.95
1.9
1.85
Age (Ga)
Savo belt + Felsic gneisses and tonalites X Basalts/amphibolites P Pyhäsalmi rhyolite K Kotajärvi diorite/lava, Pielavesi V Joroinen rhyolite Saimaa area R Rantasalmi picrites, average J Ultramafic sill, Juva
Pohjanmaa belt E Evijärvi basalts Tampere belt Sammatinjärvi dacite H Haveri basalts Uusimaa belt Pellinki lavas O Orijärvi granodiorite S Salittu picrite ▼
•
Fig. 8.2. εNd vs. age diagram for Svecofennian metavolcanic rocks in Finland. Data from Huhma (1986, 1987), Patchett and Kouvo (1986), Makkonen (1996), Lahtinen and Huhma (1997), and Vaasjoki and Huhma (1999). The Chondritic Uniform Reservoir (CHUR) and depleted mantle (DM) evolution lines after DePaolo and Wasserburg (1976) and DePaolo (1981), respectively.
Interpretation of the U-Pb age distribution of detrital zircons in Svecofennian sedimentary rocks (Figure 8.3) must be made with caution. Lahtinen et al. (2002), based on duplicate concordant analyses, estimated that true precision for reliable analytical data is ± 15 Ma. Further, a sample from the Tiirismaa quartz arenite contains two grains with interpreted ages of 1.77 Ga, one of them with a relatively low discordance of 6% and 2σ error of ±18 Ma. These ages are not geologically realistic. Therefore, when considering maximum depositional ages using individual anomalously young grains, the possibility exists that an inferred age has no clear geological 352
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explanation. In spite of these problems, the U-Pb ages of detrital zircons and Nd isotope studies show that the Svecofennian sedimentary rocks were mainly derived from ~ 2.1 Ga to 1.9 Ga sources; a prominent Archean component is present in most cases, however (Figures 8.3 and 8.4). Some of the sedimentary rocks are relatively poor in or devoid of Archean detritus (see also Table 8.2). These include, in particular, arcrelated sedimentary rocks from the Tampere and Pohjanmaa belts, which represent the upper sedimentary group of central Svecofennia in the generalized division of Lahtinen et al. (2002). In southern Svecofennia, two quartz
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youngest: 1907 ± 30 next: 1915 ± 26 1915 ± 60 1920 ± 72
4 2
A
youngest: 1926 ± 26 next: 1959 ± 26 1990 ± 40 1999 ± 34
2
A1556 Himanka
4
1925 ± 14 1933 ± 8 1944 ± 10
2
1.81 1.90 1.99
Frequency
Frequency
1899 ± 20
2.08 2.17 2.26 2.35 2.44 2.5 2.7 2.9 3.0<
B
6
1889 ± 14
2.08 2.17 2.26 2.35 2.44 2.5 2.7 2.9 3.0<
A15 Joroinen 1917 ± 10 1927 ± 6 1954 ± 6 1999 ± 28
4 2
G 1.81 1.90 1.99
A1555 Kannus
4
1895 ± 10 1994 ± 10 1995 ± 22
2
6
2.08 2.17 2.26 2.35 2.44 2.5 2.7 2.9 3.0<
A1558 Ristiina
1911 ± 16
1942 ± 10 1972 ± 20 1982 ± 30
4 2
H 2.08 2.17 2.26 2.35 2.44 2.5 2.7 2.9 3.0<
1882 ± 14
6 4
A1554 Sievi
1.81 1.90 1.99
Frequency
1.81 1.90 1.99
Frequency
1.81 1.90 1.99
Frequency
Frequency
1879 ± 10
2.08 2.17 2.26 2.35 2.44 2.5 2.7 2.9 3.0<
C
1895 ± 10 1896 ± 10 1921 ± 14
2
D
1879 ± 10
2.08 2.17 2.26 2.35 2.44 2.5 2.7 2.9 3.0<
A1557 Rantasalmi
4
1892 ± 8 1905 ± 8 1939 ± 12
2
I 8 6 4 2
1906 ± 16
2.08 2.17 2.26 2.35 2.44 2.5 2.7 2.9 3.0<
A1199 Pielavesi
1.81 1.90 1.99
Frequency
1.81 1.90 1.99
Frequency
A934 Orijärvi 4
F 1.81 1.90 1.99
E
Frequency
Frequency
A1 & A57 Tampere belt
6
1913 ± 36 1914 ± 10 1934 ± 34
J
1.81 1.90 1.99
2.08 2.17 2.26 2.35 2.44 2.5 2.7 2.9 3.0< Age (Ga)
1858 ± 8
6 4 2
2.08 2.17 2.26 2.35 2.44 2.5 2.7 2.9 3.0<
A696 Tiirismaa and A 361 Hyvinkää 1859 ± 10 1869 ± 13 1877 ± 12
1.81 1.90 1.99
2.08 2.17 2.26 2.35 2.44 2.5 2.7 2.9 3.0< Age (Ga)
Fig. 8.3. Detrital zircon ages of Svecofennian metamorphosed sandstones. The data comprise 207Pb/ 206 Pb ages from Huhma et al. (1991), Claesson et al. (1993), and Lahtinen et al. (2002). From the data of Lahtinen et al. (2002), analyses with >30% discordance were rejected, as were two unrealistically young ages in sample A696 Tiirismaa. Class interval is 0.03 Ga except 0.1 Ga for the 2.5–3.0 Ga interval. Grains older than 3.0 Ga are grouped as a single class. The vertical line with number gives the youngest interpreted concordant age (in Ma, discordance < 5%) and ±2 σ error for each sample; the interpreted ages (Lahtinen et al., 2002) use a lower intercept value of 250 Ma. The additional three to four numbers give the next to youngest concordant ages. Underlined numbers indicate two ages from one grain. (A) and (B) show graywackes representing the lower sedimentary group of central Svecofennia (Lahtinen et al., 2002). (C) is a cross-bedded arkose from the Pohjanmaa belt. Its stratigraphic position is unknown, but Lahtinen et al. (2002) interpreted it as a molasse deposit. (D) is from the Pohjanmaa belt and represents the sandstones of the upper sedimentary group of central Svecofennia (Lahtinen et al., 2002). A maximum deposition age ~1.86 Ga implied by a slightly discordant (7%) grain with an interpreted age of 1852 ± 8 Ma is not geologically realistic. (E) is a paragneiss inclusion in a ~1.925 Ga tonalite at Pielavesi, Savo belt. (F) through (J) are from southern Svecofennia. (F) is an immature graywacke from Orijärvi, Uusimaa belt. (G) and (H) are immature graywackes from the Saimaa area. (I) is a mature graywacke from the Saimaa area. (J) includes two quartz arenites. CHAPTER
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A
Savo belt Paragneiss inclusion in tonalite, Pielavesi
DM 2
CHUR
0
εNd -2 -4
2.3
2.0
Pohjanmaa belt Sandstone, upper sedimentary group of CS, Sievi Possibly a molasse arkose, Kannus Lower sedimentary group of CS Graywackes and mudrock, Himanka Graywacke and mudrock,Vimpeli Mica schist, Alajärvi
1.7
Age Ga
B
DM 2
εNd
CHUR
0 -2
Tampere belt Two arenites (from Mauri) and a mudrock (from Ylöjärvi),upper sedimentary group of CS Graywacke-slate with unknown stratigraphic position, data from Miller et al. (1986) Lower sedimentary group of CS Graywackes from Siivikkala Graywackes from Nokia and Ahvenlammi
-4
2.3
2.0
1.7
Age Ga
C
DM 2
εNd
0
Häme and Uusimaa belts and Saimaa area Quartz arenites, Tiirismaa and Hyvinkää Immature graywacke, Orijärvi Immature graywacke, Haukivuori Immature graywacke and mudrock, Ristiina Immature graywacke, Joroinen Mature graywacke and mudrock, Rantasalmi
CHUR
-2 -4
2.3
2.0
1.7
Age Ga
Fig. 8.4. Nd isotope evolution diagrams for Svecofennian metasedimentary rocks. Data from Miller et al. (1986), Huhma (1987), and Lahtinen et al. (2002). The CHUR and DM evolution lines as in Figure 8.2. CS–central Svecofennia.
arenites are characterized by 1.93–1.86 Ga grains and deviate from the other Svecofennian sedimentary rocks in having a maximum deposition age of ~1.87–1.86 Ga.
3. Geochemical and tectonomagmatic characterization of the volcanic rocks In this paper, the classification of volcanic rocks is largely based on silica and alkali
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Table 8.2. Ages of plutonic clasts in Svecofennian conglomerates. Sample/Locality Pohjanmaa belt Haapajärvi Tampere belt A1203 Ahvenlammi
Rock type
Age ± 2σ (Ma)
Reference
Comments
granodiorite
1888 ± 7 Ma
Marttila (1987)
ref. in Vaasjoki and Sakko (1988)
tonalite
2556 and 2536 Ma
Kähkönen and Huhma (1993)
A26 Vähä-Lima, Tampere granitoid A90 Iso-Kartano, Tampere granitoid A144 Välimäki, Lavia tonalite Saimaa area Haukivuori
granitoid
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Pb/206Pb ages,
1884 ± 3 Ma 1890 ± 3 Ma 1888 ± 3 Ma
Nironen (1989) Nironen (1989) Nironen (1989) Kähkönen and Huhma (1993): εNd(at 1.90 Ga) +0.1
1885 ± 6 Ma
Korsman et al. (1988)
contents. This approach can be problematic because of the possible mobility of alkalies. Because P is significantly less mobile than K, P contents are in some cases discussed to further assess the character of the rocks (low-K, medium-K, high-K, and shoshonitic types). According to Ewart (1982) and Pearce (1982), low-K basalts have in general lower P2O5 contents (~0.1–0.15 wt.%) than mediumK basalts (~0.2–0.3 wt.%). Similarly, P2O5 contents in high-K and shoshonitic basalts tend to be higher (~0.3–0.7 wt.%) than those in medium-K basalts. Interpretations on the tectonomagmatic affinities of the Svecofennian volcanic rocks are largely based on proportions of mafic, intermediate, and felsic rocks, Ti vs. Zr variation, and Ti contents of mafic rocks. Where analyses made by the INAA or ICP-MS methods are available, the relatively immobile trace elements Th, Ta, Nb, and REE can be used for tectonomagmatic discrimination (see Pearce, 1982, 1983). In these cases the data are shown on chondrite- and NMORB-normalized diagrams. Note, however, that mafic volcanic rocks of some parts of continental flood basalt provinces and oceanic plateaus tend to have relatively low Ti contents and thus fall in the arc field in the classic Ti vs. Zr diagram of Pearce (1982), rather the field of within-plate lavas (e.g., Marsh, 1987; Neal et al., 1997; Tejada et al., 2002). CHAPTER
207
εNd(at 1.90 Ga) –11.7
4. Savo belt 4.1. General The Savo belt (Figures 8.1 and 8.5) is bounded and penetrated by numerous faults and shear zones. In general, the belt consists of migmatized mica gneisses (Figure 8.6A) that are mainly of graywacke and mudrock origin. Quartz-feldspar schists and amphibolites (both of volcanic and volcaniclastic origin) as well as graphite schists (black shales) are locally abundant. The volcanic rocks have a limited areal extent but are economically important due to massive sulfide deposits. Locally preserved pillow structures (Figure 8.6B) indicate that the eruptions were at least in part subaqueous. The Savo belt differs from the other Finnish Svecofennian belts in including ~1.93–1.92 Ga arc-type volcanic rocks and tonalitic gneisses with εNd (T) typically on the order of ~ +2 to +4 (Table 8.1; Figure 8.2). The Pb isotope composition of the sulfide deposits is also relatively primitive (Vaasjoki, 1981). The eastern parts of the Savo belt in particular, are characterized by fault-bounded blocks with variable metamorphic and structural histories. Metamorphic grade varies from medium-T amphibolite facies (550–600 oC) to granulite facies (800–880 oC) at pressures of 5 ± 1 kb (Hölttä, 1988, 1995). The metamorphic evolution culminated close to a main phase of
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granitoid magmatism at ~1885 Ma. Bedding planes, gneissose banding, and a dominant schistosity in the Savo belt have northwesterly or northerly strikes, but vary somewhat due to polyphase deformation. There is some variation in the characterization of the earliest deformation phases (Ekdahl, 1993; Kärki et al., 1993; Luukas, 1997; Korsman et al., 1999) but the following generalization seems plausible. The two earliest folding phases produced isoclinal to tight recumbent folds with east to northeast vergence; a regionally pervasive foliation or gneissose banding also developed during these phases. A third folding phase turned the flat-lying structures subvertical and resulted in approximately Norientated elongate antiforms and synforms, which govern the structure of the Savo belt. Open curving of the axial planes of these folds is largely due to a fourth phase. N-striking dextral shear zones were related to the third phase, whereas a system of NE-striking sinistral and S-striking dextral zones characterize the fourth deformation phase. The 1.93–1.92 Ga arc magmatism in the Savo belt was possibly related to SW-directed subduction (Ward, 1987; Lahtinen, 1994) towards the Keitele microcontinent. The age of the earliest tectonothermal events in the Savo belt is not well constrained. However, the newly formed Keitele–Savo entity possibly collided with the Archean craton at 1.91–1.90 Ga (Lahtinen, 1994; Nironen, 1997; Chapter 11), and this event may have involved extensive deformation and metamorphism.
4.2. Pielavesi–Pyhäsalmi region The Pielavesi–Pyhäsalmi region (Figures 8.1 and 8.5) is relatively rich in volcanic rocks and massive sulfide deposits. Kousa et al. (1994) divided the volcanic rocks at Pyhäsalmi into two major units: the (older) Eastern volcanic sequence (EVS), the host of the Pyhäsalmi massive sulfide deposit, and the (younger) Western volcanic sequence (WVS). The EVS 356
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consists of volcanic rocks and rare graphiteand sulfide-bearing, calcareous, and tuffaceous rocks (see also Kousa and Lundqvist, 2000). The EVS volcanic rocks comprise a bimodal association of low-K basalts, basaltic andesites, and rhyolites with arc affinity (Figure 8.7; Kousa et al., 1994). They show a slight LREE enrichment (Figure 8.8), and a 1.92 Ga rhyolite and 1.93 Ga tonalitic gneisses have a juvenile Nd isotope signature (Table 8.1; Figure 8.2; Lahtinen, 1994). The LREE enrichment and the abundance of felsic rocks indicate that the 1.93–1.92 Ga arc, although relatively immature, was not as primitive as true primitive arcs characterized by LREE-depleted mafic rocks of island arc tholeiitic series (Jakeš and Gill, 1970). The felsic EVS rocks were possibly formed by melting of newly formed, evidently ~1.95 Ga Paleoproterozoic crust (Lahtinen and Huhma, 1997). The WVS represents the volcanic rocks of the Ylivieska field (Pohjanmaa belt) and consists of medium- to high-K calc-alkaline basalts to dacites. These are rich in andesites, have arc affinities, and show pronounced enrichments in the LREE (Kousa et al., 1994; Figures 8.7, 8.8). The rocks probably represent extensions of the Kuusaa Formation ~35 km northwest of Pyhäsalmi (Figure 8.5) and were emplaced in an evolved arc milieu. The supracrustal sequence at Pielavesi (Figure 8.1) was divided by Ekdahl (1993) into four major units, of which the oldest has mafic volcanic rocks at the base but is otherwise characterized by migmatitized mica gneisses (graywackes) with graphitic and volcanic intercalations. The second stratigraphic unit consists of sedimentary carbonates, calc-silicate rocks, felsic volcanic rocks, cherts, minor iron formations and black shales, a uranium-phosphorus horizon, and a conglomerate (in the uppermost part). This shelf-type unit was deposited in a relatively shallow sea. In the third unit, known as the Säviä suite, intermediate tuffites and tuffs prevail but its uppermost volcanic rocks have a bimodal mafic-felsic character and host
SVECOFENNIAN
S U P R AC RU S TA L
ROCKS
Raahe
of Bo th nia
39
75
Vihanti
ulf
33
G
N
53
Oulujärvi
16
20 km Himanka
35
50
26
Ylivieska
32 76
Sievi
Kannus
Nivala 28
Kokkola 25
Kuusaa 46
Haapajärvi 36
41
Pyhäsalmi 58 30
38
Evijärvi
15
Pihtipudas
33
Svecofennian domain 15 = granodiorites, tonalites, quartz 38 = gneissic tonalites and granodiorites diorites, granites (~1.89–1.87 Ga) (~1.93–1.91 Ga)(Savo belt) 16 = gabbros, diorites, peridotites Karelian domain (~1.89–1.87 Ga) 39 = granites and granodiorites 25 = mica schists with intercalated are(~1.80 Ga) nites and conglomerates (Ylivieska field) 41 = granites and granodiorites 26 = mafic, intermediate and felsic vol(~1.89–1.86 Ga) canic rocks with sedimentary intercala46 = gabbros tions (1.90–1.88 Ga) (Ylivieska field) 50 = gneissic alkaline granite 28 = mica schists and mica gneisses with (~1.95–1.96 Ga) black schist intercalations (Evijärvi field) 30 = mafic volcanic rocks (Evijärvi field) 53 = Karelian mica schists, black schists, conglomerates, and arenites 32 = pyroxene granitoids (~1.885 Ga) 58 = Karelian quartz arenites 33 = mafic, intermediate, and felsic volcanic rocks with sedimentary intercala- Archean rocks tions (~1.90–1.88 Ga) (Savo belt) 75 = paragneisses 35 = mica gneisses and mica schists with intercalated carbonate rocks (Savo belt) 76 = tonalite–trondhjemite–granodiorite gneisses and migmatites 36 = felsic, intermediate, and mafic volcanic rocks (~1.92 Ga)(Savo belt) Fig. 8.5. Geologic map of the northern and central parts of the Savo and Pohjanmaa belts. Simplified from Korsman et al. (1997).
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B
A
Fig. 8.6. Structures of supracrustal rocks of the Savo belt. (A) Migmatitic mica gneiss, Ruukki, close to Raahe. The label is 10 cm long. (B) Pillow lava, Eastern volcanic sequence, Tetrinmäki, Pyhäsalmi. The pen is 13 cm long. Photos: Jukka Kousa (A),Yrjö Kähkönen (B).
base metal mineralizations. The youngest unit consists predominantly of graywackes with mafic volcanic intercalations. The Säviä suite is probably an equivalent of the 1.93–1.92 Ga volcanic rocks that host the Pyhäsalmi sulfide deposit, whereas the age of the deposition of the two lowermost Pielavesi units is not well constrained. However, δ13C values of ~ 0.3 to –3 from sedimentary carbonates in the shelf-type second unit show that these rocks were not deposited contemporaneously with the 2.2–2.1 Ga Jatulian carbonates (cf. Ekdahl, 1993; see also Karhu, 1993, and Chapter 16). The volcanic rocks of the uppermost unit are probably 1.90–1.88 Ga (Korsman et al., 1997); the massive Kotajärvi lava or diorite within this unit is ~1882 Ma (Salli, 1983; Ekdahl, 1993; Lahtinen, 1994; Table 8.1). Geochemical characterization of the Pielavesi volcanic rocks is problematic because up to 40% of them may have been hydrothermally altered (Ekdahl, 1993). However, the rocks point to a volcanic arc rather than intraplate or ocean floor environments; some rocks with fairly high Ti might indicate episodes of rift-related volcanism (Figure 8.7; see also Lahtinen, 1994). The arc-type signatures are supported by Ti/Zr/Y and Ti/Mn/P variations (Ekdahl, 1993). The volcanic rocks of the two 358
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oldest units are mostly medium-K, whereas those of the Säviä suite are of low-K type. The bulk of the rocks of the youngest unit, especially those of the Kotajärvi unit, have high-K, high-P trachyandesitic compositions and represent arc-type magmatism at a late stage of evolution of the Savo belt.
4.3. Rautalampi region The supracrustal rocks at Rautalampi include volcanic and sedimentary rocks also formed in a Paleoproterozoic immature arc (Lahtinen, 1994). In general, they are correlated with the three earliest units at Pielavesi but primary features have not been preserved as well as at Pielavesi and Pyhäsalmi. The volcanic rocks at Rautalampi are mainly mafic to felsic gneisses; however, andesitic rocks are fairly common, too. In general, the mafic volcanic rocks are of low-K type with arc affinity and show slight enrichments in the LREE (Figures 8.7, 8.8; Lahtinen, 1994).
4.4.Volcanic rocks of the Virtasalmi region The Virtasalmi region (Figure 8.1) resembles structurally the Savo belt as planar structural elements dominantly strike northwesterly to
SVECOFENNIAN
S U P R AC RU S TA L
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A
B Oldest unit Second to oldest Säviä suite
10000
Eastern volcanic sequence Western volcanic sequence
10000
Youngest unit
Ti ppm
Ti ppm
Pielavesi 1000 10
Pyhäsalmi 1000 100
Zr ppm
10
500
C
Zr ppm
100
500
100
500
D Virtasalmi field Haukivuori field Joroinen field Lawrie averages
RB
O
WITHIN PLATE LAVAS
M
10000
Ti ppm
10000
Ti ppm
ARC LAVAS
Rautalampi
Virtasalmi region
1000
1000 10
Zr ppm
100
10
500
Zr ppm
Fig. 8.7. Ti vs. Zr diagrams of volcanic rocks of the Savo belt: (A) Pielavesi, (B) Pyhäsalmi, (C) Rautalampi, (D) Virtasalmi region. Data from Lawrie (1992), Ekdahl (1993), Kousa et al. (1994), Lahtinen (1994), and Pekkarinen (2002). The fields of MORB (mid-ocean ridge basalts), within-plate lavas, and arc lavas according to Pearce (1982). The cross in (C) shows the average of Malaitan type C-G basalts from the Ontong Java Plateau (Neal et al., 1997).
northerly, opposite to the mainly easterly to northeasterly trends in the Saimaa area and Pirkanmaa belt. Korsman et al. (1997) included large parts of the region into southern Svecofennia (their Accretionary arc complex of southern Finland) and considered only the volcanic-dominated part (Virtasalmi field, see below) to belong to the Savo belt (their Primitive arc complex of central Finland). In this chapter, the Virtasalmi region is divided into three fields: (1) the Virtasalmi field dominated by amphibolitic mafic volcanic rocks; (2) the Haukivuori field dominated by mica gneisses and migmatites (but locally CHAPTER
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with well-preserved primary features) of the Saimaa area; and (3) the Joroinen field, east of Virtasalmi, characterized by mica gneisses and mica schists but also containing relatively abundant felsic to mafic volcanic rocks. The boundary between the Virtasalmi and Joroinen fields is not clear, and Pekkarinen (2002) suggested that the volcanogenic rocks of these areas are closely related. In the following, I will mainly discuss the volcanic rocks, the sedimentary rocks will be treated below in Section 8.4. The Virtasalmi field consists of amphibolites, mica gneisses and mica schists, graphite-
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A
ROCK / CHONDRITE
ROCK / CHONDRITE
B Pyhäsalmi eastern volcanic sequence Mafic rocks
100
100
10
1
Pyhäsalmi eastern volcanic sequence Felsic rocks
La Ce
Nd
Sm Eu
Tb
1
Yb Lu
C
10
La Ce
Nd
Sm Eu
Tb
Yb Lu
D Rautalampi volcanic rocks
Intermediate rocks Mafic rocks
100
ROCK / CHONDRITE
ROCK / CHONDRITE
Pyhäsalmi western volcanic sequence
100
10
1
La Ce
Nd
Sm Eu
Mafic Felsic
Tb
10
1
Yb Lu
La Ce
Nd
Sm Eu
Tb
Yb Lu
Fig. 8.8. Chondrite-normalized rare earth element (REE) patterns of volcanic rocks of the Savo belt. (A) Pyhäsalmi EVS mafic (SiO2 <57%) volcanic rocks; (B) Pyhäsalmi EVS felsic (SiO2 >67%) volcanic rocks; (C) Pyhäsalmi WVS mafic (SiO2 <57%) and intermediate (SiO2>57%) volcanic rocks; (D) Rautalampi mafic and felsic volcanic rocks. Data from Kousa et al. (1994) and Lahtinen (1994). Chondrite values from Boynton (1984).
bearing and black schists, sedimentary carbonates, calc-silicate rocks, cherts, Fe-rich strata, U-P-bearing horizons, minor felsic volcanic rocks, and Cu and Zn deposits (Lawrie, 1992; Reinikainen, 2001; Pekkarinen, 2002). Recrystallization and deformation have largely destroyed primary structures, but such features as massive and pillow lavas, tuffs, flow top breccias, lava tubes, and amygdules have been preserved in low-strain zones. Highly vesicular pillow structures indicate that the mafic volcanic rocks of the Virtasalmi field were erupted in relatively shallow water. The volcanic rocks of the Virtasalmi field are subalkaline, mainly medium-K tholeiitic basalts and andesites (Lawrie, 1992; Pekkarinen, 2002), and disperse into the three fields in Figure 8.7. Non-arc affinities are 360
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supported by the relatively high Nb values (Lawrie, 1992). The contents of P2O5 and K2O (~0.3–0.6 wt.% and 0.4–1.0 wt.%, respectively) and La/Y ratios are higher than in typical NMORB. Overall, the rocks resemble EMORB, WPL or oceanic plateau lavas rather than NMORB or arc basalts. The mafic volcanic rocks of the Virtasalmi field are probably ~1.92–1.905 Ga in age (Korsman et al., 1997; Pekkarinen, 2002). A rhyolite close to the boundary of the Virtasalmi and Joroinen fields has an age of 1906 Ma (Pekkarinen, 2002; Table 8.1). The rhyolite and a nearby 1903 Ma gneissic tonalite have εNd (T) values typical of the tonalitic and volcanic rocks of the Savo belt (Huhma, 1986; Lahtinen and Huhma, 1997; Figure 8.2) although they are slightly younger than the latter rocks.
SVECOFENNIAN
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ROCKS
Geographically, the Virtasalmi field is an offshoot from the bulk of the Savo belt. It also differs geologically, because sedimentary carbonates are relatively abundant and the mafic volcanic rocks are not of arc type. The mafic rocks were possibly emplaced in a riftrelated or marginal basin setting, but an origin as part of an oceanic plateau is also feasible. The Virtasalmi field would then represent an exotic terrane within the Savo belt. The volcanic rocks of the Haukivuori field occur as thin and short lenses but show relatively well-preserved pillow and lava breccia structures. Pekkarinen (2002) suggested them to be considerably younger than the mafic lavas of the Virtasalmi field. The Haukivuori volcanic rocks are mainly low-K mafic or ultramafic and tend to have lower K2O and P2O5 contents (typically 0.07–0.13 wt.% and 0.12–0.14 wt.%, respectively) than the mafic lavas of the Virtasalmi field (Pekkarinen, 2002). This and the fact that they are in part high in Ti and Zr (Figure 8.7D) suggest that, as a group, they resemble N- or transitional (T-) MORB rather than EMORB, WPL or island arc tholeiites.
5. Pohjanmaa belt 5.1. General Supracrustal rocks in the Pohjanmaa belt are divided into two fields (Figures 8.1, 8.5). The rocks of the Evijärvi field are dominated by turbiditic graywackes and mudrocks but also contain units of mafic lavas with MORB to WPL affinities as well as associated black shales, carbonate and calc-silicate rocks, and cherts (Figures 8.9 through 8.11). The rocks of the Ylivieska field are probably younger than these and are composed of ~1.90–1.88 Ga arc-type volcanic and related clastic sedimentary rocks. In the generalised division of Lahtinen et al. (2002), the sedimentary rocks of the Evijärvi field belong to the lower CHAPTER
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sedimentary group of central Svecofennia, whereas those of the Ylivieska field represent the upper sedimentary group of this terrane. Lehtonen et al. (2003) reported a similar system of arc-type volcanic rocks in the east and MORB- to WPL-like rocks in the west, both enveloped by mica gneisses of turbidite origin near Siipyy, southernmost Pohjanmaa. Here, however, sedimentary carbonates are relatively common in the east. Regional metamorphism in the Pohjanmaa belt peaked at 1.89–1.88 Ga (Mäkitie, 1999, 2000). The supracrustal rocks of the Evijärvi field were metamorphosed to schists, gneisses, and locally migmatites, but they occasionally show well-preserved primary structures (Figure 8.9). In general, the metamorphic conditions in the Evijärvi field increased, from medium-T amphibolite facies in the northeast to lower-T granulite facies in the southwest (Vaarma, 1990; Mäkitie and Lahti, 1991; Vaarma and Pipping, 1997; Mäkitie, 1999). The rocks of the Ylivieska field were metamorphosed at low-T to medium-T amphibolite facies, and well-preserved primary structures are abundant. At Evijärvi, an early phase of deformation created tight to isoclinal recumbent folds and thrusts with east to northeast vergence (Vaarma and Pipping, 1997). Subsequent deformation included open folds with E-trending subvertical axial planes and open to tight folds with axial planes striking northwest. In the Seinäjoki area (Figure 8.1), Mäkitie (1999, 2000) presented a similar sequence of deformation but distinguished a weak schistosity formed before the early isoclinal recumbent folding. Lahtinen et al. (Chapter 11) suggest that the early thrusts with east to northeast vergence in the Evijärvi field were related to the ~1.90 Ga collision of an inferred Bothnia microcontinent with the Keitele microcontinent in the southwest. Deformation in the Ylivieska field seems less prominent than in the Evijärvi field. Around Ylivieska, Salli (1964) indentified
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several, mostly isoclinal early-stage synclines and anticlines with NW-trending subvertical axial planes and subhorizontal fold axes. In the study area of Strand (2002), supracrustal rocks form a wide, relatively tight, northwestplunging syncline. The ~1.89 Ga arc magmatism of the Ylivieska field was probably related to subduction directed approximately to the northeast (Nironen, 1997). This idea is in part based on a N-dipping mantle reflector under the Gulf of Bothnia (BABEL Working Group, 1990), i.e., beneath the western extensions of the Pohjanmaa belt. The change from the N-directed subduction at Tampere (see Section 6.1) to the NE-directed subduction beneath the Pohjanmaa belt might have been caused by pinning of the subduction zone in the Tampere area (Nironen, 1997).
5.2. Evijärvi field The mafic volcanic rocks of the Evijärvi field mainly occur as relatively thin, less than 1km-wide formations and are mostly lavas frequently showing pillow structures (Figure 8.9). The carbonate rocks, black shales, and cherts of the Evijärvi field are associated with these volcanic formations (Vaarma and Pipping, 1997). On a Ti vs. Zr diagram (Figure 8.10), the volcanic rocks fall into the MORB, WPL, and arc fields but, in general, do not have arc affinity. Besides this, Ti, Zr, Cr, V, and Sr variations (figures not shown) indicate MORB and WPL affinities rather than an arc setting (Vaarma and Kähkönen, 1994). Chondritenormalized REE patterns vary from LREE depleted to LREE enriched (Figure 8.11). The LREE-depleted, NMORB-like Evijärvi basalts have εNd (at 1.9 Ga) values of ~ +3 to +4 (Figure 8.2) and were derived from depleted mantle sources. Based on Ti and Zr contents, the partly pillowed basalts at Nurmo and Vittinki close to Seinäjoki (Figure 8.1) resemble those at Evijärvi (Figure 8.10). Felsic volcanic rocks are rare in the Evijär362
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vi field and have only been found near Kokkola (Figure 8.5). These rocks host U-bearing phosphorite layers and are stratigraphically below skarn and carbonate rocks, black shales, and graywackes (Kousa and Lundqvist, 2000). Near Seinäjoki, plagioclase and uralite-plagioclase porphyritic sills or dikes, 1886 Ma in age (Table 8.1), are possibly related to the arc-type volcanic rocks of the Ylivieska field. The age distribution of detrital zircon in a graywacke from the Evijärvi field (Figure 8.3B) is like those in the two graywackes from the stratigraphically low levels of the Tampere belt; together these samples represent the lower sedimentary group of central Svecofennia. Based on the youngest concordant grains, the maximum deposition age of these graywackes is ~1.92 Ga. Mixing of Archean and 2.0–1.92 Ga Proterozoic components is also supported by the strongly to slightly negative εNd (at 1.9 Ga) values (Figure 8.4). The 2.0–1.92 Ga detritus was possibly derived from the 1.93–1.92 Ga Savo arc and the inferred Keitele microcontinent (Lahtinen et al., 2002). The relations of the mafic lavas of the Evijärvi field to the enveloping graywackes and mudrocks are not known well because of poor exposure and lack of isotope ages. In general, however, thrusts with east to northeast vergence are evident (see above) and, probably, the Evijärvi field consists of allochthonous slices. Based on the scatter from NMORB-like to WPL-like basalts, the mafic lava units of the slices represent variable non-arc tectonic settings.
5.3.Ylivieska field Rocks of the Ylivieska field occur as scattered complexes of volcanic and sedimentary rocks (Figure 8.5). The former are of both pyroclastic to volcaniclastic and lava origin, and a significant part of them were emplaced in shallowwater or subaerial environments (Figure 8.9; Kousa and Lundqvist, 2000; Strand, 2002). They range from basalts to K-rhyolites and
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A
B
C
D
E
F
Fig. 8.9. Structural features of supracrustal rocks of the Pohjanmaa belt. (A) Pillow lava, Evijärvi field; (B) Pillow breccia, Evijärvi field; (C) Pelite and graywacke strata with concretions and porphyroblasts of sillimanite, Evijärvi field; (D) Volcaniclastic conglomerate, Antinneva Formation, Pyhäjoki (~20 km south of Raahe), Ylivieska field; (E) Cross-stratified volcaniclastic sedimentary rock with few felsic pebbles, Antinneva Formation, Pyhäjoki (~20 km south of Raahe), Ylivieska field; (F) Synsedimentary deformation in felsic sedimentary rock, Alavieska (~20 km northwest of Ylivieska),Ylivieska field.The pen is ~13 cm and the compass ~12 cm long. Photos: Yrjö Kähkönen (A, F), Markus Vaarma (B, C), and Jukka Kousa (D, E).
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363
ROCK / CHONDRITE
Evijärvi field mafic volcanic rocks RBX
O
10000
M
100
WITHIN PLATE LAVAS
X Ti ppm ARC LAVAS
10
1 1000 10
100
Zr ppm
500
Fig. 8.10. Ti vs. Zr diagrams of mafic volcanic rocks from the Evijärvi field. The symbols with different colors indicate separate formations; crosses are averages from the Nurmo–Vittinki area near Seinäjoki. Data from Mäkitie and Lahti (1991),Vaarma and Kähkönen (1994), and Vaarma and Pipping (1997). The fields of MORB (mid-ocean ridge basalts), within-plate lavas, and arc lavas are according to Pearce (1982).
have calc-alkaline, mature island-arc affinity (Figures 8.7, 8.8). The sedimentary rocks are characterized by sandstones, conglomerates, and silty mudrocks with volcanic provenance (see also Kousa, 1997); they represent the upper sedimentary group of central Svecofennia. These sedimentary rocks were mainly deposited in fluvial or shallow-water environments and the fact that the volcanic–sedimentary complexes enveloped by turbidites suggest that related deep-water deposits are probably also present. The supracrustal successions of the Ylivieska field evidently represent separate volcanoes and intervening basins. The volcanoes were partly exposed above sea level, and their erosion resulted in fluvial and shallowwater deposits on and close to the flanks of the volcanic aprons as well as in deep-water turbidites more distal to the volcanic centers. There are no published ages on the volcanic rocks of the Ylivieska field, but Korsman et al. (1997) and Kousa and Lundqvist (2000) considered these to be 1.89–1.88 Ga old. The three youngest concordant grains in sediment 364
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La Ce
Nd
Sm Eu
Tb
Yb
Fig. 8.11. Chondrite-normalized rare earth element (REE) patterns of mafic volcanic rocks from the Evijärvi field. Colors of the patterns indicate the same formations as those in Figure 8.10. Data from Vaarma and Kähkönen (1994). Chondrite values from Boynton (1984).
sample A1554 (Figure 8.3D) fall in the range of 1882 ± 14 Ma to 1896 ± 10 Ma. Because the provenance of the sample was dominated by local arc-derived material, it can be assumed that the ages of arc-related volcanic and plutonic rocks of the Ylivieska field vary from ~1.88 Ga to ~1.90 Ga, rather than from ~1.88 Ga to ~1.89 Ga. Based on the interpreted ages of concordant grains, the sedimentary rocks of the Ylivieska field have a maximum deposition age of ~1.89 Ga (Figure 8.3D; see also Lahtinen et al., 2002). This is supported by a 1888 Ma granitoid cobble in a related conglomerate (Table 8.2). The scarcity of Archean grains in sample A1554 and the mixture of Archean and Proterozoic detritus in sample A1555 agree with the εNd evolution lines (Figure 8.4). The detailed stratigraphic position of the latter sample is not quite clear, but Lahtinen et al. (2002) tentatively suggested it to be a molasse deposit related to a ~1.89 Ga collision.
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6. Tampere and Pirkanmaa belts 6.1. General Since the studies of Sederholm (1897), Seitsaari (1951), Simonen and Kouvo (1951), and Simonen (1953), the well-preserved medium-grade Tampere belt has been a key area of the Svecofennian supracrustal successions in Finland. The belt is primarily composed of turbiditic graywackes and mudrocks as well as ~1.905–1.89 Ga arc-type volcanic and related sedimentary rocks (Ojakangas, 1986; Kähkönen, 1987, 1989, 1994, 1999; Table 8.1; Figures 8.12 through 8.17), and it serves as a good example of a Proterozoic greenstone belt. In the north, the belt is bounded by the ~1.88 Ga Central Finland granitoid complex and in the south by the high-grade Pirkanmaa belt. The belt is dominated by migmatites and gneisses of turbiditic origin, but is also relatively rich in black shales and includes mafic and ultramafic volcanic rocks with MORB to WPL affinities. The boundary between the two belts is in places gradual, in other places a fault zone. The bulk of the Tampere belt belongs to a volcanic arc, whereas the Pirkanmaa belt mainly represents the subduction zone complex of this arc system; the latter might also contain Paleoproterozoic rocks considerably older than 1.9 Ga (Kähkönen, 1999). Metamorphism in the Tampere and Pirkanmaa belts was of low-P type and took place at low-T amphibolite to greenschist/amphibolite facies (Tampere belt) and at high-T amphibolite facies (Pirkanmaa belt) (Kilpeläinen et al., 1994; Kilpeläinen, 1998). Metamorphism in the latter culminated at 1.88 Ga (Mouri et al., 1999). The structure of the central part of the Tampere belt has been regarded as a syncline (e.g., Kähkönen, 1989; Nironen, 1989a), but a U-Pb zircon age of ~1892 Ma from Tesoma in the southernmost part of the belt (Figure 8.12) is not compatible with this (Table 8.1). Instead, Kähkönen et al. (2004) considered CHAPTER
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a folded early thrust as a probable explanation. The earliest folding produced a major E-trending synform with subvertical axial planes, subhorizontal fold axes, and subvertical stretching lineations. The folds are mainly isoclinal but in a few places open. Folding after this phase was concentrated in subvertical zones with variously trending axial surfaces (Nironen, 1989a). Kilpeläinen (1998) suggested that the tectonic and metamorphic evolution started earlier in the Pirkanmaa belt than in the Tampere belt. In the Pirkanmaa belt, an early recumbent folding produced subhorizontal schistosity parallel with bedding and compositional banding (Kilpeläinen et al., 1994; Kilpeläinen, 1998). The F2 folds had E-trending subvertical axial planes and subhorizontal fold axes. Overall, deformation resulted in complex fold interference patterns (Figure 8.13L; Arkimaa et al., 2000). Recently, Kilpeläinen and Nironen (2002) described relics of foliation that preceded the D1 event. Evidently, the earliest deformation in the Pirkanmaa belt involved more than one phase of recumbent folding. Based on a N-dipping mantle reflector, the 1905–1890 Ma arc-type volcanism in the Tampere belt was probably related to subduction to the north under the Keitele microcontinent (e.g., Chapter 11). The interpretation of the Tampere belt as a volcanic arc and the Pirkanmaa belt as the associated subduction zone complex is in line with this idea. The bulk of metamorphic and deformational events in these belts were probably related to accretion or collision of southern Svecofennia with central Svecofennia at 1.88–1.87 Ga.
6.2. Central Tampere belt According to the generalized stratigraphic scheme, the Haveri Formation is probably the oldest supracrustal unit in the central Tampere belt (Figures 8.12, 8.14). It is characterized by basaltic lavas that are commonly pillowed and have EMORB affinity (Figures 8.13A, 8.15
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365
500
470
Haveri
Viljakkala
Osara
Central Finland granitoid complex
Lake Näsijärvi
Koskuenjärvi
Pohtola
Harhala
Valkjärvi
Sammatinjärvi Oksijärvi
PSZ
Takamaa
Hämeenkyrö batholith
6830
Veittijärvi Ylöjärvi
Lempiänniemi Tervakivi Pylsynlahti Pirttiniemi
Pulesjärvi Kolunkylä
Värmälä stock
Orivesi
Ahvenlammi Viinaränninnotko
Myllyniemi Siivikkala Mauri
Nokia batholith
6820
Tesoma Kalkku
Nokia
TAMPERE
10 km
ESZ
Metagraywackes and metapelites Metaconglomerates
Metavolcanic rocks in general Felsic metavolcanic and metasedimentary rocks
Gneisses and migmatites
Pukala porphyry
Granitoids with minor diorites and gabbros Hinge zone of major synform Shear zones; PSZ = Paarlahti shear zone ESZ = Epilä shear zone
Fault, interpreted Viljakkala, Ylöjärvi, Näsijärvi E, and Pulesjärvi sections
Fig. 8.12. Lithological map of the central Tampere belt with a simplified structural interpretation. Based on Kähkönen (1999) and references therein. The dotted lines indicate approximate positions of the Viljakkala,Ylöjärvi, Näsijärvi E, and Pulesjärvi sections.
through 8.17; see also Mäkelä, 1980; Kähkönen and Nironen, 1994). The Haveri formation is overlain by the turbidite-dominated Osara and Myllyniemi Formations, which are probably lateral counterparts and represent the lower sedimentary group of central Svecofennia. These are succeeded by units rich in arc-type volcanic and related sedimentary rocks, of which the Pulesjärvi–Kolunkylä complex is the most prominent. The arc-type volcanic rocks include both pyroclastic units and lavas and were formed in a highly evolved arc (Kähkönen, 1987, 1989, 1994; Lahtinen, 1996). The related sedimentary rocks have an overwhelmingly volcanic provenance and include turbiditic graywackes, mudrocks, and conglomerates as well as fluvial or shallowwater sandstones and conglomerates (Rautio, 1986; Leveinen, 1990; Kähkönen, 1999); they belong to the upper sedimentary group 366
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of central Svecofennia. At Ylöjärvi, a mainly mafic volcanic unit with transitional arc to WPL affinity (Takamaa Formation) overlies the arc-related sedimentary rocks. The EMORB-like basaltic lavas of the Haveri Formation are occasionally amygdaloidal. They have interlayers of and are overlain by tuffs, cherts, sedimentary carbonates, and skarns. The lava-dominated part further grades into tuffs, sulfide-rich fine-grained tuffaceous rocks, and black shales. The εNd (at 1.9 Ga) values of the Haveri basalts are slightly positive (Figure 8.2), probably indicating mantle source enrichment relative to depleted mantle some time before the emplacement of the lavas. Pb isotopes in the Haveri sulfides are consistent with mantle sources (Vaasjoki and Huhma, 1999). Pb isotope compositions in the other sulfide deposits of the Tampere belt, and in central Svecofennia in general, are less
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primitive than at Haveri and in the Savo belt (Vaasjoki, 1981). The turbidites of the Myllyniemi Formation show westerly paleocurrents and are largely mid-fan deposits of submarine fans (Ojakangas, 1986), but also contain conglomerates deposited in submarine fan channels (Leveinen, 1990; Figures 8.13B through 8.13D). Black shales and mudrocks relatively rich in C and S are not common but occur at Nokia, at Ahvenlammi (Myllyniemi Formation), and in low horizons in the Osara Formation (Figures 8.12 and 8.14), i.e., in the lowermost parts of the thick turbidite unit of the central Tampere belt. The Pulesjärvi–Kolunkylä complex is characterized by intermediate to mafic lava or lava-like and pyroclastic rocks as well as sedimentary rocks with a volcanic provenance (Figures 8.13F–I). The sandstones and conglomerates are largely fluvial or shallow-water deposits, but the sedimentary rocks also include mudrocks, graded graywackes (see the cover picture of this chapter), and conglomerates probably deposited from turbidity currents in relatively deep water. The complex possibly comprises emergent volcano(es) with subaerial to submarine sedimentary aprons. The arc-type volcanic rocks of the central Tampere belt range from basalts to rhyolites and are most commonly dacites and andesites (Kähkönen, 1987, 1989, 1994). They are mainly of high-K and medium-K character but shoshonitic and trachytic or high-K rhyolitic types are also relatively abundant. Low-K arc tholeiitic rocks have not been identified. The chondrite-normalized REE patterns of the arc-type volcanic rocks show moderate to pronounced LREE enrichment (Figure 8.15; see also Kähkönen, 1994; Lahtinen, 1996). In general, the rocks have pronounced Nb depletions and positive anomalies of La, P, and Sm (Figure 8.17) indicating enriched mantle sources with a subduction component. These features point to an evolved arc setting. The LREE enrichment in the Takamaa Formation CHAPTER
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tends to be less pronounced than in the arctype volcanic rocks and may, together with the transitional character of this unit (Figure 8.15A), indicate incipient rifting at ~1.89 Ga. The 1904 ± 4 Ma Koskuenjärvi Formation (Table 8.1) in the northern limb of the synform at Orivesi (Figure 8.12) is the oldest dated unit of the Tampere belt but, based on regional correlation, it is younger than the Haveri Formation. It seems to have deposited after the Myllyniemi and Nokia graywackes that have detrital zircons of ~1.92 Ga as the youngest population. The 1889 ± 5 Ma Takamaa Formation is the youngest dated unit and is, within error limits, coeval with the 1892 ± 3 Ma Tervakivi rhyolite (Figure 8.14). A plagioclase porphyry has an even younger age of 1880 ± 7 Ma (Table 8.1), but it may be related to the enveloping granitoids. Granitoid cobbles in the conglomerates of the Pulesjärvi–Kolunkylä complex have ages of ~1890–1885 Ma (Table 8.2) and suggest a maximum deposition age of ~1.89 Ga. Within error limits these rocks are coeval with the Takamaa Formation and some plutonic rocks that intrude the belt. Evidently, the orogenic evolution was rapid. Zircon in the Mauri arenites, ~20–40 km west–southwest of Tampere (Figure 8.12) is ~1.90 Ga (Matisto, 1977). Thus these arenites were deposited at ~1.90–1.89 Ga at the earliest and were mainly derived from ~1.90–1.89 Ga sources. They are possibly coeval with the Pulesjärvi–Kolunkylä complex (Kähkönen, 1999) and represent the upper sedimentary group of central Svecofennia. The 1898 ± 4 Ma Sammatinjärvi dacite from Ylöjärvi has an εNd (at 1.9 Ga) value of –0.7 (Figures 8.3, 8.12; Table 8.1). Similar near-zero values are registered by a mudrock from Ylöjärvi and the Mauri arenites (Figure 8.4) that represent the upper sedimentary group of central Svecofennia with a major volcanic provenance. These values are similar to those of the Central Finland granitoid complex
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367
A
B
C
D
E
F
Fig. 8.13. Structures of supracrustal rocks of the Tampere (A to J) and Pirkanmaa (K to L) belts. (A) Pillow lava, Haveri Formation,Viljakkala. (B) Conglomerate, Ahvenlammi Member, Myllyniemi Formation, Kangasala. (C) Graded bedding in turbidite, Myllyniemi Formation, Siivikkala,Ylöjärvi. The pen is ~15 cm long. (D) Load casts in turbidite, Myllyniemi Formation, Alasenlahti, Tampere. The pen is 17 cm long. (E) Small-scale cross lamination in the Tuuliniemi Formation, Tampere. The vertical dimension corresponds 8 cm in nature. (F) Dark fiamme in the Sileäkallio ignimbrite, Pulesjärvi–Kolunkylä complex, Tampere. (G) Reverse to normal grading in a mafic volcaniclastic stratum, Pulesjärvi–Kolunkylä complex, Tampere. (H) Trough-type cross bedding in fluvial sandstone with volcanic provenance, Pulesjärvi–Kolunkylä complex, Tampere. (I) Volcanic conglomerate with an interbed of sandstone in the center, Pulesjärvi–Kolunkylä complex, Tampere. (J) Pillow lava, Hoivasvuori, Suodenniemi. (K) Pillow lava,Vähä-Kassari, Kylmäkoski (commune just north of Urjala). (L) Polyphase folding,Vammala. In (F) through (L) the pen is 13 cm long. Photos:Yrjö Kähkönen (A, B, E through L), Maunu Härme (C), and Ragnar Törnroos (D).
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G
H
I
J
K
L
and they probably indicate derivation from a mafic source separated from the mantle during the Paleoproterozoic (Lahtinen and Huhma, 1997; Rämö et al., 2001). The Myllyniemi and Nokia turbidites received detrital material mainly from ~2.05– 1.92 Ga sources but also have a marked Archean component (Figures 8.3, 8.12, 8.14; Table 8.2). The Nd isotope evolution lines agree with the idea of a mixture of Archean and Proterozoic source material and further
suggest that the proportion of Archean detritus is more significant in the lowermost graywackes (Nokia and Ahvenlammi in Figure 8.4B) than in the Siivikkala graywackes in the stratigraphically middle parts of the Myllyniemi Formation.
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6.3. Western and eastern Tampere belt In the western Tampere belt, the volcanic rocks at Suodenniemi (Figures 8.1, 8.13J) are domi-
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Ylöjärvi profile Takamaa Fm. 1889 ± 5 Ma Veittijärvi conglomerate 6 km
Näsijärvi E profile Pulesjärvi–Kolunkylä complex
Pulesjärvi profile 5
Pohtola 2 km W
2 km E 4 km
Pulesjärvi–Kolunkylä complex
4 km
Tervakivi Fm. 1892 ± 3 Ma Pylsynlahti Fm. 3
3 Tuuliniemi Fm.
2
Pirttiniemi Fm. 1898 ± 8 Ma
2
1
Myllyniemi Fm.
1
Tervakivi Fm.
Multivuori
Viljakkala profile Myllyniemi Fm.
Harhala Fm. Osara Fm.
Ahvenlammi Mb. thrust?
Haveri Fm. 1 2 3
4 5 6
7 8 9
10 11 12
Viinaränninotko 13 14 15
Fig. 8.14. Generalized stratigraphic columns of the Tampere belt near Lake Näsijärvi based on Kähkönen (1999) and references therein. The illustration schematically shows variation in bed thickness, grain size, and sand to mud ratios; individual bed thicknesses for example in the Myllyniemi Formation are not to true scale. Legend: 1–mudrocks; 2–graywackes; 3–conglomerates; 4–felsic to intermediate volcanic and sedimentary rocks, variations in grain size are preliminary estimates; 5–trachytes and high-K rhyolites; 6–intermediate to felsic crystal and lithic tuffs and lapilli tuffs; 7–mainly andesitic lavas and lava-like rocks, in part pyro- or volcaniclastic; 8–basaltic to andesitic tuff breccias, agglomerates and lapilli tuffs, in part lava-like rocks; 9–mafic tuffs; 10–basaltic to andesitic matrix-supported tuff breccias or debris flow deposits; 11–Sileäkallio ignimbrite; 12–plagioclase phenocrysts; 13–clinopyroxene phenocrysts and clasts (presently hornblende/uralite); 14–sills and subvolcanic intrusions; 15–mylonite. The ages are from Table 8.1.
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SVECOFENNIAN
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ROCKS
A
B Takamaa Fm. Other units
10000
Suodenniemi Vihteljärvi, Kankaanpää
10000
Haveri Fm.
Ti ppm
Ti ppm
Central Tampere belt 1000 10
Western Tampere belt 1000
Zr ppm
100
10
500
C
100
Zr ppm
500
D
Kiikoinen
10000
10000
WITHIN RB
PLATE
O
Ti ppm
M
Ti ppm
ARC LAVAS Pirkanmaa belt
Pirkanmaa belt
1000
1000 10
Zr ppm
100
10
500
Zr ppm
100
500
Fig. 8.15. Ti vs. Zr diagram for volcanic rocks of the Tampere (A, B) and Pirkanmaa (C, D) belts. Data from Kähkönen (1989, 1994), Kähkönen and Nironen (1994), Lahtinen (1996), and from unpublished files of Markku Tiainen and the author. In (C), the symbols with different colors indicate samples from separate units. The fields of MORB (mid-ocean ridge basalts), within-plate lavas, and arc lavas are according to Pearce (1982).
nated by basaltic andesites and andesites with a medium-K and high-K character (Kähkönen, 1987). At Kankaanpää and Ikaalinen, the volcanic rocks range from basalts to dacites and are mostly of medium-K or high-K type (Kähkönen, 1987; Yli-Kyyny, 1990). The Ti vs. Zr diagrams and spidergrams mainly indicate an arc setting (Figures 8.15 and 8.17D) although some of the mafic rocks at Kankaanpää have relatively high Ti contents. At Luhanka, in the eastern part of the Tampere belt, the volcanic rocks are dominated by medium-K basalts to andesites and high-K rhyolites (Ikävalko, 1981; Luukkonen, 1994). High-K basalts and andesites are relatively CHAPTER
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common as well. The rocks have considerable depletions in Ta and Nb relative to Th (figures not shown) and were probably formed in an evolved arc.
6.4. Pirkanmaa belt In addition to turbidites and black shales, the sedimentary rocks of the Pirkanmaa belt include arenites and some conglomerates, but sedimentary carbonates are rare. Besides these, Hytönen (1999) reported cherts in glacial boulders. Some of the black shales seem to be spatially associated with the mafic volcanic rocks, and comparison of bedrock and
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371
B
Koskuenjärvi Fm. dacites and a rhyolite
66
100 63 10
66 1
La Ce
C
Nd
Sm Eu Tb
Ylöjärvi:
ROCK / CHONDRITE
68
100
10
57
63 53
1
La Ce
Nd
63
10 62
1
La Ce
D 58
Sm Eu Tb
Sm Eu Tb
Ylöjärvi: 62 59
Takamaa Fm.
Yb
63 55
10
54 53 57
1
Yb
E
51
Nd
100
55
Subalkaline rocks Shoshonitic rocks Trachytic rocks
70
Yb
Volcanic rocks stratigraphically below the Takamaa Fm.
61
Valkjärvi Fm.
100
74
65
Orivesi:
ROCK / CHONDRITE
ROCK / CHONDRITE
Orivesi:
ROCK / CHONDRITE
A
La Ce
Nd
56
Sm Eu Tb
Yb
Haveri Fm. mafic lavas
ROCK / CHONDRITE
ROCK / CHONDRITE
F
100
Pirkanmaa belt mafic lavas
100
10
1
10
1 La Ce
Nd
Sm Eu Tb
Yb
La Ce PrNd
Sm Eu Gd Tb Dy Ho Er Tm Yb Lu
Fig. 8.16. Chondrite-normalized rare earth element (REE) patterns of representative volcanic rocks of the Tampere (A to E) and Pirkanmaa (F) belts. The numbers in (A) to (D) give the SiO2 contents of the rocks; in (B) just for the subalkaline rocks. The samples in (E) and (F) are mafic. In (F), colors of the patterns indicate the same units as those in Figure 8.15C. Data from Kähkönen and Nironen (1994) and from unpublished files of Markku Tiainen and the author. Chondrite values from Boynton (1984).
low-altitude geophysical maps suggests the black shales are more abundant than the mafic volcanic rocks (see Korsman et al., 1997; Arkimaa et al., 2000). Whether this is a real feature or a biased view due to poor exposure is not known. In chemical composition, the bulk of 372
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the sedimentary rocks of the Pirkanmaa belt resemble those of the Myllyniemi and Osara Formations in the Tampere belt (Lahtinen, 1996); all these rocks belong to the lower sedimentary group of central Svecofennia. Furthermore, black shales are abundant compared to their rarity among the typical turbidites of
SVECOFENNIAN
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ROCKS
A
B 100
100
10 59 1
0.1
Ylöjärvi:
Orivesi–Tampere– 57 Ylöjärvi area ROCK / MORB
ROCK / MORB
55
58 53
10
1
0.1 Sr K Rb Ba Th Ta Nb La P Zr
Sm Ti Y
Sc Cr
Orivesi: Subalkaline basalt Sr K Rb Ba Th Ta Nb La P Zr
ROCK / MORB
ROCK / MORB
10
10
1
1
0.1 Sr K Rb Ba Th Ta Nb Ce P ZrHf Sm Ti Y Yb Sc Cr
Sr K
Ba Th Ta Nb Ce P Zr Hf Sm Ti Y Yb Sc Cr
F 100
100
Pirkanmaa belt 10
10
ROCK / MORB
ROCK / MORB
Pirkanmaa belt
1
0.1
Sc Cr
Haveri Fm. lavas
Hoivasvuori, Suodenniemi
E
Sm Ti Y
D100
C 100
0.1
Takamaa Fm. mafic volcanic rocks Mafic volcanic rocks below the Takamaa Fm.
1
0.1 Sr K Rb Ba Th
Nb La P Zr
Ti Y
Sc Cr
Sr K Rb Ba
NbCe P Zr
Sm Ti Y Yb Sc Cr
Fig. 8.17. Mid-ocean ridge basalt-normalized trace element patterns of representative volcanic rocks of the Tampere (A to D) and Pirkanmaa (E to F) belts. The numbers in (A) give the SiO2 contents of the rocks; the rocks in the other figures are mafic. In (F), colors of the patterns indicate the same units as those in Figure 8.15C. Data from Kähkönen and Nironen (1994), Lahtinen (1996), and from unpublished files of Markku Tiainen and the author. Normalizing values mostly from Pearce (1982), La (3.8 ppm) from Lahtinen (1996).
the Tampere belt. In this respect, a significant part of the Pirkanmaa belt sedimentary rocks resembles the lowest parts of the turbiditedominated units in the Tampere belt, where black shales are relatively common. However, a part of the Pirkanmaa belt sedimentary rocks were probably derived from the arc-type CHAPTER
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volcanic units of the Tampere belt (Lahtinen, 1996), and they evidently represent fore-arc sediments of the Tampere arc system. Compared to the Tampere belt, volcanic rocks in the Pirkanmaa belt are not abundant and they are mainly mafic to ultramafic lavas with rare pillow structures (Figure 8.13K).
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The basalts resemble MORB or WPL, although some units are transitional between MORB/WPL and arc-type rocks (Figures 8.15 through 8.17). The rocks display both flat and slightly LREE-enriched REE patterns. Picrites with trace element patterns similar to those of transitional MORB have also been described (Peltonen, 1995). Volcanic rocks of MORB affinity are relatively abundant in the southern Pirkanmaa belt and may indicate a suture between central and southern Svecofennia (Lahtinen, 1994, 1996). The mafic lava units of the Pirkanmaa belt possibly represent varying non-arc settings, and the belt may consist of several allochthonous slices.
7. Supracrustal belts within the Central Finland granitoid complex Several fragmentary arc-type successions occur within the Central Finland granitoid complex (Figure 8.1; Korsman et al., 1997). The sedimentary rocks in these belts are mainly mudrocks and resemble those related to the arc-type volcanic rocks of the Tampere belt and Ylivieska field. In the northern parts of the Central Finland granitoid complex, the volcanic rocks at Pihtipudas are felsic to intermediate and have been dated at 1883 ± 20 Ma (Table 8.1). The significance of this age is questionable, because it is a pooled age from seven granitoids, three felsic porphyries, and a dacite. The supracrustal rocks at Perho, 70–80 km west of Pihtipudas, resemble those of the Ylivieska field in age and overall character (Korsman et al., 1997). In the central parts of the Central Finland granitoid complex at Saarijärvi quartz-feldspar porphyritic rocks are abundant. They are mainly hypabyssal rocks probably related to the surrounding granitoids and have been dated at 1887 ± 2 Ma (Table 8.1). The plagioclase-porphyritic intermediate rocks at Haapamäki, ~50 km southwest of Saarijärvi, 374
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comprise another wide field and are predominantly subvolcanic, in places pyroclastic (Nironen, 2003; observations by the author). In the southwestern parts of the complex, the supracrustal formation at Parkano consists of mafic to felsic volcanic rocks intruded by dikes and granitoids (Vaasjoki and Lahti, 1991). A felsic, possibly volcanic rock has been dated at 1907 ± 13 Ma (Table 8.1) but, because the zircons seem to be in part inherited, the significance of this age is equivocal (Nironen, 2003). A granite with a U-Pb zircon age of 1893 ± 6 Ma is cut by a 1883 ± 15 Ma plagioclase porphyritic dike, which represents synkinematic magmatism of the Central Finland granitoid complex (Nironen, 2003). A 1872 ± 12 Ma rhyolitic dike cross-cuts the bedding of the mafic volcanic rocks and is possibly related to the postkinematic granitoids of the Central Finland granitoid complex. Farther to the east, the volcanogenic belt at Kuru (Figure 8.1) consists of intermediate to felsic rocks (Tiainen and Kähkönen, 1994). The intermediate rocks are, in general, homogeneous; in places stratified and fragmental structures indicate fallout and pyroclastic flow deposits. The felsic rocks are mainly massive and porphyritic, possibly of subvolcanic origin, but they also include fine-grained stratified tuffs; one of the latter has been dated at 1897 ± 2 Ma (Table 8.1). The abundance of andesites to rhyolites, Ti vs. Zr variation, and the considerable enrichment in LREE [(La/Yb)N ranges from 5 to 11, diagrams not shown] indicate that the volcanic rocks of this belt resemble the arc-type volcanic units of the Tampere belt and were probably formed in the same evolved arc system. The Hirsilä belt north of the Tampere belt (Figure 8.1) is mainly composed of mica gneisses, veined gneisses, amphibolites, and felsic gneisses. Blastoclastic textures, amygdules, and pumice have been sporadically preserved (Lahtinen, 1996). In general, the Hirsilä volcanic rocks show pronounced Nb depletion and clear positive La, P, and Sm anomalies
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indicating enriched mantle sources with a subduction component (figures not shown). Thus they resemble the volcanic rocks of the Tampere belt and were probably emplaced in the same evolved arc setting (Lahtinen, 1996). The sedimentary rocks at Hirsilä differ from the Myllyniemi-type turbidites of the Tampere belt and resemble those higher in the stratigraphic succession; they were derived from ~1.905–1.89 Ga arc-type sources. The Hirsilä belt is a close analogue to the Tampere belt in its original character and tectonic setting.
8. Häme belt and Saimaa area 8.1. General The Häme belt and the Saimaa area (Figure 8.1) are treated together because their mutual boundary east–northeast of Lahti has not been delineated. Bedding planes, banding, and dominant schistosity strike east–northeast to east, except at and around Virtasalmi where strikes are mainly north–northwesterly. The sedimentary rocks are characterized by mudrocks rather than graywackes (Korsman et al., 1997; Nironen et al., 2002). Black shales are rare in the Häme belt and in the bulk of the Saimaa area, but relatively abundant in the Virtasalmi region and in places in the southeastern Saimaa area (Arkimaa et al., 2000; Pekkarinen, 2002). Quartz arenites, sedimentary carbonates (relatively common at Virtasalmi, see p. 360), as well as Fe-sulfide and Fe-oxide formations are scarce. The Häme belt is essentially composed of volcanic rocks emplaced in a volcanic arc (Hakkarainen, 1994; Lahtinen, 1996), whereas in the Saimaa area these are less common and show variable tectonomagmatic affinities. The supracrustal rocks of the Häme belt and Saimaa area were metamorphosed mainly at amphibolite facies conditions 1.88–1.86 Ga ago (Nironen, 1999; Väisänen, 2002; Väisänen et al., 2002). At 1.83–1.81 Ga, they largely experienced a high-T event as the late SvecoCHAPTER
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fennian granite-migmatite zone of southern Finland was formed. Väisänen et al. (2002) found this event to have peaked at 1824 ± 5 Ma. The youngest detrital zircons in the sillimanite-bearing Tiirismaa quartz arenite near Lahti (Figures 8.1 and 8.3) indicate deposition at ~1.87–1.86 Ga at the earliest, i.e., after or during the 1.87–1.86 Ga metamorphic peak. Thus these arenites do not belong to the typical supracrustal association of the Häme belt that was deposited before the peak of metamorphism. Korsman et al. (1997) showed two other small quartz arenite occurrences in the Häme belt and Saimaa area. A fourth quartz arenite is found at Hyvinkää in the northern Uusimaa belt (Figures 8.1 and 8.3) and it resembles the Tiirismaa arenite in the age pattern of detrital zircon (see also Lahtinen et al., 2002). Muddy counterparts of these relatively young arenites may be common in southern Finland and could provide an explanation for the high relative abundance of pelitic rocks in southern Svecofennia. Lahtinen (1994, 1996) and Nironen (1997) suggested that the subduction related to the ~1.89 Ga arc magmatism in the Häme belt was directed to the south (see also Chapter 11). Instead, Väisänen and Mänttäri (2002) considered it probable that the subduction was directed to the north. The concept of Sdirected subduction is supported, though not proven, by a south-dipping mantle reflector found beneath the western extensions of the Häme belt (Chapter 11). See Section 10.3 for further discussion.
8.2.Volcanic rocks of the Häme belt The Häme belt contains the largest volcanicdominated area in the Svecofennian domain in Finland (Figure 8.1). Hakkarainen (1994) studied the belt around Hämeenlinna and Forssa and identified several early synclines and anticlines with approximately ENE- to E-trending vertical axial planes and subhori-
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B
A
Fig. 8.18. Structures of volcanic rocks of the Häme belt. (A) Tuff breccia, Forssa Group, Forssa; (B) Coherent to autobrecciated lava, Häme Group, Lautaporras, Tammela. The pen is 13 cm long. Photos: Yrjö Kähkönen.
zontal fold axes. The intensity of subsequent deformation varies, but folds with NE-trending subvertical axial planes are the most prominent feature. The Häme belt includes lavas and pyroclastic rocks (Figure 8.18), which range from basalts to rhyolites, are mainly medium-K type, and show arc affinity (Figure 8.19). Hakkarainen (1994) divided the volcanic rocks into two units, the (older) Forssa Group of separate stratovolcanoes and the (younger) Häme Group related to a linear E-trending fissure system. An age of 1888 ± 11 Ma has been published on the volcanic rocks of the Häme belt (Table 8.1). A plagioclase porphyry from Hyvinkää, close to the boundary to the Uusimaa belt, has an age of 1880 ± 3 Ma (Table 8.1). This rock is probably cogenetic with the nearby 1880 ± 5 Ma (Patchett and Kouvo, 1986) gabbros. The ~1.88 Ga rocks at Hyvinkää are probably closer to the Häme Group than to the Forssa Group in age and setting. The Forssa Group includes both volcanic (Figure 8.18A) and sedimentary (mainly pelitic) rocks. The basement is unkown (Hakkarainen, 1994) or may consist of graywackes (Lahtinen, 1996). The volcanic rocks range from basalts to rhyolites with andesite as the most common type. The lower parts of the 376
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group contain andesitic pillow lavas and minor sedimentary carbonates. The bulk of the volcanic rocks are related to stratovolcanoes, which fed pyroclastic material and detritus into the basins between. The strata between the Forssa and Häme Groups include mudrocks, graywackes, and minor conglomerates, all with volcanic provenance. Thin Fe-oxide–chert layers and Fe-sulfide formations are also present; these may be associated with the early stages of the Häme Group fissure eruptions. The Häme Group is characterized by uralite and plagioclase porphyritic basaltic lavas, which occasionally grade into andesites. They are mainly coherent but locally autoclastic (Figure 8.18B) or pillowed (Hakkarainen, 1994). Pyroclastic interbeds are present but mudrocks or other sedimentary rocks have not been found. Felsic volcanic rocks are scarce according to Hakkarainen (1994). However, based on the data of Aulis Kinnunen and the author (Figure 8.20), they seem to be relatively abundant and might indicate bimodal affinity. Basaltic feeder dikes are found throughout the Häme Group and provide evidence for E-trending fissure eruptions reflecting extension of the arc. The basalts, andesites, and rhyolites of
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O
10000
Ti ppm
M
WITHIN PLATE LAVAS
ARC LAVAS
ROCK / CHONDRITE
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58 74 10 74 72 1
Zr ppm
100
500
Fig. 8.19. Ti vs. Zr diagram of the Häme belt volcanic rocks. Data from Hakkarainen (1994), Lahtinen (1996), and from unpublished files of Aulis Kinnunen and the author. The fields of MORB (mid-ocean ridge basalts), within-plate lavas, and arc lavas are according to Pearce (1982).
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the Forssa Group are medium-K calc-alkaline rocks. In general, they have tectonomagmatic affinities indicating a relatively mature island arc. The pronounced Nb depletion typical of arc magmas and relatively steep trend from Zr to Y (Figure 8.21) indicate an enriched mantle source with a subduction component. Lavas with a less clear Nb depletion are also found and they probably represent volcanism contemporaneous with the Häme Group (Lahtinen, 1996). The volcanic rocks of the Häme Group are mainly tholeiitic medium-K basalts. In general, their Nb depletion is less pronounced than in the Forssa Group (Figure 8.21) and indicates a setting of advanced arc rifting for the Häme Group. No Nd isotope data are available for the Häme belt volcanic rocks. The 1886 ± 14 Ma granodiorite near Hämeenlinna (Table 8.1) has an εNd (at 1.89 Ga) value of ~ +2 (Patchett and Kouvo, 1986), and the flat and slightly LREEdepleted patterns in Figure 8.20 might indicate similar, relatively juvenile sources. All in all, the volcanic rocks of the Häme belt seem to represent a setting less evolved than that of
La Ce
B ROCK / CHONDRITE
1000 10
58
76 71
La Ce
Nd
64
Sm Eu
Fig. 8.20. Chondrite-normalized rare earth element (REE) patterns of volcanic rocks of the Häme belt. The samples are from the area of the Häme Group in Hakkarainen (1994). Colors indicate samples from unnamed individual units. The numbers give the SiO2 contents of the rocks. Unpublished data of Aulis Kinnunen and the author. Chondrite values from Boynton (1984).
the Tampere belt.
8.3.Volcanic rocks of the Saimaa area In the Saimaa area, volcanic rocks, in addition to those in the Virtasalmi region, are found particularly at Rantasalmi, Parikkala, and Punkaharju (Figure 8.1). Relatively wide volcanic fields are also found in Heinola, ~40 km northeast of Lahti, and in the southern Saimaa area. The 1906 Ma rhyolite from Joroinen (Table 8.1) is the only dated volcanic rock in or close to the area. The volcanic rocks at Rantasalmi are mafic to ultramafic in composition and pillow
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A
100
B 100 Andesite ROCK / MORB
ROCK / MORB
Basalt 10 Basaltic andesite 1
Basalt
0.1
10
1
0.1 Sr K Rb Ba Th Ta Nb Ce P Zr
Sm Ti Y
Sc Cr
10
10
ROCK / MORB
D 100
ROCK / MORB
C 100
Basalts 1
Sr K Rb Ba Th Ta Nb Ce P Zr
Sm Ti Y
Sc Cr
Sr K Rb Ba Th Ta Nb Ce P Zr
Sm Ti Y
Sc Cr
1
0.1
0.1 Sr K Rb Ba Th Ta Nb Ce P Zr
Sm Ti Y
Sc Cr
Fig. 8.21. Mid-ocean ridge basalt-normalized trace element patterns for representative volcanic rocks of the Häme belt. (A) Basalt to andesite, Forssa Group; (B) Andesites, Forssa Group; (C) Häme Group -type basalts and basaltic andesites from the area of the Forssa Group; (D) Basalts, Häme Group. Data from Lahtinen (1996). Normalizing values are from Pearce (1982).
lavas are common (Kousa, 1985; Viluksela, 1988). The Ti vs. Zr, Ti vs. Cr, Ti vs. V, Ti/Zr/ Y, and Ti/Cr vs. Ni relations as well as slight enrichments in LREE indicate similarity with transitional MORB (Viluksela, 1988). This is supported by the relatively low K2O and P2O5 contents of the mafic rocks (0.25 and 0.12 wt.% on the average, respectively). Lahtinen and Huhma (1997) referred to picrites at Rantasalmi that have average εNd (at 1.9 Ga) of +3.3 (Figure 8.2), indicating a depleted mantle source. The volcanic rocks at Parikkala and Punkaharju differ drastically from those at Rantasalmi. They range from basalts to rhyolites with dacites as the most common type; minor trachytic and trachyandesitic rocks also occur (Viluksela, 1988, 1994). The volcanic rocks are mostly of medium-K or high-K type and 378
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calc-alkaline rather than tholeiitic. In chondrite-normalized diagrams, they show fair LREE enrichments. Relative depletions in Ta and Nb indicate arc-type affinity (figures not shown). Viluksela (1994) suggested that the volcanic rocks at Parikkala and Punkaharju were emplaced in a mature arc setting similar to recent active continental margins.
8.4. Sedimentary rocks of the Saimaa area The sedimentary rocks of the Saimaa area are migmatized and mainly represent mudrocks and graywackes. Some sedimentary rocks in the Haukivuori–Virtasalmi–Rantasalmi district (Figure 8.1) are relatively well preserved, whereas those in the south have been extensively migmatitized at 1.83–1.81 Ga (Figure
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B
A
Fig. 8.22. (A) Migmatitic gneiss from the Saimaa area, Ristiina; (B) Haukivuori conglomerate with a 1885 Ma granitoid clast. The compass is ~12 cm long. Photos: Jukka Kousa.
8.22). The studied sedimentary rocks of the Saimaa area show variable characteristics from immature to mature types (Lahtinen, 2000) and also vary in the age distribution of detrital zircon (Figure 8.3). Rocks of the immature group envelope the Virtasalmi volcanic field and, at Haukivuori, have thin interbeds of MORB-like pillow lavas. They are mainly graywackes with chemical index of alteration (CIA; Nesbitt and Young, 1982) values less than 57 (Lahtinen, 2000). The youngest concordant zircons at Joroinen and a 1885 Ma granitoid clast in a conglomerate interbedded with the immature graywackes at Haukivuori indicate a maximum deposition age of ~1.89 Ga (Figures 8.3, 8.22B; Table 8.2). Thus these graywackes were deposited after the Upper Kaleva psammites and lower sedimentary group of central Svecofennia. Compared to the Upper Kaleva graywackes, they are typically enriched in the LREE, Rb, Ba, Th, and U. Some of the rocks have high Cr/Sc ratios indicative of mafic to ultramafic components in the source. Evidently, these sedimentary rocks were largely derived from island arc or active continental margin rocks and in part also from ultramafic rocks (Lahtinen, 2000). The graywacke from Joroinen has a significant Archean component, whereas the Haukivuori graywacke is lower in Archean detritus (Figures 8.3G, 8.4). CHAPTER
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The immature sedimentary rocks at Ristiina, 60–70 km south of Virtasalmi, have a maximum deposition age of ~1.92–1.91 Ga and an εNd (at 1.9 Ga) of ~ –2. They resemble the lower sedimentary group of central Svecofennia in these respects (Figures 8.3, 8.4). Sedimentary rocks of the mature group are most abundant at Juva but they also occur east and southeast of the Virtasalmi volcanic field (Lahtinen, 2000). In spite of pronounced migmatization, graywackes and mudrocks can in places be distinguished. The rocks have CIA values of 62 to 65 and thus they differ from the immature group discussed above as well as from the Upper Kaleva psammites and pelites. They have characteristically high Zn and low Co, and Lahtinen (2000) suggested that their sources were dominated by alkaline within-plate granitoids. Moreover, high Cr and Cr/Sc in some of the mature rocks indicate significant ultramafic components. Considering the age distribution of detrital zircon and the low proportion of Archean grains in a mature psammite (Figure 8.3I), the bulk of the sources were probably 2.1–2.0 Ga old. The psammites were deposited at ~1.89 Ga at the earliest.
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9. Uusimaa belt 9.1. General aspects The E-trending Uusimaa belt consists mainly of mudrocks, graywackes, volcanic rocks, felsic supracrustal rocks, and sedimentary carbonates. The belt also contains massive sulfide deposits and banded iron-formations. Black shales have only been reported from the Salittu Formation of the Orijärvi area. Quartz arenites are rare, the only example is at Hyvinkää and resembles the Tiirismaa arenite (Figures 8.1, 8.3, 8.4). The abundance of felsic supracrustal rocks, sedimentary carbonates, and massive sulfide deposits distinguish the bulk of the Uusimaa belt from most other Svecofennian supracrustal belts in Finland. Furthermore, Pb in the sulfide deposits of the Uusimaa belt tends to be less primitive than in central Svecofennia and the Savo belt (Vaasjoki, 1981). In the Uusimaa belt, the early tectonometamorphic evolution peaked at 1.87–1.86 Ga, whereas the subsequent major metamorphic event related to the late Svecofennian granitemigmatite zone of southern Finland peaked at ~1825 Ma (Pajunen et al., 2002; Väisänen and Mänttäri, 2002; Väisänen, 2002). The metamorphic conditions and degree of deformation varied widely. For instance, the Orijärvi area (Figure 8.23) with well-preserved primary structures (Figure 8.24) has been a type locality of amphibolite facies rocks since the study of Eskola (1914), but some 15 km northeast of Orijärvi the amphibolite facies rocks change abruptly to low-P granulites (Schreurs and Westra, 1986; Korsman et al., 1997). The 1.83–1.81 Ga event has also largely obscured features of the preceding 1.87–1.86 Ga orogenic events; for instance, in the otherwise well-preserved Orijärvi area a ~1878 Ma dacite (Table 8.1) has concordant 1797 ± 3 Ma titanite (Väisänen et al., 2002). On the other hand, titanites from migmatites on the southern coast of Finland yield ages of 380
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1.88–1.86 Ga (Hopgood et al., 1983) indicating that here the 1.83–1.81 Ga event did not exceed the blocking temperature of titanite. The volcanic rocks of the Uusimaa belt range from picrites to rhyolites and show EMORB or WPL to arc affinities (Figures 8.25 through 8.27). Their U-Pb zircon ages range from 1900–1895 Ma to ~1878 Ma (Table 8.1). Three parts of the Uusimaa belt are discussed here: (1) the Kemiö–Järvenpää field in the center is the widest and shows the typical lithological association of the Uusimaa belt; (2) the Nauvo–Korppoo field in the west has fewer sedimentary carbonates and its mafic volcanic rocks are like WPL; and (3) the Pellinki field in the southeast is also relatively poor in sedimentary carbonates but rich in arc-type basalts and andesites.
9.2. Kemiö–Järvenpää field The Kemiö–Järvenpää field is characterized by felsic supracrustal rocks, which largely represent volcanic and volcanogenic sedimentary deposits, and sedimentary carbonates (Figure 8.1; Kähkönen, 1998; Reinikainen, 2001). Massive sulfide deposits, related alteration products, and minor banded iron-formations are abundant particularly in the west (Latvalahti, 1979; Mäkelä, 1989). Picritic to intermediate lavas (in part with pillow structures) and pyroclastic rocks as well as intermediate subvolcanic intrusions are also relatively abundant (Latvalahti, 1979; Schreurs et al., 1986; Colley and Westra, 1987; Mäkelä, 1989; Väisänen and Mänttäri, 2002). Among the volcanogenic rocks, felsic to intermediate types predominate. The volcanic rocks are mostly subalkaline and include both calc-alkaline and tholeiitic associations. The frequency distribution of silica is bimodal in the vicinity of massive sulfide deposits but, regionally, this feature is less conspicious (Mäkelä, 1989). The chondrite-normalized REE patterns of the volcanic rocks show mostly moderate enrich-
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ment of LREE (Figure 8.26; see also Mäkelä, 1989), and relative depletions in Nb and Ta are pronounced to absent (Figure 8.27). Stratigraphic concepts of the Kemiö–Järvenpää field are mainly based on the well-preserved rocks of the Orijärvi triangle (Ploegsma and Westra, 1990; Figure 8.23). The views of Latvalahti (1979), Colley and Westra (1987), and Mäkelä (1989) show some differences and are not discussed here. Väisänen and Mänttäri (2002), supported by U-Pb zircon ages, distinguished three major units: (1) the ~1895 Ma Orijärvi Formation, (2) the ~1878 Ma Kisko Formation, and (3) the Salittu Formation (youngest). Outside the Orijärvi triangle, a fourth unit, the Toija Formation, probably underlies the Salittu Formation. The volcanic formations are separated by sedimentary units and truncated by a shear zone. The Orijärvi Formation comprises a bimodal association of mainly medium-K basalts and medium-K to high-K dacites and rhyolites interbedded with sedimentary carbonates, iron formations, and intensively altered supracrustal rocks hosting Cu-Zn-Pb deposits. Pillow lavas (Figure 8.24A) and sedimentary intercalations indicate subaqueous eruptions. The 1.90–1.89 Ga Orijärvi granodiorite (Table 8.1) is cogenetic with this unit. The Orijärvi Formation is separated from the Kisko Formation by a unit of graywackes and mudrocks. The graywackes are relatively rich in ~2.1–2.0 Ga detrital zircon, whereas younger and Archean grains are not abundant (Figure 8.3F; see also Claesson et al., 1993). The youngest zircons are ~1.93–1.92 Ga and thus older than the ~1895 Ma Orijärvi Formation beneath. Väisänen and Mänttäri (2002) explained the absence of 1.90–1.89 Ga zircons in the graywackes by the subaqueous character of the Orijärvi Formation. The scarcity of Archean grains is in agreement with the positive εNd (at ~1.9 Ga) value of ~ +1 (Figure 8.4C). The volcanic rocks of the Kisko Formation range from medium-K/high-K basalts to CHAPTER
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rhyolites with basaltic andesites and andesites as the most common rock types. The upper part of the formation also includes a picritic interlayer with felsic fragments. According to Väisänen and Mänttäri (2002), the Salittu Formation is younger than the Kisko Formation, and these units are separated by mica schists (mudrocks and graywackes) with occasional andalusite porphyroblasts. The Salittu Formation consists mainly of picritic (>12 wt.% MgO) and basaltic lavas, which are in places brecciated or pillowed and contain interbeds of graphite-bearing mica gneisses (black shales) and minor sedimentary carbonates. It mainly lies outside the Orijärvi triangle and differs from the Orijärvi and Kisko Formations because migmatites are common and gently dipping early structures are present, as opposed to the mainly upright structures of the latter two units. The Toija Formation west of the Kisko shear zone lies outside the Orijärvi triangle (Figure 8.23). It resembles the Salittu Formation in structural style and apparently underlies it. The Toija Formation includes mafic pillow lavas and rhyolites with intercalations of marbles and mica schists. A picritic volcanic horizon is found close to the contact with the Salittu Formation. The basalts, basaltic andesites, and picrite of the Orijärvi and Kisko Formations plot into the arc field in the Ti vs. Zr diagram (Figure 8.25). The LREE enrichments are moderate, LILE/HFSE ratios high, and Nb depletions pronounced (Figures 8.26, 8.27; see also Väisänen and Mänttäri, 2002). The Nb, Zr, and Sm contents in the least evolved rocks of the Orijärvi Formation are below those in average MORB and indicate a depleted rather than enriched mantle source. The felsic and intermediate rocks in these units also have subduction-related characteristics (Figure 8.25; see also Väisänen and Mänttäri, 2002). The slightly negative εNd (at 1.9 Ga) value of the Orijärvi granodiorite, which is probably comagmatic with the Orijärvi Formation, sug-
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ed Sw
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Turku
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Rapakivi granite
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sko
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u Fm
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sk o
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Granodiorite
Mafic–intermediate volcanic rock
Marble
Altered rock
Gabbro
Felsic volcanic rock
Picrite and basalt
Metapelite
Granite
Iron formation
Road
Shear zone (SZ)
Fig. 8.23. Lithological map of the Orijärvi area. B−Bergslagen, CFGC−Central Finland granitoid complex, CS−central Svecofennia, K−Kemiö, NK−Nauvo−Korppoo, OM−Orijärvi mine, SS−southern Svecofennia. Slightly modified from Väisänen and Mänttäri (2002).
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Fig. 8.24. Structures of supracrustal rocks of the Uusimaa belt. (A) Pillow lava, Orijärvi Formation, Hyypiänmäki, Kisko. The pen is 13 cm long; (B) Tuff breccia, Orijärvi Formation, Multsilta, Kisko. The compass is 6 cm wide; (C) Dacitic volcanic breccia, locality of the dated sample of the Kisko Formation, Kisko. The pen is 12 cm long; (D) Picritic volcanic breccia, Salittu Formation, Kisko. The pen is 12 cm long; (E) Pillow lava, Toija Formation, Kisko. The scale bar is 10 cm long; (F) Erosional scour filled by volcaniclastic material, subvertical section looking north, Pellinki Group, Suur-Pellinki, Porvoo. The scale bar is 12 cm long. Photos: Yrjö Kähkönen (A, B, and E), Markku Väisänen (C, and D), and Matti Laitala (F).
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RB
O
Ti ppm
M
Ti ppm
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ARC LAVAS
1000 10
100
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Fig. 8.25. Ti vs. Zr diagram of representative Uusimaa belt volcanic rocks. Data from Ehlers et al. (1986), Lindroos and Ehlers (1994), and Väisänen and Mänttäri (2002). (A) Orijärvi area; (B) Vestlax (Kemiö) and Nauvo–Korppoo (mafic middle unit). The fields of MORB (mid-ocean ridge basalts), within-plate lavas, and arc lavas are according to Pearce (1982). One felsic sample from Kemiö has 960 ppm Ti and 250 ppm Zr. Two felsic samples from the Toija Formation have Ti below 1000 ppm and 110–180 ppm Zr.
gests that the source of the felsic rocks of this formation had a significant contribution from older Proterozoic crust (Figure 8.2). In the basaltic and picritic rocks of the Salittu Formation, the LILE elements, particularly Ba, are variable (at least in part due to alteration) but, in general, enrichments in LILE and LREE are moderate to minor. Depletions in Nb relative to Th are mostly absent. Overall, the Salittu Formation volcanic rocks resemble T- or EMORB rather than WPL. However, the εNd (at 1.9 Ga) value of ~ +3 of the Salittu picrite (Figure 8.2) indicates a depleted mantle source. The stratigraphically high picrite in the Toija Formation resembles those of the Salittu Formation (Figure 8.27). The mafic rocks at Toija show depletion in Nb and are like the arc-type rocks of the Kisko and Orijärvi Formations. These similarities support the view of Väisänen and Mänttäri (2002) that the Salittu Formation overlies the Kisko Formation and that the Toija Formation represents a transition from the Orijärvi and Kisko Formations to the 384
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Salittu Formation. The most mafic rocks of the Orijärvi Formation tend to be lower in Zr, Hf, P, LREE, and LILE, and higher in Cr and Ni than the mafic volcanic rocks in the Kisko Formation. The most primitive lavas are found in the lower parts of the Orijärvi Formation. Thus evolution from a less mature arc (lowermost parts of the Orijärvi Formation) to a more mature arc (Kisko Formation) is obvious. The mafic rocks of the Orijärvi Formation resemble the classic island arc tholeiites of Jakeš and Gill (1970) in their low Zr contents but are distinct due to their LREE enrichment. The abundance of felsic volcanic rocks and the initial εNd value of ~ –1 of the Orijärvi granodiorite (Figure 8.3) further indicate that the Orijärvi Formation was not emplaced in a true primitive oceanic arc. Väisänen and Mänttäri (2002) suggested that the T- or EMORB-like rocks of the Salittu Formation were emplaced in a backarc or intra-arc basin during arc extension and that the Toija Formation represents an
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initial stage of the opening of that basin. The clearly positive εNd (T) value in the Salittu Formation contrasts with the negative εNd (T) of the Orijärvi granodiorite and could imply marked extension. On the other hand, according to my view, the strong contrast in εNd (T) values, structural style, and metamorphic grade between the Salittu Formation and the Orijärvi and Kisko Formations suggest that the stratigraphic scheme may be more complex than currently perceived. Supracrustal rocks at Vestlax, Kemiö, approximately halfway between Orijärvi and Nauvo–Korppoo (Figure 8.1), display a mushroom-shaped fold interference pattern with subvertical E-striking axial plane of the second fold event (Lindroos and Ehlers, 1994). The interpreted stratigraphic sequence of the supracrustal rocks begins with felsic ignimbrites. These are covered by pillow-basalts, mafic volcanic breccias, and fine-grained felsic volcaniclastic rocks. The Vestlax mafic volcanic rocks fall in all three fields in the Ti vs. Zr diagram, have flat to slightly LREE enriched REE patterns, and show relative depletions in Ta (Figures 8.25 through 8.27). Thus they differ from the mafic volcanic rocks of the Orijärvi area and the Nauvo–Korppoo field (see below).
9.3. Nauvo–Korppoo field In the westernmost part of the Uusimaa belt, the volcanic rocks at Nauvo–Korppoo (Figure 8.1) comprise a <1-km-thick formation enveloped by turbidites. Ehlers et al. (1986) divided the volcanic rocks into three units: (1) subvolcanic gabbro sills, (2) pillowed mafic lavas and pyroclastic volcanic rocks, and (3) ultramafic to mafic lavas. A thin horizon of sedimentary carbonates is found between the second unit and the enveloping cordierite- and garnet-bearing mica gneisses. The ultramafic to mafic lavas are found as an interlayer in the gneisses. The mafic volcanic rocks at Nauvo– Korppoo are within-plate basalts (Figures 8.25 CHAPTER
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through 8.27). The Ti/Zr/Y diagram (figure not shown) suggests that the ultramafic rocks also are of WPL type. Ehlers et al. (1986) suggested that the mafic to ultramafic volcanic rocks represent a stage of rift-related volcanism that preceded the ~1.9 Ga Svecofennian arc volcanism. Accordingly, the setting of the Nauvo–Korppoo volcanic units seems to differ from that of the Salittu Formation. Assuming that the stratigraphic concepts of Ehlers et al. (1986) and Väisänen and Mänttäri (2002) are both correct, ultramafic to mafic volcanic rocks are found at two stratigraphic levels in the Uusimaa belt.
9.4. Pellinki field The supracrustal rocks at Pellinki in the southeastern part of the Uusimaa belt (Figure 8.1) show several major E-trending syncline-anticline pairs with subvertical axial planes and subhorizontal fold axes (Laitala, 1972). These are curved due to subsequent deformation mainly in ENE-trending zones. The succession at Pellinki includes more mafic volcanic rocks than the Uusimaa belt in general and has only minor sedimentary carbonates. It is ~7 km thick and is here divided into (1) a stratigraphically lower unit of mainly felsic schists and gneisses in the north (Tirmo Group) and (2) an upper unit of mafic to intermediate volcanic rocks in the south (Pellinki Group). In the Tirmo Group, the felsic schists and gneisses are mainly sedimentary rocks and have intercalations of silty and clayey mudrocks, intermediate to mafic volcanic and volcaniclastic rocks, conglomerates, and rare sedimentary carbonates (Laitala, 1972; Korsman et al., 1997). The felsic rocks grade from generally stratified and fine-grained types in the north to coarser-grained, poorly stratified types in the south. The volcanic rocks of the Pellinki Group show well-preserved primary structures (Figure 8.24F). The lower part of the unit consists of subaqueous rocks with pillow
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B Basaltic andesite Basalt
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Yb Lu Rhyolite Dacite Andesites Basalts and basaltic andesites Picrite
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Nauvo–Korppoo Mafic volcanic rocks
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La Ce
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Sm Eu
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Yb Lu
La Ce
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Sm Eu
Fig. 8.26. Chondrite-normalized rare earth element (REE) patterns of representative volcanic rocks of the Uusimaa belt. (A, B) Orijärvi Formation; (C) Kisko Formation; (D) Salittu Formation; (E) Vestlax (Kemiö) and Nauvo–Korppoo (mafic middle unit); (F) Pellinki area. Data from Koljonen and Rosenberg (1975), Ehlers et al. (1986), Lindroos and Ehlers (1994), and Väisänen and Mänttäri (2002). Chondrite values from Boynton (1984).
lavas, pyroclastic deposits, and associated volcaniclastic sedimentary rocks. These are cut by subvolcanic mafic to dacitic intrusions and sills, which are in part cryptodomes with peperitic contacts (Strandman and Fröjdö, 2002a,b). They are overlain by extensive, 386
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mainly basaltic to andesitic massive lavas. A pooled U-Pb zircon date of 1887 ± 14 Ma (Table 8.1) yields an approximate age for the Pellinki Group. The volcanic rocks of the Pellinki Group show a continuous trend from basalts to da-
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Kisko Fm. picrite, basalt, basaltic andesites ROCK / MORB
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C
0.1 Sr K Rb Ba Th
Sr K Rb Ba Th Ta Nb Ce P Zr Hf Sm Ti Y Yb Sc Cr
NbCe P Zr Hf Sm Ti Y Yb Sc Cr
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Salittu Fm. mafic volcanic rocks and an andesite
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Salittu Fm. picrites 10
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10 Andesite
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0.1 Sr K Rb Ba Th Ta NbCe P Zr Hf Sm Ti Y Yb Sc Cr
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Sr K Rb Ba Th Ta NbCe P Zr Hf Sm Ti Y Yb Sc Cr
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Toija Fm. 10
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Vestlax, Kemiö Mafic rocks Picrite
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1 Nauvo–Korppoo
0.1
0.1 Sr K Rb Ba Th Ta NbCe P Zr Hf Sm Ti Y Yb Sc Cr
Sr K Rb Ba Th Ta NbCe P Zr Hf Sm Ti Y Yb Sc Cr
Fig. 8.27. Mid-ocean ridge basalt-normalized trace element patterns of representative volcanic rocks of the Uusimaa belt. (A) Least evolved samples, Orijärvi Formation; the arrow indicates some samples had Nb contents below detection limit; (B) A picrite, basalt and basaltic andesites, Kisko Formation; (C) Picrites, Salittu Formation; (D) Mafic volcanic rocks (basalts, basaltic andesites) and an andesite, Salittu Formation; (E) A picrite and two mafic volcanic rocks, Toija Formation; (F) Mafic volcanic rocks from Vestlax (Kemiö) and Nauvo–Korppoo (middle unit). Data from Ehlers et al. (1986), Lindroos and Ehlers (1994), and Väisänen and Mänttäri (2002). Normalizing values are from Pearce (1982).
cites, are mostly medium-K rocks, and exhibit moderate enrichment of LREE (Figure 8.26; see also Laitala, 1972; Koljonen and Rosenberg, 1975; Strandman and Fröjdö, 2002a, CHAPTER
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b). The low Ti contents and, in particular, clear depletions in Ta and Nb indicate an arc setting (Strandman and Fröjdö, 2002a, b). The slightly to moderately positive initial εNd
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(at 1.9 Ga) values (Figure 8.2) are relatively high, but the variation in these values may imply heterogeneous sources. In general, the volcanic rocks of the Pellinki Group resemble those of evolved arcs rather than those of true primitive arcs.
9.5. Sedimentary carbonates of the Uusimaa belt Sedimentary carbonates of the Uusimaa belt are commonly associated with volcanic rocks and locally also with chemical precipitates (Ehlers et al., 1986; Lindroos, 1990; Reinikainen, 2001). Lindroos (1990) suggested that the sedimentary carbonates and the associated cherts and iron formations represent a marker horizon between the (older) MORB- or WPLtype mafic lavas and the (younger) arc-type volcanic rocks. In the Orijärvi area, however, sedimentary carbonates are found at several stratigraphic levels (Mäkelä, 1989; Väisänen and Mänttäri, 2002). Sedimentary carbonates in the Bergslagen field in south-central Sweden, a lateral extension of the Uusimaa belt (see section 10.2), have also been deposited at different periods (Allen et al., 1996a). Some of these show stromatolite structures (Lundqvist, 1979, p. 27) and are thus shallow-water deposits. In the Uusimaa belt the sedimentary carbonates were in part deposited in relatively deep water (Reinikainen, 2001).
10. Discussion 10.1. Correlation of the Pohjanmaa belt to northern Sweden In northern Sweden, the Skellefte field with abundant ~1.89–1.88 Ga arc-type volcanic rocks (volcanic arc) and the area of turbidite-derived migmatitic gneisses to the south (subduction zone complex) probably belong to lateral extensions of the arc system that in Finland is represented by the Pohjanmaa, 388
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Tampere, and Pirkanmaa belts. Volcanism in the Skellefte field evidently took place in an extensional continental margin arc rather than in an oceanic island arc (Allen et al., 1996b). Subduction was directed towards the present northeast, and the arc system at Skellefte was thrust onto an Archean craton in the north (BABEL Working Group, 1990). The supracrustal rocks of the Skellefte field resemble those of the Ylivieska field in some respects but there are also differences. Both are characterized by well-preserved arctype volcanic and related sedimentary rocks. The volcanic and synvolcanic plutonic rocks of the Skellefte field show a range in age from ~1890 Ma to ~1875 Ma (Wilson et al., 1987; Billström and Weihed, 1996) but Kousa and Lundqvist (2000), referring to a 1907 ± 12 Ma tonalite intrusive into the Skellefte supracrustal rocks, suggested that the Skellefte volcanism may have commenced before that time. The typically submarine rhyolitic and dacitic volcanic rocks of the Skellefte field are overlain by sedimentary rocks, mafic lavas, and felsic subaerial ~1876 ± 3 Ma (Skiöld et al., 1993) volcanic rocks of the Arvidsjaur field. The subduction-related volcanic activity in the Skellefte field thus seems to have lasted longer than in the Ylivieska field and Tampere belt. The volcanic rocks of the Skellefte field are, although locally rich in andesites, characterized by calc-alkaline or subalkaline submarine rhyolites and dacites (Claesson, 1985; Vivallo and Claesson, 1987; Allen et al., 1996b). Instead, the Ylivieska field tends to be more abundant in intermediate and mafic rocks and generally displays features indicative of a shallow-water or subaerial environment (Kousa et al., 1994; Kousa and Lundqvist, 2000; Koistinen et al., 2001). Furthermore, counterparts of the mafic volcanic rocks overlying the typical felsic volcanic rocks of the Skellefte field and rocks similar to the felsic volcanic rocks of the Arvidsjaur field have not been reported in the Ylivieska field.
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In the Skellefte field, εNd (at 1.88 Ga) values of the 1.89–1.87 Ga arc-related volcanic and plutonic rocks range mostly from ~ +1 to ~ +3 (Wilson et al., 1985, 1987; Billström and Weihed, 1996). No εNd values have been published on the volcanic rocks of the Ylivieska field but the possibly arc-related 1.90–1.88 Ga plutonic rocks in the northwestern part of the Pohjanmaa belt also have relatively high εNd (T) values from ~ +2.5 to ~ +4 (Huhma, 1986; Patchett and Kouvo, 1986; Lahtinen and Huhma, 1997). It will be interesting to see if the volcanic rocks of the Ylivieska field in general have similar values or if they resemble more rocks of the Central Finland granitoid complex – the latter have εNd (at 1.88 Ga) close to zero. At Luleå, some 100 km north of Skellefte, subalkaline andesites to rhyolites with arc affinity differ from the typical volcanic rocks of the Skellefte field (Perdahl, 1995). Based on the age of intrusive granitoids they seem to be older than 1.90–1.89 Ga. The εNd values of the felsic rocks are clearly negative and indicate substantial involvement of Archean crust (Öhlander et al., 1993). The Luleå volcanic rocks were emplaced in an active continental margin arc and closer to the Archean craton than those of the Skellefte field. Similar rocks are unknown in the Pohjanmaa belt. The Skellefte field is rich in submarine felsic volcanic rocks comprising domes or cryptodomes, and its VHMS deposits occur almost entirely in below-wave base facies associations (Allen et al., 1996b). Submarine volcanism with abundant rhyolitic domes and cryptodomes seems to have been more common in the Skellefte field than in the Ylivieska field. Mafic volcanism, similar to that overlying the felsic volcanic series in the Skellefte field, is not known in the Ylivieska field. These features may indicate less pronounced extension in the Ylivieska arc and could possibly explain the difference in the amount of VHMS deposits in the two fields. In the Skellefte field about 15 deposits have been mined (Nordin et CHAPTER
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al., 1997), whereas no economic VHMS ore deposits are known in the Ylivieska field. The Bothnian basin south of the Skellefte field is dominated by gneissose graywackes to mudrocks. The northern part of the basin, here called the Umeå field, is rich in black shales. The Umeå and Evijärvi fields are probably lateral counterparts, because both are characterized by gneissose turbidites with relatively abundant black shales and belong to the zone of crustal conductors that curves from the Pirkanmaa belt in southern Finland via the Evijärvi field to the Umeå field (Korja, 1993; Korja, 1995; Korja and Hjelt, 1993; Korja et al., 1993; Eriksson and Henkel, 1993; Arkimaa et al., 2000; Rutland et al., 2001). They also contain mafic volcanic rocks claimed to have MORB-like geochemical features. The Umeå field includes mafic volcanic rocks, graywackes, black shales, and cherts intruded by ~1.95 Ga granitoids (Wasström, 1990, 1993, 1996). These volcaniclastic and pillow basalts have been suggested to show evolution from a of mid-ocean ridge to an island-arc setting (ibid.), but in the absence high-quality data on key immobile trace elements and Nd isotopes, I consider these interpretations provisional. In any case, the pillow basalts have low contents of K2O and P2O5 (0.11–0.19 wt.% and 0.05–0.11 wt.%, respectively) suggesting an immature system. These rocks might resemble the NMORB-like pillow basalts of the Evijärvi field but 1.95 Ga granitoids are not known in Pohjanmaa. The typical volcanic rocks of the Skellefte field are overlain by, in addition to mafic volcanic rocks, subaerial to shallow-water sedimentary rocks that laterally grade into finer-grained turbiditic graywackes. These graywackes and intercalated black shales and mafic volcanic rocks are equivalents of the upper part of the sedimentary rocks of the Umeå field (Weihed et al., 2002). Accordingly, a part of the black shales and graywackes of the Umeå field are younger than the bulk of the volcanic rocks of the Skellefte field. Thus
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the Umeå field, like the Pirkanmaa belt, may contain fore-arc sedimentary rocks derived from the volcanic arc to the north.
10.2. Correlation of the Uusimaa belt to the Bergslagen field The Bergslagen field in south-central Sweden resembles the Uusimaa belt, particularly the Kemiö–Järvenpää field, in many respects. The two areas are strikingly different from the other Svecofennian supracrustal belts and are commonly regarded as lateral counterparts. Lahtinen et al. (Chapter 11) suggest that they originally belonged to the Bergslagen microcontinent that was divided into two parts by a NW-striking shear zone. The Bergslagen field is characterized by 1.90–1.88 Ga felsic volcanic rocks and is also rich in sedimentary carbonates and massive sulfide deposits (e.g., Lundström, 1987; Allen et al., 1996a). Besides these, it locally includes units of fluvial arkoses (up to 5 km thick; Kumpulainen et al., 1996) and quartz arenites of deltaic to tidal flat environments (Gavelin and Russell, 1967; Claesson et al., 1993). The Bergslagen field was interpreted as an extensional, probably back-arc active continental margin magmatic region by Allen et al. (1996a). In general, the Bergslagen field is dominated by felsic volcanic rocks. However, mafic volcanic rocks and subvolcanic intrusions are also common in the stratigraphically upper parts (e.g., Valbracht et al., 1991; Kumpulainen et al., 1996). Lundström et al. (1998) scrutinized the reliability and stratigraphic significance of U-Pb zircon ages on volcanic rocks from the Bergslagen field and found a 1904 ± 4 Ma rhyolite from the eastern parts to mark the onset of volcanism in the area. Another reliable age from a stratigraphically low level in western Bergslagen is 1891 ± 4 Ma. The other, less reliable ages evaluated by these authors mainly cluster between 1.90 Ga and 1.88 Ga and mostly have analytical errors of ±10 to ±25 390
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Ma, consequently the duration of volcanism remains unclear. The earliest dated volcanic events in the Bergslagen field and Uusimaa belt took place at ~1.90 Ga, but it is not clear if volcanism in the Bergslagen field lasted to ~1.88 Ga as in the Uusimaa belt. The age distribution of detrital zircons in a sandstone from the Bergslagen field resembles that of the Orijärvi graywacke (Claesson et al., 1993; Lahtinen et al., 2002) and renders these sedimentary rocks unique in this respect within the Svecofennian domain. A quartz arenite in the southeastern part of the field differs from the former both in older ages of the Proterozoic population and higher proportion of Archean grains (ibid.), thus emphasizing the need for further studies. The εNd (T) values in most igneous rocks of the Uusimaa belt and Bergslagen field show largely similar features but the data from the volcanic rocks of the Uusimaa belt are scarce. In addition to the Pellinki lavas with εNd (at 1.89 Ga) values of ~ +1 to +2.5, a picrite from the Salittu Formation with an εNd value of ~ +3 is the only volcanic rock sample studied so far for Nd isotopes (Figure 8.2). The Salittu picrite, in particular, deviates from the ~1.89–1.88 Ga mafic to felsic plutonic rocks of the belt because the latter have εNd (T) values of ~0 to –0.7 (Huhma, 1986; Patchett and Kouvo, 1986). Among them, the Orijärvi granodiorite (Figure 8.2) exemplifies the excepted εNd values of the arc-type volcanic rocks and is probably cogenetic with the volcanic rocks of the Orijärvi Formation. In the Bergslagen field, εNd (T) values for ~1.89–1.88 Ga felsic volcanic rocks and subvolcanic intrusions are close to zero. The εNd (at 1.88 Ga) values in the least-altered and albitized samples vary from ~ –1 to +2.5 and are mostly between –0.5 and +1 (Valbracht et al., 1994; Kumpulainen et al., 1996). Similar variation is shown by the Bergslagen data of Patchett et al. (1987). Mafic subvolcanic intrusions have εNd (at 1.9 Ga) from ~ +1.5 to +4.5 (Kumpulainen et al., 1996), and a similar
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range (mostly from +1 to +2) is seen in the mafic volcanic rocks studied by Valbracht (1991). The εNd values in the felsic rocks indicate the presence of relatively evolved Proterozoic crust at ~2.0–2.1 Ga. The mafic rocks, which mainly occur relatively high in the succession, show evidence for an increase in more primitive components with advancing rifting. A similar increase with time, from εNd (T) values of ~ –2.5 to +1, in the sedimentary rocks in Bergslagen is probably caused by increasing proportion of young, relatively juvenile detritus (Kumpulainen et al., 1996). The values in the lowermost sedimentary rocks suggest a significant involvement of Archean crustal material. The εNd (T) values in the Orijärvi granodiorite and Salittu picrite and the stratigraphy of the Orijärvi area (Väisänen and Mänttäri, 2002) suggest similarities between the Kemiö–Järvenpää and Bergslagen fields. The Orijärvi graywacke with slightly positive εNd (at 1.9 Ga) value (Figure 8.4) resembles, in this respect, the sedimentary rocks relatively high in the succession (cf. Kumpulainen et al. 1996). There are also differences between the Uusimaa belt and the Bergslagen field. Metamorphism in the Uusimaa belt was mainly of higher grade than in Bergslagen; the latter shows very low grade metamorphic conditions in the west. In general, the volcanic rocks tend to be more mafic and contain more pillow lavas in the Uusimaa belt than in Bergslagen, and thus have a less mature arc character and indicate a deeper water environment. This concerns, in particular, the eastern parts of the Uusimaa belt. The depositional environments are also in part different. Quartz arenites and fluvial arkoses are locally thick in Bergslagen but, excluding the relatively young Tiirismaa-type quartz arenite at Hyvinkää, they have not been reported in the Uusimaa belt. Stromatolitic carbonate sedimentary rocks are known in Bergslagen (Lundqvist, 1979; Allen et al., 1996a) but not CHAPTER
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in the Uusimaa belt. This might be due to the more pronounced deformation and higher degree of metamorphism in Uusimaa, but it may also indicate deposition in deeper water in Uusimaa (Reinikainen, 2001). Note also that cross-bedding has been preserved well in some quartz arenites of Bergslagen although metamorphic cordierite and sillimanite in the associated pelitic rocks indicate a fairly high grade of metamorphism (Gavelin and Russell, 1967). Finally, isotope data of Patchett et al. (1987), Claesson et al. (1993), Kumpulainen et al. (1996), and Lahtinen et al. (2002) imply a more significant Archean component in the sedimentary rocks in Bergslagen. The differences between the two areas possibly resulted from along-arc lateral changes from a more evolved and more continental setting in Bergslagen to a less evolved setting in Uusimaa. The situation largely resembles the change from the microcontinental Taupo volcanic system of New Zealand to the more oceanic arc setting in the north (e.g., Cole, 1982).
10.3. Correlation of the Häme and Uusimaa belts Arc volcanism in the Häme and Uusimaa belts may have been approximately coeval, but data from Häme are scarce. The geochemical (including Nd isotopic) character of the volcanic and plutonic rocks of the Häme belt indicate that the ~1.90–1.88 Ga magmatism in the belt took place in an arc setting that was less evolved than that of the Tampere belt and the bulk of the Uusimaa belt. The boundary between the Häme and Uusimaa belts is difficult to demarcate with certainty because of prominent faults and the effects of the late Svecofennian granite-migmatite zone of southern Finland (see the maps of Koistinen, 1994, and Korsman et al., 1997). Moreover, it is not known for sure if the two belts were attached at 1.89 Ga. The volcanic rocks at Pellinki, southeastern Uusimaa belt, resemble those of the Häme
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belt in many respects. This would support the view that the two belts were attached and emphasizes the eastward change from a more continental setting in Bergslagen to a more oceanic setting in the east. On the other hand, the Salittu Formation at Kisko and its possible counterparts in the east might be considered as a remnant of oceanic crust between the Häme and Uusimaa belts. At the moment, there are two models with contrasting subduction directions for the evolution of the 1.90–1.88 Ga volcanism in the Häme and Uusimaa belts. Lahtinen (1994, 1996) and Lahtinen et al. (Chapter 11) suggest that the 1.89–1.88 Ga arc volcanism in the Häme belt was caused by south-directed subduction beneath Paleoproterozoic protocrust (the Bergslagen microcontinent) and the attached pre-1.92 Ga arc crust of the Häme belt. This is in line with, though not proven by, the S-dipping mantle reflector beneath the Häme and Uusimaa belts. In this scenario, the Häme belt is characterized by arc-type volcanism whereas the 1.90–1.88 Ga magmatism in the Bergslagen field and Uusimaa belt represents a back-arc setting. According to my view, this can be the case in Bergslagen but the Orijärvi area poses problems, because the abundant subduction-related andesites of the 1878 Ma Kisko Formation were emplaced in a proper arc rather than in a back arc. Väisänen and Mänttäri (2002) and Väisänen (2002) suggested subduction to the present north and emphasized that subduction-related volcanism was still active in the Uusimaa belt at ~1.88 Ga when collision was ongoing between southern and central Svecofennia. In their model, the Häme and Uusimaa belts once belonged to the same arc system that was rifted apart, the Salittu Formation represents the back-arc basin, and the Häme belt was a remnant arc at ~1.88 Ga. The model is in line with the ordinary spatial arrangement of active arcs, back-arc basins, and remnant arcs in relation to subduction zones in recent arc systems. In the modern arc systems, the volcanically 392
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active arc occurs above shallower parts of the subduction zone, the associated remnant arc is above deeper active parts of the subduction zone, and the back-arc basin lies between the two arcs. The S-dipping mantle reflector (cf. Chapter 11), however, cannot be explained by subduction to the north. Finally, the Häme belt could be speculated to have been originally exotic in relation to the Uusimaa belt because it shows a less evolved arc character than the typical parts of the Uusimaa belt and the Bergslagen field. This could be attributed to along-arc lateral variation, but the scarcity of sedimentary carbonates in the Häme belt might also indicate that the Häme belt and the Kemiö–Järvenpää field were exotic and were accreted shortly after the 1.90–1.88 Ga arc magmatism. In this model, the Salittu Formation would represent pre-1.90 Ga oceanic crust.
10.4. Tiirismaa-type quartz arenites The ~1.86 Ga detrital zircons in the Tiirismaa-type quartz arenites from southern Finland (Figure 8.3) suggest that very mature sandstones were deposited after or during the 1.88–1.86 Ga tectonometamorphic events. Similar rocks are known at other localities in southern Finland (Figure 8.1; Korsman et al., 1997), but the character and distribution of probably associated fine-grained sedimentary material is unknown. Their sporadic occurrence, the absence of Archean zircon grains, and prismatic 1.87–1.86 Ga crystals with rounded terminations suggest short transportation and deposition in small isolated basins (Lahtinen et al., 2002). In central Sweden, a low-grade quartz arenite at Los, close to the southern margin of the Bothnian basin, has a similar relatively young zircon population (Claesson et al., 1993; see also Lahtinen et al., 2002). The youngest zircons in the Los arenite were probably in part derived from nearby 1867 ± 9 Ma (Welin, 1987) felsic volcanic rocks. In Finland,
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proper volcanic rocks of this age are unkown, but the ~1.87 Ga porphyritic dikes and their possible volcanic counterparts might have been a potential source. A further possibility are the 1.87–1.86 Ga granitoids, which are relatively abundant in southern and central Finland. In central Sweden, the peak of regional metamorphism occurred at ~1.86–1.82 Ga (Billström and Weihed, 1996; Kousa and Lundqvist, 2000), and the Swedish Los-type quartz arenites may have deposited before this. In Finland, the Tiirismaa-type quartz arenites were evidently deposited after or during the 1.88–1.86 Ga tectonometamorphic events. A problem is that angular unconformities have not been identified so far. A further problem is the generally pronounced weathering implied by quartz-rich arenites. It is not easy to understand the origin of the Tiirismaa-type quartz arenites from volcanic, plutonic or metamorphic sources during the apparently short time of erosion in an active orogenic setting. A viable explanation is that the major tectonometamorphic event preceding the 1.83–1.81 Ga events in the late Svecofennian granitemigmatite zone of southern Finland took place at 1.86–1.85 Ga, not 1.87–1.86 Ga. This might be indicated by the collision-related tonalite for which Väisänen et al. (2002) presented an age of 1854 ± 18 Ma. However, this rock may also have been emplaced at 1.87–1.86 Ga (see Van Duin, 1992).
10.5. Angular unconformities? No clear angular unconformities have been identified within the Svecofennian domain in Finland even though several collisional or accretional orogenic events are evident. In the Lachlan Fold Belt of eastern Australia, for instance, the orogenic framework consists of four main orogenic pulses between 440 Ma to 340 Ma, originally delineated by unconformities (Gray et al., 1997). Three angular unconformities seem posCHAPTER
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sible in the Svecofennian domain in Finland. An angular unconformity (1.87–1.86 Ga) probably exists between the bulk of the supracrustal rocks and those associated with the Tiirismaa-type quartz arenites. The 1.93– 1.92 Ga Savo arc, together with the Keitele microcontinent, possibly collided with the Archean craton at 1.91–1.90 Ga (Lahtinen, 1994; Nironen, 1997; Chapter 11). Therefore, the 1.90–1.88 Ga volcanic rocks at Pielavesi should be underlain by an unconformity. The third case may be speculated for the Pirkanmaa and Pohjanmaa belts, where an unconformity could exist between rocks clearly older than 1.91 Ga and rocks deposited after ~1.92 Ga. As referred to by Eskola (1941), J.J. Sederholm tried to find the discordance between the supracrustal rocks of the Tampere belt that he called Bothnian schists and the Svionian rocks that he considered to compose the depositional basement of the Bothnian schists. Sederholm was very eager in this topic and stated: “Ich frage nach der Unterlage ebenso hartnäckig wie Franzose fragt: où est la femme?” Considering the established multiple orogenic events it seems that we should ask instead: Où sont les femmes?
11. Summary The Svecofennian domain of Finland is a prime example of Proterozoic accretionary orogens. Its supracrustal lithological units are dominated by ~1.9 Ga turbiditic sedimentary and ~1.90–1.88 Ga arc-type volcanic rocks (Table 8.3). Black shales, MORB- and WPLlike, often pillowed basalts, arc-related fluvial to turbiditic sedimentary rocks, felsic rocks (largely volcanic or volcanogenic sedimentary deposits), and sedimentary carbonates are locally common, whereas quartz arenites and cherts are rare. Three major terranes are discerned: the Savo belt, central Svecofennia, and southern Svecofennia. The rocks record two major orogenic periods: 1.89–1.86 Ga
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events throughout the area and a 1.83–1.81 Ga event in the south. In addition, a 1.91–1.90 Ga event in the Savo belt seems plausible. HighT, low-P amphibolite facies metamorphism dominated, but there was a variation from the greenschist/amphibolite facies boundary to the granulite facies. The tectonothermal events were mainly related to accretions and collisions of Paleoproterozoic microcontinents and newly formed arc systems as well as to their collision with the Archean craton in the northeast, and they were preceded by or partly coeval with subduction events at 1.93–1.92 Ga and 1.905–1.88 Ga. The Savo belt, close to the Archean craton, differs from the other Svecofennian supracrustal belts, because it is characterized by 1.93–1.92 Ga bimodal volcanic arc rocks with juvenile Nd isotope composition, and thus represents a relatively immature arc, and because it includes numerous massive sulfide deposits. The EMORB- or WPL-like mafic volcanic rocks of the Virtasalmi field are distinct and might represent a rift, back-arc basin or oceanic plateau setting. The 1.93–1.92 Ga arc magmatism in the Savo belt was evidently related to subduction under a 1.95 Ga crustal segment close to the inferred ~2.1–2.0 Ga Keitele microcontinent in the southwest. This newly formed Keitele–Savo entity possibly collided with the Archean craton at 1.91–1.90 Ga. Within central Svecofennia, the 1.905 Ga to 1.88 Ga volcanic and related sedimentary rocks of the Tampere belt, Ylivieska field, and the Central Finland granitoid complex represent an arc, whereas the Pirkanmaa belt and Evijärvi field compose the subduction zone complex of this evolved arc system. The system was formed close to the western and southern margins of the Keitele microcontinent, and the associated subduction zone was curved with N-directed subduction at Tampere and NE-directed subduction at Pohjanmaa. The arc-type volcanic rocks range from basalts to rhyolites, are rich in andesites and 394
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dacites, and largely of high-K and mediumK character, but shoshonitic and trachytic or high-K rhyolitic types are also relatively abundant. Like the 1.89–1.87 Ga plutonic rocks of central Svecofennia in general, the arc volcanic units seem to have εNd (T) close to zero. The fluvial to turbiditic sedimentary rocks derived from the arc also show εNd (T) close to zero. In the northwestern part of the Ylivieska field, however, possibly arc-related plutonic rocks have high εNd (T) of ~ +3 and thus indicate evolution in a less mature system, evidently outside the realm of the Keitele microcontinent. The Haveri basalts (Tampere belt) are largely pillow lavas and were probably emplaced before the 1.905–1.89 Ga arc volcanism. They show EMORB affinity, have mantle-type Pb isotopes, and show slightly positive εNd (at 1.9 Ga) values indicating mantle enrichment (relative to depleted mantle) some time before 1.90 Ga. The Haveri Formation is overlain by a thick unit of turbidites (Osara and Myllyniemi Formations), which have a significant Archean component and are dominated by 2.0–1.92 Ga detritus derived from the Savo belt and the Keitele microcontinent. Largely similar turbidites dominate in the Pirkanmaa–Evijärvi subduction zone complex, but this is relatively rich in black shales and, therefore, geophysically appears as a curved zone of crustal conductors. The subduction zone complex also includes NMORB- to WPL-like basaltic pillow lavas, minor cherts, and sedimentary carbonates as well as sedimentary rocks probably derived from the 1.905–1.88 Ga arc. The NMORB-like basalts (from Evijärvi) show clearly positive εNd (at 1.9 Ga) values indicating depleted mantle sources. In general, the basalts were formed in various non-arc settings, and the subduction zone complex consists of allochthonous slices. The Tampere–Ylivieska arc system probably continues as the Skellefte and Umeå fields (arc and subduction zone complex, respectively) in northern Sweden but the slightly to clearly positive εNd (T) values in the arc volcanic rocks
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Table 8.3. General characteristics of Svecofennian supracrustal belts in Finland. Savo belt in black, central Svecofennia in blue, southern Svecofennia in red.
Belt/area/field
Characteristic primary rock types
Volcanic age / εNd (T) maximum deposition age in Ga
Environment of deposition
Tectonic setting
Pyhäsalmi– Pielavesi region
low-K arc basalts– rhyolites (bimodal); black shales, dolomites; graywackes, mudrocks EMORB or WPB pillow lavas
1.93– 1.92
subaqueous
immature arc
shelf submarine fan shallow subaqueous
rift or marginal basin
Virtasalmi field Tampere belt and Ylivieska field lower parts at Tampere Pirkanmaa belt and Evijärvi field Häme belt
Saimaa area
Uusimaa belt Kemiö– Järvenpää field
Nauvo– Korppoo Pellinki
volcanic sandstones, mudrocks, conglomerates; medium-K to shoshonitic arc volcanics; graywackes, mudrocks; black shales, cherts; EMORB pillow lavas
1.90 1.905– 1.88 1.92
graywackes, mudrocks; black shales, cherts; NMORB to WPB pillow lavas
1.92
medium-K basalts– andesites; medium-K arc basalts– andesites–rhyolites
1.89– 1.88 1.89– 1.88
mudrocks, graywackes; TMORB pillow lavas; arc dacites
1.9
T-EMORB pillow lavas; medium-K to high-K arc basalts–andesites– rhyolites (in part bimodal), felsic volcanic and sedimentary rocks and sedimentary carbonates; graywackes, mudrocks graywackes, mudrocks; WPB pillow lavas medium-K arc basalts– andesites–dacites; felsic sedimentary rocks
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+0.5–0 0±1 +2±1
fluvial to submarine fan subaqueous to subaerial submarine fan
+0.5
subaqueous
–2±1
submarine fan
+3
subaqueous subaerial to subaqueous subaerial to subaqueous
0±1 +3
arc related highly evolved arc rift or marginal basin subduction zone complex various non arc rifting arc relatively evolved arc
subaqueous evolved arc
+3 –1
1.90– 1.88
subaqueous subaqueous
1.93
+2
1.89
at Skellefte indicate a less evolved setting than what is typical at Tampere and Pohjanmaa. In addition, the arc-type volcanic activity seems to have lasted longer in the Skellefte field. In southern Svecofennia, the sedimentary rocks have diverse sources and tend to be more pelitic than those of central Svecofennia. In the bulk of this terrane, pronounced metamorphism and migmatization as well as abundant 1.83–1.81 Ga potassic granites make identifiCHAPTER
+3
+2
relatively shallow water submarine fan submarine fan subaqueous subaqueous
marginal basin evolving arc, rifting arc in early stage
rifted crust relatively evolved arc
cation of primary features largely impossible. The terrane is characterized by two approximately coeval, E-trending volcanic belts with different characteristics: the Häme belt in the north and the Uusimaa belt in the south. The Häme belt consists mainly of 1.89– 1.88 Ga medium-K basaltic to rhyolitic volcanic arc rocks, which comprised separate stratovolcanoes, and overlying mafic lavas; the latter were associated with a stage of arc
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rifting. Compared to the Tampere belt and the bulk of the Uusimaa belt, this arc represents a less mature setting. In the northeast, the predominantly sedimentary rocks of the Saimaa area have diverse sources, show variation in the degree of maturity, and have variable ages of deposition. Volcanic rocks are not common in this area but range from the TMORB-like mafic to ultramafic pillow lavas with clearly positive εNd (at 1.9 Ga) at Rantasalami to the arc-type medium-K and high-K, mainly intermediate rocks at Parikkala and Punkaharju. The Uusimaa belt consists mainly of mudrocks, graywackes, 1.90–1.88 Ga arc-type volcanic rocks, and EMORB- and WPL-like basalts and picrites. Overall, it is relatively rich in felsic volcanic to volcanogenic sedimentary rocks and sedimentary carbonates. The mainly medium-K arc-type volcanic and related plutonic rocks of the Uusimaa belt show a variation in εNd (T) from slightly negative values (at Orijärvi) to moderately positive values (at Pellinki), while the T- or EMORBlike Salittu picrites have a clearly positive εNd (T) value. The EMORB- to WPL-like mafic and ultramafic volcanic rocks of the Uusimaa belt were possibly emplaced both before and after the 1.90–1.88 Ga arc volcanism. The belt is a lateral counterpart of the Bergslagen field of Sweden. The supracrustal rocks were deposited in a fairly evolved system within or close to the ~2.1–2.0 Ga Bergslagen microcontinent, but the environment in Uusimaa was less continental than in Bergslagen. Compared to the Kemiö–Järvenpää field in the west, the arc setting in the eastern Uusimaa belt (at Pellinki) was even less mature. Both S- and N-directed subduction have been suggested for the arc volcanism in the Häme and Uusimaa belts. The original spatial relation of the belts remains to be determined. Recent observations on the Tiirismaa-type quartz arenites from southern Svecofennia with ~1.86 Ga detrital zircons indicate deposition after or during the 1.87–1.86 Ga tectonothermal peak, but associated mudrocks and 396
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angular unconformities have not been identified. Similar sequences are known in central Sweden where they are partly derived from the nearby 1.87–1.86 Ga volcanic rocks. Such age volcanic rocks are, however, unknown in Finland.
Acknowledgments First, I would like to thank numerous colleagues for discussions, information, and suggestions before and during preparation of the manuscript: Carl Ehlers, Hannu Huhma, Juha Karhu, Jarmo Kohonen, Timo Kilpeläinen, Aulis Kinnunen, Tapio Koistinen, Kalevi Korsman, Jukka Kousa, Raimo Lahtinen, Alf Lindroos, Mikko Nironen, Lauri Pekkarinen, Fredrik Strandman, Ragnar Törnroos, Markus Vaarma, and Markku Väisänen, among others. Sincere thanks go also to Gerhard Hakkarainen, Raimo Lahtinen, Lauri Pekkarinen, and Markku Väisänen for providing access to data files. Aulis Kinnunen and Markku Tiainen kindly accepted the use of the Häme belt data and the Pirkanmaa belt data, respectively, that we had collected together. Annakaisa Korja, Raimo Lahtinen, and Mikko Nironen shared unpublished information from Chapter 11. Within the framework of GGT-project, Kalevi Korsman provided an opportunity for field and geochemical studies in the Tampere–Vammala area. I also gladly recall the times of the project led by the late Tapio Koljonen sponsoring REE analyses on the Häme, Pohjanmaa, and Tampere belts. Excursion together with Asko Käpyaho and Jukka Reinikainen to the Uusimaa belt opened new views. Mirjam Ajlani kindly processed Figure 8.5. Carl Ehlers and Raimo Lahtinen reviewed the manuscript and made critical and helpful comments. The efforts and suggestions by Martti Lehtinen and Tapani Rämö were invaluable during preparation of this article. Hugh O’Brien checked the English and also made significant comments and proposals.
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Wasström, A., 1990. Knaftenområdet - en primitiv tidig Proterozoisk vulkaniska öbåge söder om Skellftefältet, norra Sverige. M.Sc. Thesis, Åbo Akademi University, Finland. 1–128. (in Swedish) Wasström, A., 1993. The Knaften granitoids of Västerbotten County, northern Sweden. In: T. Lundqvist (Ed.), Radiometric dating results. Geol. Surv. Sweden, Res. Pap. C 823, 60–64. Wasström, A., 1996. U–Pb zircon dating of a quartz-feldspar porphyritic dyke in the Knaften area, Västerbotten County, northern Sweden. In: T. Lundqvist (Ed.), Radiometric dating results 2. SGU series C 828, 34–40. Weihed, P., Billström, K., Persson, P.-O., Bergman Weihed, J., 2002. Relationship between 1.90–1.85 Ga accretionary processes and 1.82–1.80 Ga oblique subduction at the Karelian craton margin, Fennoscandian Shield. GFF 124, 163–180. Welin, E., 1987. The depositional evolution of the Svecofennian supracrustal sequence in Finland and Sweden. In: G. Gaál, R. Gorbatschev (Eds.), Precambrian geology and evolution of the central Baltic Shield. Precambrian Res. 35, 95–113. Wilson, M.R., Hamilton, P.J., Fallick, A.E., Aftalion, M., Michard, A., 1985. Granites and early Proterozoic crustal evolution in Sweden: evidence from Sm–Nd, U–Pb and O isotope systematics. Earth Planet. Sci. Lett. 72, 376–388. Wilson, M.R., Claesson, L.-Å., Sehlstedt, S., Smellie, J.A.T., Aftalion, M., Hamilton, P.J., Fallick, A.E., 1987. Jörn: An early Proterozoic intrusive complex in a volcanic arc environment. Precambrian Res. 36, 201–225. Windley, B.F., 1992. Proterozoic collisional and accretionary orogens. In: K.C. Condie (Ed.), Proterozoic Crustal Evolution. Elsevier, Amsterdam, 419–446. Yli-Kyyny, K., 1990. Geology and geochemistry of volcanic rocks in Viljakkala, Ikaalinen and Kankaanpää. In: Y. Kähkönen (Ed.), Symposium Proterozoic Geochemistry, Helsinki ’90. December 13–14, 1990, Helsinki, Finland. Abstracts. Dept. Geol., Univ. Helsinki, p. 46.
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Chapter 9
SVECOFENNIAN MAFIC–ULTRAMAFIC INTRUSIONS
P. Peltonen
Cover page: Orbicular peridotite from the Kylmäkoski Ni-Cu deposit. Photo: Jari Väätäinen.
Peltonen, P., 2005. Svecofennian mafic–ultramafic intrusions. 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. 407–442. © 2005 Elsevier B.V. All rights reserved.
Three types of mafic–ultramafic intrusions (Groups I, II, and III) were emplaced during the Svecofennian orogeny at ~1.89–1.87 Ga. Altogether, this magmatism represents a significant fraction of the Paleoproterozoic crustal growth of the Fennoscandian Shield and it also had a major influence on its metamorphic evolution. Most of the Group I intrusions were emplaced close to the peak of the Svecofennian orogeny (~1.89 Ga) and were derived from hydrous arc-type basalts. They bear striking geochemical, mineralogical, and structural similarities to the mafic–ultramafic complexes exposed in younger deeply eroded oceanic and continental arcs (e.g., the Aleutians and Andes). The Group I intrusions were emplaced over a protracted period during the amalgamation of the Svecofennian arc collage; some of them represent conduits of arc basalts or were emplaced within an accretionary wedge, others were emplaced during the Svecofennian arc–Archean craton collision along transtensional shear zones. The Group I intrusions show evidence for syncrystallization deformation and assimilation of country rocks. They have a high potential for magmatic Ni-Cu sulfide deposits and have been the main source of Ni in Finland. The Group II intrusions are large synvolcanic layered gabbro complexes located in the Southern Finland arc complex. They represent low-pressure crystallization products of relatively juvenile subalkalic tholeiitic basalts within an oceanic arc and are not spatially associated with the Group I bodies. This suggests that the southern Finland oceanic arc terrain was amalgamated to the Western Finland arc complex only after the emplacement of the Group I and Group II intrusions. The latter have low potential for magmatic sulfide and oxide deposits. The Group III intrusions are Ti-Fe-P-rich gabbros within the Central Finland granitoid complex region. They share the geochemical similarities with anorogenic gabbros and probably do not all have a common origin. Several of these intrusions are genetically related to K-rich granitoid plutons and form, together with the granites, a bimodal magmatic suite that was generated by magmatic underplating at the postkinematic stage of the Svecofennian orogeny. However, a few of the Ti-Fe-P gabbros yield synorogenic crystallization ages and may actually represent evolved Group I magmas. Some of the Group III intrusions host important Ti ore reserves.
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1. Introduction Mafic–ultramafic plutonic rocks emplaced during Precambrian orogenic episodes are a poorly defined category of intrusions. This is not only because of the long tectonometamorphic history of such terrains and scarcity of young analogies (deeply eroded oceanic or continental arcs), but also because orogenic emplacement of magmas is often accompanied by subsequent breakup or boudinage of the intrusions, syncrystallization deformation, metamorphism, assimilation of country rock material, and prolonged thermal re-equilibration with the country rocks. These processes cause lithologic and textural diversity, which hampers their characterization and exact timing of emplacement relative to the major stages of the orogeny. Proper characterization of orogenic mafic plutonism is of prime importance for a number of reasons. First, their parental melts have ultimately been generated in the upper mantle, and these rocks thus provide information on the nature of the mantle source beneath convergent plate margins. Often, however, the composition of the parental melts becomes strongly modified by assimilation of crustal material at lower crust–upper mantle boundary region or during magma ascent towards higher crustal levels (Hildreth and Moorbath, 1988). Second, orogenic mafic–ultramafic plutonic rocks may constitute a significant part of the new crust generated at convergent tectonic settings (Robins and Gardner, 1974; Boyd and Mathieson, 1979; Snoke et al., 1982; Thompson, 1984; Burns, 1985; Butler, 1989; DeBari and Coleman, 1989; Grissom et al., 1991; Kepezhinskas et al., 1993; Skirrow and Sims, 1999; Schersten, 2001). Third, mafic plutonism provides a mechanism to transport heat upwards within the crust and thus has a strong impact on the metamorphic evolution of the lower and middle crust (Komatsu et al., 1994). As any other mafic–ultramafic magmas, orogenic melts can also lead to accumulation 410
of valuable magmatic Ni-Cu-PGE sulfides, Ti-Fe-V oxides or apatite. Clearly, proper understanding of both the magmatic and subsolidus history of orogenic mafic–ultramafic intrusions is a prerequisite for the modeling of crustal evolution and development of successful exploration strategies. This chapter provides a review of the mafic–ultramafic intrusions of the Svecofennian orogen (Figure 9.1). Distribution of various intrusion types in the Svecofennian domain will be outlined and several case studies will be described in detail. In most cases, the cumulus terminology of Irvine (1982) is followed unless only conventional rock names were used in the original descriptions and corresponding cumulus names cannot be deduced. The mutual relationship of the intrusion types, timing of their emplacement relative to the tectonic evolution, and significance for regional studies of the Svecofennides will be scrutinized. Lahtinen et al. (Chapter 11) describe in detail the geodynamic context of the Svecofennian magmatism and give necessary tectonic information that is not repeated here.
2. Classification of the intrusions Mafic–ultramafic plutonic rocks can be classified according to their isotope ages, petrology, and geochemistry or tectonic setting. According to the classification scheme of Naldrett (1989), all Svecofennian intrusions belong, in a broad sense, to Category IV – “Intrusions emplaced in an active orogenic belt”. Such bodies are characterized by syndeformational intrusion resulting in fragmentation and boudinage, partial metamorphism, and presence of primary hydrous phases that, at least in some cases, indicate origin above active subduction zone. Also typical are complex contact phenomena and emplacement at relatively deep crustal levels where high ambient temperatures result in assimilation of country rock material. For the Svecofennian intrusions, however, a more specific classification scheme
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Tyypekinlampi Saarisenjärvi Kotalahti Laukunkangas +0.1
Kauhajärvi Perämaa –0.7– +0.6 +0.2
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+1.2 Vammala
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Primitive (1.93–1.91 Ga) arc complex Arc complex of western and southern Finland (1.89–1.87 Ga) Pirkanmaa belt Tampere belt Synorogenic (Groups Ia, Ib) intrusions
Postkinematic granitoid plutons Ti-Fe-P gabbros (Group III) +2.3
Initial Nd value Archean Karelian craton Southwestern margin of the craton
Proterozoic metasedimentary cover of the Archean craton Rapakivi granites and minor sedimentary rocks Fig. 9.1. Generalized geological map of central and southern Finland (modified from Korsman et al., 1997, and Rämö et al., 2001) show the occurrences of different types of mafic–ultramafic intrusions.The synorogenic Group I bodies are found both within the Primitive and Western Finland arc complexes and also intrude the Archean basement gneisses (and their metasedimentary cover), but are absent from the Häme belt. Group I intrusions have variable and relatively non-depleted initial Nd isotope composition. Much of the Häme belt consists of Group II synvolcanic intrusions and associated volcanic formations with depleted Nd isotope signatures. Group II bodies are absent from the Pirkanmaa belt suggesting a major tectonic boundary between Häme and Pirkanmaa belts. Group III Ti-Fe-P gabbros are found within the peripheral zones of the Central Finland granitoid complex, some of them closely associated with postkinematic granitoid plutons. Initial εNd values after Huhma (1986; unpubl.), Patchett and Kouvo (1986), Makkonen (1996), Rämö et al. (2001). Häme belt Synvolcanic mafic intrusions (Group II)
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is desirable. U-Pb zircon ages of the Svecofennian mafic–ultramafic intrusions form a rather continuous spectrum spanning the synorogenic (~1.87–1.89 Ga) stage of the Svecofennian orogeny (Figure 9.2). Some intrusions record slightly older ages that, however, tend to be associated with large errors, probably indicative of heterogeneous zircon populations. Most clearly, this may be the case for Lapinlahti, which is a rare example of Svecofennian intrusions emplaced into the Archean crust (Figure 9.1). Whether some of the mafic–ultramafic intrusions within the Svecofennian orogen were emplaced at the early orogenic stage (~1.9 Ga) remains uncertain. Similarly, the samples yielding the youngest ages (~1.87 Ga) have rather large errors. The Soukkio gabbro, for example, is a bimodal mafic–felsic igneous complex (Eerola et al., 2001; Huhma, 1986) in which mixing of older mafic intrusive rocks with younger granite remains a possibility. Therefore, the actual range of emplacement ages for the Svecofennian mafic–ultramafic intrusions is probably somewhat smaller than the range indicated by the minimum and maximum ages of the chronogram (Figure 9.2). Most of the mafic–ultramafic intrusions record ages between 1875 Ma and 1885 Ma, which corresponds to the peak of the synorogenic stage of the orogeny. Two small flexures on the chronogram divide this population into three subgroups, 1885, 1880, and 1875 Ma. All these contain samples from both southern and central Finland and from various geotectonic and lithologic units and correlations between age and geographic location are absent. One kind of relationship is, however, evident: all intrusions that contain significant magmatic Ni-Cu sulfide deposits belong to the 1880 Ma age group. Because age data cannot be used to divide the intrusions into meaningful lithologic groups, a geotectonic domain concept is used. Basically, the terminology for the geotectonic units follows that presented elsewhere (Kors412
man et al., 1997). Below, a general outline of the intrusions and some detailed case studies will be provided for the following subgroups of mafic intrusions: Group I: Intrusions of the Arc complex of western Finland Group Ia: Intrusions close to the Archean craton margin Group Ib: Intrusions of the Tampere and Pirkanmaa belts Group II: Synvolcanic intrusions of the Arc complex of southern Finland Group III: Ti-Fe-P gabbros of the Central Finland granitoid complex.
3. Intrusions close to the craton margin (Group Ia) The craton margin environment in central Finland shows a marked concentration of mafic–ultramafic intrusions, several of which host small magmatic Ni-Cu sulfide occurrences or deposits (e.g., Papunen and Gorbunov, 1985). Within this domain the areal distribution of the intrusions is not restricted to any major geotectonic unit (Figure 9.1). They are found both in the older (~1.92 Ga) Primitive arc complex and the younger (1.89–1.87 Ga) Arc complex of western Finland. Some of them are found east of the (subsurface) Archean–Proterozoic boundary and are intrusive both to the Archean basement gneisses and the overlying metasediments (e.g., the Lapinlahti gabbro–anorthosite). Traditionally, the emplacement of these intrusions has been related to development of subvertical D3 wrench lineaments (e.g., Gaál, 1972). Spatial association of shear zones and intrusions is especially evident adjacent to the suture zone in the southeast (Figure 9.1). Most of the intrusions are found within a broad belt outside these shear zones and an unambiguous genetic relationship between them has not been established. Structural analysis implies
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Laukunkangas (Ni-Cu) Hyvinkää layered intrusion Koivusaarenneva Ti-Fe-P gabbro Kaipola layered intrusion
Kotalahti (Ni-Cu)
Vammala (Ni-Cu)
Groups Ia and Ib Group III Unclassified
Mean
Group II Lapinlahti
Fig. 9.2. Chronogram of U-Pb zircon ages of the Svecofennian mafic–ultramafic plutonic rocks. Data sources for targets discussed in the text: Soukkio (Huhma, 1986), Hitura (Isohanni et al., 1985), Laukunkangas (Huhma, 1986), Hyvinkää (Patchett and Kouvo, 1986; Suominen 1988), Koivusaarenneva (Kärkkäinen, 1999b), Kotalahti (Gaál, 1980), Vammala (Häkli et al., 1979), Lapinlahti (Paavola, 1988), Perämaa (Rämö et al., 2001), Saarisenjärvi and Tyypekinlampi (Ekdahl, 1993).The remaining non-labeled data from Helovuori (1979), Honkamo (1988), Hopgood et al. (1983), Marttila (1981), Nurmi et al. (1984), Nykänen (1983), Suominen (1991),Vaasjoki (1989),Vaasjoki et al. (1988, 1996) and from the unpublished database of the Geological Survey of Finland. C H A P T E R 9 • S V E C O F E N N I A N M A F I C – U LT R A M A F I C I N T RU S I O N S •
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that the intrusions were emplaced already during early D2 deformation and were deformed by recumbent D2 folding that predates the D3 lineament formation (Jokela, 1991; Koistinen, 1996). The present author favors a model in which plutonism occurred over a wide zone due to the westward subduction during the final stages of the closure of the basin between the Primitive arc complex and the Arhean craton. Synchronous or subsequent to the amalgamation, transtensional shear systems developed at the continental margin locally facilitating the ascent of melts along subvertical shear zones. This model also explains the more primitive composition and higher Ni potential of the shear zone-associated intrusions compared to intrusions elsewhere. Within shear zones, magmas are expected to rise faster and undergo less fractionation during emplacement thus also retaining higher potential to saturate nickeliferous sulfides. During the D3 phase these zones were further reactivated, deforming and brecciating the intrusions. Mafic magmatism continued for some time after the amalgamation as evidenced by intrusions (e.g., Saarisenjärvi, Tyypekinlampi) that yield postkinematic crystallization ages of ~1875 Ma (Figure 9.2; Ekdahl, 1993). Three representative intrusions, Laukunkangas, Kotalahti, and Lapinlahti, of the craton margin environment are described in more detail. Two of these, Laukunkangas and Kotalahti, hosted economic magmatic Ni-Cu sulfide deposits.
3.1. Laukunkangas The Laukunkangas mafic intrusion is a small mafic–ultramafic body within a zone of intense transcurrent faulting adjacent to the southwestern margin of the Archean craton (Figure 9.1). It is enclosed by high-grade Svecofennian graphite-bearing migmatites that, close to the intrusion margins, may contain garnet, cordierite, and orthopyroxene porphyroblasts. The associated magmatic Ni-Cu sulfide deposit 414
was one of the largest in the Svecofennian intrusions and yielded 6.6 Mt ore (0.78 wt.% Ni and 0.22 wt.% Cu) in 1984–1994 (Puustinen et al., 1995). The Laukunkangas intrusion is elongated, pipe-shaped, approximately 1 km long, 200 m wide, and more than 800 m deep. The mineralized eastern part of the body is well known because of extensive drilling. Breccia structures and graphite-rich gneiss xenoliths are common throughout the body, and are considered indications of tectonic disturbance during intrusion and solidification (Grundström, 1980). Laukunkangas intrusion can be divided into marginal and layered series (Figure 9.3). The marginal series is heterogeneous, noritic and shows reverse fractionation. The layered series comprises two distinct zones: peridotite and norite. The peridotite zone is located in the eastern tip of the intrusion and consists of olivine, olivine-plagioclase, and olivineorthopyroxene cumulates. The peridotite zone is overlain by the norite zone that is more than 200 m thick, comprising rhythmically layered orthopyroxene-plagioclase and plagioclaseorthopyroxene cumulates. On the basis of the Ni content of silicates and sulfides the norite zone can be divided into three subzones (Figure 9.3). Generally, the norite subzone 1 overlays the peridotite zone but is locally missing, whereas subzone 2 lies directly on the peridotite zone. Subzone 3 is above subzone 2 and consists of evolved plagioclase-rich cumulates. Subzone 1 has the highest Ni content, subzone 2 is intermediate, and subzone 3 is most depleted (Pertti Lamberg, pers. comm., 2001). Clinopyroxene, plagioclase, magmatic amphibole, phlogopite, and Ni-Cu sulfides dominate as intercumulus minerals. The Ni-Cu mineralization is associated with the peridotite zone close to intrusion margin. The mineralization includes both disseminated, massive, and breccia-textured ore types. Sulfide breccias and sulfide veins, which are confined to the contact zone between the intrusion and country rocks, consist of massive sulfides and
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Laukunkangas intrusion
Layered series
Marginal series
Peridotite zone
Norite zone Norite subzone 3 Ni(opx) <50 ppm
Norite subzone 2 Ni(opx) 50–250 ppm
Norite subzone 1 Ni(opx) >250 ppm
Fig. 9.3. Subdivision of the Laukunkangas intrusion (Pertti Lamberg, pers. comm., 2001). Opx–orthopyroxene.
contain abundant country rock fragments. Pyrrhotite, pentlandite, and chalcopyrite are the major ore-forming minerals and sphalerite, gersdorffite, violarite, ilmenite, magnetite, rutile, graphite, and molybdenite are found as minor constituents (Grundström, 1980). Fractional crystallization and sulfide saturation modeling suggest that crustal contamination of a relatively primitive parental magma resulted both in a shift of the melt composition from the olivine field to the orthopyroxene field and sulfide saturation (Pertti Lamberg, pers. comm., 2001).
3.2. Kotalahti The Kotalahti intrusion is one of several mafic–ultramafic intrusions within the NWrunning tectonic shear zone, the Kotalahti Ni belt (Gaál, 1972). It hosted the largest Svecofennian magmatic Ni-Cu sulfide deposit with total production of 12.3 Mt of ore (0.66 wt.% Ni and 0.26 wt.% Cu) in 1957–1987 (Puustinen et al., 1995). According to Gaál (1980) the magma was emplaced along subvertical axial plane of a NNW-trending synform in the Archean bedrock. Next deformation phase
was associated with high strain, refoliation, and generation of the NW-trending Kotalahti Ni belt shear zone. The Kotalahti intrusion is a subvertical plate that is ~1300 m long and 200 m wide at maximum. The nothern part of the body is steeply dipping and shows normal order of fractionation form footwall peridotites towards hanging-wall gabbros (Figure 9.4). The central part of the plate is characterized by “upside-down” structure with pyroxenitic and peridotitic cumulates in the upper parts of the body. Some structures indicate that the emplacement of the ultramafic rocks postdates that of the gabbros. The southern part of the Kotalahti intrusion consists of a mineralized ultramafic pipe that does not display internal fractionation. Heterogeneous gabbros are abundant in the lower parts and at the marginal zones of the intrusion. They include olivine gabbros, olivine norites, norites, gabbros and hornblende gabbros. The most fractionated rock types are diorites and quartz diorites at the bottom of the complex (Papunen and Koskinen, 1985). Breccia textures are common between the peridotitic, pyroxenitic, and gabbroic units and indicate polyphase intrusion. The sulfides can be classified as disseminated (interstitial), breccia, and massive ore veins (Papunen, 1970). They are associated with the ultramafic cumulate units and a separate breccia-textured ore body (Jussi Ore) within the graphitic gneisses outside the ultramafic intrusion proper. Such offset ores, also present in most of the other Svecofennian Ni-Cu deposits, are of high-grade and form the economic backbone of these otherwise rather low-grade deposits. The mineralogical composition of the Kotalahti ore is simple: the main minerals are monoclinic and hexagonal pyrrhotite, rare troilite, pentlandite, and chalcopyrite. The Jussi Ore also contains pyrite, millerite, and bornite (Papunen, 1970).
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100 m
Mertakoski ore body
100 m
200 m
200 m
300 m
300 m Välimalmi ore body
400 m
400 m
500 m
500 m
600 m
600 m
Gabbro
Mica gneiss
Pyroxenite
Amphibolite
Peridotite
Granite
Ni-Cu sulfide ore Fig. 9.4. Two representative cross-sections from the northern part of the Kotalahti intrusion (after Papunen and Koskinen, 1985).
3.3. Lapinlahti gabbro–anorthosite The Lapinlahti gabbro–anorthosite was emplaced into the Archean crust close to the craton margin (Figure 9.1). A coarse-grained gabbro dike from the central part of the body yielded a zircon age of 1895 ± 15 Ma (Paavola, 1988) implying that also this gabbro belongs 416
to the major phase of Svecofennian mafic plutonism (Figure 9.2). Lapinlahti intrusion is a subrounded body with a narrow 8-km-long “tail” protruding towards southwest from the main body, and has a total areal coverage of 44 km2 (Figure 9.5). The intrusion is completely enclosed by Archean banded tonalite–trondhjemite migmatites and has a concentric struc-
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ture with steeply (60–90°) dipping layering and a strike parallel to the intrusion margins. Magmatic layering proceeds from olivine gabbronorites and gabbronorites within the outer rim of the body towards more evolved leucogabbros, anorthosites, and hornblende gabbros in the central parts (Figure 9.6A, B). Minor ultramafic rocks include olivine websterites, websterites, and their hornblendebearing varieties. Crystallization started with plagioclase and olivine and was followed by orthopyroxene and clinopyroxene. Apparently, the parental melt had high volatile content as not only calcic amphibole but also biotite is present as large poikilitic intercumulus grains throughout the crystallization sequence. In the more evolved rock types hornblende and apatite appear as cumulus minerals. Kerkkonen (1985) argued that the parental melt of the Lapinlahti was probably high-Al basalt. The internal structure of the Lapinlahti gabbro–anorthosite is well displayed on aeromagnetic and ground gravity survey anomaly maps. The Bouguer anomaly (Figure 9.5B) is tightly restricted within the exposed area of the gabbro showing maxima in the center of the body. These features are consistent with the intrusion being an almost vertical funnel-shaped pipe (Kukkonen, 1981). The second vertical derivate of the gravity data visualizes the concentric structure of the body (Figure 9.5C). Ultramafic cumulates, olivine gabbro-norites, and hornblende gabbros outcome as dense outer and inner layers while the anorthosite-dominated middle layers appear as a gravity low. Gray-tone and obliquely illuminated low-altitude aeromagnetic maps bring out more subtle features of magmatic layering (Figure 9.5D, E).
4. Intrusions of the Tampere and Pirkanmaa belts (Group Ib) The supracrustal belt between the Central Finland granitoid complex and the Häme belt
consists of two distinct lithological domains, the high-grade Pirkanmaa belt and the medium-grade Tampere belt (Figure 9.1). The Pirkanmaa belt consists of high-grade and polydeformed tonalitic migmatites derived from psammitic protoliths (Koistinen, 1996). The Tampere belt is a narrow volcano-sedimentary sequence of basaltic to rhyolitic mature arc-type rocks and turbiditic graywackes (Kähkönen, 1989). The boundary between the Pirkanmaa and the Tampere belts is located within a major shear zone that is not a major terrain boundary but a thrust or reverse fault (Nironen, 1989). These belts are believed to represent the upper and middle crustal expressions, respectively, of the same volcanic arc–accretionary wedge terrain. The mafic–ultramafic intrusions of the Pirkanmaa belt are, in general, more mafic compared to those of the Tampere belt and have high potential for magmatic Ni-Cu sulfide deposits (Lamberg, 1990; Papunen and Gorbunov, 1995; Peltonen, 1995a). Within these belts a broad correlation exists between the nature of the Group Ib intrusions and the metamorphic grade of their country rocks. Intrusions within high-grade domains tend to be more metamorphosed and deformed, smaller, and more mafic (and Ni-ore potential) than their counterparts in the lower-grade crustal domains. This is a typical feature of synorogenic intrusions and suggests that the depth of their emplacement corresponds to the pressure determined from the metamorphic mineral assemblage of the enclosing supracrustal rocks (Peltonen, 1995a). A complete layered series is not preserved in any of the intrusions but a cumulate pseudostratigraphy – obtained by combining petrographic data from over 50 bodies – illustrates some salient features of their fractionation. Figure 9.7 shows an idealized layered series that would result from closed-system crystallization of the parental melt, as well as the extent of layered series in some example intrusions. The layered series and country rocks are sepa-
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Fig. 9.5. Geological (A), Bouguer gravity anomaly (B), second vertical derivative of Bouguer anomaly (C), gray-tone low-altitude aeromagnetic (D), and obliquely-illuminated aeromagnetic anomaly (E) maps of the Lapinlahti gabbro–anorthosite. Geology modified after Kerkkonen (1985) and Paavola (1988). Geophysical data from the Geological Survey of Finland (processed by Seppo Elo).
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rated by a hybrid zone of magma–sediment mingling and by a more extensive marginal zone. The marginal zone shows a reverse trend of differentiation dominated by two-pyroxene cumulates. At the top, direction of differentiation changes to normal and marginal contact zone gives way to the cumulates of the layered series. Peridotite zone is distinguished by the presence of cumulus olivine. Earliest cumulates are composed of cotectic proportions of olivine and chromite and intercumulus material (locally including abundant magmatic sulfides). Olivine-chromite cumulates are followed by olivine-two-pyroxene cumulates until the cumulus termination of olivine marks the top of the peridotite zone. Pyroxenite zone is dominated by two-pyroxene cumulates with plagioclase as a common intercumulus phase; the amount of plagioclase increases with stratigraphic height. In some bodies of the Pirkanmaa belt plagioclase is absent and magmatic amphibole is the major intercumulus phase. Beginning of the gabbro zone is marked by the appearance of cumulus plagioclase. Lower gabbro zone is dominated by clinopyroxeneplagioclase and clinopyroxene-orthopyroxene-plagioclase cumulates. The most evolved rock types are plagioclase-orthopyroxene rocks and plagioclase cumulates; these may contain abundant euhedral apatite, ilmenite, and magnetite. Upper contact zones are only sporadically exposed and are composed of cognate gabbro xenoliths in a matrix of hybrid gabbro–metapelite rocks.
4.1. Ultramafic intrusions of the Vammala Ni province More than 50 small (100–1000 m long) ultramafic cumulate bodies are concentrated within a roundish crustal block ~15 km in diameter near Vammala (Figure 9.8). This area – the Vammala Ni province – is associated with a moderate gravity anomaly maximum that is not explicable by the small and extensively serpentinized cumulates exposed in the area,
but rather indicates the presence of voluminous mafic–ultramafic massifs at subsurface levels (Elo, 1992). Most of the ultramafic bodies can be depicted as boudins or lenses that “float” in polydeformed, medium- to high-grade paragneisses. The intrusions consist mainly of olivine-chromite and olivineclinopyroxene cumulates with clinopyroxene, orthopyroxene, magmatic amphibole, and phlogopite as intercumulus phases (Figure 9.9A). The entry of plagioclase was delayed, probably as a consequence of crystallization at relatively high confining pressure; moderate crustal pressures are also indicated by relatively aluminous pyroxenes and chromian spinels (Peltonen, 1995a). Metamorphism, deformation, subsolidus equilibration, and geochronology suggest that the emplacement of these cumulatetextured bodies coincided with the peak of regional meta morphism and deformation (Figure 9.6C; Jokela, 1991; Peltonen, 1995a; Marshall et al., 1995). Internal structures of the cumulate lenses and absence of strong penetrative tectonic fabric (in spite of their deformed large-scale morphology) may indicate that the cumulate-textured sills were boudinaged while not completely solidified. Partial recrystallization of cumulate bodies is also consistent with their synkinematic intrusion. Lack of prograde reactions in cumulate cores indicates that, between igneous crystallization and regional metamorphism, the cumulate bodies became hydrated only close to their margins. The metamorphic conditions reached upper amphibolite–lower granulite facies, i.e., 600–700 °C and 5–6 kbar (Peltonen, 1990). Folding of the host migmatites probably resulted in flexing and melt-facilitated fracturing of the ultramafic bodies and their veining by neosome material (Marshall et al., 1995) – a feature common also in the Hitura and Kotalahti intrusions (Papunen, 1970). The slow cooling of cumulates from peak conditions is evidenced by extensive subsolidus reequilibration of olivine and chromian spinel
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(Peltonen, 1995c) and redistribution of Ca between pyroxenes. Slow cooling rates and the absence of significant contact aureoles may reflect a low temperature gradient between the cooling intrusions and their metapelite surroundings undergoing migmatization. Later, Fe-Mg silicates became extensive replaced by pseudomorphic lizardite. Although ultramafic in bulk composition, these cumulates were not derived from ultramafic melt as suggested by the relatively low and uniform forsterite content of olivine (Fo = 77.0–82.4 mol.%). Instead, they either crystallized in an open system or represent fragments of much larger intrusions (Figure 9.8). However, because only minor amounts of more evolved cumulates are found in the region, the bodies cannot represent the basal units of fragmented layered intrusions, neither do they represent ophiolitic cumulates or the products of ultramafic magmas. They have been interpreted as remnants of synorogenic conduits for tholeiitic arc-type magmas that became choked by cumulus crystals and were boudinaged into small lenses and fragments by concomitant tectonic movements (Peltonen, 1995a, b). Several of the Vammala Ni province intrusions are mineralized (Figure 9.6D). The largest magmatic Ni-Cu sulfide deposit was hosted by the Vammala intrusion (Häkli et al., 1979) – it yielded 7.6 Mt of ore (0.68 wt.% Ni, 0.42 wt.% Cu) in 1974–1994 (Puustinen et al., 1995). Most of the sulfides are interstitial,
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indicative of early formation of immiscible sulfide liquid in the magma (Figure 9.9B). Minor remobilization of the interstitial ore occurred during metamorphism and deformation resulting in formation of thin massive sulfide
Fig. 9.6. (facing page) (A) Layered leucogabbro, Lapinlahti gabbro–anorthosite. (B) “Mottled” anorthosite, Lapinlahti. (C) Pyroxenitic dike in polydeformed graywacke–slate migmatite in the proximity of the Piimäsjärvi intrusion,Vammala Ni province. Such features can be applied to constrain the timing of the magmatism relative to the regional deformation. (D) Orbicular peridotite from the Kylmäkoski Ni-Cu deposit. Several origins have proposed for such a texture, e.g., the orbicules could represent rounded cumulate fragments (autoliths) that settled to the base of the magma chamber together with the immiscible sulfide liquid or be products of rapid olivine crystallization due to supercooling. (E) Pothole structure in the layered series of the Hyvinkää intrusion. (F) Layered ultramafic cumulates, Hyvinkää. (G) Rhytmically layered gabbronorite cumulate layers, Hyvinkää. (H) Fragmental unit with gabbro autoliths embedded in fine-grained gabbro, Hyvinkää. Photos: Petri Peltonen (A, B, G), Markku Tiainen (C), Jari Väätäinen (D) and Riku Raitala (E, F, H). C H A P T E R 9 • S V E C O F E N N I A N M A F I C – U LT R A M A F I C I N T RU S I O N S •
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veins. The dominating sulfide assemblage is monoclinic pyrrhotite + pentlandite + chalcopyrite ± cubanite ± mackinawite ± valleriite. Less common sulfides include gersdorffite, niccoline, various tellurides, sphalerite, galena, molybdenite, gold, silver, and PGM (Peltonen, 1995b). The petrogenesis of these magmatic sulfide ores will be discussed in Section 7.
4.2. Porrasniemi layered gabbro Porrasniemi layered gabbro is a typical example of mafic–ultramafic intrusions within the high-grade Pirkanmaa belt. The internal structure of Porrasniemi is well-preserved and enables detailed study of magmatic evolution and primary structures. Porrasniemi exhibits an extensive and complete cumulate sequence and has some special interest because of its apparent potential for magmatic sulfide deposits at depth. The Porrasniemi layered gabbro comprises three tectonic blocks separated from each other by migmatites (Figure 9.10). Originally, the gabbro was probably emplaced as a continuous 2-km-long and 400-m-thick stratiform sheet but became boudinaged by tectonic movements. In the subsequent deformation, the three blocks were rotated relative to each other and the magmatic layering was tilted close to vertical (Lamberg, 1990). Each block can be subdivided into layered series and marginal series. Marginal series is located between the layered series and the footwall contact of the intrusion and shows a reverse trend of fractionation: olivine content increases and plagioclase content decreases upwards in the sequence towards the peridotite zone. The marginal series is an up to 75-m-thick heterogeneous unit with abundant country rock xenoliths and poorly developed cumulus textures. Lamberg (1990) distinguished a country rock xenolith-rich “fragmental unit” within the marginal series of the Kärki block. Transition from the marginal series to the layered series 422
is relatively abrupt and is marked by a change from reverse to normal fractionation. The layered series can be divided into peridotite, pyroxenite, and gabbro zones. The peridotite zone is relatively thin (~50 m) and consists of peridotites and olivine websterites at the base and wehrlites at the top. The overlaying pyroxenite zone is approximately 250 m thick and consists of two-pyroxene cumulates (Figure 9.9C). The modal amount of intercumulus plagioclase gradually increases upwards in the pyroxenite zone. The peridotite and pyroxenite zones consist of ~ 40-cm-thick layers of uniform composition. At the base of the peridotite zone modal rhytmic layering is present: 15-cm-thick augite-bronzite-olivine cumulate layers are frequently separated by 1–2-cm-thick olivine-rich laminae. In the gabbro zone, melano- and leucocratic laminae alternate and, at the highest stratigraphic levels, 1–3-cm-thick pyroxene laminae alternate with plagioclase-rich layers. This type of layering resembles the schlieren-lamination of Irvine (1982). The ~100-m-thick gabbro zone is characterized by cumulus plagioclase gabbros and norites. The contacts between peridotite, pyroxenite, and gabbro zones are phase boundaries: beginning of the pyroxenite zone is marked by disappearance of cumulus olivine, and the start of the gabbro zone by appearance of cumulus plagioclase. Smooth geochemical and mineral chemical trends suggest that Porrasniemi crystallized from a single pulse of magma. Because chilled margins are not exposed and the rocks are cumulates, the composition of parental magma is difficult to estimate. However, back-calculation from cumulus mineral compositions implies that the parental magma was close to tholeiitic basalt. Ubiquitous hydrous intercumulus minerals suggests that it had relatively high volatile content. Mass balance calculations have shown that the cumulus sequence is incomplete also in Porrasniemi. The most primitive olivinecumulates are missing from the intrusion, and
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olivine may have saturated already within the feeder system. Also, the most evolved plagioclase-titanian magnetite cumulates, predicted to occur at the top, have not been found; they may remain unexposed above the norites or the most evolved melt may have escaped the magma chamber (Lamberg, 1990).
4.3. Kaipola layered intrusion The Kaipola layered intrusion (Figure 9.11) is located close to the easternmost tip of the Tampere belt close to the boundary of the Central Finland granitoid complex (Sandholm, 1970). The intrusion is situated between two NW-trending faults about 20 km to the north of the shear zone separating lower grade rocks
of the Tampere belt (upper crustal milieu) from the strongly deformed and metamorphosed rocks of the Pirkanmaa belt (middle crustal milieu). The intrusion is a 4.6-km-long and 2.2-km-wide oval gabbro–diorite body, which is beautifully displayed in the aeromagnetic map (Figure 9.11A). Its residual gravity anomaly is about 13 mGal and, according to gravity modeling, the body dips to the north–northwest with an average depth extent of about 1.4 km. The associated volcanic rocks give rise to a magnetic maximum on the northwest side of the intrusion (Peltonen and Elo, 1999). Unlike most other Svecofennian mafic–ultramafic intrusions, Kaipola is not enclosed by metasedimentary rocks but by syn- and postkinematic granitoids. Gabbro–granite
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D C Fig. 9.9. (A) Olivine-clinopyroxene-orthopyroxene cumulate with magmatic amphibole occupying the intercumulus space. Note the resorbed outlines of the cumulus clinopyroxe crystals indicative of the peritectic reaction clinopyroxene + intercumulus melt = amphibole. Vammala Ni province. (B) Olivine + chromite cumulate with interstitial Ni-Cu-Fe sulfides consisting of pentlandite (light yellow), chalcopyrite (yellow), and pyrrhotite (light brown).Vammala Ni-Cu deposit. (C) Clinopyroxene-orthopyroxene cumulate with intercumulus plagioclase. Porrasniemi layered gabbro. (D) Clinopyroxene-orthopyroxene-plagioclase cumulate with brown poikilitic intercumulus amphibole (pargasite). Note the resorbed outlines of the cumulus crystals due to their peritectic reaction with hydrous intercumulus melt. Kaipola layered intrusion. Width of the images corresponds to 2.5 mm.
relationships are well exposed at a road cut close to the western tip of the intrusion where the contact zone is characterized by gabbro enclaves in the granite and small mica-rich clots in the gabbro – both features implying immiscibility of two melts. Two distinct types of granitoid dikes intrude the gabbro: older fine-grained, 5–50-cm-wide, and deformed dikes with smooth and irregular boundaries and younger coarse-grained pegmatite dikes that sharply cut both the gabbro and the finegrained granite dikes. The relationship of the gabbro with the older dikes gives an impres424
sion of immiscibility and co-existence of felsic and mafic magmas, whereas the younger dikes clearly postdate the solidification of the gabbro. The older dikes are interpreted to be coeval with the synkinematic granitoids enclosing the gabbro and the younger dikes are related to the emplacement of the somewhat younger, postkinematic, Kaipola granitoid pluton nearby (cf. Nironen et al., 2000). The Kaipola layered intrusion is characterized by a well-preserved layered series showing distinctive large-scale repetitive layering with at least seven zones (Figure 9.11B). Thin
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Fig. 9.10. Geological map of the Porrasniemi layered gabbro (modified after Lamberg, 1990).The originally sill-type intrusion that crystallized from a single pulse of magma is interpreted to have boudinaged into three roundish fragments rotated relative to each other during regional deformation.
olivine-bearing (augite-bronzite-olivine cumulates, aboC) and pyroxenitic layers (augite or augite-bronzite cumulates, aC/abC) represent the most primitive fractionation products and have been observed at the base of some of the zones. Most of the cumulus sequence is composed of more evolved leucocratic–mesocratic plagioclase-dominated orthocumulates (paC/pabC/pbC/pC) with apatite and titanian magnetite as minor cumulus phases. A striking feature of the intrusion are large postcumulus oikocrysts of green and brown amphibole, implying that the parental melt of the intru-
sion was relatively hydrous (arc-type basalt?). Some of the poikilitic amphibole was formed by interstitial crystallization, but most of it was produced by peritectic replacement of cumulus pyroxenes and plagioclase (Figure 9.9D). Other intercumulus minerals include ilmenite, apatite, phlogopite, quartz, zircon, and plagioclase. Sulfides are uncommon which, together with low PGE contents, implies that Kaipola layered intrusion has low potential for magmatic sulfide deposits.
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5. Synvolcanic intrusions of the Arc complex of southern Finland (Group II) The Häme belt is a major volcanic-dominated terrain in southern Finland bounded by the Pirkanmaa belt in the north and high-grade metamorphic gneiss complexes in the south (Figure 9.1). Hakkarainen (1994) divided the supracrustal formations into the Forssa Group and stratigraphically younger Häme Group. The volcanic rocks of the Forssa Group are mainly medium-K, calc-alkaline pyroclastic andesites probably related to mature arc stratovolcanoes with intervening sedimentary basins. The Häme Group volcanic rocks are medium-K tholeiitic basalts and basaltic andesites believed to represent fissure eruptions during late intra-arc rifting. These volcanic formations are spatially associated with mafic intrusions, the most extensive of these are the Forssa gabbro and the Hyvinkää layered intrusion (Figure 9.1). Age data from the Häme belt and associated mafic intrusions are few. An andesitic lava close to the western margin of the Forssa gabbro yielded an age of 1888 ± 11 Ma (Vaasjoki, 1994). This felsic lava unit represents the uppermost units of the Häme belt (Hakkarainen, 1994) and thus provides a minimum age for the Forssa gabbro. Patchett and Kouvo (1986) reported an age of 1880 ± 5 Ma for a gabbro pegmatoid of the Hyvinkää layered intrusion. Plagioclase porphyrite close to the eastern margin of the intrusion yielded an age of 1880 ± 3 Ma and is considered to be indicative of the cogenetic origin of the gabbro and volcanic rocks (Suominen, 1988). A somewhat younger age, 1870 ± 8 Ma, was yielded by a hornblende gabbro from the Soukkio complex ~20 km east of Hyvinkää (Huhma, 1986). In the light of these data the possibility remains that the mafic magmatism gets younger from west to east: i.e., Forssa gabbro (>1888 Ma) > Hyvinkää layered intrusion (~1880 Ma) > Soukkio complex (~1870 Ma). 426
5.1. Forssa gabbro The Forssa gabbro consists of medium-grained amphibole and pyroxene gabbros, diorites, and quartz diorites. Ultramafic and anorthositic varieties are uncommon. Aeromagnetic low-altitude map (not shown) brings out the concentric structure of the gabbro with gabbroic rocks in the core and diorite in the margins. In many places the diorite brecciates the gabbro. This, together with the lack of fine-scaled magmatic layering, is indicative of a dynamic environment of crystallization. In the northwest, the plutonic rocks grade into a plagioclase-phyric and weakly ophitic hypabyssal rock type that gradually grades into hornblende-plagioclase porphyry (Neuvonen, 1956). All these features, also supported by the U-Pb zircon ages (Figure 9.2), are indicative of a synvolcanic nature of the gabbros and a comagmatic origin of the spatially associated volcanic formations.
5.2. Hyvinkää layered intrusion The synvolcanic Hyvinkää layered intrusion in the Häme belt (Figure 9.12) has an areal extent of ~120 km2 and is one of the largest Svecofennian layered gabbro complexes. It is an oval lopolithic body consisting of layered peridotites, pyroxenites, olivine gabbros, gabbronorites, non-layered isotropic gabbros, and granophyre. According to Raitala (1997), the body has been slightly tilted from its primary position so that the western part exposes deeper levels of the igneous stratigraphy and thus more primitive cumulates than those exposed in the east. In the east, the gabbros are rich in hornblende, biotite, and magnetite which, together with some quartz and alkali feldspar, imply proximity of roof. The layered series of the Hyvinkää intrusion has been studied in detail by Raitala (1997). The outermost shell is a hybrid zone between the country rocks and the first marginal series cumulates. The hybrid zone may reach
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Fig. 9.11. (A) Low-altitude aeromagnetic anomaly map of the Kaipola layered intrusion and surroundings. The intrusion is located between two parallel NW-trending faults. The lower part of the image illustrates the sharp boundary between the high-grade Pirkanmaa belt and lower-grade Tampere belt (after Peltonen and Elo, 1999; data from the Geological Survey of Finland). (B) Geology of the Kaipola layered intrusion. Mineral abbreviations for the cumulate names as in Figure 9.7. C H A P T E R 9 • S V E C O F E N N I A N M A F I C – U LT R A M A F I C I N T RU S I O N S •
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600 m in thickness but locally may be only a few meters. It consists of a variety of rock types that formed through mixing between the magma and country rocks. Frequently, the hybrid rock types bear evidence for strong interaction with late magmatic fluids. The presence of some disseminated sulfides and tellurides, including some Pt group minerals, has raised economic interest for the hybrid zone (Raitala, 2000). The marginal series is believed to have formed from cumulus crystals that nucleated and grew higher in the magma chamber but which, due to density contrast or convective currents, settled at the sidewalls of the magma chamber. The rock types of the marginal series are cumulates with clinopyroxene, orthopyroxene, and olivine as cumulus phases and plagioclase and clinopyroxene as the most common intercumulus minerals. The cumulus paragenesis alternates in a random manner and is distinct for example from the marginal series of the Porrasniemi intrusion (above) in which the order of crystallization was reverse to that of the layered series. The layers are frequently graded and 1 to 5 m thick. Magmatic erosion, slumping, gliding as well as textures implying filter pressing are common (Figure 9.6E). Erosional discordance separates the marginal series from the overlying layered series. Most of the layered series consists of rhytmically layered gabbro to gabbronorite cumulates (Figure 9.6F, G). The layers are generally 0.5–30 cm thick. A characteristic feature are country rock xenoliths and autolithic fragments (Figure 9.6H). These have settled parallel to the magmatic layering. A late magmatic dunite pipe, which consists of cumulus olivine and chromite and intercumulus clinopyroxene, intrudes the layered series (Figure 9.12). Later, the Hyvinkää layered intrusion was intruded by K-rich granite and diabase (Raitala, 1997).
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6. Ti-Fe-P gabbros of the Central Finland granitoid complex (Group III) Several gabbroic intrusions, some of them hosting important magmatic oxide deposits, are found within the Central Finland granitoid complex (Figure 9.1). Most of them are located in the peripheral areas of the complex and are confined to two clusters: the Kauhajärvi and Koivusaarenneva gabbro provinces. Although intrusions within these two provinces share a number of features, distinct origins have been proposed. Rämö et al. (2001) argued that the Kauhajärvi gabbros are genetically related to the ~1.87 Ga postkinematic Lauhanvuori granite and form a bimodal magmatic suite. In contrast, according to Kärkkäinen (1999a), the Koivusaarenneva gabbros are Group I intrusions emplaced during the synorogenic stage of the orogeny.
6.1. Kauhajärvi gabbro province The Kauhajärvi gabbro province (Figure 9.1) consists of five, 2–10-km-long and 1–3-kmwide gabbro intrusions. These gabbros are located between the postkinematic (1867 ± 6 Ma) metaluminous to peraluminous Lauhanvuori granite in the west and a foliated synkinematic (1886 ± 11 Ma) granodiorite in the east (Rämö et al., 2001). Field observations imply that gabbros are younger than the synkinematic granitoids but are intruded by the Lauhanvuori granite that postdates the main stage of the Svecofennian orogeny (Rämö, 1986). The Perämaa gabbro has yielded a similar crystallization age (1874 ± 14 Ma; Rämö et al., 2001) and has been related to the same geotectonic event as the Lauhanvuori granite. The intimate association of the Ti-Fe-P gabbros and latekinematic granites imply an mature, postorogenic type setting for the magmatism. Nironen et al. (2000) interpreted the
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contact features of the postkinematic granites and spatially associated gabbros as a result of mixing and mingling of coeval felsic and mafic magmas. They disregarded the possibility of a single parental magma and presented a two-stage model for the coexistence of the felsic and mafic magmatic suites. According to this model, compositionally variable and anhydrous lower crust was first produced as a result of extraction of synkinematic magmas. This granulitic residue was then melted due to
the mafic underplating. Both anatectic silicic melts and underplating mafic magmas contributed to the bimodal postkinematic magmatism that took place in response to extensional or transtensional events that modified the tectonically thickened Svecofennian crust.
Kauhajärvi gabbro The Kauhajärvi gabbro consists of two chemically distinct zones: relatively thin (~50 m), poorly layered basal zone and thicker (>400
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m) well-layered main zone (Figure 9.13). The basal zone consists of gabbro, gabbronorite, and olivine gabbronorite and layering is only weakly developed. Ilmenite and apatite are strongly concentrated in the uppermost part of the basal zone. The main zone is texturally and modally layered and extends from peridotite to anorthosite. Layers are 0.2–20 m thick. The chemical compositions of the basal zone and main zone are distinct – the main zone is enriched in Fe, Ti, and P and depleted in Mg, Cr, Si, and Al relative to the basal zone. Ilmenite was saturated early in the main zone and was concentrated into the most primitive Mg-rich cumulates. The main zone contains ~1500 ppm F, which is believed to reflect interaction of basaltic parental melt with coeval granitic magma. The compositional variation of both the basal zone and main zone can be explained by a closed-system fractional crystallization of a single pulse of tholeiitic magma. However, redox conditions during their crystallization were variable. The basal zone crystallized under low fO2, which resulted in strong enrichment of Ti, Fe, and P at the top of the zone. The main zone crystallized under relatively high oxygen fugacity resulting in co-precipitation of ilmenite and apatite throughout the layered series, thus preventing formation of massive oxide ore layers (Kärkkäinen, 1999a).
Perämaa gabbro The Perämaa intrusion is found in the same tectonic setting as the Kauhajärvi gabbro (Figure 9.1). Intermediate differentiates dominate but ultramafic–gabbroic rocks make approximately one third of the total volume of the intrusion (Figure 9.14). This mafic part of the body consists of rhytmically layered or massive cumulus-textured peridotites, olivine gabbronorites, gabbronorites, and gabbros, derived from a tholeiitic parental magma (Rämö, 1986). Plagioclase (An32–59), olivine (Fo35–70), titanian magnetite, ilmenite, apatite, and clinopyroxene are cumulus phases, plagioclase and orthopyroxene are intercumulus. 430
Rämö (1986) described rhytmic layering where individual 10–15-cm-thick layers consist of thin seam of magnetite in the bottom followed first by gabbro and then leucogabbro on the top. As was in the case of Kauhajärvi intrusion, also in Perämaa the Fe, Ti, and P abundances are highest in the melanocratic cumulates. The Perämaa intrusion shears many features in common with anorogenic intrusions; this reflects its emplacement into stabilized crust during the latest stages of the Svecofennian orogeny (Nironen et al., 2000). Before the final emplacement the primary melt for the Perämaa gabbro evolved under low fO2 at almost closed system, which resulted in a Ti-Fe-P-rich parental magma for the intrusion (Rämö, 1986). During crystallization, however, fO2 increased as evidenced by the gravitative accumulation of oxides at the base of individual cumulate layers. Importantly, the H2O content of the melt was low. This is in marked contrast to the synorogenic intrusions such as Kaipola that crystallized from hydrous (arc-type) magma.
6.2. Koivusaarenneva layered intrusion The Koivusaarenneva layered intrusion – the host for a major magmatic ilmenite deposit – is located ~170 km north–northeast of the Perämaa gabbro (Figure 9.1). It is an elongated, 0.5–1-km-wide and 3-km-long sill-like intrusion belonging to a suite of several similar intrusions that are found adjacent to the intersection of SW- and SE-trending fault zones. The intrusion itself bears some similarities with the intrusions of the Kauhajärvi gabbro province – it is intrusive to the Central Finland granitoid complex and shares some compositional features of anorogenic mafic plutonism. However, the Koivusaarenneva intrusion is not associated with postkinematic granites and has yielded an older zircon age of 1881 ± 6 Ma (Kärkkäinen, 1999b). Kärkkäinen (1999a) interpreted it to belong to the synorogenic group and to share common origin with other
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MAIN ZONE Anorthosite, anorthositic gabbro Gabbro, gabbronorite, olivine gabbro Peridotite BASAL ZONE Gabbro, gabbronorite, olivine gabbro Fault Lauhanvuori granite (late orogenic) Granodiorite (synorogenic)
400 m
Fig. 9.13. Geology of the Kauhajärvi gabbro (after Kärkkäinen, 1999a).
craton margin intrusions. On the basis of modal mineral variations, Kärkkäinen (1999a) divided the Koivusaarenneva intrusion into lower, middle, and upper zones characterized by titanian magnetite, ilmenite, and apatite, respectively (Figure 9.15). The 150–500-m-thick lower zone consists of layered gabbro and gabbronorite (plagioclase-pyroxene cumulates). Thin (0.2–1 m) pyroxenitic layers are spatially associated with oxide-rich layers. The lower zone hosts disseminated (8–18 vol.% ilmenite) to semimassive titanian magnetite-ilmenite ore that
is characterized by low TiO2/Fe2O3. The lower zone is, however, of only minor economic importance. The beginning of the middle zone is marked by an abrupt increase in ilmenite and magnetite and normative pyroxenes coupled with decrease in plagioclase. The middle zone hosts a 5–20-m-thick layer of massive, magnetite-poor, ilmenite ore (18–48 vol.% ilmenite) overlain by up to 40 m of disseminated ilmenite. The massive ore layers are associated with both pyroxenitic and gabbro layers, the remainder of the middle zone consists of gabbro and gabbronorite. The middle zone is the
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A
B
C
mGal
mGal/km2
-18
+10
-24
+0
-30
-10
1 km PERÄMAA GABBRO Diorite, quartz diorite, granodiorite Gabbro, ultramafic rocks
D
Lauhanvuori granite (lateorogenic) Granodiorite (synorogenic)
E
nT +4000 +2000 0
Fig. 9.14. Geological (A), Bouguer gravity anomaly (B), second vertical derivative of Bouguer anomaly (C), gray-tone low-altitude aeromagnetic (D), and obliquely illuminated aeromagnetic (E) anomaly maps of the Perämaa gabbro.Ti-rich oxide layers give rise to pronounced aeromagnetic maxima in the central part of the body. Geology after Rämö (1986). Geophysical data from the Geological Survey of Finland, processed by Seppo Elo.
economically most important ore horizon. The 300–800-m-thick upper zone is only weakly layered and is dominated by fluorapatite-rich gabbro with minor ilmenite.
7. Chemical and isotope composition of the mafic–ultramafic intrusions The chemical compositions of mafic–ultramafic intrusions are determined by initial composition of the parental magma, assimilation of country rock material, possible segregation of immiscible sulfide liquid, and by cumulus processes. As all Svecofennian intrusions are composed of various types of cumulates and fresh chilled margins are not well-preserved, the composition of the parental magma has to 432
be determined indirectly. Because of their potential for Ni-Cu ± PGE deposits, several attempts have been made to determine the parental magma composition of the Group I intrusions (Mäkinen, 1987; Lamberg, 1990; Peltonen, 1995a; Makkonen, 1996). The Mg content of the primary magma was close to 12 wt.% MgO as backcalculated from the forsterite content of the most magnesian cumulus olivine (Lamberg, 1990; Makkonen, 1996). This implies that the ultramafic cumulates were not derived from ultramafic magma but represent cumulates from more evolved melts. Peltonen (1995a) argued that the incompatible trace element composition of the most primitive olivine cumulates of the Vammala Ni province closely approximates that of the parental magma. This approach – based on the fact that olivine has
• C H A P T E R 9 • S V E C O F E N N I A N M A F I C – U LT R A M A F I C I N T RU S I O N S
Country rock (tonalite) UPPER ZONE Gabbro/gabbronorite Ilmenite and apatite
200–500 m
50–200 m Ilmenite > titanian magnetite
15–50 m 10–30 m 2–3 m
Ilmenite
Titanian magnetite > ilmenite
50–100 m
10–70 m
50 m
Leucogabbro Chilled upper zone/diabase? MIDDLE ZONE Gabbro/gabbronorite Low-grade ilmenite ore High-grade ilmenite ore Pyroxenite High-grade ilmenite ore (+magnetite) LOWER ZONE Gabbro/gabbronorite Cyclic zone: low-grade ilmenite ore gabbro pyroxenite semimassive and disseminated oxides Gabbro, coarse-grained Country rock (tonalite)
Fig. 9.15. Igneous stratigraphy of the Koivusaarenneva layered intrusion (after Kärkkäinen, 1999a).
extremely low mineral/melt partition coefficients for incompatible trace elements and that trapped melt thus determines the trace element composition of such cumulates – implied that the incompatible trace element composition of the magma was similar to that of the enclosing metaturbidites. This, together with lower than depleted mantle εNd (at 1.9 Ga) values (Fig. 7.1; Huhma, 1986, unpubl.; Makkonen, 1996), low Se/S of the sulfides (Peltonen, 1995b), and presence of graphite in the ores (Peltonen et al., 1995) shows that the trace element composition of parental magma for the Group I intrusions was strongly modified during emplacement through the Svecofennian crust. Thus it bears no unequivocal information of its mantle source. Primary magmatic intercumulus amphiboles are particularly common in the Group I bodies (Lamberg, 1990; Makkonen, 1996;
Peltonen and Elo, 1999) indicating high water content of the parental magma. This is characteristic for arc cumulates of Phanerozoic terrains (Snoke et al., 1982; Regan, 1985; Butler, 1989; DeBari and Coleman, 1989; Kepezhinskas et al., 1993; Skirrow and Sims, 1999) and led Peltonen (1995a) to suggest that the Group I bodies are cumulates of Svecofennian arc basalts. Arc-type volcanic rocks are, however, uncommon in the vicinity of the Pirkanmaa belt intrusions. The belt is characterized by metaturbidites and fragments of primitive oceanic crust and this led Lahtinen (1994) to favor an accretionary wedge setting for the intrusions of the Pirkanmaa belt. Major element ternary plots (Figure 9.16) are used to illustrate fractional crystallization of the intrusions. The crystallization started with olivine in Vammala, Porrasniemi, Laukunkangas, and Kotalahti. In the ultramafic cu-
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mulate bodies of Vammala Ni province (black circles) the extent of fractional crystallization is restricted to olivine and olivine-pyroxene cumulates. In all bodies olivine is followed by pyroxenes in the cumulate sequence. In this respect, intrusions of the craton margin environment and Pirkanmaa belt show contrasting behavior. Generally, clinopyroxene saturated after olivine in the Pirkanmaa belt intrusions (such as Porrasniemi), shifting the cumulate compositions towards the CaO-apex in the CMA-ternary (Figure 9.16). Figure 9.16 also shows that the most primitive olivinebearing cumulates are absent in Porrasniemi. This implies early segregation of significant amounts of olivine (± chromite) during ascent or intermediate magma storage – a process that significantly decreased the potential of the Por rasniemi intrusion to host magmatic sulfide deposits. Orthopyroxene dominates in the craton margin intrusions resulting in Laukunkangas-type trends with low and ~constant CaO abundances (Figure 9.16). This feature was recognized by Mäkinen (1987), who divided intrusions into Vammala-type (cpx-dominated) and Kotalahti-type (opx-dominated) and interpreted the division to reflect higher degree of mantle melting for Kotalahti-type magmas. However, crustal contamination provides an alternative explanation: extensive assimilation of country-rock sediments would increase the silica content of the melt resulting in early crystallization of orthopyroxene instead of clinopyroxene (Haughton et al., 1974). This is supported by the Nd isotope composition of the intrusions. In the Juva area (Figure 9.1), where Kotalahti- and Vammala-type bodies coexist in the same region, two Kotalahti-type intrusions yield an average εNd (at 1.9 Ga) of +0.7, whereas three Vammala-type bodies yield an average εNd(at 1.9 Ga) of +1.7. This suggests higher amount of crustal material in the Kotalahti-type magmas (Makkonen, 1996). The Laukunkangas intrusion, which is a typical Kotalahti-type intrusion and hosts a 434
major magmatic Ni-Cu sulfide deposit, has also a relatively low εNd (at 1.9 Ga) value of +0.2 ± 0.5 (Huhma, 1986). It seems reasonable to conclude that the crustal environment of emplacement had a profound effect on the chemical evolution of the melts and thus on the mineralogy of the forming cumulates. Intrusions that were emplaced through the thick sialic Archean crust or the Primitive arc complex were more likely to become contaminated by SiO2 and crystallize orthopyroxene. Intrusions in western Finland became contaminated as well, but the main contaminant was carbonaceous and calcareous sulfidic black schist material resulting in early sulfide and clinopyroxene saturation in the melt. The Group II synvolcanic intrusions of the arc complex of southern Finland have somewhat different chemical and isotope characteristics compared to Group I intrusions (Huhma, 1986; Patchett and Kouvo, 1986; Vuori, 1999). The Hyvinkää layered intrusion, for example, crystallized from a subalkalic tholeiitic magma (Vuori, 1999). Low Nb and Zr, together with slight LREE enrichment, high Sr and Ba/Rb suggest an island arc setting. In the AFM ternary diagram the Hyvinkää layered intrusion follows a typical tholeiitic trend with local production of Ti-Fe-V-rich residual melts (Figure 9.16). The compositional trend in the CMA ternary implies that the cumulate compositions are largely related to orthopyroxene fractionation – this is consistent with the noritic bulk composition of Hyvinkää. Vuori (1999) concluded that, although crustal xenoliths are locally common in the cumulate sequence of the Hyvinkää layered intrusion, the effect of contamination on the melt composition was relatively small. This is consistent with the high εNd (at 1.9 Ga) value of +2.7 (Patchett and Kouvo, 1986). The Group III Perämaa intrusion records εNd (at 1875 Ma) of –0.1 (average of five) and is less juvenile compared to the intrusions of the Arc complex of southern Finland but is similar
• C H A P T E R 9 • S V E C O F E N N I A N M A F I C – U LT R A M A F I C I N T RU S I O N S
FeOtot
MgO
Al2O3
CaO GROUP Ia
Na2O+K2O GROUP Ib
Laukunkangas (Ni-Cu)
Porrasniemi
Kotalahti (Ni-Cu)
Vammala (Ni-Cu)
Lapinlahti
Kaipola
GROUP II Hyvinkää
MgO
GROUP III Perämaa Koivusaarenneva
Fig. 9.16. Major element CMA and AFM ternary plots illustrating the trends of the fractional crystallization for selected Group I, II, and III mafic intrusions. Boundary between the calc-alkaline and tholeiitic fields in the AFM diagram is after Irvine and Baragar (1971).
to other mafic–ultramafic intrusions outside Central Finland granitoid complex (Rämö et al., 2001). The time-corrected Pb isotope ratios for Perämaa gabbro and diorite average at 206Pb/204Pb = 15.64 and 207Pb/204Pb = 15.28 (Rämö et al., 2001); these are close to those of the adjacent postkinematic granitoid plutons. They both plot close to the composition of average crustal Pb. Rämö et al. (2001) favored these values to reflect an enriched subcontinental lithospheric mantle source. However, they could also indicate pervasive interaction
of depleted mantle derived magmas and crust during the genesis of postkinematic granite plutons and associated mafic intrusions.
8. Economic aspects and petrogenesis of the ores Svecofennian mafic–ultramafic intrusions show high potential for both magmatic NiCu ± PGE sulfide deposits (Papunen and Gorbunov, 1985) and ilmenitic (FeTiO3) Ti
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ores (Kärkkäinen et al., 1997). Magmatic Ni sulfide deposits are restricted to the orogenic Group Ia and Ib bodies. The potential of the synvolcanic Group II intrusions has also been extensively evaluated, but the results have been discouraging (Häkli, 1970; Raitala, 1997; Vuori, 1999). Since the start up the production of the Makola deposit in 1941, altogether nine deposit have been exploited from Group I bodies. The total production has been 41 Mt ore with an average (weighted) grade of 0.67 wt.% Ni and 0.28 wt.% Cu. Most of the deposits have low abundances of Pt group elements. Rare examples of PGE mineralized Group I intrusions are Ekojoki (Peltonen et al., 1995) and Uudiskorhola (Papunen, 1989). Group I intrusions have been the most important source rock for Ni in Finland and their production has far exceeded that from other type of formations (e.g., Archean komatiites in eastern Finland ). Sulfide mineralogy and ore textures indicate that the Ni-Cu deposits originated as concentrations of immiscible sulfide liquid (Figures 9.6D, 9.9B). The mineralized zones are frequently located within the most primitive cumulates at the stratigraphic base of the intrusions (Papunen and Gorbunov, 1985; Mäkinen, 1987; Peltonen, 1995a; Makkonen, 1996). In addition, deformation has resulted in remobilization of primary sulfides and formation of economically important high-grade offsets in several intrusions (Kotalahti, Laukunkangas, Telkkälä). Chemical, isotope, and mineral composition of the ores require that assimilation of sedimentary rocks by the magma, combined with decreasing temperature, was the ultimate cause for the sulfide saturation and formation of immiscible nickeliferous sulfide liquid. For the deposits of the Vammala Ni province (Group Ib bodies), Peltonen (1995b) suggested that in the metamorphic environment, H2S-bearing C-O-H-S fluids were continuously produced in the surrounding schists through the conversion of pyrite to pyrrhotite in the presence of graphite. The 436
circulation of such fluids enabled selective transfer of large quantities of S and Zn from black schists into the cooling magma. In the craton margin intrusions (Group Ia) the ore genesis was basically similar, but black schist sulfides were less important contaminants. In this case, the magmas gained excess SiO2 from their country rocks, which promoted orthopyroxene crystallization and sulfide saturation (Makkonen, 1996). The Group III intrusions host important ilmenite resources. The Koivusaarenneva layered intrusion is estimated to contain 44 Mt ore down to 150 m with 15% ilmenite and 6% vanadiferous magnetite (Kärkkäinen et al., 1997). The Koivusaarenneva intrusion is currently under feasibility study. The whole Koivusaarenneva gabbro province (Figure 9.1) has exploration potential as also several other intrusions host Ti deposits. Kärkkäinen (1999a) proposed a two-stage model for the Koivusaarenneva intrusion and the oxide mineralizations. At the first stage, a primary tholeiitic arc basalt magma underwent fractional crystallization in a deep magma chamber. This took place at very low fO2 to prevent early saturation of Ti-rich oxides. In the second stage, the modified residual magma – enriched in TiO2 up to 3.3 wt.% – was emplaced at higher crustal levels to form the intrusion. The average chemical compositions of the lower, middle, and upper zones are not consistent with the origin from a single batch of magma, but requires successive magma pulses from the same source. The lower zone crystallized from a single magma pulse and fractional crystallization of ilmenite and ferroan titanian spinel led to the stratification. The middle zone and the ilmenite ore were probably formed through open-system fractional crystallization in a dynamic system as the amount of ilmenite ore exceeds what might be expected to saturate from the volume of the magma represented by the middle zone. The upper zone crystallized from a more fractionated and P-rich pulse of the parental magma.
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9. Concluding remarks The mafic–ultramafic intrusions of the Svecofennian orogen are classified into three major groups. Their areal distribution, petrology, and geochemical and isotope characteristics provide important constraints for crustal evolution. Group I intrusions include all synorogenic bodies that were emplaced either within the Archean craton margin, the Primitive arc complex or the Arc complex of western Finland (Figure 9.1). This implies that their emplacement was coeval with or slightly postdated the amalgamation of these arc complexes. Although Group I intrusions may have diverse country rocks, shapes, and to some extent crystallization order of minerals, they share a number of common features. Their U-Pb zircon systematics invariably record ages between 1.89–1.87 Ga – the time of the synorogenic stage of the orogeny. Structural analysis indicates that the bodies were emplaced during early orogenic stage into short-lived extensional structures within the arc crust. Ubiquitous boudinage and fragmentation, partial metamorphism, and extensive fluid-driven contamination are all indicative of synkinematic intrusion. Group I intrusions have high potential for magmatic Ni-Cu sulfide deposits. All intrusions within the Arc complex of southern Finland are Group II bodies. They differ from Group I intrusions in being much larger and spatially associated with metavolcanic rocks. In contrast to the Group I intrusions – which show evidence for crystallization at high confining pressures (Peltonen, 1995 a) – the Group II intrusions crystallized at low pressure. The boundary of the Häme and Pirkanmaa belts is a major tectonic boundary, no Group I intrusions are found south of it and no Group II intrusions north of it. This suggests that the emplacement of both types of mafic intrusions preceded amalgamation of these two arc complexes. This is also supported by the distinct initial εNd values for
mafic–ultramafic intrusions across the boundary. These imply more a depleted source or less crustal contamination for the Group II intrusions (Figure 9.1). The Group II intrusions – apparently because of their more evolved compositions and less dynamic crystallization regime (Maier et al., 2001) – have significantly lower potential for magmatic sulfide deposits than the Group I intrusions. Group III comprises evolved gabbroic intrusions in the Central Finland granitoid complex region. This is a relatively poorly characterized suite of bodies including gabbros gene tically related to postkinematic (~1.87 Ga) granites (e.g., Perämaa) and intrusions that share the characteristics of evolved Group I intrusions (e.g., Koivusaarenneva). It is, however, important to notice that postkinematic gabbros are not strictly restricted to the Central Finland granitoid complex. Saarisenjärvi and Tyypekinlampi (Figure 9.1) are examples of intrusions that were formed within the Primitive arc complex and have postkinematic crystallization ages (Figure 9.2; Ekdahl, 1993).
Acknowledgments Numerous individuals are acknowledged for putting their expertise, unpublished data or photographs at the author’s disposal. Thanks go to Seppo Elo, Niilo Kärkkäinen, Markku Tiainen (GTK), Pertti Lamberg (Outokumpu Research), and Riku Raitala (University of Helsinki). Heikku Papunen (University of Turku), Hannu Makkonen (GTK), and the volume editors carefully reviewed the manuscript and suggested numerous improvements.
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Burns, L.E., 1985. The Border Ranges ultramafic and mafic complex, south-central Alaska: cumulate fractionates of island-arc volcanics. Can. J. Earth Sci. 22, 1020–1038. Butler, J.R., 1989. Review and classification of ultramafic bodies in the Piedmont of the Carolinas. Geol. Soc. Am., Spec. Pap. 231, 19–31. DeBari, S., Coleman, R.G., 1989. Examination of the deep levels of an island arc: Evidence from the Tonsina ultramafic-mafic assemblage, Tonsina, Alaska. J. Geophys. Res. 94, 4373–4391. Eerola, T., Törnroos, R., Lallukka, H., Kärkkäinen, N., Valli, T., Elo, S., Tiainen, M., 2001. The Svecofennian layered mafic–ultramafic intrusions in Mäntsälä, southern Finland. In: S. Autio (Ed.), Current Research 1999–2000, Geol. Surv. Finland, Spec. Pap. 31, 17–23. Ekdahl, E., 1993. Early Proterozoic Karelian and Svecofennian formations and the evolution of the Raahe–Ladoga Ore Zone, based on the Pielavesi area, central Finland. Geol. Surv. Finland, Bull. 373, 1–137. Elo, S., 1992. Gravity anomaly maps. In: T. Koljonen (Ed.), The Geochemical Atlas of Finland. Part 2., Geol. Surv. Finland, Espoo, pp. 70–75. Gaál, G., 1972. Tectonic control of some Ni-Cu deposits in Finland. 24th Int. Geol. Congr., Montreal 4, 215–224. Gaál, G., 1980. Geological setting and intrusion tectonics of the Kotalahti nickel-copper deposit, Finland. Bull. Geol. Soc. Finland 52, 101–128. Grissom, G.C., DeBari, S.M., Page, S.P., Page, R.F.N., Villar, L.M., Coleman, R.G., de Ramirez, M.V., 1991. The deep crust of an early Paleozoic arc; the Sierra de Fiambala, northwestern Argentina. In: R.S. Harmon, C.W. Rapela (Eds.), Andean magmatism and its tectonic setting. Geol. Soc. Am., Spec. Pap. 265, 189–200. Grundström, L., 1980. The Laukunkangas nickelcopper occurrence in southeastern Finland. Bull. Geol. Soc. Finland 52, 23–53. Hakkarainen, G., 1994. Geology and geochemistry of the Hämeenlinna-Somero volcanic belt, southwestern Finland: a Paleoproterozoic island arc. In: M. Nironen, Y. Kähkönen
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(Eds.), Geochemistry of Proterozoic supracrustal rocks in Finland. Geol. Surv. Finland., Spec. Pap. 19, 85–100. Häkli, T.A., 1970. Factor analysis of the sulphide phase in mafic - ultramafic rocks in Finland. Bull. Geol. Soc. Finland 42, 109–118. Häkli, T.A., Vormisto, K., Hänninen, E., 1979. Vammala, a nickel deposit in layered ultramafite, southwest Finland. Econ. Geol. 74, 1166–1182. Haughton, D.R., Roeder, P.L., Skinner, B.J., 1974. Solubility of sulfur in mafic magmas. Econ. Geol. 69, 451–467. Helovuori, O., 1979. Geology of the Pyhäsalmi ore deposit, Finland. Econ. Geol. 74, 1084– 1101. Hildreth, W., Moorbath, S., 1988. Crustal contributions to arc magmatism in the Andes of Central Chile. Contrib. Mineral. Petrol. 98, 455–489. Honkamo, M., 1988. Explanation to the map of rocks. Geological map of Finland 1:100 000, Sheets 2533 (Haukipudas) and 3511 (Kiiminki), Geol. Surv. Finland, Espoo. Hopgood, A.M., Bowes, D.R., Kouvo, O., Halliday, A.N., 1983. U-Pb and Rb-Sr isotopic study of polyphase deformed migmatites in the Svecokarelides, southern Finland. In: M.P. Atherton, C.D. Gribble (Eds.), Migmatites, melting and metamorphism. Shiva Geology Series, 80–92. 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. Irvine, T.N., 1982. Terminology for layered intrusions. J. Petrol. 23, 127–162. Irvine, T.N., Baragar, W.R.A., 1971. A guide to the chemical classification of the common volcanic rocks. Can. J. Earth Sci. 8, 523–548. Isohanni, M., Ohenoja, V., Papunen, H., 1985. Geology and nickel-copper ores of the Nivala area. Geol. Surv. Finland, Bull. 333, 211–228. Jokela, J., 1991. Valkeakosken alueen kallioperän integroitu rakennetulkinta. M.Sc. Thesis, Univ. of Turku, Finland. (in Finnish) Kähkönen, Y., 1989. Geochemistry and petrology of the metavolcanic rocks of the early Proterozoic Tampere Schist Belt, southern Finland.
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Geol. Surv. Finland, Bull. 345, 1–107. Kärkkäinen, N., 1999a. Titanium ore potential of small mafic intrusions based on two examples in western Finland. Ph. D. Thesis, Houghton, Michigan Technol. Univ., U.S.A. Kärkkäinen, N., 1999b. The age of the Koivusaarenneva ilmenite gabbro, western Finland. In: S. Autio, Ed.), Current Research 1997–1998. Geol. Surv. Finland Spec. Pap. 27, 35–37. Kärkkäinen, N., Sarapää, O., Huuskonen, M., Koistinen, E., Lehtimäki, J., 1997. Ilmenite exploration in western Finland, and the mineral resources of the Kälviä deposit. In: S. Autio (Ed.), Current Research 1995–1996. Geol. Surv. Finland, Spec. Pap. 23, 15–24. Kepezhinskas, P.K., Reuber, I., Tanaka, H., Miyashita, S., 1993. Zoned calc-alkaline plutons in Northeastern Kamchatka, Russia: Implications for the crustal growth in magmatic arcs. Mineral. Petrol. 49, 147–174. Kerkkonen, K., 1985. Lapinlahden gabron petrografiasta, rakenteesta ja kemismistä. M.Sc. Thesis, Univ. of Oulu, Finland. (in Finnish) Koistinen, T.J. (Ed.), 1996. Explanation to the map of Precambrian basement of the Gulf of Finland and surrounding area 1:1 mill. Geol. Surv. Finland, Spec. Pap. 21. Komatsu, M., Toyoshima, T., Osanai, Y., Arai, M., 1994. Prograde and anatectic reactions in the deep arc crust exposed in the Hidaka metamorphic belt, Hokkaido, Japan. Lithos 33, 31–49. Korsman, K., Koistinen, T., Kohonen, J., Wennerström, M., Ekdahl, E., Honkamo, M., Idman, H., Pekkala, Y. (Eds.), 1997. Bedrock map of Finland 1:1 000 000. Geol. Surv. Finland, Espoo. Kukkonen, R., 1981. Painovoima-anomalioiden atktulkintaohjelmisto kolmedimensionaalisten geologisten rakenteiden tulkitsemiseksi. M.Sc. Thesis, Helsinki Univ. of Technology, Espoo, Finland. (in Finnish) Lahtinen, R., 1994. Crustal evolution of the Svecofennian and Karelian domains during 2.1–1.79 Ga, with special emphasis on the geochemistry and origin of 1.93–1.91 Ga gneissic tonalites and associated supracrustal rocks in the Rautalampi area, central
Finland. Geol. Surv. Finland, Bull. 378, 1–128 Lamberg, P., 1990. Porrasniemen intruusio, sen rakenne ja petrologia – 1.9 Ga-magmatismi ja Ni-malmit. M.Sc. Thesis, Univ. of Oulu, Finland. (in Finnish) Maier, W.D., Li, C., de Waal, S.A., 2001. Why are there no major Ni-Cu sulfide deposits in large layered mafic - ultramafic intrusions? Can. Min. 39, 547–556. Mäkinen, J., 1987. Geochemical characteristics of Svecokarelidic mafic–ultramafic intrusions associated with Ni-Cu occurrences in Finland. Geol. Surv. Finland, Bull. 342, 1–109. Makkonen, H.V., 1996. 1.9 Ga tholeiitic magmatism and related Ni-Cu deposition in the Juva area, SE Finland. Geol. Surv. Finland, Bull. 386, 1–101. Marshall, B., Smith, J.V., Mancini, F., 1995. Emplacement and implications of peridotitehosted leucocratic dykes, Vammala Mine, Finland. GFF 117, 199–205. Marttila, E., 1981. Explanation to the map of rocks. Geological map of Finland 1:100 000, Sheet 3323 (Kiuruvesi), Kiuruveden kartta-alueen kallioperä. Geol. Surv. Finland, Espoo. Naldrett, A.J., 1989. Magmatic sulfide deposits. Oxford University Press. 186 p. Neuvonen, K.J., 1956. Kallioperäkartan selitys. Summary: Explanation to the map of rocks. Geological map of Finland 1:100 000, sheet 2113 (Forssa), Geol. Surv. Finland, Helsinki. Nironen, M., 1989. Emplacement and structural setting of granitoids in the early Proterozoic Tampere and Savo Schist Belts, Finland – implications for contrasting crustal evolution. Geol. Surv. Finland, Bull. 346, 1–83. Nironen, M., Elliott, B.A., Rämö, O.T., 2000. 1.88–1.87 Ga post-kinematic intrusions of the Central Finland Granitoid Complex: a shift from C-type to A-type magmatism during lithospheric convergence. Lithos 53, 37–58. Nurmi, P.A., Front, K., Lampio, E., Nironen, M., 1984. Etelä-Suomen svekokarjalaiset porfyyrityyppiset molybdeeni- ja kupariesiintymät, niiden granitoidi-isäntäkivet ja litogeokemiallinen etsintä. Summary: Svecokarelian porphyry-type molybdenum and
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copper occurrences in southern Finland: their granitoid host rocks and lithogeochemical exploration. Geol. Surv. Finland, Tutkimusrap., Rep. Invest. 67, 1–88. Nykänen, O., 1983. Explanation to the map of rocks. Geological map of Finland 1:100 000, Sheets 4124 + 4142 (Punkaharju) and 4123 + 4114 (Parikkala), Geol. Surv. Finland, Espoo. Paavola, J., 1988. Lapinlahden kartta-alueen kallioperä. Summary: Pre-Quaternary rocks of the Lapinlahti map-sheet area. Geological map of Finland 1:100 000, sheet 3332 (Lapinlahti), Geol. Surv. Finland, Espoo. Papunen, H., 1970. Sulfide mineralogy of the Kotalahti and Hitura nickel-copper ores, Finland. Ann. Acad. Sci. Fenn., Ser. A III 109, 1–74. Papunen, H., 1989. Platinum-group elements in metamorphosed Ni-Cu deposits in Finland. In: M.D. Prendergast, M.J. Jones (Eds.), Magmatic sulphides – the Zimbabwe volume. London: The Institution of Mining and Metallurgy, pp. 165–176. Papunen, H., Gorbunov, G. I. (Eds.), 1985. Nickelcopper deposits of the Baltic Shield and Scandinavian Caledonides. Geol. Surv. Finland, Bull. 333, 1–394. Papunen, H., Koskinen, J., 1985. Geology of the Kotalahti nickel-copper ore. In: H. Papunen, G.I. Gorbunov (Eds.), Nickel-copper deposits of the Baltic Shield and Scandinavian Caledonides. Geol. Surv. Finland, Bull. 333, 228–240. Patchett, J., Kouvo, O., 1986. Origin of continental crust of 1.9–1.7 Ga age: Nd isotopes and U-Pb zircon ages in the Svecokarelian terrain of South Finland. Contrib. Mineral. Petrol. 92, 1–12. Peltonen, P., 1990. Metamorphic olivine in picritic metavolcanics from Southern Finland. Bull. Geol. Soc. Finland 62, 99–114. Peltonen, P., 1995a. Petrogenesis of ultramafic rocks in the Vammala Nickel Belt: Implications for crustal evolution of the early Proterozoic Svecofennian arc terrane. Lithos 34, 253–274. Peltonen, P., 1995b. Magma-country rock interaction and the genesis of Ni-Cu deposits in the Vammala Nickel Belt, SW Finland.
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Mineral. Petrol. 52, 1–24. Peltonen, P., 1995c. Crystallization and re-equilibration of zoned chromite in ultramafic cumulates, Vammala Ni-belt, southwestern Finland. Can. Min. 33, 521–535. Peltonen, P., Elo, S., 1999. Petrology of the Kaipola layered intrusion, southern Finland. In: S. Autio (Ed.), Current Research 1997–1998. Geol. Surv. Finland, Spec. Pap. 27, 21–24. Peltonen, P., Pakkanen, L., Johanson, B., 1995. Re-Mo-Cu-Os sulphide from the Ekojoki Ni-Cu deposit, SW Finland. Mineral. Petrol. 52, 257–264. Puustinen, K., Saltikoff, B., Tontti, M., 1995. Distribution and metallogenic types of nickel deposits in Finland. Geol. Surv. Finland, Rep. Invest. 132, 1–38. Raitala, R., 1997. Hyvinkään emäksinen kerrosintruusio. M.Sc. Thesis, Univ. of Helsinki, Finland. (in Finnish) Raitala, R. (Ed.), 2000. Hyvinkään-Mäntsälän malmiprojektin loppuraportti. Report, Ministry of Trade and Industry, Helsinki, Finland. (in Finnish) Rämö, T., 1986. Honkajoen Perämaan emäksinen intruusio – erityisesti sen gabro-osien petrografia, mineralogia ja petrologia. M.Sc. Thesis, Univ. of Helsinki, Finland. (in Finnish) Rämö, O.T., Vaasjoki, M., Mänttäri, I., Elliott, B.A., Nironen, M., 2001. Petrogenesis of the postkinematic magmatism of the Central Finland Granitoid Complex I; Radiogenic isotope constraints and implications for crustal evolution. J. Petrol. 42, 1971–1993. Regan, P.F., 1985. The early basic intrusions. In: W.S. Pitcher, R.S. Middleton (Eds.), Magmatism at Plate Edge, The Peruvian Andes. Wiley, New York, pp. 72–89. Robins, B., Gardner, P.M., 1974. Synorogenic layered basic intrusions in Seiland province, Finnmark. Norg. Geol. Unders. 312, 91–130. Sandholm, G., 1970. Om gabbroplutonen i Kaipola, Jämsä. M.Sc. Thesis, Univ. of Helsinki, Finland. (in Swedish) Schersten, A., 2001. Mafic intrusions in SW Sweden. Ph. D. Thesis, Earth Sciences Center, Göteborg University, Sweden. Skirrow, R.G., Sims, J.P., 1999. Genesis and setting
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of intrusion-hosted Ni-Cu mineralisation at Las Aguilas, San Luis Province, Argentina: implications for exploration of an Ordovician arc. Expor. Mining Geol. 8, 1–20. Snoke, A.W., Sharp, W.D., Wright, J.E., Saleeby, J.B., 1982. Significance of mid-Mesozoic peridotitic to dioritic intrusive complexes, Klamath Mountains—western Sierra Nevada, California. Geology 10, 160–166. Suominen, V., 1988. Radiometric ages on zircons from a cogenetic gabbro and plagioclase porphyrite suite in Hyvinkää, southern Finland. Bull. Geol. Soc. Finland 60, 135–140. Suominen, V., 1991. The chronostratigraphy of southwestern Finland with special reference to Postjotnian and Subjotnian diabases. Geol. Surv. Finland, Bull. 356, 1–100. Thompson, J.F.H., 1984. Acadian synorogenic mafic intrusions in the Maine Appalachians. Am. J. Sci. 284, 462–483. Vaasjoki, M., 1989. Pb and S isotopic studies on the Rauhala Zn-Cu-Pb sulphide deposit and its
environment. In: K. Kojonen (Ed.), The early Proterozoic Zn-Cu-Pb sulphide deposit of Rauhala in Ylivieska, western Finland. Geol. Surv. Finland, Spec. Pap. 11, 59–65. Vaasjoki, M., 1994. Valijärven hapan vulkaniitti: minimi Hämeen liuskejakson iäksi. Summary: Radiometric age of a meta-andesite at Valijärvi, Häme schist zone, southern Finland. Geologi 46, 91–92. Vaasjoki, M., Sakko, M., 1988. The evolution of the Raahe–Ladoga zone in Finland: isotopic constraints. In: K. Korsman (Ed.), Tectonometamorphic evolution of the Raahe–Ladoga zone. Geol. Surv. Finland, Bull. 343, 7–32. Vaasjoki, M., Pietikäinen, K., Vaarma, M., 1996. U-Pb zircon determinations from the Keikyä breccia and other sites in the Svecofennides: indications of a Svecokarelian protocrust. Bull. Geol. Soc. Finland 68, 3–10. Vuori, S., 1999. Hyvinkään emäksisen kerrosintruusion länsiosan rakenteesta ja geokemiasta. M.Sc. Thesis, Univ. of Helsinki, Finland. 1–78. (in Finnish)
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Cover page: Postkinematic quartz monzonite crosscutting mafic monzodiorite. Aplite dikes crosscut the quartz monzonite and monzodiorite. Distance between the two horizontal dikes is 15 cm. Paloinen, Toivakka, central Finland. Photo: Mikko Nironen.
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Nironen, M., 2005. Proterozoic orogenic granitoid 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. 443–480. © 2005 Elsevier B.V. All rights reserved.
The classification of the Paleoproterozoic granitoid rocks of Finland has been renewed to conform to the evolution of the Svecofennian orogen. Preorogenic granitoids (1.95–1.91 Ga) are located in the Primitive arc complex of central Finland and in northernmost Finland. They differ from the orogenic ones in having higher Fe, Mg, and Ca contents typical of primitive arc environments. Emplacement and deformation of the synorogenic rocks in southern and central Finland relate to accretion of two arc complexes at 1.91–1.89 Ga. These rocks are divided into synkinematic (1.89–1.87 Ga) and postkinematic (1.88–1.86 Ga). Both show high K contents; the postkinematic rocks also have high Fe/Mg and an overall alkaline character. The synorogenic (1.89–1.86 Ga) rocks of northern Finland comprise high-Na monzonitic and tonalitic–granodioritic suites. The lateorogenic granites of southern Finland (1.84–1.81 Ga) are peraluminous S-type granites with abundant garnet. They are associated with migmatizing veins and were formed by partial melting of Paleoproterozoic metasedimentary rocks. The postorogenic plutons of southern Finland (1.81–1.77 Ga) were emplaced at high crustal levels. They have a wide compositional range, shoshonitic affinity, and they were probably derived from an enriched lithospheric mantle source. The lateorogenic (1.84–1.80 Ga) granites of northern Finland are associated with migmatizing veins and often contain fragments of older rocks. The postorogenic granites of northern Finland (1.80–1.77 Ga) are high-level intrusions with abundant quartz porphyry and aplite dikes. The lateorogenic and postorogenic granites of northern Finland are peraluminous and alkaline. Initial εNd values are ~ +3 in the preorogenic rocks, –1.5 to +1 and ~ +3 in the synorogenic rocks, –1 to +0.5 in the postkinematic rocks of the Central Finland granitoid complex, 0 to +2 in the lateorogenic and postorogenic granites of southern Finland, and strongly negative in the lateorogenic and postorogenic rocks in areas underlain by Archean crust. These values reflect a shift from juvenile arc magmatism to mixed sources involving Paleoproterozoic and Archean components. Orogenic magmatism in central Finland started with calc-alkaline synkinematic magmatism and ended in alkaline postkinematic magmatism. A younger cycle in southern Finland commenced with peraluminous lateorogenic magmatism and ended in shoshonitic postorogenic magmatism. The first evolution line resulted from accretion of two arc complexes and subsequent stabilization of the crust in central Finland. The other reflects prolonged deformation, magmatism, and metamorphism along the southern collision zone, possibly as the result of orogenic collapse.
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1. Classification of plutonic rocks The classification of plutonic rocks with respect to orogeny originated in Finland. Pentti Eskola (1932) classified plutonic rocks in southern Finland as synkinematic, late kinematic, and postkinematic using the degree of deformation and compositional differences as criteria. In the same year, J. J. Sederholm divided the Precambrian rocks of Fennoscandia into four sedimentation cycles, each with distinct plutonism. The plutonic rocks of the first and second cycles mainly correspond to the synkinematic rocks of Eskola (1932). Some granite stocks in southern Finland and Lapland (northern Finland) belong to the third cycle, and the rapakivi granites in southern Finland (Chapter 12) are part of the fourth cycle. Later, Simonen (1960) distinguished synkinematic granodiorite, trond hjemite, charnockite, and granite provinces as well as a belt of latekinematic microcline granites. In 1980, Simonen divided the Svecofennian plutonic rocks into synorogenic, lateorogenic, and postorogenic and considered that the rapakivi granites are anorogenic. Along with the acceptance of the plate tectonic theory the magmatectonic classification of plutonic rocks was renewed to conform to the new concepts of crustal evolution. An attempt was focused to unravel which areas are remnants of ancient ocean floor or magmatic arc and which are part of an older continental crust onto which arcs collided. The following
classification of magmatism corresponds to a plate tectonic application of the classic (geosynclinal) orogenic hypothesis: (1) island arc magmatism preceding onset of accretion or collision (preorogenic stage); (2) magmatism associated with accretion of arc(s) to a continent or collision of two continents (synorogenic stage); (3) magmatism associated with large horizontal movements and extension of orogenically thickened crust (late- and postorogenic stages). As mid-ocean ridges and island arcs ultimately collide with continents, virtually all plutonic rocks except intracontinental suites will be deformed in crustal movements. This leads to a general problem of the tectonic setting of plutonic rocks: the style and degree of deformation does not necessarily indicate tectonic regime. Contemporaneous plutons may appear syntectonic or posttectonic with respect to a specific deformation, depending on their emplacement level (e.g., Brown, 1994). The classification according to geological environment (e.g., Pitcher, 1982) was aimed for Phanerozoic granitoid rocks but it has also been applied to the Precambrian. The tectonic classification based on trace elements (Pearce et al., 1984) has been widely used, but the complex history of granitoid rocks (e.g., geochemical inheritage in the source area,
Table 10.1. Classifications of Proterozoic orogenic granitoid rocks in Finland. Simonen (1960)
Simonen (1980)
This study
synkinematic (1.85–1.75 Ga)
synorogenic (1.90–1.86 Ga)
latekinematic
lateorogenic postorogenic (~1.80 Ga)
preorogenic (1.95–1.91 Ga) synorogenic (1.89–1.86 Ga) synkinematic (1.89–1.87 Ga) postkinematic (1.88–1.86 Ga) lateorogenic (1.84–1.80 Ga) postorogenic (1.81–1.77 Ga)
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Inari area
Hetta complex
Central Lapland granitoid complex
Preorogenic rocks Synorogenic rocks Lateorogenic rocks AWF Central Finland granitoid complex
Postorogenic rocks
PA
Rapakivi granite Supracrustal rocks Archean rocks
ASF
100 km Helsinki
Fig. 10.1. Paleoproterozoic felsic and intermediate plutonic rocks in Finland (modified from Korsman et al., 1997, and Koistinen et al., 2001). Terrane boundaries (in blue) between the Arc complex of southern Finland (ASF), Arc complex of western Finland (AWF), and Primitive arc complex (PA) are from Korsman et al. (1997).
mixing of magmas) hamper, often severely, the interpretation of discrimination diagrams. There are indications that the Svecofennian bedrock grew by sequential accretion of arcs and probably includes several collision zones and remnants of marine basins (Lahtinen, 1994; Nironen, 1997). Therefore, different parts of the Svecofennian bedrock
underwent orogenic processes at different times: arc magmatism took place in one area and initial accretion in another area, and the crust was solidified and subject to isostatic uplift in one area, while elsewhere folding and magmatism were still in operation. A renewed classification of the orogenic granitoid rocks of Finland (including quartz
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Preorogenic rocks Synkinematic rocks Synorogenic rocks (northern Finland) Postkinematic rocks Lateorogenic rocks Postorogenic rocks
1750
1800
1850
1900
1950 Ma
Fig. 10.2. Age histogram of the Paleoproterozoic granitoid rocks in Finland (Geological Survey of Finland database). Each individual age is shown as a rectangle; question marks denote ages with two sigma error larger than ±10 Ma. Numbers in the 1880–1890 Ma column indicate number of individual analyses not shown.
monzonite, quartz monzodiorite, and quartz diorite) is shown in Table 10.1. This is based on new geochronological data and recent concepts of the evolution of the Svecofennian but retains the general terminology of Simonen (1980). The distribution of the rocks is shown in Figure 10.1, and the age data that form the basis of the classification are shown in Figure 10.2. The preorogenic rocks were generated in an island arc environment and were placed in their present location during the Svecofennian orogeny. Zircons dated from volcanic rocks, granitoids, and mafic plutonic rocks of southern Finland cluster in the age range of 1.89–1.87 Ga (Vaasjoki, 1996). This time interval also includes prominent tectonic activity in the Finnish Svecofennian. The emplacement and deformation of the synorogenic rocks is assigned to the accretion of three arc com448
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plexes to the Archean craton: the Primitive arc complex and the Arc complex of western Finland 1.91 Ga ago and the Arc complex of southern Finland 1.89 Ga ago (Figure 10.1; Lahtinen, 1994; Nironen, 1997). Field studies have shown that some synorogenic rocks in central Finland crosscut their host (synorogenic) plutonic rocks. These rocks have been divided into syn- and postkinematic groups with respect to prominent deformation within the area in question. Recent geochemical and isotope studies have confirmed this distinction (Nironen et al., 2000; Rämö et al., 2001). The lateorogenic granites of southern Finland are located in the collision zone between the two southern arc complexes and are associated with metamorphism and low-angle crustal movements. The postorogenic plutons of southern Finland are located within the same
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zone but, in contrast to the lateorogenic granites, were emplaced along faults and shear zones within a largely stabilized crust.
2. Preorogenic rocks 2.1. Preorogenic rocks of central Finland (1.93–1.91 Ga) The Primitive arc complex adjacent to the southwestern margin of the Archean bedrock (Figure 10.1) contains 1.93–1.91 Ga gneissic tonalites, which are associated with a bimodal volcanic sequence with numerous interlayered Zn-Cu deposits and graywacke-type mica gneisses (Ekdahl, 1993). They are often found in the center of domal structures and are surrounded by low-K tholeiitic to calc-alkaline basalts, andesites, and low-K felsic volcanic rocks with arc affinity. These gneissic tonalites are the oldest plutonic rocks so far identified in the Finnish Svecofennian. A 1921 ± 2 Ma rhyolite (Kousa et al., 1994) indicates that the 1.93–1.91 Ga tonalitic plutonism was associated with volcanism. The gneissic tonalites are pervasively foliated and range from tonalite to granodiorite. They vary in appearance from felsic, homogeneous, and intrusive-like to migmatitic. The main minerals are plagioclase, quartz, biotite, and occasionally hornblende, with orthopyroxene in some of the tonalites. The tonalites contain abundant mafic and felsic enclaves. Some of these are disrupted dikes, whereas others appear to be country rock xenoliths. In some domal areas the composition of the mafic enclaves matches the composition of the surrounding mafic volcanic rocks. Ekdahl (1993) considered the enclaves intrusive equivalents of the overlying volcanic rocks. This leads to a yet unresolved problem regarding the geologic setting of the gneissic tonalites: do they represent the basement of the supracrustal rocks (Ekdahl, 1993) or are they sills that intruded the supracrustal rocks (Lahtinen, 1994)? An-
other problem is the tectonic setting of the rocks: were they formed at the Archean craton margin as the result of accretion of the Arc complex of western Finland (Ekdahl, 1993), or did the Primitive arc complex unite with the Arc complex of western Finland elsewhere and dock on the craton margin at 1.91 Ga (Lahtinen, 1994; Nironen, 1997)?
2.2. Preorogenic rocks of northern Finland (1.95–1.91 Ga) In northernmost Finland, within the Inari area, (Figure 10.1) there are foliated plutonic bodies that usually have concordant contacts with the surrounding Paleoproterozoic gneisses. The composition of the bodies varies from peridotite through gabbro to granodiorite and they probably consist of several intrusive units. The felsic rocks of the group have yielded zircon ages of 1.95–1.93 Ga (Meriläinen, 1976). As the tectonic character of the Inari area is not well constrained, little can be said about the tectonic setting of these rocks. The Hetta complex, consisting of tonalitic–granitic intrusive rocks as well as remnants of supracrustal rocks, is enigmatic in age and tectonic setting. A group of flat-lying tonalite and granodiorite bodies with subhorizontal foliation are located within the allochtonous metavolcanic rocks to the east of the Hetta complex and it is unknown whether also the intrusive rocks are allochtonous. A tonaliticgranitic body of this group yielded an age of 1.91 Ga (Rastas et al., 2001), and a granite of the Hetta complex was dated at 1.81 Ga (Huhma, 1986). Korsman et al. (1997) assigned these rocks to the ~1.8 Ga age group, whereas Lehtonen et al. (1998) considered the Hetta complex to be synorogenic (~1.88 Ga); here the rocks are considered preorogenic. In the recent lithological map of Fennoscandia (Koistinen et al., 2001) the Hetta complex is assigned to the 1.96–1.91 Ga age group.
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Oulu
lmi Iisa x ple com
Veteli
Vaasa
Central Finland granitoid complex 10 Puruvesi
7
8
Tampere
9
6
Kalanti Oripää
4 3
5
Turku
100 km
2
1
Helsinki
Hanko
Preorogenic rocks
Lateorogenic rocks
Supracrustal rocks
Synkinematic rocks
Postorogenic rocks
Archean rocks
Postkinematic rocks
Rapakivi granite
Fig. 10.3. Paleoproterozoic felsic and intermediate plutonic rocks in southern Finland (modified from Korsman et al., 1997). The western margin of the Archean craton (paleosuture) is shown by a dotted line. Postkinematic microtonalite dikes are found west of the dashed line. Blue line as in Figure 10.1. Postorogenic intrusions: 1–Lemland, 2–Mosshaga, 3–Seglinge, 4–Åva, 5–Turku, 6–Renko, 7–Parkkila, 8–Luonteri, 9–Eräjärvi, 10–Pirilä.
3. Synorogenic rocks Synorogenic plutonic rocks, 1.89–1.87 Ga in age, are found in all the three arc complexes of southern and central Finland (Figure 10.3). They are especially abundant in the ~40,000 km2 Central Finland granitoid complex that 450
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covers most of the Arc complex of western Finland. The synorogenic rocks are subdivided into synkinematic (1.89–1.87 Ga) and postkinematic (1.88–1.86 Ga) with reference to prominent deformation in the area in question. In central Finland the prominent deformational event is dated at 1885–1880 Ma (Nironen,
PROTEROZOIC
OROGENIC
GRANITOID
ROCKS
A
B
C
D
Fig. 10.4. Synkinematic rock types of southern and central Finland. (A) Gneissic, folded and sheared granodiorite of the Central Finland granitoid complex. Length of code bar 12 cm. (B) Coarse-porphyritic granodiorite of the Central Finland granitoid complex. Length of code bar 10 cm. (C) Mafic enclaves in a granodiorite of the Central Finland granitoid complex. Diameter of lens cap 5.5 cm. (D) Strongly foliated tonalite near Turku, southwestern Finland. Length of code bar 10 cm. Photos: Mikko Nironen.
1989; Hölttä, 1995; Mäkitie, 1999; Mouri et al., 1999). The overlapping ages of the two groups show that, at 1.88–1.87 Ga, some parts of the Svecofennian crust were subject to penetrative deformation and synkinematic magmatism, whereas in other areas the postkinematic stage had already been reached.
3.1. Synkinematic rocks of southern and central Finland (1.89–1.87 Ga) Synkinematic rocks of the Central Finland granitoid complex are typically mediumgrained granodiorites and granites; mediumgrained tonalites are found especially along the western margin of the complex (Figure
10.3). They exhibit a foliation that varies from slight fabric to pervasive gneissic foliation (Figure 10.4A). Coarse-porphyritic granodiorite and granite with abundant subhedral to anhedral alkali feldspar megacrysts 1–4 cm in diameter are located in the central and northern parts of the complex (Figure 10.4B). The granites contain biotite as the only main mafic mineral; accessory minerals are hornblende, titanite, apatite, magnetite, and zircon. Both hornblende and biotite are present in the tonalites and granodiorites, and the tonalites may also contain clinopyroxene or orthopyroxene. Typical accessory minerals are titanite, apatite, magnetite, and zircon. The minor mafic rocks mostly consist of quartz diorite
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and diorite. The main mafic minerals in the dioritic rocks are biotite and hornblende. The dioritic rocks usually contain so much feldspar that they have a monzonitic affinity. Further accessory minerals include clinopyroxene, quartz, orthopyroxene, magnetite, apatite, titanite, and zircon. Abundant mafic enclaves in the tonalites and granodiorites (Figure 10.4C) suggest that mixing and mingling processes were important; in the granites mafic enclaves are sparse. The felsic to intermediate synkinematic intrusive rocks grade in places into subvolcanic quartz-feldspar and feldspar porphyries (Nironen, 2003). Synkinematic rocks of the supracrustaldominated belts surrounding the Central Finland granitoid complex are generally granodiorites and tonalites. They are usually oval or roundish intrusions and exhibit a foliation that is locally conformable with the country rock contacts. In places the orientation extends from the plutonic rock into the country rock – a typical feature for syntectonic intrusions (Figure 10.5). Most of the plutons are multiphase and show normal zoning with tonalitic marginal parts and granodioritic to granitic central parts; reversely zoned plutons are also found (Nironen, 1989; Nironen and Bateman, 1989). Biotite is generally more common than hornblende, and accessory clinopyroxene is some times present. Other accessory minerals are titanite, magnetite, apatite, and zircon. The granitoids contain mafic enclaves, and may also contain quartz dioritic or dioritic phases. These features suggest that mafic mantle-derived magma and felsic crustal magma mixed producing a variety of hybrid rocks (Nironen and Bateman, 1989; Lahtinen, 1996). Abundant mafic and, less frequently, intermediate dikes are found among migmatitic mica gneisses that fringe the Central Finland granitoid complex in the south and east (Figure 10.3). These dikes crosscut the earliest deformation structures but have been broken or boudinaged during subsequent deformation, hence they may be considered synkinematic dikes. 452
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In southwestern Finland intensely foliated tonalites are found as concordant, sheetlike intrusions within the supracrustal rocks (Figure 10.4D). There are also tonalites and granodiorites varying in age from 1.87 Ga to 1.84 Ga (Suominen, 1991; van Duin, 1992; Nironen, 1999). These rocks become pyroxene-bearing as the grade of metamorphism in the host rocks increases. The considerable age variation may arise from growth of secondary zircon during a high grade metamorphic event 1.82 Ga ago (Väisänen et al., 2002), and thus the crystallization age of these rocks is probably ~1.87 Ga.
3.2. Postkinematic rocks of central Finland (1.88–1.86 Ga) Within the Central Finland granitoid complex there is a suite of 1.88–1.87 Ga quartz monzonitic, granodioritic, and granitic plutons that crosscut the 1.89–1.88 Ga synkinematic rocks (Elliott et al., 1998; Nironen, 2003). These are considered postkinematic because they are usually unoriented or only slightly oriented and truncate the foliation of the synkinematic rocks. They are found all over the Central Finland granitoid complex and also outside the complex – however, not in the Arc complex of southern Finland (Figure 10.3). A postkinematic sequence, consisting of hypersthene granite, monzogranite, and coarse-porphyritic granite, was emplaced at 1886–1880 Ma in the northeastern part of the complex (Nironen and Front, 1992). The postkinematic plutonic magmatism shifted from northeast toward west during 1885–1870 Ma (Rämö et al., 2001). The postkinematic plutons are located at or close to major crustal shear zones. These plutons were emplaced within an extensional or transtensional environment and shear zones controlled their emplacement. A couple of small monzodioritic to gabbroic bodies are associated with the postkinematic plutons of the Central Finland granitoid complex. Magma mingling textures indicate that
PROTEROZOIC
OROGENIC
GRANITOID
ROCKS
A
Central Finland granitoid complex Lake Näsijärvi
H
V Tampere
B
10 km
C
3 km
D
3 km
E
Erosional surface
>
> >
3 km
Diorite
Granodiorite
Granitoid rocks
Quartz monzodiorite
Granite
Metavolcanic rocks
Tonalite
Feldspar porphyry
Metasedimentary rocks
Fig. 10.5. Two synkinematic plutons near the southern margin of the Central Finland granitoid complex. (A) The location of the Hämeenkyrö (H) and Värmälä (V) plutons. (B) Lithology of the Hämeenkyrö pluton. (C) Lithology of the Värmälä pluton. (D) LS texture in the Värmälä pluton and trend lines of foliation (local S1) within and outside the Värmälä pluton. The continuation of foliation into the pluton and other contact phenomena (Nironen, 1989) indicate that the pluton is syntectonic with respect to local D1 deformation. (E) Interpretation of the emplacement of the Värmälä pluton (Nironen, 1989). Plagioclase accumulated at the base of a magma chamber in deeper crust. The dioritic body was emplaced first, then the granitic and granodioritic phases. The quartz monzodioritic magma, containing plagioclase cumulates, was emplaced as the last phase in the center of the pluton. CHAPTER 10 • PROTEROZOIC OROGENIC GRANITOID ROCKS •
453
A
B
C
D
Fig. 10.6. Postkinematic rock types of the Central Finland granitoid complex. (A) A monzodioritic enclave in quartz monzonite. Note the reaction rim and the felsic clots in the monzodiorite, suggesting mingling between a monzodioritic and a quartz monzonitic magma. Diameter of coin 2.2 cm. (B) Coarse-grained granite. Diameter of coin 2.4 cm. (C) Pyroxene-bearing marginal variety of a porphyritic quartz monzonite. Length of code bar 10 cm. (D) Postkinematic porphyritic quartz monzonite. Note the alkali feldspar crystals mantled by plagioclase. Length of code bar 10 cm. Photos: Mikko Nironen.
the mafic rocks are coeval with the felsic rocks (Figure 10.6A). The postkinematic plutons are multiphase intrusions that can be divided into three types (Elliott et al., 1998). Type 1 plutons are biotite granodiorites and granites that are found along the southern margin of the Central Finland granitoid complex. They are coarseporphyritic rocks with abundant orthoclase megacrysts. Accessory minerals are fluorite, zircon, apatite, and ilmenite. Type 2 plutons are coarse-porphyritic or equigranular granites that vary in grain size from medium to coarse (Figure 10.6B). The main mafic mineral is biotite but some plutons contain amphibole; the mafic minerals are generally interstitial and 454
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thus rather late in the crystallization sequence. Fluorite is a characteristic accessory mineral of the Type 2 plutons. Other accessory minerals include apatite, zircon, titanite, and allanite. Type 3 plutons either have a pyroxene-bearing margin (Figure 10.6C) or contain pyroxene through out. Large orthoclase megacrysts mantled by plagioclase are typical of the Type 3 quartz monzonites (Figure 10.6C, D). Especially the Type 2 granites in the western part of the granitoid complex resemble the rapakivi granites in their mineralogy and geochemical characteristics (Nironen et al., 2000). Typical structural features of the postkinematic plutons are shown in Figure 10.7. A group of 1.87–1.86 Ga quartz diorites,
PROTEROZOIC
OROGENIC
GRANITOID
ROCKS
+
Porphyritic granite Even-grained granite Pyroxene-bearing quartz monzonite Gabbro
Location of Figure 10.6C
Synkinematic granitoid rocks Intermediate and felsic metavolcanic rocks Mica gneiss Foliation form line Shear zone Shear sense 5 km
Fig.10.7. Two postkinematic granitoid plutons in the southern part of the Central Finland granitoid complex (modified from Nironen et al., 2000). The northern one with pyroxene-bearing margin is a Type 3 pluton, and the southern one is Type 2.
granodiorites, and granites is found east of the suture that marks the boundary of the Archean craton (Figure 10.3; Huhma, 1986; Ruotoistenmäki et al., 2001). The rocks are late in the intrusion sequence and some are associated with wrench faults (Halden, 1982). These features are typical of postkinematic plutons and the rocks are thus considered postkinematic. A swarm of dikes is found in a broad zone in central Finland, extending from the Archean Iisalmi terrain to the northeastern part of the Central Finland granitoid complex (Figure 10.3). These dikes mark a prominent plutonic event in the zone and are especially abundant in the east. The dikes are generally fine-grained and quartz dioritic to tonalitic; hence they have been called microtonalite dikes (Huhma, 1981). The dike swarm is complex in composition, origin, and age. The dikes are generally deformed and they crosscut sedimentary rocks, igneous rocks, and each
other. They were emplaced at 1.89–1.85 Ga, probably in several periods. They may be divided into homogeneous (Figure 10.8A) and composite (Figure 10.8B) (Rautiainen, 2000). The former probably resulted from mixing and the latter from mingling of mafic and felsic magmas.
3.3. Synorogenic rocks of northern Finland (1.89–1.86 Ga) In western Lapland (Figure 10.9), there are granitoids that are broadly coeval with the synorogenic granites of central and southern Finland. These rocks consitute the Haaparanta (Haparanda) suite and they are found on both sides of the Finland–Sweden border (e.g., Lehtonen et al., 1998; Bergman et al., 2001). The synorogenic rocks in western Lapland may be divided into two groups: (1) monzonites, consisting of gabbro, quartz monzodiorite, monzodiorite, monzonite, and quartz
CHAPTER 10 • PROTEROZOIC OROGENIC GRANITOID ROCKS •
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B A Fig. 10.8. Microtonalite dikes from central Finland. (A) Homogenous, fine-grained microtonalite crosscuts Archean gneiss and is crosscut by another homogeneous microtonalite dike. Width of the older dike 30 cm. (B) Composite dike, consisting of microtonalite with mafic encaves, crosscuts mica gneiss. Length of code bar 10 cm. Photos: Olli Äikäs.
monzonite; and (2) trondhjemites, consisting of quartz diorite, tonalite, trondhjemite, and granodiorite (Lehtonen, 1984). The monzonites are darkish, medium- to coarse-grained, and foliated and contain mafic enclaves. They have conformal or tectonic contacts with host supracrustal rocks. The most common mafic mineral is hornblende. Titanite, epidote, magnetite, apatite, and zircon are accessory minerals. The trondhjemites are pale, medium-grained and generally weakly foliated rocks that consist mainly of plagioclase and quartz; the amount of alkali feldspar varies. Hornblende, clinopyroxene, and biotite are found as primary minor minerals, and magnetite, titanite, apatite, and zircon are accessory minerals. Age determinations of the diorites and monzonites have yielded Paleoproterozoic ages (1.89–1.86 Ga), whereas the zircons in the trondhjemites are mixed populations with an Archean component (Hiltunen, 1982; Lehtonen, 1984; Huhma, 1986; Perttunen, 1991; Väänänen, 1998; Perttunen and Vaasjoki, 2001; Väänänen and Lehtonen, 2001). A monzonite with an εNd (at 1880 Ma) value of –3.2 and depleted mantle model age of 2.42 Ga (Chapter 4, Table 4.1.) suggests an Archean component also in the monzonites. 456
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4. Lateorogenic granites 4.1. Lateorogenic granites of southern Finland (1.84–1.81 Ga) In southern Finland, the lateorogenic granites constitute an ENE-trending, roughly 100-kmwide zone (Figure 10.3). The zone is characterized by migmatization of supracrustal rocks, and the plutonic rocks are mainly granites containing garnet and cordierite (e.g., Nurmi and Haapala, 1986). The paucity of intermediate and mafic plutonic rocks is characteristic. The mode of occurrence of the granites varies: they may be homogeneous, coarse-grained, locally porphyritic rocks, small intrusions of equigranular granite, or may grade from fairly homogeneous garnet-bearing rocks to granite vein systems that migmatize supracrustal rocks. Many localities within the zone exhibit gradation from a schist or gneiss cut by granite dikes to a heterogeneous diatexitic rock with remnants of older rock (Figure 10.10A), and further into a homogeneous plutonic rock. Lateorogenic granites may also brecciate synorogenic plutonic rocks (Figure 10.10B). The lateorogenic granites are generally found as flat-lying sheets. Tectonic activity continued within the zone of lateorogenic
PROTEROZOIC
OROGENIC
GRANITOID
ROCKS
Inari area Vainospää
Preorogenic rocks Hetta complex
Synorogenic rocks
Nattanen Tepasto
Riestovaara Pomovaara
Lateorogenic rocks Postorogenic rocks Supracrustal rocks
Central Lapland granitoid complex
Archean rocks
Rovaniemi
Rovaniemi
100 km Fig. 10.9. Paleoproterozoic felsic and intermediate plutonic rocks in northern Finland. Some of the pre- and postorogenic plutons are indicated by their names. (Modified from Korsman et al., 1997).
granites longer than elsewhere in southern and central Finland. This is indicated by the apparent association of the granites with overthrust structures. Ehlers et al. (1993) suggested that the granites intruded as sheets between older rocks along subhorizontal thrust planes during transpressional deformation and were locally folded into vertical position by subsequent open folding. In line with this, Selonen et al. (1996) proposed strike-slip dilatacy pumping for the emplacement mechanism of a flat-lying granite sheet in southernmost Finland. In contrast, Korja and Heikkinen (1995) presented a model favoring emplacement during crustal extension. Nurmi and Haapala (1986) concluded that the lateorogenic granites of southern Finland
were formed when Proterozoic sedimentary rocks were partially melted in the continental crust. Thus the granites migmatize rocks similar to those they originated from. As the grade of metamorphism rose, the melt collected into pressure minima such as fractures and shear zones (Figure 10.10C). These melts may have remained in the shear zones or escaped to form a granite magma. Granite dikes that alternate with layers of garnet-bearing country rock represent granite that has been segregated from the original site of melting. The melt and unmelted rocks also rose together to form migmatite domes (Korsman et al., 1984; Bleeker and Westra, 1987). Some granites contain felsic layers that probably represent melts that accumulated in shear planes (Figure 10.10D).
CHAPTER 10 • PROTEROZOIC OROGENIC GRANITOID ROCKS •
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A
B
C
D
Fig. 10.10. Various types of lateorogenic granites in southern Finland. (A) A diatexitic granite with ghost-like remnants of older gneiss. Diameter of lens cap 5.5 cm. (B) Lateorogenic granite brecciates synkinematic granodiorite. (C) Migmatitic garnet-cordierite gneiss in which melt has accumulated into shear zones during syn-anatexis deformation. (D) Banded, garnet-bearing granite. Photos: Mikko Nironen.
The ages of the lateorogenic granites in the western part of the zone vary between 1.84 and 1.83 Ga (Huhma, 1986; Suominen, 1991). A study of a granite–monzodiorite association at Turku yielded a ~1815 Ma age (Väisänen et al., 2000). The age indicates that the Turku magmatism is postorogenic (see below) but geochemically the granite is similar to the lateorogenic granites.
4.2. Lateorogenic granites of northern Finland (1.84–1.80 Ga) Mapping of the Central Lapland granitoid complex (Figure 10.9) dates back to the beginning of the 20th century, and the rock distribution of the region is poorly known. 458
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The dominant rock type is a coarse-grained, heterogeneous granite with abundant fragments of supracrustal rocks. Available zircon ages vary rather susbtantially (2.14–1.77 Ga; Lauerma, 1982; Huhma, 1986). The youngest zircon age (1.77 Ga) was obtained from the Rovaniemi granite that probably belongs to the postorogenic group (see below). Recent monazite and titanite ages (Corfu and Evins, 2002) show that granite magmatism, associated with regional deformation and metamorphism, occurred as late as 1.77 Ga ago in the Central Lapland granitoid complex. The microcline granites in western Lapland are found both as plutons crosscutting all other rocks and as heterogeneous, migmatizing veins associated with regional meta-
PROTEROZOIC
OROGENIC
GRANITOID
ROCKS
B
A
C Fig. 10.11. Rock types in the postorogenic Åva ring complex. (A) Porphyritic Åva granite crosscuts supracrustal gneiss. Length of code bar 12 cm. (B) Lamprophyre dike crosscuts Åva granite. The dike is ~50 cm wide. (C) Åva granite, containing a gneiss xenolith, crosscuts a monzonitic member. Length of pen 13.5 cm. Photos: Mikko Nironen (A) and Veli Suominen (B, C).
morphism (Lehtonen, 1984; Perttunen et al., 1996; Väänänen, 1998). The plutons consist of reddish, coarse-grained granite, typically with fragments of supracrustal rocks. Biotite is commonly the only mafic mineral, and magnetite is an abundant accessory phase; some plutons contain also tourmaline. A zircon age of 1.81 Ga was obtained from such a granite, and an age of 1778 Ma from a pegmatitic granite (Lehtonen, 1984; Lehtonen, 1988; Väänänen and Lehtonen, 2001). In Sweden, corresponding plutons are considered to belong to the 1.81–1.78 Ga Lina granite–pegmatite association (Bergman et al., 2001). The ~1.81 Ga porphyritic granite southeast of Oulu and granites crosscutting the Archean complex (Figure 10.3) presumably belong to the same 1.8 Ga age group as the
Central Lapland granitoid complex (see also Korsman et al., 1997).
5. Postorogenic rocks 5.1. Postorogenic rocks of southern Finland (1.81–1.77 Ga) The postorogenic rocks of southern Finland are found as relatively small intrusions that roughly follow the northern boundary of the of lateorogenic granite zone to Russia (Figure 10.3; Eklund et al., 1998). Their ages range between 1815 Ma and 1770 Ma (Vaasjoki and Sakko, 1988; Suominen, 1991; Vaasjoki, 1996; Väisänen et al., 2000). The intrusions are generally rounded (diameter 2–15 km);
CHAPTER 10 • PROTEROZOIC OROGENIC GRANITOID ROCKS •
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B
A Monzonite
Lateorogenic granite
Granite
Gneiss
Lamprophyre dike
Fig. 10.12. (A) Postorogenic Åva ring complex (after Ehlers and Bergman, 1984). (B) Emplacement mechanism of the Åva monzonite and granite as suggested by Bergman (1986). The emplacement of a lateorogenic granite plastically deformed the host gneisses. The monzonite intruded the lateorogenic granite, and pearced upward by stoping and wedging. Volatiles and excess magma pressure caused fracturing and uplifting of the roof. The granite followed the hot trail of monzonitic precursors, subsequently widening the funnel laterally. The hatched line shows the present level of erosion.
the Åva, Seglinge and Mosshaga intrusions are ring complexes and also the Lemland intrusion has concentric compositional and structural features. In contrast to the other intrusions, Parkkila and Eräjärvi are dike-like bodies with lengths of several kilometers. All intrusions sharply crosscut their host rocks (Figure 10.11A). Large compositional variation (monzodiorite to granite) is present in the intrusions of the Åland islands and at Luonteri whereas the others are more homogeneous (Parkkila is granodioritic and Pirilä and Eräjärvi are granitic). The Luonteri intrusion is a funnel460
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shaped multiphase pluton that consists of an older tonalitic phase followed by granodiorite and granite (Pitkänen, 1985). Biotite is the main mafic mineral, hornblende is an abundant mineral in some granodiorite varietes. Typical accessory minerals are titanite, apatite, magnetite, zircon, and allanite. Chlorite and fluorite are found as alteration products of biotite, and primary fluorite is present in the late dike rocks. Lamprophyric dikes are associated with the Åva and Seglinge ring complexes (Figure 10.11B). Bimodal lamprophyre–granite magmatism has resulted in frequent magma
PROTEROZOIC
OROGENIC
GRANITOID
ROCKS
Coarse-porphyritic granite
Location of Figures 13B and 13C
Porphyritic granite Even-grained granite Granite porphyry Rhyolite dike Granite gneiss Mafic and ultramafic metavolcanic rock Quartzite and arkosite Mica gneiss 0
3
6 km
A 6 mm
B
C
Fig. 10.13. (A) Map of the postorogenic Riestovaara pluton (modified from Front et al., 1989). Part of the Nattanen pluton is seen in the upper right corner. A possible unexposed postorogenic pluton is shown as dashed line on the basis of a positive magnetic anomaly. (B) A boulder close to the margin of the postorogenic Nattanen pluton, consisting of spherulitic rhyolite dike (right) and host mica schist (left). Diameter of coin is 2.6 cm. (C) Photomicrograph of the spherulitic rhyolite in Figure 10.13B. Photos: Kai Front.
mixing and mingling structures (Hubbard and Branigan, 1987; Branigan, 1989; Eklund et al., 1998; Figure 10.11C). Lindberg and Eklund (1988) compared the geochemical features and contact relationships of mafic, intermediate, and granitic rocks in the Lemland area and
considered that both chemical and mechanical mixing occurred at several stages in a zoned magma chamber during upward movement. According to Bergman (1986), the Åva monzonite intruded a lateorogenic granite as branching concentric dikes by stoping, and
CHAPTER 10 • PROTEROZOIC OROGENIC GRANITOID ROCKS •
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the subsequent granite widened the funnel laterally (Figure 10.12). The intrusion mechanism of the Åva ring complex and tuffisites at Seglinge (Hubbard and Branigan, 1987) imply emplacement of these ~1.80 Ma rocks into rigid country rocks at a shallow depth (a few kilometers). In contrast, a thermobarometric study of the 1815 Ma postorogenic rocks in the Turku area indicates an emplacement pressure of 4.1 kbar, corresponding to a minimum depth of 14–15 km (Väisänen et al., 2000).
5.2. Postorogenic granites of northern Finland (1.80–1.77 Ga) The postorogenic granites of northern Finland (the Nattanen type granites; Mikkola 1928, 1941) extend from western Lapland to Murmansk in Russia. In Finland, the Tepasto, Pomovaara, Riestovaara, Nattanen, and Vainospää plutons (Figure 10.9) belong to this group. Zircon age from the postorogenic granites range from 1.80 Ga to 1.77 Ga (Huhma, 1986; Rastas et al., 2001); hence they are broadly coeval with the lateorogenic granites of northern Finland and the postorogenic rocks of southern Finland. The postorogenic granites are generally seen in aeromagnetic maps as positive anomalies. Also the 1.77 Ga Rovaniemi granite (Figure 10.9) is seen as a positive anomaly distinct from the migmatitic lateorogenic granites of the Central Lapland granitoid complex. The Rovaniemi granite crosscuts migmatitic structures (Perttunen et al., 1996) and is thus considered postorogenic. The postorogenic granite plutons of northern Finland are rounded or elongated, unmetamorphosed and generally undeformed and sharply crosscut their host rocks (Haapala et al., 1987; Front et al., 1989). Individual plutons are multiphase and show a zonal structure with the oldest phases dominating the area of intrusion and the younger phases generally in the center of the plutons. Abundant quartz porphyry and aplite dikes are characteristic of 462
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the postorogenic granites. A typical intrusion sequence is seen in the Riestovaara pluton (Figure 10.13A): a coarse-porphyritic granite was emplaced first, then porphyritic granite, and finally an even-grained granite. Granite porphyry and aplite dikes are the latest phases. Spherulitic rhyolite dikes (Figure 10.13B, C) are found adjacent to the Nattanen pluton. These dikes indicate high-level emplacement and may be associated to an unexposed postorogenic pluton between the Riestovaara and Nattanen plutons (hatched line in Figure 10.13A). Alkali feldspar megacrysts mantled by plagioclase are rather common in the early granitic phases. The granites generally contain biotite as the only mafic mineral. Magnetite and titanite are typical minor phases, and zircon, apatite, ilmenite, monazite, and fluorite are found as accessory minerals.
6. Geochemical comparison and petrogenetic implications The main geochemical differences between the granitoid groups and their petrogenetic implications are reviewed. The geochemical data have been collected from published sources (Eklund et al., 1998; Lahtinen, 1994; Lehtonen, 1988; Nironen and Bateman, 1989; Väänänen, 1998; Väisänen et al., 2000), the Geological Survey of Finland databank, and unpublished personal datasets. Analyses from clearly altered and/or mineralized rocks and extremely fractionated granites were excluded. The fact that the analyses were made in different laboratories and at different times causes unavoidable scatter, especially in minor and trace elements, which hampers identification of fractionation trends.
6.1. Preorogenic rocks The preorogenic tonalites of the Primitive arc complex are charcterized by high Fe (and
PROTEROZOIC
OROGENIC
GRANITOID
ROCKS
TiO2 Al2O3 FeOtot CaO SiO2 SiO2 Fig. 10.14. Variation diagrams (pp. 463–465) for the Paleoproterozoic felsic and intermediate plutonic rocks in Finland. Total Fe is expressed as FeOtot. Fields in K2O vs. SiO2 diagrams after Rickwood (1989), ca–calc-alkaline, sh–shoshonitic. Boundary of I-type and S-type granites in molecular A/CNK [Al2O3/(CaO+Na2O+K2O)] vs. SiO2 diagrams after Chappell and White (1974). CHAPTER 10 • PROTEROZOIC OROGENIC GRANITOID ROCKS •
463
Na2O K2O Ba Rb
SiO2
SiO2
Fig. 10.14. (continued)
464
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PROTEROZOIC
OROGENIC
GRANITOID
ROCKS
Zr FeOtot/MgO A/CNK
SiO2
+
Synkinematic rock of southern and central Finland Synkinematic rock of Central Finland granitoid complex Synorogenic rock of northern Finland
SiO2 Preorogenic rock of central Finland Postkinematic rock of central Finland Postkinematic rock of Central Finland granitoid complex Lateorogenic rock of southern Finland Lateorogenic rock of northern Finland Postorogenic rock of southern Finland Postorogenic rock of northern Finland
Fig. 10.14. (continued)
CHAPTER 10 • PROTEROZOIC OROGENIC GRANITOID ROCKS •
465
A peraluminous
A/NK
metaluminous
peralkaline
A/CNK
A/CNK
Al2O3+CaO FeOtot+Na2O+K2O
B
ca+sp
alkaline 100 x
MgO+FeOtot+TiO2 SiO2
Fig.10.15. Chemical composition of the Paleoproterozoic felsic and intermediate plutonic rocks of Finland in (A) molecular A/NK[Al2O3/(Na2O+K2O)] vs. A/CNK [Al2O3/(CaO+Na2O+K2O)] diagrams (Maniar and Piccoli, 1989); and (B) a discrimination diagram for granites (SiO2 >68%) (Sylvester, 1989). ca+sp is the field for calc-alkaline and strongly peraluminous granitoid rocks.
Mg) and Ca as well as low K and Rb (Figure 10.14). In the A/CNK vs. SiO2 diagram the preorogenic tonalites plot in the I-type field. In the A/NK vs. A/CNK diagram (Figure 10.15A) they plot on both sides of the boundary between the metaluminous and peraluminous fields. In Figure 10.15B the tonalites plot clearly in the field of calc-alkaline and strongly peraluminous granites. According to Lahtinen (1994), the low K and Al contents of the gneissic tonalites 466
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indicate low degrees (10–15 wt.%) of melting of low-K tholeiitic island arc basalts in an immature arc where thick crust had not yet been developed. The high Fe, Mg, and Ca contents conform to a mafic source and the low incompatible element abundances to an overall immature source area. The εNd (at 1930 Ma) values of the gneissic tonalites vary from +1.1 to +4.4 with a cluster around +3, and the depleted mantle model ages vary from 2.33 Ga to 1.86 Ga with most values
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A
DM
εNd
CHUR
Mean 3.1 St. dev. 0.49 Age (Ga)
B
DM
C
CHUR
CHUR
Mean 3.0 –0,2 St. dev. 0.24 0.99
Mean –0,2 St. dev. 0.47 Age (Ga)
D
DM
CHUR
DM
εNd
Age (Ga)
E
DM
CHUR
Age (Ga)
εNd
Age (Ga)
Fig. 10.16. εNd vs. age diagram for the Paleoproterozoic felsic and intermediate plutonic rocks in Finland (data from Huhma, 1986; Patchett and Kouvo, 1986; Lahtinen and Huhma, 1997; Rämö and Nironen, 2001; Rämö et al., 2001; Ruotoistenmäki et al., 2001). Analyses with147Sm/144Nd less than 0.08 and over 0.145 have been excluded. Analyses from rocks of the Archean craton area are shown in blue. (A) Preorogenic rocks of central Finland; (B) Synkinematic rocks of southern and central Finland; (C) Postkinematic rocks of central Finland; (D) Lateorogenic rocks; (E) Postorogenic rocks. DM is depleted mantle (DePaolo, 1981), CHUR is the Chondritic Uniform Reservoir (DePaolo and Wasserburg, 1975). Gray vertical bars denote approximate age of magmatism.
around 2.0 Ga (Figure 10.16A; Lahtinen and Huhma, 1997). The clearly positive initial εNd values suggest that the tonalites do not contain an Archean component, although the Primitive arc complex is located adjacent to the Archean craton. An ion microprobe study of zircons confirmed the absence of an Archean
component and showed that 1.93–1.92 Ga is the actual crystallization age of the tonalites (Vaasjoki et al., 1998).
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6.2. Synorogenic rocks Synkinematic rocks of southern and central Finland The synkinematic rocks of the supracrustal belts (red crosses in Figure 10.14) have generally higher Al and Na and lower K, Ba, and Rb contents than the synkinematic rocks of the Central Finland granitoid complex (blue crosses in Figure 10.14). These differences indicate the tonalitic–granodioritic character of the former and the granodioritic–granitic character of the latter group. In the K2O vs. SiO2 diagram the rocks of the supracrustal belts plot mainly in the calc-alkaline field, whereas the rocks of the Central Finland granitoid complex plot in the high-K calc-alkaline and shoshonitic fields. The several trends in the CaO versus SiO2 diagram do not result from analytical scatter but probably reflect source heterogeneity. The synkinematic rocks show a typical shift from the I-type to the Stype field with increasing SiO2, but the rocks of the supracrustal belts show, in general, a stronger S-type character than the rocks of the Central Finland granitoid complex. The synkinematic rocks of the supracrustal belts are also more peraluminous than the rocks of the complex (Figure 10.15A). In Figure 10.15B the synkinematic granites mainly plot in the calc-alkaline and strongly peraluminous granite field. The synkinematic rocks of southern and central Finland resemble modern calc-alkaline andesites and have I-type granitoid characteristics (Front and Nurmi, 1987), indicating a prominent arc-related source component. The peraluminous affinity of the synkinematic rocks of southern Finland as well as their high Al contents suggest a sedimentary component in the source area. Lahtinen (1996) studied the rocks of the supracrustal belt immediately south of the Central Finland granitoid complex and concluded that some strongly peraluminous granitoids are the result of mantlederived magma that melted sedimentary rocks. 468
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He also proposed three end-members for the most common calc-alkaline granitoids of the supracrustal belt: mantle-derived magma of within plate affinity, calc-alkaline intermediate melt, and silicic crustal melt. The synkinematic granitoids may be divided into two groups by their εNd (at 1880 Ma) values (Figure 10.16B; Huhma, 1986; Patchett and Kouvo, 1986; Lahtinen and Huhma, 1997; Rämö et al., 2001). The rocks adjacent to the Archean craton margin and the Veteli and Kalanti intrusions in western Finland (Figure 10.3) form a group with εNd (at 1880 Ma) values from +2.7 to +3.4 and depleted mantle model ages between 1.96 Ga and 1.91 Ga. The granitoids of the Central Finland granitoid complex and granitoids of the supracrustal belts in southernmost Finland have somewhat lower εNd values (from –1.5 to +1.1) and higher depleted mantle model ages (from 2.41 Ga to 2.05 Ga), indicating that these areas are less juvenile and probably contain an older (~2.0 Ga) nucleus (Lahtinen and Huhma, 1997; Rämö et al., 2001). Nironen et al. (2000) proposed that K-rich calc-alkaline volcanic rocks were subducted beneath the older nucleus and that partial melting of these rocks in the lower crust, triggered by heat and magmatic addition from the mantle, produced the synkinematic magmatism in the Central Finland granitoid complex.
Postkinematic rocks of central Finland There is no marked difference between the postkinematic rocks of the Central Finland granitoid complex and those of the supracrustal belts. The postkinematic rocks are in general higher in Fe, K, Ba, Zr, and Fe/Mg and lower in Mg and Ca than the synkinematic rocks at similar SiO2 contents (Figure 10.14). In the A/CNK vs. SiO2 diagram the postkinematic rocks show a trend from the I-type field towards the S-type field with increasing SiO 2 , and also a change from marginally peraluminous to marginally metalu-
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ROCKS
minous (Figure 10.15A). In Figure 10.15B the postkinematic granites show a slightly more alkaline character than the synkinematic ones. Overall, the postkinematic rocks are more alkaline and evolved than the synkinematic rocks. The composition and similarity to rapakivi granites of the postkinematic plutons indicate A-type affinity and suggest crystallization of relatively dry magmas derived from deep crust (Elliott et al., 1998; Nironen et al., 2000). The εNd (at 1880 Ma) values of the postkinematic granitoids within the Central Finland granitoid complex constitute a tight range of –1.1 to +0.5 (Figure 10.16C; Rämö et al., 2001). Depleted mantle model ages of the rocks vary from 2.27 to 2.11 Ga. These values do not differ remarkably from the values of the synkinematic granitoids. The homogeneity of the Nd isotope composition implies a homogeneous source over the entire granitoid complex and that the crust-forming process was quite rapid (Rämö et al., 2001). Nironen et al. (2000) proposed that the postkinematic rocks were derived from a granulitic residue left in the lower crust after the extraction of the synkinematic granitoids. Heat from the mafic underplate triggered partial melting of the granulite, and mafic magmatic addition also contributed to the generation of the postkinematic magmas. According to Elliott (2003), the postkinematic magmas may consist of partial melts of evolved mafic (ferrodioritic) lower crust with minor incorporation of restitic material from lower crust. The εNd (at 1880 Ma) values of the postkinematic granitoids east of the paleosuture (Figure 10.3) are from –6.6 to –1.9 (Figure 10.16C; Huhma, 1986; Ruotoistenmäki et al., 2001), indicating major Archean component in these rocks. The depleted mantle model ages for these rocks are 2.28 Ga to 2.20 Ga.
Synorogenic rocks of northern Finland The synorogenic rocks of northern Finland
have clearly higher Na and somewhat lower Ti, Al, and Fe contents than the synkinematic rocks of southern and central Finland (Figure 10.14). The monzonites (SiO2 less than 67 wt.%) and trondhjemites (SiO2 over 67 wt.%) do not constitute different trends in the discrimination diagrams, except for K and Rb that are lower in the trondhjemites. The synorogenic rocks plot in the I-type field. The monzonites are mostly metaluminous but the trondhjemites plot around the boundary between the metaluminous and peraluminous fields (Figure 10.15A). The I-type characteristics of the synorogenic rocks of northern Finland suggest arc affinity. However, the high Na and K contents of the monzonites also point to alkaline affinity; high Na contents, in particular, are not typical for arc-related rocks. Although clearly albitized rocks with Na2O >7 wt.% were excluded from the presented data, it is possible that the overall high Na values reflect, at least some, Na metasomatism.
6.3. Lateorogenic granites The lateorogenic granites have high K and Rb contents compared to synorogenic rocks within their limited (70–77 wt.%) SiO2 range (Figure 10.14). Moreover, the granites of southern Finland have somewhat higher Al and lower Zr contents and the granites of northern Finland have elevated Na contents. The lateorogenic granites of southern Finland plot in the S-type field whereas the granites of northern Finland plot in the I-type field. The lateorogenic granites are clearly peraluminous (Figure 10.15A). In Figure 10.15B the granites of northern Finland show a more alkaline character than the southern Finland granites. The peraluminous character of the lateorogenic granites in southern Finland is in line with the interpretation that they are S-type granites derived by partial melting of sedimentary rocks. The low Zr contents in these rocks suggest low melting temperatures. The
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lateorogenic granites of northern Finland are also peraluminous but they have higher Zr contents and a more alkaline character than the lateorogenic granites of southern Finland. There are few Nd isotope studies on the lateorogenic granites. The εNd (at 1830 Ma) values of the Hanko and Oripää granites in southwestern Finland (Figure 10.3) are +0.1 and +2.2, and the depleted mantle model ages are 2.08 Ga and 1.94 Ga, respectively (Figure 10.16D; Huhma, 1986; Rämö and Nironen, 2001). The Puruvesi granite close to the paleosuture yielded an age of ~1.81 Ga (Nykänen, 1983; Huhma, 1986). The strongly negative εNd (at 1810 Ma) value of this granite (–6.8) indicates an Archean component in the source. Nd isotope data from a lateorogenic granite within the Hetta complex in northern Finland with εNd (at 1810 Ma) value of –8.0 (Figure 10.16D; Huhma, 1986) also indicates the presence of an Archean component.
6.4. Postorogenic rocks The postorogenic rocks of southern and northern Finland differ from each other in their SiO2 contents: the southern Finland rocks cover a wide range from 32 to 78 wt.% SiO2 (Eklund et al., 1998), the northern Finland rocks have a more limited range at the felsic end (68–78 wt.% SiO2). The postorogenic rocks plot in the highK and shoshonitic fields in the K2O vs. SiO2 diagram (Figure 10.14). The rocks of southern Finland have high Ti and Ba contents, whereas the granites of northern Finland have elevated Rb contents. In the A/CNK versus SiO2 diagram the postorogenic rocks mainly plot in the I-type field. The rocks of southern Finland have a trend from metaluminous to marginally peraluminous, whereas the rocks of northern Finland are mainly peraluminous, approaching the peralkaline field (Figure 10.15A). In Figure 10.15B, the postorogenic rocks show a pronounced alkaline affinity. The rocks of southern Finland are strongly enriched in the 470
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LREE (Nurmi and Haapala, 1986; Eklund et al., 1998). Pitkänen (1985) concluded that fractional crystallization of a F-rich magma produced the phases in the Luonteri intrusion, and that F was confined into the early phases, dominantly biotite. Eklund et al. (1998) showed that the postorogenic rocks of southern Finland are highly enriched in the incompatible elements and that the more mafic rocks are particularly enriched in F, P, Ba, Sr, and LREE. They classified the postorogenic rocks of southern Finland as shoshonitic by their high K, Ba, and Sr contents. However, the Ti contents of these rocks are high for typical shoshonitic rocks and indeed high compared to other orogenic rocks (Figure 10.14). According to Lahtinen (1994), geochemical features of all the postorogenic rocks of southern Finland are in line with fractional crystallization of mantle-derived alkali basaltic magmas which assimilated crustal rocks. He proposed an enriched subcontinental lithospheric mantle source for the mafic magmas. Eklund et al. (1998) suggested carbonate metasomatism as the reason for enrichment in the lithospheric mantle and concluded that metasomatism was more extensive in the east. Front et al. (1989) showed that the postorogenic granites of northern Finland have an alkaline affinity with their high SiO2 and total alkali values as well as high Fe/Mg ratios. In Figures 10.15A and B these granites plot in the peraluminous and alkaline fields, respectively, and are tightly associated with the lateorogenic granites of northern Finland. The postorogenic granites of northern Finland are F-rich and show decreasing F with differentiation of the magma, i.e., a trend similar to the one in the postorogenic Luonteri intrusion of southern Finland (Kai Front, pers. comm., 2001). The εNd (at 1800 Ma) value of a granite of the Åva ring complex is +0.2, and the depleted mantle model age is 2.02 Ga (Figure 10.16E; Patchett and Kouvo, 1986). The εNd (at 1800
PROTEROZOIC
OROGENIC
GRANITOID
ROCKS
Table 10.2. Geochemical and Nd isotope characteristics of the Proterozoic orogenic granitoid rocks of Finland. The synkinematic rocks serve as the reference group in geochemical comparison. The source for each group is also assessed. See text for details. Preorogenic rocks Synorogenic rocks (central Finland) Synkinematic rocks Postkinematic rocks northern Finland
Lateorogenic rocks
Postorogenic rocks
SiO2 72–76%
SiO2 71–75% (south), 70–76% (north) high Al, K, Rb
SiO2 50–74% (south), 68–77% (north) high K (+ Ti, Ba, Sr; south)
SiO2 57–77%
high Fe, Mg, Ca low K, Rb I-type
I-type/S-type
SiO2 56–77%
SiO2 47–75%
high Fe, K, Ba, Zr
high Na
low Mg, Ca I-type/S-type
I-type
met/peraluminous met/peraluminous met/peraluminous metaluminous
low Zr (south) S-type (south), I-type (north) peraluminous
I-type met/peraluminous
Nd isotopes
εNd (at 1930 Ma) +1.1 to +4.4
εNd (at 1880 Ma) –1.5 to +3.4
εNd (at 1880 Ma) –6.6 to +0.5
εNd (at 1880 Ma) –3.2
TDM 1.86–2.33 Ga TDM 1.93–2.41 Ga TDM 2.11–2.88 Ga TDM 2.42 Ga
Source primitive arc
mature arc + mantle + sedimentary rocks
mature arc
Ma) values of the Parkkila granodiorite and the Pirilä granite (Patchett and Kouvo, 1986; Lahtinen and Huhma, 1997) are +0.5 and +0.7, respectively. The εNd (at 1770 Ma) values of the Nattanen, Rovaniemi, and Vainospää granites are strongly negative (from –9.2 to –6.2) and the depleted mantle model ages are from 2.55 to 2.41 Ga (Huhma, 1986). The Nattanen and Vainospää granites yielded initial εHf values of –12 and –10 and plot well below the chondritic curve (Patchett et al., 1981). Both the Nd and Hf isotope results indicate a substantial Archean component in the granites.
enriched mantle + crust
εNd (at 1830 Ma) +0.1,+2.2 (south), –7.7 to –6.5 (north) TDM 1.94 Ga, 2.08 Ga (south), 2.41–2.51 Ga (north)
εNd (at 1800 Ma) +0.2 to +0.7(south), –8.7 to –5.8 (north) TDM 1.97–2.02 Ga (south), 2.41–2.55 Ga
mainly sedimentary rocks (south)
enriched mantle + crust
7. Discussion The classification presented here is based on orogenic evolution in the Finnish Svecofennian and does not apply to other Paleoproterozoic parts of the Fennoscandian Shield; hence also the orogenic concept differs from the model presented in Chapter 11. The term postcollisional has been applied to postorogenic rocks of southern Finland (Eklund et al., 1998). Because also the postkinematic and lateorogenic rocks are postcollisional with respect to the accretionary events, the term has not been adopted here. The geochemical and Nd isotope characteristics of the plutonic groups are summed up in Table 10.2. Figure 10.16 shows that the initial εNd values are around +3 in the preoro-
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genic rocks of central Finland, –1.5 to +1 and around +3 in the synorogenic rocks, –1 to +0.5 in the postkinematic rocks of the Central Finland granitoid complex, 0 to +2.2 in the lateorogenic and postorogenic granites of southwestern Finland, and strongly negative in the postkinematic, lateorogenic, and postorogenic rocks that are located in areas underlain by Archean crust. These changes reflect a shift from juvenile arc magmatism to mixed sources with an older (Paleoproterozoic) component and further to a predominant Archean component. The positive (+0.2 to +3.4) initial εNd values of the synkinematic, lateorogenic, and postorogenic rocks of southwestern Finland suggest a primitive arc component in the source of these rocks. Recently Väisänen et al. (2002) proposed that the newly dated ~1900 Ma bimodal plutonic and volcanic rocks in southernmost Finland represent pre-collisional island arc magmatism. It is quite possible that, in addition to the Primitive arc complex, also the southern Finland and western Finland arc complexes contain remains of preorogenic rocks. The new ages and interpretations would thus extend the age range of preorogenic rocks. The Hetta compex in northern Finland, considered here as preorogenic, may actually consist of a variety of intrusive rocks ranging in age from 1.95 Ga to 1.80 Ga. Front and Nurmi (1987) concluded that the synorogenic granitoids of the Central Finland granitoid complex with their high Fe/Mg ratios indicate a tholeiitic character in contrast to the calc-alkaline, synkinematic rocks of the supracrustal belts. Their study covered both synkinematic and postkinematic plutons. The Fe enrichment in the postkinematic rocks, coupled with high K contents, indicates alkaline rather than tholeiitic affinity. Front and Nurmi (1987) also concluded that the Central Finland granitoid complex area represents a thicker crust than the adjacent supracrustaldominated areas. Moreover, the generally high-K magmatism in the granitoid complex 472
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and the fact that the initial εNd values in this area are slightly lower than those of the synkinematic granitoids in the supracrustal belts indicate that the Central Finland granitoid complex area is an evolved Paleoproterozoic terrane. The similar initial εNd values of the synkinematic and postkinematic granitoids in the Central Finland granitoid complex indicate thorough recycling and homogenization of the lower crust (cf. Rämö et al., 2001). The age of postkinematic magmatism is ~1885 Ma in the area of the Primitive arc complex, and decreases westwards with a concomitant shift in composition toward fluorite-bearing granites with A-type affinity. Hence, postkinematic magmatism initiated at the zone where the Primitive arc complex and the Arc complex of western Finland had accreted against the Archean craton, and then spread out to the west (until 1.87 Ga) and east (until 1.86 Ga). The microtonalite dikes in eastern Finland are a further expression of the postkinematic magmatism. The shift from granites and granodiorites to microtonalite dikes with magma mixing and mingling structures shows an increasing mantle component in extensional postkinematic magmatism at the Archean craton margin. The fact that postkinematic magmatism has not been found in the area of the Arc complex of southern Finland and that lateorogenic and postorogenic magmatisms of southern Finland are largely confined to this area (Figure 10.3) implies, together with Nd isotope data (Rämö et al., 2001), that the Arc complex of southern Finland is a terrain with lithospheric evolution distinct from the evolution in the Arc complex of western Finland. The details of the difference will be a matter of future studies. At this stage the orogenic, accretional magmatism of southern and central Finland may be grouped into two evolution lines: one starting with calc-alkaline synkinematic magmatism and ending in alkaline postkinematic magmatism in the Central Finland granitoid complex area, and another in
PROTEROZOIC
OROGENIC
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ROCKS
southern Finland, starting with synkinematic calc-alkaline magmatism followed by peraluminous lateorogenic magmatism and ending in shoshonitic postorogenic magmatism. The first evolution line results from the accretion of the Arc complex of southern Finland against the Arc complex of western Finland, and subsequent stabilization of the crust in central Finland. The other evolution line reflects prolonged deformation, magmatism, and metamorphism along the southern collisional zone. The origin of magmatism in southern Finland (melting of Svecofennian sedimentary rocks) as well as the flat-lying and in places domal structures bear similarities to core complexes that are generally assigned to extensional collapse of the orogen. Extensional collapse may have taken place in southern Finland but the association of at least some of the lateorogenic intrusions with transpressional structures implies that the tectonic setting of this granite group is still unclear. The fairly large compositional variation in the synorogenic rocks of northern Finland suggests a mafic component in the source. On the other hand, the negative initial εNd value of a monzonite of this group implies an Archean component. The negative initial εNd values obtained from the lateorogenic and postorogenic rocks of northern Finland, and the fact that many of them contain Archean zircons, indicate a substantial Archean contribution in the formation of the magmas. Such an Archean component may have been: (1) Proterozoic sedimentary rocks with detrital Archean material; (2) Archean crust that was remelted or assimilated by magmas rising from the mantle; or (3) Archean lithospheric mantle. The first possibility may be considered with the lateorogenic granites, which are found both as migmatizing veins and as plutons that contain fragments of older rocks. The lateorogenic and postorogenic granites of northern Finland are very similar both in their major and trace element characteristics and could thus represent one granite group.
However, the former are granites within migmatitic complexes, whereas the latter are high-level intrusive plutons. The lack of mafic rocks in association with the postorogenic granites suggests melting of Archean crust rather than magmatic input from the mantle. As the lateorogenic and postorogenic rocks are geochemically similar, both groups probably result from remelting of Archean crust. This implies that the entire northern Finland crust contains Archean rocks that are partly buried by the Proterozoic cover sequence. Another implication is that the orogenic classification of granitoid rocks is not as readily applicable for northern Finland as for southern Finland. The orogenic rocks of northern Finland have an alkaline affinity and especially the synorogenic rocks are Na-rich. The reason for Na enrichment is problematic. Scapolite, albite, carbonate, and tourmaline are common alteration products in supracrustal and igneous rocks of northern Finland and Sweden. Frietsch et al. (1997) assigned this alteration to evaporitic sediments that were deposited in an intracratonic rift basin. They concluded that metasomatizing components were removed from the evaporites during low- to medium-grade regional metamorphism and that large-scale metasomatism was connected with intrusion of granitoids at 1.89–1.77 Ga. Another explanation for the Na enrichment of the synorogenic rocks would be melting of relatively Na-rich Archean lower crust. The high F contents of the postorogenic rocks in northern Finland bring up another problem – the source of the F-rich fluids. The postorogenic rocks in southern Finland with similarly high F contents were explained to be derived from an enriched (metasomatic) lithospheric mantle (Eklund et al., 1998). It is possible that the lithospheric mantle in northern Finland, with a long enrichment history from the Archean to the Proterozoic, was the source of the F-rich fluids. The tectonic regime for postorogenic magmatism with high-level emplacement and influx of F-rich volatiles
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would have been large-scale extension that affected the crust from the top to the base.
8. Summary The Proterozoic orogenic granitoid rocks of Finland may be divided into preorogenic (1.95–1.91 Ga), synorogenic (1.89–1.86 Ga), lateorogenic (1.84–1.80 Ga), and postorogenic (1.81–1.77 Ga). The synorogenic rocks of southern and central Finland may be further divided into synkinematic (1.89–1.87 Ga) and postkinematic (1.88–1.86 Ga). The preorogenic rocks are located in the Primitive arc complex in central Finland and in northernmost Finland. Their high Fe, Mg, and Ca, low K contents, and positive initial εNd values (around +3) are consistent with juvenile primitive arc magmatism. The synkinematic rocks of southern Finland are peraluminous and represent typical calc-alkaline magmatism, whereas the synkinematic and postkinematic rocks of central Finland have higher K, Ba, and Rb contents; the postkinematic rocks also have high Fe/ Mg ratios and an overall alkaline character. The synorogenic rocks of northern Finland are chartacterized by high Na contents. The initial εNd values of the synorogenic rocks in southern and central Finland are from –1.5 to +3.4 and suggest a slightly evolved Paleoproterozoic source. The strongly negative initial εNd values of the synorogenic rocks in northern Finland imply a substantial Archean source component. The lateorogenic rocks are peraluminous S-type granites with migmatizing veins. The postorogenic rocks of southern Finland are metaluminous, have a wide compositional range (monzonite–granite) and shoshonitic affinity, whereas the postorogenic rocks of northern Finland are peraluminous, alkaline granites. The initial εNd values of the lateorogenic and postorogenic rocks of range from 0 to +2.2 in southwestern Finland and are strongly negative in areas underlain by Ar474
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chean crust. The orogenic, accretional magmatism of southern and central Finland consists of two evolution lines: (1) calc-alkaline synkinematic magmatism in the Central Finland granitoid complex area, and subsequent alkaline postkine matic magmatism that spread to the west and east of the area of the Primitive arc complex; and (2) calc-alkaline synkinematic magmatism in southern Finland, followed by peraluminous lateorogenic magmatism and, subsequently by shoshonitic postorogenic magmatism. The first resulted from accretion of the Arc complex of southern Finland against the Arc complex of western Finland, and subsequent stabilization of the crust in central Finland. The second reflects prolonged deformation, magmatism, and metamorphism along the southern collisional zone, possibly as a result of orogenic collapse.
References Bergman, L., 1986. Structure and mechanism of intrusion of postorogenic granites in the archipelago of southwestern Finland. Acta Acad. Aboensis, Ser. B, 46, 1–74. Bergman, S., Kubler, L., Martinsson, O., 2001. Description of regional geological and geophysical maps of northern Norrbotten county (east of the Caledonian orogen). Geol. Surv. Sweden, Ser Ba 56, 1–110. Bleeker, W., Westra, L., 1987. The evolution of the Mustio gneiss dome, Svecofennides of SW Finland. Precambrian Res. 36, 227–240. Branigan, N.P., 1989. Hybridisation in middle Proterozoic high-level ring complexes, Åland, SW Finland. Precambrian Res. 45, 83–95. Brown, M., 1994. The generation, segregation, ascent and emplacement of granite magma: the migmatite-to-crustally-derived granite connection in thickened orogens. Earth-Sci. Rev. 36, 83–130. Chappell, B.W., White, A.J.R., 1974. Two contrasting granite types. Pacific Geology 8, 173–174. Corfu, F., Evins, P.M., 2002. Late Palaeoproterozoic
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monazite and titanite U–Pb ages in the Archaean Suomujärvi Complex, N-Finland. Precambrian Res. 116, 171–181. DePaolo, D.J., 1981. Neodymium isotopes in the Colorado Front Range and crust-mantle evolution in the Proterozoic. Nature 291, 193–196. DePaolo, D.J., Wasserburg, G.J., 1976. Nd isotopic variations and petrogenetic models. Geophys. Res. Lett. 3, 249–252. Ehlers, C., Bergman, L., 1984. Structure and mechanism of intrusion of two postorogenic granite massifs, southwestern Finland. In: A. Kröner, R. Greiling (Eds.), Precambrian Tectonics Illustrated. E. Schweizerbart’sche Verlagsbuchhandlung, Stuttgart, 173–190. Ehlers, C., Lindroos, A., Selonen, O., 1993. The late Svecofennian granite-migmatite zone of southern Finland – a belt of transpressive deformation and granite emplacement. Precambrian Res. 64, 295–309. Ekdahl, E., 1993. Early Proterozoic Karelian and Svecofennian formations and the Evolution of the Raahe-Ladoga Ore Zone, based on the Pielavesi area, central Finland. Geol. Surv. Finland, Bull. 373, 1–137. Eklund, O., Konopelko, D., Rutanen, H., Fröjdö, S., Shebanov, A.D., 1998. 1.8 Ga Svecofennian post-collisional shoshonitic magmatism in the Fennoscandian shield. Lithos 45, 87–108. Elliott, B.A., 2003. Petrogenesis of the post-kinematic magmatism of the Central Finland Granitoid Complex II; sources and magmatic evolution. J. Petrol. 44, 1681–1701. Elliott, B.A., Rämö, O.T., Nironen, M., 1998. Mineral chemistry constraints on the evolution of the 1.88-1.87 Ga post-kinematic granite plutons in the central Finland granitoid complex. Lithos 45, 109–129. Eskola, P., 1932. On the origin of granitic magmas. Mitt. Mineral. Petrogr. 42, 455–481. Frietsch, R., Tuisku, P., Martinson, O., Perdahl, J.A., 1997. Early Proterozoic Cu-(Au) and Fe ore deposits associated with regional Na-Cl metasomatism in northern Fennoscandia. Ore Geol. Rev. 12, 1–34. Front, K., Nurmi, P.A., 1987. Characteristics and
geological setting of synkinematic Svecokarelian granitoids in southern Finland. Precambrian Res. 35, 207–224. Front, K., Vaarma, M., Rantala, E., Luukkonen, A., 1989. Keski-Lapin varhaisproterotsooiset Nattas-tyypin graniittikompleksit, niiden kivilajit, geokemia ja mineralisaatiot. Summary: Early Proterozoic Nattanentype granite complexes in central Finnish Lapland: rock types, geochemistry and mineralizations. Geol. Surv. Finland, Rep. Invest. 85, 1–77. Haapala, I., Front, K., Rantala, E., Vaarma. M., 1987. Petrology of Nattanen-type granite complexes, northern Finland. Precambrian Res. 35, 225–240. Halden, N.M., 1982. Structural, metamorphic and igneous history of migmatites in the deep levels of a wrench fault regime, Savonranta, eastern Finland. Trans. Roy. Soc. Edinburgh Earth Sci. 73, 17–30. Hiltunen, A., 1982. The Precambrian geology and skarn iron ores of the Rautuvaara area, northern Finland. Geol. Surv. Finland, Bull. 318, 1–133. Hölttä, P., 1995. Contact metamorphism of the Vaaraslahti pyroxene granitoid intrusion in Pielavesi, Central Finland. In: P. Hölttä (Ed.), Relationship of granitoids, structures and metamorphism at the eastern margin of the Central Finland Granitoid Complex. Geol. Surv. Finland, Bull. 382, 27–79. Hubbard, F., Branigan, N., 1987. Late Svecofennian magmatism and tectonism, Åland, southwest Finland. Precambrian Res. 35, 241–256. Huhma, A., 1981. Youngest Precambrian dyke rocks in North Karelia, East Finland. Bull. Geol. Soc. Finland 53, 67–82. 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. Irvine, T.N., Baragar, W.R.A., 1971. A guide to the chemical classification of the common volcanic rocks. Can. J. Earth Sci. 8, 523–548. Koistinen, T., Stephens, M.B., Bogatchev, V., Nord-
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gulen, Ø., Wennerström, M., Korhonen, J., 2001. Geological map of the Fennoscandian Shield, scale 1:2 000 000. Geol. Surveys of Finland, Norway and Sweden and the North-West Dept. of Nat. Res. Russia. Korja, A., Heikkinen, P., 1995. Proterozoic extensional tectonics of the central Fennoscandian Shield: results from the Baltic and Bothnian echoes from the lithosphere experiment. Tectonics 14, 504–517. Korsman, K., Hölttä, P., Hautala, T., Wasenius, P., 1984. Metamorphism as an indicator of evolution and structure of the crust in Eastern Finland. Geol. Surv. Finland, Bull. 328, 1–40. Korsman, K., Koistinen, T., Kohonen, J., Wennerström, M., Ekdahl, E., Honkamo, M., Idman, H., Pekkala, Y. (Eds.), 1997. Bedrock map of Finland 1:1 000 000. Geol. Surv. Finland, Espoo, Finland. Kousa, J., Marttila, E., Vaasjoki, M., 1994. Petrology, geochemistry and dating of Paleoproterozoic metavolcanic rocks in the Pyhäjärvi area, central Finland. In: M. Nironen, Y. Kähkönen (Eds.), Geochemistry of Proterozoic supracrustal rocks in Finland. Geol. Surv. Finland, Spec. Pap. 19, 7–27. Lahtinen, R., 1994. Crustal evolution of the Svecofennian and Karelian domains during 2.1-1.79 Ga, with special emphasis on the geochemistry and origin of 1.93-1.91 Ga gneissic tonalites ans associated supracrustal rocks in the Rautalampi area, central Finland. Geol. Surv. Finland, Bull. 378, 1–128. Lahtinen, R., 1996. Geochemistry of Palaeoproterozoic supracrustal and plutonic rocks in the Tampere–Hämeenlinna area, southern Finland. Geol. Surv. Finland, Bull. 389, 1–113. Lahtinen, R., Huhma, H., 1997. Isotopic and geochemical constraints on the evolution of the 1.93-1.79 Ga Svecofennian crust and mantle. Precambrian Res. 82, 13–34. Lauerma, R., 1982. On the ages of some granitoid and schist complexes in northern Finland. Bull. Geol. Soc. Finland 54, 85–100.
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Lehtonen, M., 1984. Geological map of Finland 1:100 000, Explanation to the Pre-Quaternary rocks of the Muonio map-sheet area. Geol. Surv. Finland. (in Finnish, with an English summary) Lehtonen, M., 1988. Muonion-Kihlangin alueen geologiasta ja granitoidien petrokemiasta. Phil. Lic. Thesis, Univ. Helsinki. (in Finnish) Lehtonen, M., Airo, M.-L., Eilu, P., Hanski, E., Kor telainen, V., Lanne, E., Manninen, T., Rastas, P., Räsänen, J., Virransalo, P., 1998. Kittilän vihreäkivialueen geologia. Summary: The stratigraphy, petrology and geochemistry of the Kittilä greenstone area, northern Finland. A report of the Lapland Volcanite Project. Geol. Surv. Finland, Rep. Invest. 140, 1–144. Lindberg, B., Eklund, O., 1988. Interactions between basaltic and granitic magmas in a Svecofennian postorogenic granitoid intrusion, Åland, southwest Finland. Lithos 22, 13–23. Mäkitie, H., 1999. Structural analysis and metamorphism of Palaeoproterozoic metapelites in the Seinäjoki–Ilmajoki area, western Finland. Bull. Geol. Soc. Finland 71, 305–328. Maniar, P.D., Piccoli, P.M., 1989. Tectonic discrimination of granitoids. Geol. Soc. Am. Bull. 101, 635–643. Meriläinen, K., 1976. The granulite complex and adjacent rocks in Lapland, northern Finland. Geol. Surv. Finland, Bull. 281, 1–129. Mikkola, E., 1928. Über den Nattanengranit im finnischen Lapplande. Fennia 50, N:o 12, 1–24. Mikkola, E., 1941. Kivilajikartan selitys. Lehdet B7-C7-D7, Muonio–Sodankylä–Tuntsajoki. English summary: Explanation to the map of rocks, Sheets B7–C7–D7. The general geological map of Finland. Geologinen toimikunta, Helsinki, 1–286. Miyashiro, A., 1978. Nature of alkalic volcanic rock series. Contrib. Mineral. Petrol. 66, 91–104. Mouri, H., Korsman, K., Huhma, H. 1999. Tectono-metamorphic evolution and timing of
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the melting processes in the Svecofennian Tonalite-Trondhjemite Migmatite Belt: An example from Luopioinen, Tampere area, southern Finland. Bull. Geol. Soc. Finland 71, 31–56. Nironen, M., 1989. Emplacement and structural setting of granitoids in the early Proterozoic Tampere and Savo Schist Belts, Finland - implications for contrasting crustal evolution. Geol. Surv. Finland, Bull. 346, 1–83. Nironen, M., 1995. Block boundary at the southeastern margin of the Paleoproterozoic central Finland Granitoid Complex. In: P. Hölttä (Ed.), Relationship of granitoids, structures and metamorphism at the eastern margin of the Central Finland Granitoid Complex. Geol. Surv. Finland, Bull. 382, 91–115. Nironen, M., 1997. The Svecofennian orogen: a tectonic model. Precambrian Res. 86, 21–44. Nironen, M., 1999. Structural and magmatic evolution in the Loimaa area, southwestern Finland. Bull. Geol. Soc. Finland 71, 57–71. Nironen, M., 2003. Keski-Suomen granitoidikompleksi – kallioperäkartan selitys. Central Finland Granitoid Gomplex – Explanation to the bedrock map. Geol. Surv. Finland. Rep. Invest. 157, 1–45. Nironen, M., Bateman, R., 1989. Petrogenesis and syntectonic emplacement in the early Proterozoic south-central Finland: a reversely zoned diorite-granodiorite and a granite. Geol. Rundsch. 78, 617–631. Nironen, M., Front, K., 1992. The 1.88 Ga old Mäntylä complex central Finland: emplacement and deformation of mafic to felsic plutonic rocks and associated Mo mineralization. Bull. Geol. Soc. Finland 64, 75–90. Nironen, M., Elliott, B.A., Rämö, O.T., 2000. 1.88-1.87 Ga post-kinematic intrusions of the Central Finland Granitoid Complex: a shift from C-type to A-type magmatism during lithospheric convergence. Lithos 53, 37–58. Nurmi, P.A., Haapala, I., 1986. The Proterozoic granitoids of Finland: Granite types, metallogeny and relation to crustal evolution.
Geol. Soc. Finland, Bull. 58, 203–233. Nykänen, O., 1983. Geological map of Finland 1:100 000, Pre-Quaternary rocks of the Punkaharju and Parikkala map-sheet areas. Geol. Surv. Finland. (in Finnish, with an English summary) Patchett, P.J., Kouvo, O., 1986. Origin of continental crust of 1.9-1.7 Ga age: Nd isotopes and U-Pb ages in the Svecokarelian Terrain of South Finland. Contrib. Mineral. Petrol. 92, 1–12. Patchett, P.J., Kouvo, O., Itedge, C.E., Tatsumoto, M., 1981. Evolution of continental crust and mantle heterogenity: evidence from Hf isotopes. Contrib. Mineral. Petrol. 78, 279–297. Pearce, J.A., Harris, N.B.W., Tindle, A.G., 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. J. Petrol. 25, 956–983. Perttunen, V., 1991. Geological map of Finland 1:100 000, Explanation to the Pre-Quaternary rocks of the Kemi, Karunki, Simo and Runkaus map-sheet areas. Geol. Surv. Finland. (in Finnish, with an English summary) Perttunen, V., Vaasjoki, M., 2001. U-Pb geochronology of the Peräpohja Schist Belt, northwestern Finland. In: M. Vaasjoki (Ed.), Radio metric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Pap. 33, 45–84. Perttunen, V., Hanski, E., Väänänen, J., Eilu, P., Lappalainen, M., 1996. Geological map of Finland 1:100 000, Explanation to the Pre-Quaternary rocks of the Rovaniemi map-sheet area. Geol. Surv. Finland. (in Finnish, with an English summary) Pitcher, W.S., 1982. Granite type and tectonic environment. In: K.J. Hsu (Ed.) Mountain Building Processes. Academic Press, London, 19–40. Pitkänen, P., 1985. Anttolan Luonterin postorogeenisen intruusion petrologia ja geokemia. M.Sc. Thesis, Univ. Helsinki. (in Finnish) Rämö, T., Nironen, M., 2001. The Oripää granite,
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SW Finland: Characterization and significance in terms of Svecofennian crustal evolution. Bull. Geol. Soc. Finland 73, 103–109. Rämö, O.T., Vaasjoki, M., Mänttäri, I., Elliott, B.A., Nironen, M., 2001. Petrogenesis of the postkinematic magmatism of the central Finland Granitoid Complex I; radiogenic isotope constraints and implications for crustal evolution. J. Petrol. 42, 1971–1993. Rastas, P., Huhma, H., Hanski, E., Lehtonen, M.I., Härkönen, I., Kortelainen, V., Mänttäri, I., Paakkola, J., 2001. U-Pb isotopic studies on the Kittilä greenstone area, central Lapland, Finland. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Pap. 33, 95–141. Rautiainen, J., 2000. Arkeeisen kratonin reunalla esiintyvät intermediääriset juonet Iisalmen, Juankosken ja Siilinjärven alueilla. M.Sc. Thesis, Univ. Helsinki. (in Finnish) Rickwood, P.C., 1989. Boundary lines within petrologic diagrams which use oxides of major and minor elements. Lithos 22, 247–263. Ruotoistenmäki, T., Mänttäri, I., Paavola, J., 2001. Characteristics of Proterozoic late-/ postcollisional intrusives in Archaean crust in Iisalmi-Lapinlahti area, central Finland. In: S. Autio (Ed.), Geological Survey of Finland, Current Research 1999–2000. Geol. Surv. Finland, Spec. Pap. 31, 105–115. Sederholm, J.J., 1932. On the geology of Fennoscandia with special reference to the preCambrian. Explanatory notes to accompany a general geological map of Fennoscandia. Bull. Comm. géol. Finlande 98, 1–30. Selonen, O., Ehlers, C., Lindroos, A., 1996. Structural features and emplacement of the late Svecofennian Perniö granite sheet in southern Finland. Bull. Geol. Soc. Finland 68, 5–17. Simonen, A., 1960. Plutonic rocks of the Svecofennides in Finland. Bull. Comm. géol. Finlande 189, 1–101. Simonen, A., 1980. The Precambrian in Finland. Geol. Surv. Finland, Bull. 304, 1–58.
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Suominen, V., 1991. The chronostratigraphy of southwestern Finland with special reference to Postjotnian and Subjotnian diabases. Geol. Surv. Finland, Bull. 356, 1–100. Sylvester, P.J., 1989. Post-collisional alkaline granites. J. Geol. 97, 261–280. Väänänen, J., 1998. Geological map of Finland 1:100 000, Explanation to the maps of Pre-Quaternary rocks, sheets 2713 and 2731. Kolarin ja Kurtakon kartta-alueiden kallioperä. Geol. Surv. Finland. (in Finnish, with English summary) Väänänen, J., Lehtonen, M.I., 2001. U-Pb isotopic age determinations from the Kolari-Muonio area, western Finnish Lapland. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Pap. 33, 85–93. Vaasjoki, M., 1977. Rapakivi granites and other postorogenic rocks in Finland: their age and the lead isotopic composition of certain associated galena mineralizations. Geol Surv. Finland, Bull. 294, 1–64. Vaasjoki, M., 1996. Explanation to the geochronological map of southern Finland: The development of the continental crust with special reference to the Svecofennian orogeny. Geol. Surv. Finland, Rep. Invest. 135, 1–30. Vaasjoki. M., Kontoniemi, O., 1991. Isotopic studies from the Proterozoic Osikonmäki gold prospect at Rantasalmi, southeastern Finland. In: S. Autio (Ed.), Geological Survey of Finland, Current Research 1989–1990. Geol. Surv. Finland, Spec. Pap. 12, 53–57. Vaasjoki, M., Sakko, M., 1988. The evolution of the Raahe–Ladoga zone in Finland: isotopic constraints. In: K. Korsman (Ed.), Tec tono-metamorphic evolution of the Raahe–Ladoga zone, eastern Finland. Geol. Surv. Finland, Bull. 343, 7–32. Vaasjoki, M., Huhma, H., Lahtinen, R., Vaarma, M., Vestin, J., 1998. The protoliths of Svecofennian granitoids in light of U-Pb ion microprobe measurements on the NORD-
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SIM instrument. In: J. R. Wilson (Ed.), 23. Nordiske Geologiske Vintermøde, Århus 13-16 January 1998: Abstract volume. Århus: University of Aarhus, 308. Väisänen, M., Mänttäri, I., Kriegsman, L.M., Hölttä, P., 2000. Tectonic setting of postcollisional magmatism in the Palaeoproterozoic Svecofennian orogen, SW Finland. Lithos 54, 63–81.
Väisänen, M., Mänttäri, I., Hölttä, P., 2002. Svecofennian magmatic and metamorfic evolution in soutwestern Finland as revealed by U-Pb zircon SIMS geochronology. Precambrian Res. 116, 111–127. van Duin, J.A., 1992. The Turku granulite area, SW Finland: a fluid-absent Svecofennian granulite occurrence. Ph. D. Thesis, Free Univ., Amsterdam.
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Chapter 11
PALEOPROTEROZOIC TECTONIC EVOLUTION
R. Lahtinen, A. Korja, M. Nironen CHAPTER
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Cover page: A frozen Proterozoic subduction zone in the Gulf of Bothnia (BABEL Working Group, 1990). Smoothed instantaneous amplitude of the migrated section of BABEL 4 reflection profile (image processed by Pekka Heikkinen, Institute of Seismology, University of Helsinki). The image comprises a 35-km-wide section of the crust and mantle at depths between 20 km and 70 km. Red and yellow colors image highly reflective crust, and pale blue denotes poorly reflective mantle.
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Lahtinen, R., Korja, A., Nironen, M., 2005. Paleoproterozoic tectonic evolution. 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. 481–532. © 2005 Elsevier B.V. All rights reserved.
An integrated study of petrological, geochronological, potential-field, deep seismic reflection and refraction, and geoelectric data leads to a new tectonic model for the Paleoproterozoic of the Fennoscandian Shield. The key issue in the model is the amalgamation of several microcontinents and island arcs with the following pre-1.92 Ga components: Karelian, Kola, and Norrbotten cratons (Archean); Keitele, Bergslagen, and Bothnia microcontinents (> 2.0 Ga); Kittilä island arc (~2.0 Ga); and Savo, Knaften, Inari, and Tersk island arcs (~1.95 Ga). We reject the concept of a semi-continuous Svecofennian (or Svecokarelian) orogeny and propose five orogenies for the period 1.92–1.79 Ga; these overlap partly in time and space and have different structural grains. The orogenic evolution is divided into (1) microcontinent accretion stage (1.92–1.87 Ga), (2) continental extension stage (1.86–1.84 Ga), (3) continent–continent collision stage (1.84–1.79 Ga), and (4) orogenic collapse and stabilization stage (1.79–1.77 Ga). Paleoproterozoic tectonic evolution of the Karelian craton, the Archean nucleus of the shield, involved a long period of rifting at 2.5–2.1 Ga that finally led to continental breakup at ~2.06 Ga. The microcontinent accretion stage included collision of the Kola and Karelian cratons (Lapland–Kola orogeny), collision of the Karelian craton with both the Norrbotten craton and the Keitele microcontinent, and docking of the Bothnia microcontinent (Lapland–Savo orogeny). The collision of the Bergslagen microcontinent with the newly-formed Archean–Paleoproterozoic complex led to the Fennian orogeny. During the continent–continent collision stage, two subduction zones, in the south and west, were active at 1.86–1.81 Ga. Extension of hot crust in the hinterlands of the subduction zones began at ~1.86 Ga. This was followed by oblique collision of Fennoscandia with Sarmatia at 1.84–1.80 Ga (Svecobaltic orogeny). A crustal-scale shear zone divided the Svecobaltic orogen into two different convergent areas, one related to a retreating Andean-type subduction zone in the southwest and another related to a transpressional regime in the southeast. Collision of Amazonia with Fennoscandia modified the central and northern parts of the western edge of the Fennoscandian Shield at 1.82–1.80 Ga (Nordic orogeny). Orogenic collapse and stabilization of the Fennoscandian Shield at 1.79–1.77 Ga were followed by the Gothian orogeny at the southwestern margin of the shield at 1.73–1.55 Ga.
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1. Introduction The East European craton (Figure 11.1A) is composed of the Fennoscandian, Sarmatian, and Volgo–Uralian crustal segments (Gorbatschev and Bogdanova, 1993) of which the latter two are mainly covered by Phanerozoic platform sediments. The northeastern part of the Fennoscandian Shield is composed of Archean rocks and the central part of Paleoproterozoic rocks, forming an area traditionally called the Svecofennian domain (Figure 11.1B). Initial correlations between the Svecofennian and younger orogens were made when Wegmann (1928) noted the similarities between Alpine ophiolites and the ultramafic rocks of eastern Finland. Simonen (1953) was the first to compare Svecofennian rocks of the Tampere schist belt (TB in Figure 11.2) to modern island arcs. Hietanen (1975) presented the first plate tectonic interpretation of the Svecofennian domain based on the correlation between the western Cordillera of North America and the Svecofennian rocks. Accumulation of geochronological data (mainly U-Pb on zircon) and structural mapping in the 1970’s allowed new plate tectonic interpretations and reconstructions, focusing mainly on the nature of the Archean–Proterozoic boundary in central Finland. Bowes and Gaál (1981) and Koistinen (1981) introduced the term geosuture for this boundary. Gaál (1982) presented a plate tectonic model with subduction toward the east-northeast and collision ~1.9 Ga ago. Park (1983) proposed a model with the development of a continental back-arc, the Outokumpu assemblage, at 1.97 Ga. Later Park et al. (1984) and Park (1985) proposed a Western Cordillera-type model with exotic terranes juxtaposed along the Archean craton margin. Gaál (1986) presented a comprehensive model for the evolution of the Fennoscandian Shield (cf. Gaál and Gorbatschev, 1987). The basic concept for the Svecofennian orogeny at 1.92 Ga was an ENE-directed subduction 484
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zone and the development of an associated back-arc basin along the craton margin. At the same time, a new subduction zone developed farther southwest, leading to the formation of Svecofennian crust in an island arc setting (cf. Baker et al., 1988). The scarcity of subduction-related magmatism in the Archean craton margin and the easterly-directed tectonic transport prompted Ward (1987) to suggest westerly-directed subduction before collision at the Archean–Proterozoic boundary. Gaál (1990) adopted the idea of Ward (see also Tuisku and Laajoki, 1990) and included a subduction reversal in his model to account for the voluminous magmatism in central Finland. Geophysical investigations (especially seismic refraction and reflection studies) in the 1980’s allowed the study of the crust in the vertical dimension and also inspired correlations with modern analogues (BABEL Working Group, 1990). In an integrated geological-geophysical study Korja et al. (1993; see also Korja and Hjelt, 1993) attempted to locate the sutures and terrane boundaries of the Svecofennian domain and proposed a mantle-underplating model to account for the thick crust in central Finland. Gorbatschev and Bogdanova (1993) extended the tectonic discussion, by the use of drill-core and geophysical data, to the extensions of the shield that are covered by Phanerozoic rocks. Ekdahl (1993) and Ruotoistenmäki (1996) adopted the model of Gaál (1986) with subduction toward the northeast. Lahtinen (1994) presented a plate tectonic model for the Svecofennian domain of Finland involving several accretionary units and three collisional stages, at 1.91–1.90 Ga, 1.89–1.88 Ga, and 1.86–1.84 Ga. Kohonen (1995) studied the extended eastern part of the craton and its cover in eastern Finland and proposed a model including arc–continent collision that began in the north and continued to the south. Korja (1995) introduced the concept of orogenic collapse to account for the variation in crustal thickness in
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Transscandinavian granite-porphyry belt Fig. 11.1. (A) Crustal segments of the East European craton (modified after Gorbatschev and Bogdanova, 1993). FS – Fennoscandian Shield; US – Ukrainian Shield. (B) Major tectonic units of the Fennoscandian Shield modified after Gaál and Gorbatschev (1987). CHAPTER
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21°E Barents Sea Murmansk
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Archean Igneous rocks and gneisses (3.20–2.50 Ga) Supracrustal rocks (3.20–2.75 Ga)
Mesoproterozoic Rapakivi granite association (1.65–1.47 Ga) Sedimentary rocks (1.50–1.27 Ga)
Paleoproterozoic Supracrustal rocks (2.50–1.96 Ga)
Neoproterozoic Sveconorwegian orogenic belt (1.10– 0.92 Ga), partly reworked Paleo- to Mesoproterozoic rocks
Mafic intrusive rocks (2.50–1.96 Ga) Granulite belt
(>1.90 Ga)
Supracrustal rocks (1.96–1.84 Ga) Igneous rocks
(1.96–1.84 Ga)
Granite and migmatite (1.85–1.75 Ga)
Phanerozoic Caledonian orogenic belt (0.51– 0.40 Ga) Alkaline intrusions Sedimentary rocks
Igneous rocks, TIB1 (1.85–1.76 Ga) Igneous rocks, TIB2 (1.71–1.66 Ga)
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21°E
9°E
33°E
Fa(DGRF65 (nT) 600
68°N
500 400 300 200 100 0 -100
64°N
-200 -300 -400
60°N
Magnetic anomaly reduced to the pole
56°N
21°E
Continued upwards 5 km
Fig. 11.3. Magnetic anomaly map of the Fennoscandian Shield after Korhonen et al. (2002). Total intensity anomaly is reduced to the pole, continued upward to 5 km above ground, scale 1:15,000,000. Horizontal gradients are emphasized by vertical illumination of total intensity. Fig. 11.2. (facing page) Simplified geological map of the Fennoscandian Shield based on Koistinen et al. (2001). Archean units: Norrbotten craton, Kola craton, and Karelian craton, including Belomorian mobile belt. Karelian craton: Ka1 – Central Karelian complex; Ka2 – Iisalmi complex; Ka3 – Pudasjärvi complex. Paleoproterozoic units in Kola Peninsula: IA – Inari area; PeB – Pechenga belt; IVB – Imandra– Varzuga belt; UGT – Umba granulite terrane; TT – Tersk terrane. Paleoproterozoic units in Finland: LGB – Lapland granulite belt; KA – Kittilä allochthon; CLGC – Central Lapland granitoid complex; PB – Peräpohja belt; KB – Kuusamo belt; SB – Savo belt; CFGC – Central Finland granitoid complex; TB – Tampere belt; HB – Häme belt; UB – Uusimaa belt. Paleoproterozoic units in Sweden: SD – Skellefte district; BB – Bothnian basin; BA – Bergslagen area; OJB – Oskarshamn–Jönköping belt; TIB – Transscandinavian igneous belt. Boundaries: BBZ – Baltic–Bothnian megashear; PRZ – Piteå–Raahe shear zone; HSZ – Hassela shear zone. Specific localities: J – Jormua; O – Outokumpu; K – Knaften. CHAPTER
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southern and central Finland. Nironen (1997) presented a kinematic model for the Svecofennian orogen: opening of an ocean basin at 1.95 Ga and progressive accretion of two arc complexes to the Karelian craton (Karelian province in Figure 11.1B) at 1.91–1.87 Ga. Korsman et al. (1999) gave the latest summary of the Svecofennian orogen along the GGT/SVEKA Transect (Figure 11.4). Different plate tectonic models have also been presented for the Lapland–Kola orogen (LGB, UGT, and TT in Figure 11.2). These involve two interpretations based on the proposed location of the subduction zone. In the first, subduction is toward the northeast and the suture zone is within the Lapland granulite belt (Barbey et al., 1984; Krill, 1985; Daly et al., 2001). In the second, subduction is toward the southwest and the suture zone is in the Imandra–Varzuga and Pechenga belts (IVB and PeB in Figure 11.2; Berthelsen and Marker, 1986a; Marker, 1990). The tectonic models above have mainly focused on the evolution of the Svecofennian domain especially in the vicinity of the Archean–Proterozoic boundary. Most commonly, accretionary-type tectonics have been proposed to account for the Proterozoic crustal growth. It has been suggested that the accretionary orogens become younger toward the west (cf. Gothian evolution at 1.75–1.55 Ga; Åhäll and Larson, 2000). These orogenic domains were reworked in the Sveconorwegian/Grenvillian orogeny at 1.2–0.9 Ga (e.g., Gorbatschev and Bogdanova, 1993; Åhäll and Larson, 2000). The model by Nironen (1997) is hitherto the only tectonic model that attempts to explain the evolution of the entire Svecofennian domain. Reinterpretation of seismic reflection and refraction data (e.g., Korja et al., 2001; Luosto and Heikkinen, 2001), new comprehensive isotope and age data (especially from the NORDSIM facility), the results from the EUROBRIDGE (e.g., Claesson et al., 2001) and SVEKALAPKO (e.g., Daly et al., 2001) 488
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projects, and the outcome of several regional to shield-scale mapping projects (Lehtonen et al., 1998; Kousa and Lundqvist, 2000; Bergman et al., 2001; Koistinen et al., 2001; Nironen et al., 2002) have produced large amounts of new data on plate and terrane boundaries. One major goal of this paper is to integrate sample-scale petrologic, chronologic, and isotope data with both regional-scale lithologic and potential-field data, crustal-scale seismic refraction and reflection results, and geoelectric data. Variable data from surface rocks are available for exposed areas, which often lack seismic reflection data. Good seismic profiles are available from the Baltic Sea and Gulf of Bothnia, without any bedrock data. Areas covered by Phanerozoic platform sediments, with sparse drilled bedrock data, often lack published seismic reflection data. Therefore, the evolution of the southeastern part of the Karelian craton and the junction between Fennoscandia and the Volgo–Uralian crustal segments are not discussed. Furthermore, there is a slight bias to Finland and to the time period between 1.93 and 1.77 Ga – the late Paleoproterozoic to Mesoproterozoic Gothian orogenic events will only be touched briefly. The wealth of new data from the Fennoscandian Shield allows an attempt of a new comprehensive tectonic model for the Paleoproterozoic evolution of the Fennoscandian Shield. We regard it as a testable model that can lead to a more thorough understanding of the processes that operated in the Fennoscandian Shield during the Paleoproterozoic. Many of the ideas in this study have evolved from the authors’ earlier publications (e.g., Korja et al., 1993; Lahtinen, 1994; Korja, 1995; Nironen, 1997; Lahtinen and Huhma, 1997; Nironen et al., 2000a; Lahtinen, 2000; Korja and Heikkinen, 2000; Korja et al., 2001), and these papers will be referred to only in special cases. We also emphasize that in order to keep the number of references to a minimum we have mostly referred only to key papers and reviews.
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2. Geologic outline The Fennoscandian crustal segment (Gorbatschev and Bogdanova, 1993) consists of the Fennoscandian Shield, its southern continuation covered by platform sediments, and the Caledonides in the west (Figure 11.2). The aeromagnetic map (Figure 11.3) shows the continuation of lithologic units under the sedimentary cover. All three crustal segments within the East European craton record both Archean and Paleoproterozoic evolution (cf. Figure 11.1). The Archean bedrock in the Fennoscandian Shield includes nuclei dispersed and reassembled during the Paleoproterozoic. We divide the Archean into the Karelian craton, including the Belomorian mobile belt, and the Kola craton (Figure 11.2) (cf. Gaál and Gorbatschev, 1987). We also introduce the term Norrbotten craton for the Archean rocks west of the megashear defined by Berthelsen and Marker (1986b). The Karelian craton consists of Archean granitoid–gneiss complexes and supracrustal rocks (e.g., greenstones) ranging in age between 3.2 Ga and 2.5 Ga (see Chapter 2). The Karelian may be divided into the Belomorian mobile belt and three complexes (Figure 11.2): the Central Karelian complex (Ka1), the Iisalmi complex (Ka2), and the Pudasjärvi complex (Ka3). The Archean rocks continue under the Central Lapland granitoid complex and farther north under the Paleoproterozoic cover where they are again exposed in areas south and west of the Lapland granulite belt. The boundary zone between the Central Karelian complex and the Belomorian mobile belt was formed by sequential accretion of island arc and continental fragments to the Karelian core in the Neoarchean (e.g., Mints et al., 2001). This boundary zone and the whole Belomorian mobile belt were strongly reactivated in the Paleoproterozoic (Gaál and Gorbatschev, 1987; Bibikova et al., 2001). The Proterozoic rocks on the Archean cratons comprise autochthonous supracrustal CHAPTER
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rocks, including rift-type basalts, deposited from 2.45 Ga onwards, allochthonous younger series, and ~1.95 Ga ophiolites (Chapters 4, 6, and 7; Melezhik and Sturt, 1994; Bergman et al., 2001). One important feature in northern Finland is the Kittilä allochthon (KA in Figure 11.2), composed partly of oceanic crust (Chapter 4). Svecofennian (1.9–1.8 Ga) plutonic rocks intrude large areas of Paleoproterozoic cover rocks in northern Sweden and northern Finland (Chapter 10; Bergman et al., 2001). The 1.95 Ga Knaften granitoid (Wasström, 1993) south of the Skellefte district in Sweden (K and SD in Figure 11.2) and the 1.92 Ga primitive island arc rocks in the Savo belt (SB, Figure 11.2; Korsman et al., 1997) adjacent to the Archean craton in Finland are the oldest documented rocks in the Paleoproterozoic Svecofennian part in the Fennoscandian Shield. However, an older source (~2.1–2.0 Ga) has been proposed as a nucleus for the Central Finland granitoid complex (Lahtinen and Huhma, 1997). New Nd isotope data support this assumption (Rämö et al., 2001). Island arc-type volcanic rocks (1.90–1.87 Ga), varying from less mature in the Skellefte district (SD in Figure 11.2) to mature in the Tampere belt (TB in Figure 11.2), and calc-alkaline granitoids are dominant in the central Fennoscandian Shield. The latest major magmatic episodes in the central part of the Fennoscandian Shield are marked by plutonic rocks (1.88–1.87 Ga), with local A-type characteristics, in central Finland and 1.80–1.78 Ga granitoids in Sweden. Migmatites with tonalite leucosome were formed at 1.89–1.88 Ga, whereas younger migmatization associated with granite leucosome and S-type granites took place at 1.84–1.82 Ga (e.g., Weihed et al., 1992; Korsman et al., 1999). The southern part of the Svecofennian domain includes the 1.90–1.89 Ga Bergslagen area (BA) and Uusimaa belt (UB), in part formed in an intra-arc basin of a mature continental arc (e.g., Kähkönen et al., 1994; Allen
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Po lar
21°E
Fen nolo
ra
4B 2
3
Balti
c
4
81
Sv ek
a
4A
Fennia
Sv ek
a9
1
1
6
60°N
C
B ic S
ea
BABEL reflection line
Balt
Refraction line
Eur ob
Reflection line rid
A
ge
0
250
500 km
Fig. 11.4. Deep seismic refraction and reflection lines of the Fennoscandian Shield referred to in this paper. Refraction lines: Fennia (Fennia Working Group, 1998), Sveka81 (Grad and Luosto, 1987), Sveka91 (Luosto et al., 1994), Baltic (Luosto et al., 1990), Polar (Luosto et al., 1989), Fennolora (Guggisberg, 1986), Eurobridge (EUROBRIDGE Seismic Working Group, 2001), Baltic Sea (Ostrovsky, 1998). Reflection lines: BABEL A-C, 1-7 (BABEL Working Group, 1990), 4B (Mints et al., 2001). The bedrock map is modified from Koistinen et al. (2001), see Figure 11.2 for details.
et al., 1996a). Less-evolved island arc volcanic rocks are found in the Häme belt (HB in Figure 11.2; Chapter 8). Paleoproterozoic metapelitedominated sedimentary rocks, quartzites, and carbonate rocks characterize the southern part 490
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of the Svecofennian domain. Plutonism in the southern part of the Svecofennian domain shows age groups of 1.89–1.85 Ga, 1.84–1.82 Ga, and 1.81–1.79 Ga. The 1.84–1.82 Ga-group granites (S-type)
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21°E
60°N
-50
Moho-depth contour in km Upper surface of dipping mantle reflector
0
250
500 km
Fig. 11.5. Moho-depth map (Luosto, 1997) and upper surfaces of dipping mantle reflectors compiled from reflection and refraction studies (BABEL Working Group, 1990; Abramovitz et al., 1997; Ostrovsky, 1998; Balling, 2000; Heikkinen and Luosto, 2000; Luosto and Heikkinen, 2001). The bedrock map is modified from Koistinen et al. (2001), see Figure 11.2 for details.
and migmatites with granite leucosomes form a belt that extends from southeastern Finland to central Sweden (e.g., Korsman et al., 1999). The 1.83 Ga Oskarshamn–Jönköping belt CHAPTER
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(OJB; Figure 11.2) south of the Bergslagen area is characterized by supracrustal rocks intruded by calc-alkaline I-type syntectonic granitoids (Mansfeld, 1996). Otherwise, the central part of southern Sweden is dominated
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21°E
60°N
Magnetic lineament Magnetic and Bouguer lineament Bouguer lineament Bouguer lineament associated with shear zone at surface 0
250
500 km
Fig. 11.6. Major aeromagnetic and Bouguer anomaly lineaments of the Fennoscandian Shield. Aeromagnetic lineaments are interpreted from Korhonen et al. (1999) and Bouguer lineaments from Elo (1992) and S. Aaro/Geological Survey of Sweden (1997). The bedrock map is modified from Koistinen et al. (2001), see Figure 11.2 for detais.
by granitoids of the Transscandinavian igneous belt (TIB in Figure 11.2; Patchett et al., 1987) forming a N–S trending belt. Three age groups of volcanic and plutonic rocks (Larson and Berglund, 1992; Åhäll and Larson, 2000) have been identified: TIB 1 (1.81–1.77 Ga), TIB 2 (~1.7 Ga), and TIB 3 (1.68–1.65 Ga), of which TIB 1 is the most voluminous. 492
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The Fennoscandian Shield becomes younger toward the west, which reflects in decreasing granitoid ages (Figure 11.2), the youngest of which belong to the Sveconorwegian orogeny (e.g., Gaál and Gorbatschev, 1987).
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21°E
60°N
Conductors at depth >20 km Conductors close to surface
0
250
500 km
Fig. 11.7. Electromagnetic conductivity anomalies after Korja et al. (2002). The bedrock map is modified from Koistinen et al. (2001), see figure 11.2 for details.
3. Pre-1.92 Ga crustal components and crustal-scale boundaries There is increasingly growing evidence for >1.92 Ga crustal growth in the Fennoscandian Shield. Geochemical and isotope data (Valbracht et al., 1994; Lahtinen and Huhma, 1997; Andersson, 1997; Rämö et al., 2001) indicate that microcontinents, now seen as CHAPTER
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crustal domains, had already started to form 2.1–2.0 Ga ago. Abundant 2.1–1.95 Ga detrital zircons also suggest that pre-1.92 Ga crustforming processes were important (Huhma et al., 1991; Claesson et al., 1993; Lahtinen et al., 2002). However, it is not clear where these domains were formed. In areas lacking seismic reflection or refraction data (Figure 11.4), major crustal-scale
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boundaries are inferred solely from lineaments on magnetic, electromagnetic, and Bouguer anomaly maps (Figures 11.3, 11.6, 11.7) and comparable trends on the Moho-depth map (Figure 11.5). Geoelectric data (Figure 11.7) and deep seismic refraction and reflection data (Figure 11.4) allow us to delineate crustalscale boundaries and locate mantle reflectors (Figure 11.5; see also cover page). A geological interpretation of the BABEL profiles from the southern Baltic Sea to the northern part of the Gulf of Bothnia includes three inferred subduction zones (Figure 11.8). Lithologic, geochemical, isotope, and geophysical data allow to disinguish the following pre-1.92 Ga components (Figure 11.9A): Archean Karelian, Kola, and Norrbotten cratons; >2.0 Ga Keitele, Bergslagen, and Bothnia microcontinents; ~2.0 Ga Kittilä island arc and oceanic crust; ~1.95 Ga Savo, Knaften, Inari, and Tersk island arcs. The Karelian and Kola cratons are well exposed, the Norrbotten craton is not. The Paleoproterozoic microcontinents – Keitele, Bergslagen, and Bothnia – have no identified surface expressions. The Kittilä and Savo arcs are partly exposed and only small slivers of the Knaften arc are found on the surface. The relationship of the Inari and Tersk arcs with the Archean crust is not well known. The exposed and hidden pre-1.92 Ga components are outlined in Figure 11.9A, whereas crust with unknown mantle separation age, and crust dominated by accreted sediments are shown in white. Major Paleoproterozoic units are shown in Figure 11.9B. Rocks south of the Bergslagen area and rocks in the Baltic countries have similar NW-striking geophysical patterns (Figures 11.3 and 11.6) and lithologic continuities, and therefore have been grouped into the Svecobaltia area. The Umeå area has been delineated as a separate unit, based on reflection seismic data (Figure 11.8 and data from line 2; Korja and Heikkinen, 2000). These show different reflecting patterns for the uppermost 15 km to 25 km of the crust compared 494
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with the middle and lower crust.
3.1. Lapland–Kola area In our model, the boundary between the Kola and Karelian cratons (Figure 11.9A) coincides with the Paleoproterozoic Imandra–Varzuga and Pechenga belts (e.g., Berthelsen and Marker, 1986a) and is associated with SWdipping electrical conductors (Figure 11.7). This belt (IVB and PeB in Figure 11.2) and its continuation in Norway form the largest riftrelated belt in the Kola craton and displays a long evolutionary history from 2.5 Ga to 1.8 Ga (Melezhik and Sturt, 1994). The Karelian side of the Kola–Karelian boundary comprises several Paleoproterozoic accreted terranes in the Lapland–Kola area (Figure 11.2): Inari area (IA), Lapland granulite belt (LGB), Umba granulite terrane (UGT), and Tersk terrane (TT) (Korsman et al., 1997; Daly et al., 2001). The Archean Karelian craton has also been affected by Paleoproterozoic processes. These include magmatism, e.g., in the Inari area, and as strong crustal reworking, e.g., in the Belomorian mobile belt (Bibikova et al., 2001). The Inari area (Figure 11.2) comprises Archean gneisses and Paleoproterozoic metavolcanic and granitoid rocks. The calc-alkaline nature of the 1.94–1.93 Ga plutonic rocks with TDM model ages of 2.47–2.07 Ga imply a moderately evolved arc environment without a strong input from the Archean crust (Barling et al., 1997). The 1.93 Ga enderbites (Meriläinen, 1976) in the Lapland granulite belt have TDM model ages of 2.47–2.17 Ga (Bernard-Griffiths et al., 1984; Daly et al., 2001), similar to those of the Inari area granitoids. The metasedimentary granulites, the khondalite series, have been interpreted as marine turbidites (Barbey et al., 1982), and they have a mixed Archean to dominantly ~2.0 Ga Paleoproterozoic source based on detrital zircon and Nd isotope data (Sorjonen-Ward et al., 1994; Tuisku and Huh-
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km
S
0
BABEL B
PA L E O P R O T E R O Z O I C T E C T O N I C
50
150
250
BB
100
MR
N SW
300 km 0
50
4
UA
50
Mafic nucleus
NE
100
SE
NW S
100
75
BA
C
3
200
50
200
B
4 1
3
50
68°
N
0 km
Mantle reflections in refraction data
Major thrust or detachment
Reflective structures
16°
BABEL C
6
4A
16°
100
Sandstone
68°
BABEL 6
150
BB
Crustal scale strike-slip shear zone
NE
250 km
N S
0
PRZ
25
HSZ
Gabbro–anorthosite
Rapakivi granite
150
C1 BABEL C
125
4A BABEL 4, 4A & 3
0 km
Mafic–intermediate intrusion
MR
200
200
Paleoproterozoic microcontinent
150
250
TIB granite
100
HB
300
BA
Archean craton
UB
350
OJB
0 10 20 30 40 50 60
Fig. 11.8. A geological interpretation of the BABEL lines B, C, 6, 1, 3&4 (see Figure 11.4) modified after Korja and Heikkinen (2000). The line drawing has no vertical exaggeration. Geologic units are as in Figures 11.2 and 11.9B. Colored lines denote reflections arising from units with different reflection properties, mafic sills are in black. Mantle reflectors (MR) are from Abramovitz et al. (1997) and Heikkinen and Luosto (2000).
BABEL 1
0 10 20 30 40 50 60 70
SW
400
TIB
km
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•
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A
B Inari
Norr- Kittilä botten arc craton
Bothnia mc
Inari
be lt
Norrbotten craton Skellefteå Umeå
Kola craton
oro ge nic
Knaften arc
500 km
Kola craton
Tersk arc Karelian Belo mo craton ria n
Ca led on ian
500 km
Keitele mc
Bothnia
Sveconorwegian orogen
Bergslagen arc
Gothian orogen
Archean crust
~2.0 Ga crust
Archean / ~2.0 Ga crust
~1.95 Ga crust
Bergslagen
Karelian craton
Keitele Tavastia
Svecobaltia
Fig. 11.9. Pre-1.92 Ga crustal components and major terranes/units in the Fennoscandian Shield. (A) Exposed and hidden pre-1.92 Ga components are outlined with a broken line; juvenile crust, crust without known mantle separation age or crust dominated by accreted sediments are shown in white. (B) Major geologic units with Paleoproterozoic and younger boundaries.
ma, 1998; Daly et al., 2001). A belt of highly sheared rocks rims the Lapland granulite belt in the southwest (Marker, 1990). The Umba granulite terrane and the Tersk terrane comprise granulite-facies paragneisses and metavolcanic and metasedimentary rocks, respectively (Daly et al., 2001). The age of the arc magmatism in the Tersk terrane is ~1.96 Ga and the TDM model ages are ~2.2 Ga for both metavolcanic and metasedimentary rocks (Daly et al., 2001). Although separated by a major shear zone seen as a magnetic lineament (Figures 11.3 and 11.5), the Umba granulite terrane and the Lapland granulite belt have been correlated (e.g., Daly et al., 2001). The Lapland granulite belt appears as a NE-dipping block in seismic refraction, reflection, and magnetotelluric models (Behrens et al., 1989; Luosto et al., 1989; Walther and Flüh, 1993; Korja et al., 1989). The NE-dipping internal structures of the block have been interpreted as thrust structures related to the final emplacement of the Lapland granulite 496
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belt (Gaál et al., 1989; Marker et al., 1990; Korja et al., 1996). Marker et al. (1990) modeled the Lapland granulite belt as a NE-dipping thrust wedge that is rooted in the middle crust beneath the Inari area. It should be noted, however, that the crust is slightly thinned under the Lapland granulite belt (Figure 11.5). Fewer data are available from the Umba–Tersk areas but the interpretation of potential-field data (cf. Daly et al., 2001) and geoelectric data (Figure 11.6) favors southward-dipping structures in the eastern part of the Lapland–Kola area.
3.2. Karelian craton The mappable boundaries of the Karelian craton in the northeast, west, and southwest (Figures 11.6 and 11.9) probably delineate Paleoproterozoic sutures. The boundaries in the east and southeast are undefined and hidden under the Phanerozoic cover (Figure 11.2). Note that the Belomorian mobile belt, strongly reworked during the Paleoproterozoic, includes an Archean suture. The Archean
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rocks in central Lapland are located mainly under the Paleoproterozoic cover and, possibly, under the Central Lapland granitoid complex (GLGC) (Figure 11.2). Numerous rifting episodes occurred within the Karelian craton at 2.5–1.95 Ga. The Paleoproterozoic cover sequence in central Lapland in Finland (Chapter 4) starts with rocks formed in an intracratonic rift environment at 2.5–2.4 Ga, followed by quartzite formations deposited in a cratonic or marginal basin setting at >2.2 Ga. During basin deepening, fine-grained sediments and mafic and ultramafic volcanic rocks were deposited between 2.2–2.0 Ga. The ~2.0 Ga Kittilä allochthon (KA in Figure 11.2) represents a block of ancient oceanic lithosphere later obducted onto the rift sequence. N–S magnetic and Bouguer anomaly lineaments are found in the Karelian craton, and E–W and NE–SW lineaments in central Lapland (Figures 11.3 and 11.6). On the Moho-depth map (Figure 11.5), the isolines strike north–south within the Karelian craton and east–west within Lapland. A Paleoproterozoic cratonic stage (2.5–2.1 Ga) with multiple rifting (e.g., Vuollo, 1994; Kohonen, 1995) also characterizes the Karelian craton in eastern Finland. The 1.95 Ga Jormua ophiolite (J; Figure 11.2) is a fragment of Red Sea-type crust comprising metabasalts derived from EMORB- to OIB-like sources (Kontinen, 1987; Peltonen et al., 1996). An age of 1972 ± 18 Ma (Huhma, 1986) has been obtained from a gabbro associated with the Outokumpu serpentinites (O: Figure 11.2), which are part of a possible mantle fragment (Chapter 6). A comparable 1965 ± 18 Ma age has been obtained from tholeiitic dikes cutting the Archean basement (Vuollo et al., 1992; Chapter 5). This part of the Karelian craton is dominated by large-scale N–S magnetic and Bouguer lineaments (Figure 11.6). In refraction models, the Outokumpu structure is imaged as a high-velocity nappe (Luosto et al., 1990).
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3.3. Norrbotten Archean nucleus and attached island arcs Archean crust in northern Sweden is exposed only locally (Bergman et al., 2001). Small windows of Archean rocks are found (e.g., Lundqvist et al., 1996) and Nd isotope data indicate the existence of Archean crust below the Paleoproterozoic cover (Öhlander et al., 1993, 1999; Mellqvist et al., 1999). More juvenile arc material, including the Skellefte district, formed on the attached Knaften arc (Figure 11.9A) to the south. The western edge of the Norrbotten craton is buried under the Caledonides and the eastern limit is considered to coincide with the Baltic–Bothnian megashear (BBZ in Figure 11.2; Berthelsen and Marker, 1986b). The eastern part of the Norrbotten craton is dominated by large-scale N–S Bouguer lineaments overprinted by NNW–SSE structures (Figures 11.3 and 11.6). In the Kittilä allochthon (KA; Figures 11.2 and 11.9A), a chain of serpentinites, interpreted as dismembered ophiolitic mantle rocks, and marine metavolcanic rocks (~2.0 Ga) comprise an oceanic island arc in which ophiolitic mantle rocks form the base (Chapter 4). The latest magmatic activity in the Kittilä allochthon took place during the initial collision at ~1.92 Ga. The Kittilä island arc (allochthon) was originally attached to the Norrbotten nucleus and was subsequently overthrust onto the Karelian craton. MORB- to island arc-type metavolcanic rocks with intercalated turbiditic and graphite-bearing argillitic metasediments are found in Knaften (K; Figure 11.2), south of the Skellefte district (SD; Wasström, 1990). The sequence has been intruded by granitoids (1954 ± 6 Ma; Wasström, 1993) and rhyolite porphyry dikes (1940 ± 14 Ma; Wasström, 1996). A comparable age (1959 ± 14 Ma) of a metadacite interlayered with metagraywackes west of the Skellefte distric (Eliasson and Sträng, 1998) verifies the wider occurrence of 1.96–1.95 Ga island arc magmatism. Although
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data are sparse on the pre-1.92 Ga evolution of the Knaften island arc, it was most likely attached to the Norrbotten Archean craton before 1.92 Ga (Figure 11.9A).
3.4. Keitele microcontinent and attached island arc The presence of a 2.1–2.0 Ga continental nucleus (Figure 11.9A), coined here as the Keitele microcontinent, has been proposed on the basis of geochemical and isotopic characteristics of the Central Finland granitoid complex (CFGC; Figure 11.2). Both mature crust and enriched subcontinental lithosphere are present (Lahtinen, 1994; Lahtinen and Huhma, 1997). More juvenile, 1.93–1.92 Ga gneissic tonalites and associated supracrustal rocks are found within the Savo belt (SB) at the Archean–Proterozoic boundary (Lahtinen, 1994; Kousa et al., 1994). These rocks form the Primitive arc complex of Korsman et al. (1997). No significant Archean component is found in these rocks (Lahtinen and Huhma, 1997), but a clear suture between Keitele and the arc complex has not been identified. Thus, we interpret the 1.93–1.92 Ga rocks of the Savo belt (SB) to have formed on an island arc crust that had been attached to the Keitele microcontinent prior to 1.92 Ga. At the southern edge of the Keitele microcontinent, the Tampere belt 1.90–1.89 Ga volcanic rocks represent evolved arc volcanism (Kähkönen, 1989; Chapter 8), probably in a fore-arc position. The lower marginal-basin (Kähkönen and Nironen, 1994) or rift-basin (Lahtinen, 1994) metalavas in the Tampere belt are older than 1.91 Ga. This relation is based on their occurrence below graywackes (Kähkönen et al., 1994) that have a maximum deposition age of about 1.91 Ga (detrital zircon data; Huhma et al., 1991; Lahtinen et al., 2002) and on a 1904 Ma zircon age (Kähkönen et al., 1989) from the overlying volcanic rocks. Mantle-like lead and εNd (at 1.9 Ga) value of +0.5 (Vaasjoki and Huhma, 1999) of 498
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these lavas suggest that an enriched (relative to depleted mantle) lithospheric mantle was present below the Keitele microcontinent at ~1.91 Ga. The hidden Keitele microcontinent and younger arc material currently have a crustal thickness exceeding 52 km (Figure 11.5). The shape of the microcontinent is imaged by geoelectric conductors (Figure 11.7) and is supported by potential-field data (Figure 11.6). Thus, the Keitele area (Figure 11.9B) comprises the Central Finland granitoid complex (evolved arc), the Tampere belt (forearc), and the adjacent migmatite belt (accretionary wedge)(see Figure 11.2).
3.5. Bothnia microcontinent and attached island arc Andersson (1997) argued for the existence of Archean lower crust in the central part of the Bothnian basin (BB in Figure 11.2) based on Nd and Pb isotope and ion microprobe zircon data from the 1.53–1.50 Ga rapakivi granites. Xenocrystic 2.7 Ga zircons from a related syenite confirm the presence of Archean material (Claesson et al., 1997). The interpretation of the isotope data (Andersson, 1997) is not straightforward and an alternative could be substantial input from a heterogeneous (2.1–2.0 Ga) subcontinental lithospheric mantle (e.g., Lahtinen and Huhma, 1997). A significant pre-1.9 Ga crustal component is also evident in the 1.8 Ga granitoids (Claesson and Lundqvist, 1995). A hidden Bothnia microcontinent below the metasedimentary rocks is needed to explain the observed isotopic features. A block of higher Bouguer anomaly values (Korhonen et al., 1999) bounded in the north by an ENE–WSW lineament (Figure 11.6) and in the south by the Hassela shear zone (HSZ; Högdahl, 2000; Högdahl and Sjöström, 2001; Korja et al., 2001) is thought to represent the hidden Bothnia microcontinent. South of the Bothnia microcontinent, within
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the southern part of the Gulf of Bothnia, the Bouguer anomaly patterns, seismic structures, and mantle reflectors strike east–north-east (Figures 11.5, 11.6, and 11.8). A mafic fragment (Figure 11.8) found offshore (Korja et al., 2001) is interpreted as the mafic core of an oceanic island arc or seamount attached to the Bothnia microcontinent. The Bothnia microcontinent is overlain by the Umeå allochthon (Figure 11.9B) having a low level of magnetization, high conductivity, and shallowly dipping reflective structures.
3.6. Bergslagen microcontinent and Tavastia island arc The southernmost part of the Uusimaa Belt (UB) (Figures 11.2 and 11.9B) in Finland shows evidence for evolved crust with a long crustal history (Lahtinen and Huhma, 1997; Rämö et al., 2001). This part of the Uusimaa belt may be correlated with the active continental margin-type crust found in Bergslagen (e.g., Valbracht et al., 1994; Allen et al., 1996a; Nironen, 1997). We call this combined evolved crustal unit with a long evolution (starting at 2.1–2.0 Ga) the Bergslagen microcontinent. Based on geochemical and isotope data, less evolved and younger (< 2.0 Ga) island arc-type crust dominates southern Finland in areas directly north of the Bergslagen microcontinent (Hakkarainen, 1994; Lahtinen, 1996; Lahtinen and Huhma, 1997; Rämö et al., 2001; Chapter 8). This arc-type crust loosely correlates with the Häme belt (HB; Figure 11.2) and it is here called the Tavastia island arc. The boundary between the Bergslagen evolved crust and the Tavastia island arc is difficult to ascertain because of later tectonic transport and thus little is known about the pre-1.90 Ga evolution of these units. The Bergslagen microcontinent is characterized by low values of magnetization (Figure 11.3) and low Bouguer anomaly values. The straightforward correlation of southern Finland and Bergslagen may be hampered CHAPTER
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by right-lateral displacement along a crustalscale shear zone interpreted from Bouguer and magnetic maps.
4. Terminology related to Paleoproterozoic tectonic evolution Gaál and Gorbatschev (1987) divided the Fennoscandian Shield into three domains – the Karelian domain (Archean), the Svecofennian domain (Paleoproterozoic), and the Southwest Scandinavian domain (Mesoproterozoic; Figure 11.1). The Proterozoic orogenies are named Svecofennian (earlier the Svecokarelian; Rankama and Welin, 1972) and Gothian; those were associated with the formation of the Svecofennian and Southwest Scandinavian domains, respectively. The Transscandinavian igneous belt (TIB; Patchett et al., 1987) is an important N–S belt on the western margin of the Svecofennian domain. It is divided by age into three groups of granites (TIB1 through 3; Larson and Berglund, 1992). Åhäll and Larson (2000) coined the 1.81–1.77 Ga TIB1 magmatism Smålandian, representing a postSvecofennian tectonic stage, and included the TIB 2 and TIB 3 age groups in the Gothian evolution. The western margin of the shield was further reworked during the Sveconorwegian–Grenvillian orogeny (Gaál and Gorbatschev, 1987). The complexity of the bedrock as well as seismic reflection and refraction and geoelectric data suggest that the Svecofennian domain contains several sutures, most of them lacking a clear surface expression. The interpretations of the structural history are variable, especially in the southern part of the Svecofennian domain (see Stålhös, 1981; Ehlers et al., 1993; Koistinen et al., 1996; Korsman et al., 1999; Nironen, 1999; Väisänen and Hölttä, 1999). We propose a complex history for the Paleoproterozoic tectonic evolution of the shield that led to a collage of several accretionary
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units. We thus reject the concept of a single long (>100 Ma) Svecokarelian/Svecofennian orogeny and propose five orogenies for the period 1.92–1.79 Ga. These orogenies were partly overlapping in age but they have separate tectonic histories. Each of them produced linear belts and mountain chains, followed by exhumation; these can thus be considered orogens sensu stricto. The following Paleoproterozoic orogens are proposed: Lapland–Kola orogen, Lapland–Savo orogen, Fennian orogen, Svecobaltic orogen, and Nordic orogen. The Gothian orogeny (e.g., Gaál and Gorbatschev, 1987) comprises a time span of ~200 Ma, including several collisional events, and we prefer the term Gothian evolution (Åhäll and Larson, 2000) for this period. Short descriptions are given below for each Paleoproterozoic orogen, which then are discussed in more detail in subsequent sections. The Lapland–Kola orogen (modified from Daly et al., 2001) comprises the following dispersed Archean terranes and Paleoproterozoic accreted terranes: Lapland granulite belt, Umba granulite terrane, and Tersk terrane in the Lapland–Kola region (e.g., Daly et al., 2001). The Belomorian mobile belt within the Karelian craton is included in the Lapland–Kola orogen. The Lapland–Savo orogen (defined in this study) comprises the accreted Norrbotten craton, Keitele microcontinent, and allochthonous Paleoproterozoic cover and reworked Archean at the western boundary of the Karelian craton. The Kittilä allochthon is included in the Lapland–Savo orogen. Collision was in an E–W direction. The collisional event between the Bothnia and both Norrbotten and Keitele microcontinents is also included in the Lapland–Savo orogeny. The Fennian orogen (defined in this study) is the result of a major N–S collision at 1.89–1.87 Ga, which accreted the Bergslagen microcontinent and Tavastia island arc to the Lapland–Savo orogen. 500
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The Svecobaltic orogen (defined in this study) comprises areas previously affected by the Fennian orogeny and new crust added to the southwestern part of the shield. The Svecobaltic orogen is the result of a new collision in the south at 1.84–1.8 Ga. It has traditionally been included in the Svecofennian orogen but a definition as a separate orogenic event is warranted by the proposed large-scale extensional collapse predating a renewed compressional stage. The Nordic orogen (defined in this study) cuts across the Lapland–Savo, Fennian, and Svecobaltic orogens in N–S and SW–NE directions. It comprises, apart from the new added crust, the reworked Paleoproterozoic crust in the western Fennoscandian Shield, the Norrbotten craton, and the western part of the Karelian craton in Lapland. Gothian evolution (orogeny) at 1.75–1.55 Ga (Gaál and Gorbatschev, 1987; Åhäll and Larson, 2000), partly Mesoproterozoic in age, affected the westernmost part of southern Sweden (Figure 11.9B) and extended northwest to western Norway. It mainly reworked crust that was formed during the Nordic orogeny.
5. Tectonic model We present a model for the evolution of the Fennoscandian Shield and adjacent areas at 2.06–1.77 Ga with the Karelian craton as the fixed plate. Five orogenies are identified and these partly overlap in time and space. The key issue is the almost simultaneous amalgamation of several microcontinents and island arcs. The orogenies operated at high angle to each other, which had important implications as discussed below.
5.1. Breakup of the Archean craton (or cratons) at 2.06 Ga At the beginning of the Paleoproterozoic, a cratonic stage (2.5–2.1 Ga) with coeval
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multiple rifting events (e.g., Vuollo, 1994; Kohonen, 1995) characterized the Karelian craton. Although the lithostratigraphic correlation of the Paleoproterozic cover rocks across the shield is debatable, it is significant that the craton-type, highly 13C-enriched, sedimentary carbonates deposited at 2.2–2.1 Ga form large mappable units over the entire Karelian craton. The subsequent trangressive shift to deepwater sedimentation at 2.1–2.06 Ga is seen as a sharp drop in δ13C values of marginal sequence carbonates at the western edge of the Karelian craton (Karhu, 1993); this may indicate a continent breakup during this time period (Karhu, 1993; Kohonen, 1995). Lack of subduction-related magmatism between 2.4 and 2.1 Ga in the Karelia, Norrbotten, and Kola cratons and rare 2.5–2.1 Ga zircons in the sedimentary record (Claesson et al., 1993; Sorjonen-Ward et al., 1994; Tuisku and Huhma, 1998; Daly et al., 2001; Lahtinen et al., 2002) also support breakdown of the precursors of the Archean craton at 2.1–2.0 Ga. Continental breakup at 2.06 Ga along the present western margin of the Karelian craton created marginal basins along the edge (Figure 11.10A). A large-scale, long-lived, failed E–W rift in Lapland, now seen as the Peräpohja (PB) and Kuusamo belts (KB; Figure 11.2), was also active during this time period. At 1.95 Ga, a Red Sea-type (Peltonen et al., 1996) marginal basin formed close to the western margin (Figure 11.10B) and acted as a zone of crustal weakness during later tectonic evolution as well.
5.2. Lapland–Kola orogen In the Lapland–Kola area there is evidence for subduction-related arc magmatism in the Tersk terrane at ~1.96 Ga (Daly et al., 2001) and in the Inari area at 1.94–1.93 Ga (Barling et al., 1997). The main difference between these two areas is that there is a substantial amount of Archean crust attached to the arc crust in the Inari area. At 1.93 Ga (Figure 11.10B) subducCHAPTER
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tion was active toward the southwest under the Lapland–Kola area, and a back-arc basin formed, now expressed as the Lapland granulite belt and the Umba granulite terrane. The granulite belt is dominantly ensialic, whereas the Umba terrane is associated with arc-type crust. Back-arc rifting also extended the crust in the Belomorian mobile belt. Shortening of these back-arc basins was caused by the collision of the Kola craton with the Karelian craton. The initiation of shortening is bracketed between 1.93 and 1.91 Ga by the age of the abundant enderbites, charnockites, and granites (Meriläinen, 1976; Glebovitsky et al., 2001). These intrusions are considered as syn- to late-tectonic relative to the early deformation and metamorphism (e.g., Daly et al., 2001). The accretion of the Tersk terrane also took place before ~1.91 Ga (Daly et al., 2001). The Lapland granulite belt is a NE-dipping thrust wedge (conductors in Figure 11.7; Marker, 1990), whereas in the Umba granulite terrane a major episode of thrusting is expressed in the gently E-dipping shearing and gently SE-plunging lineation (Glebovitsky et al., 2001). Still farther east in the Tersk terrane, the major structures dip southward (cf. Daly et al., 2001). We suggest that the Lapland granulite belt represents closure of an ensialic back-arc basin, and also propose that the SW-dipping structures within the Imandra–Varzuga–Pechenga belt (conductors in Figure 11.7; Gaál et al., 1989) image the suture between the Inari arc complex and the Kola craton (cross-section a-a’ in Figure 11.11; e.g., Berthelsen and Marker, 1986a). The structures related to the same collision are also seen in the Tersk–Umba arc/back-arc pair, and in the Belomorian mobile belt (4B in Figure 11.7; Mints et al., 2001). The peak of granulite-facies metamorphism (M1) in the Lapland granulite belt and the Umba granulite terrane occurred, at least partly, before the main shearing (Barbey and Raith, 1990; Glebovitsky et al., 2001; Daly et
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A 2.06 Ga
LG
500 km
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B 1.93 Ga Norrbotten mc
n lia re cr
IA +T UG T T
B+
n
a to
Knaften arc
Bothnian mc
C 1.91 Ga
a’
D 1.90 Ga
KA
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F 1.88 Ga
E 1.89 Ga 500 km
500 km
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b c
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al., 2001). The pre-1.91 Ga age of M1 is constrained by 1.91 Ga monazites analyzed from the Lapland granulite belt (Meriläinen, 1976) and a 1.91 Ga charnockite that postdates peak metamorphism in the Umba granulite belt (Glebovitsky et al., 2001). We interpret this early metamorphism to have been related to the shortening and thickening of a pre-heated back-arc basin. Both advective and conductive heating took place in a rapidly subsiding back-arc basin. During the collision stage, thick and hot sedimentary sequences were stacked, leading to a high-temperature and medium-pressure metamorphism. The 1.94–1.87 Ga titanites in the Belomorian (Archean) mobile belt (Bibikova et al., 2001) and the strong ~1.9 Ga thermal overprinting of the Archean gneisses in the Inari area (Hannu Huhma, pers. comm., 2001)
imply that the collision between the Kola and Karelian cratons had attained a closure stage at 1.91 Ga (Figure 11.10C) and a terminal stage at 1.90 Ga (Figures 11.10C and 11.11). The main deformation in the Umba granulite terrane took place first at constant and then at decreasing pressure (Glebovitsky et al., 2001). Metamorphic rocks south and southwest of the Lapland granulite belt (e.g., Krill, 1985) and the Umba granulite terrane (see Daly et al., 2001) show an inverted metamorphic gradient. We explain these metamorphic features by early subhorizontal stacking of a hot metamorphosed package over cold rocks. Continuous shortening produced more upright folding and thrusting. Orogenic collapse of the Lapland–Kola orogen is proposed to have occurred at 1.88–1.87 Ga (Figures 11.10C and 11.12A).
Mostly Archean crust
Active plate boundaries
Archean and Proterozoic crust 2.1–2.0 Ga crust
Active terrane boundaries Direction of relative plate motion
2.0–1.95 Ga crust
Direction of compression Direction of extension
Fig. 11.10. Tectonic model for the Fennoscandian Shield, Part I; microcontinent (mc) accretion stage at 1.93–1.88 Ga. Cross-sections a-a’, b-b’, and c-c’ as in Figure 11.11, abbreviations as in Figure 11.2. (A) Breakdown of the Archean continent at 2.06 Ga. (B) Subduction and back-arc rifting in the Lapland–Kola area, westward subduction under the Keitele mc (Savo belt) and Norrbotten mc (Kittilä), and subduction toward northeast under the Norrbotten mc (Knaften). (C) Peak of the Lapland–Kola and Lapland–Savo orogenies. Initial stage of collision of the Bothnia mc with the Norrbotten and Keitele mc. (D) Docking of the Bothnia mc with the Norrbotten and Keitele mc. Differences in relative plate motions result in a transform fault between the Keitele and Bothnia mc. Subduction reversal and the onset of subduction toward north under the Keitele mc. (E) Subduction switchover and onset of subduction toward north under the Bothnia mc. Locking of subduction under the Keitele mc. Ocean is consumed by subduction toward south under the combined Tavastia island arc and Bergslagen mc. (F) Peak of the Fennian orogeny, a strong compressional stage. The Keitele–Bergslagen collision results in strong shortening within the collision zone, overthrusting at the western margin of the Karelian craton, basin inversion in Lapland, and reactivation of the Lapland–Savo suture zone. Subduction under the Bothnia mc is still active and back-arc rifting occurs in the Skellefte district. Eastward subduction under the Norrbotten mc commences and is followed by extension. Local extensional domains in the Kola and Belomorian areas develop, see Fugure 11.2 for details. CHAPTER
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This resulted in thinning of the crust under the Lapland granulite belt, but otherwise its effects are largely unknown. A further indication of the collapse is that the compressional major D1-shearing was followed by D2-extension and decompression in the Umba granulite terrane (Glebovitsky et al., 2001). A similar two-stage model has also been suggested for the Lapland granulite belt (Daly et al., 2001), where exhumation (decompression) may have taken place at 1.90–1.87 Ga (Tuisku and Huhma, 1998; Daly et al., 2001). We suggest an orogenic collapse at 1.88–1.87 Ga as the terminal stage of the Lapland–Kola orogeny. The final emplacement of the Lapland granulite belt probably occurred later, around 1.84 Ga (Figure 11.12C).
5.3. Lapland–Savo orogen Subduction-related magmatism was still active at 1.93–1.92 Ga in the island arc (Savo belt) attached to the Keitele microcontinent (Figure 11.10B). Magmatism may also have taken place on the eastern and southwestern sides of the Norrbotten microcontinent in the Kittilä and Knaften arcs, respectively. The Karelian craton had a curved leading edge when it collided with the Norrbotten and Keitele microcontinents (Figure 11.7C). Collision started in the north with emplacement of the Kittilä allochthon at ~1.92 Ga (Chapter 4) and docking of the Norrbotten and Karelian continents (cross-section a-a’ in Figure 11.11). Based on the age data from preorogenic and synkinematic intrusions (Chapter 10), the onset of the continent–continent collision between the Karelian craton and Keitele micro-continent is bracketed between 1.92 Ga and 1.89 Ga and probably occurred at ~1.91 Ga (Figure 11.10C). The characteristic features of the initial stage of the collision are vast amounts of homogenous graywackes with a maximum deposition age of 1.94–1.92 Ga (Claesson et al., 1993). Slightly earlier onset of collision in the north could have caused vo504
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luminous deposition of turbiditic graywackes south of the rising orogen (Kohonen, 1995). Both graywackes and ocean floor (ophiolites) took part in early, thin-skinned thrusting (Koistinen, 1981). The Bothnia microcontinent approached the Norrbotten and Keitele microcontinents from the west via E- to NE-directed subduction, and collision took place at 1.90 Ga (Figure 11.10B through D). The Skellefte and Norrbotten areas comprise vast amounts of 1.90–1.87 Ga subduction-related volcanic and plutonic rocks (Allen et al., 1996b; Bergman et al., 2001). It is proposed that the early ~1.90 Ga magmatism formed on the Knaften arc and Norrbotten microcontinent was related to a waning phase of subduction and initiation of collision between the Bothnia microcontinent and the Norrbotten craton (Figure 11.10D). Early thrusting (possibly E- or NE-vergent) at the western margin of the Keitele microcontinent (Markus Vaarma, pers. comm., 2002) and the early horizontal structures south of the Skellefte area (Rutland et al., 2001a) are associated with this stage. The western boundary of the Keitele microcontinent became a transform fault during the oblique collision of Bothnia and Keitele (Figure 11.10D). Docking of the rigid Bothnia, Keitele, and Norrbotten microcontinents prevented substantial shortening in the triangle between these microcontinents, and a well-preserved paleosubduction zone (BABEL Working Group, 1990; cover page) with an associated accretionary wedge is seen as NE-dipping reflectors in Figures 11.8 and 11.11 (cross-section b-b’ in the latter). Apart from the imbrication of the accretionary wedge, the configuration of the arc system seems to have retained much of its original shape (Figure 11.8). The frozen arc geometry is preserved in the middle and lower crust, whereas the uppermost 15–25 km of the BABEL 2-4 profiles represent an allochthonous unit of migmatites and plutonic rocks (Korja and Heikkinen, 2000). Thickening of the allochthonous unit and its final emplacement
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a-a’ 1.90 Ga Norrbotten craton
Inari arc
Kola craton
50 km
Vertical exaggeration 2:1 500 km
b-b’ 1.89 Ga Bothnian mc
Knaften ac
c-c’ 1.89 Ga Bergslagen mc
Keitele mc
Karelian craton
Mostly Archean crust
Mafic magmatism
Archean and Proterozoic crust
Felsic to intermediate magmatism
2.1–2.0 Ga crust
Sedimentary rocks
2.0–1.95 Ga crust
Oceanic crust
Island arc
Remnants of oceanic crust
Fig. 11.11. Crustal cross-sections pertaining to profiles a-a’, b-b’, and c-c’ in Figure 11.10.
over the now hidden island arc were related to basin inversion during the Nordic orogeny (see Section 5.9).
5.4. Subduction reversal and switch-over: prelude to the Fennian orogeny at 1.90 Ga Accretion of the Keitele microcontinent to the Karelian craton resulted in a subduction reversal. The subduction toward the north under the mature crust of the Keitele microcontinent took place between 1.90 Ga and 1.89 Ga (Figure 11.10D). At the southern edge of the Keitele microcontinent, the subduction CHAPTER
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reversal caused extension, leading to subsidence of the continental margin. Uplift of the newly formed Lapland–Savo orogen caused rapid erosion and accumulation of eroded material in the subsiding southern margin that subsequently developed into a subductionrelated foredeep (accretionary wedge). The short-lived 1.90–1.89 Ga subduction towards the north is recorded as active continental margin volcanism and associated plutonism in the Tampere belt. Docking of the Bothnia and Norrbotten microcontinents led to subduction switch-over. A new subduction zone under either oceanic crust or the Bothnia microcontinent developed
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B 1.85 Ga
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D 1.82 Ga Laurentia
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Amazonia
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E 1.80 Ga
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at 1.89–1.87 Ga (Figure 11.10E and cross-section b-b’ in Figure 11.11). The initiation of a subduction system directed to the north and and northwest is proposed, based on combined deep seismic reflection (Figure 11.8) and refraction as well as gravity data (see discussion in Korja et al., 2001).
5.5. Fennian orogeny: a north–south accretion stage at 1.89–1.87 Ga The timing of tectonic events is different across the Gulf of Bothnia from Finland to Sweden. This apparent discrepancy is solved in our model by a major transform fault system separating the Keitele and Bothnia microcontinents. To illustrate the differences, schematic crustal sections are shown in cross-sections
b-b’ and c-c’ in Figure 11.11. At 1.90–1.88 Ga, an oceanic plate was subducted toward the south, north of the combined Tavastia island arc and Bergslagen microcontinent (Figure 11.10C). Island arc type magmatism characterizes the Häme belt, where as back-arc magmatism is typical the Bergslagen area. The stratigraphic sequences begin with sedimentary and felsic volcanic (1.90 Ga) rocks, followed by arkosic and other sedimentary rocks interfingered with voluminous ~1.89 Ga felsic volcanic rocks in the Bergslagen area (Allen et al., 1996a; Kumpulainen et al., 1996). The simultaneous deposition of basement-derived quartz-rich sandstones and mature pelites with volcanic rocks (Allen et al., 1996a; Lahtinen et al., 2002) indicates that weathered ‘basement’
Mostly Archean crust
Direction of relative plate motion
Mostly Paleoproterozoic crust
Direction of compression
Active plate boundaries
Direction of extension
Active terrane boundaries Fig. 11.12. Tectonic model for the Fennoscandian Shield, Part II; continent–continent collision stage at 1.87–1.79 Ga. The Fennoscandian continental plate, formed in the accretionary stage (cf. Figure 11.10), is divided into Archean and Paleoproterozoic parts. Cross-sections d-d’, e-e’, and f-f’ as in Figure 11.13, abbreviations as in Figure 11.2. (A) Attempted collapse of the Fennian orogen. At the western margin, the subduction zone migrates southward (dashed lines denote individual microcontinents depicted in Figure 11.10F). (B) Subduction toward southeast and northeast and NE begins at the southern margin, and large-scale extension occurs in the hinterland. (C) The Sarmatian crustal segment collides with the Fennoscandian crustal segment on the southeastern margin. This initiates the Svecobaltic orogeny expressed as basin inversion and thrusting. Subduction in the west and southwest is still active. Docking of Laurentia to Fennoscandia in the northeast leads to the emplacement of the Lapland granulite belt and reactivation of the Belomorian mobile zone. (D) Peak of the Svecobaltic orogeny and onset of the Nordic orogeny. Oblique collision of Fennoscandia with Sarmatia results in the migration of a transform fault into the continent. This develops into a crustalscale shear zone that divides the Svecobaltic orogen into two different compressional regimes. A retreating subduction zone is active in the southwest and a transpressional regime prevails in the southeast. The Nordic orogeny starts with collision of Amazonia and Fennoscandia in the northwest and crustal-scale thrusting takes place in the hinterland. (E) Amalgamation of Laurentia, Fennoscandia, Amazonia, Sarmatia, and an unknown continent in the southwest comes to its end at 1.81–1.79 Ga, and a Paleoproterozoic supercontinent is formed. (F) Orogenic collapse and lithospheric delamination stabilizes the Fennoscandian Shield between 1.79 and 1.77 Ga. CHAPTER
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was also exposed close to the active volcanic basins. Thus we infer that the Bergslagen microcontinent was partly characterized by a cratonic environment and that a passive margin existed in the southern part of the Bergslagen microcontinent. At 1.89 Ga, migration of the subduction zone towards the thick lithospheric keel of the Keitele microcontinent caused steepening and locking of the subducting slab. Subduction under the Keitele microcontinent ceased (Figures 11.10E and 11.11) and contraction caused deformation and migmatization (Kilpeläinen, 1998; Korsman et al., 1999) in the accretionary wedge. After a short extensional stage at the southern margin (Kilpeläinen, 1998), the Keitele microcontinent, now as a lower plate, collided with the Tavastia island arc. Deformation due to N–S compression continued until 1.87 Ga (Kilpeläinen, 1998; Nironen, 1999). The N–S collision also had an effect farther inland in the north, where a foreland fold and thrust belt developed (e.g., Koistinen, 1981; Ward et al., 1989; Lehtonen et al., 1998). NE–NNE thrusting at ~1.89 Ga, followed by tight folding at 1885–1880 Ma (Koistinen, 1981; Korsman et al., 1999), affected the rocks of the Karelia–Keitele boundary zone. Thrusting also involved the basement and cover of the Karelian craton (Kohonen, 1995). As the strain was partioned, strong shortening and thickening occurred in the former E–W rift basins in the northern part of the Karelian craton, but otherwise the effect on the non-stretched, rigid Archean crust and its cover was minimal. For example, the Lapland failed rift in the Karelian craton (PB and KB in Figure 11.2) was closed by basin inversion and the crust was thickened at 1.89–1.88 Ga. Continuous convergence activated the craton margin and strike-slip shear zones developed in the suture and within the craton (Figure 11.7F; e.g., Kärki et al., 1993). Subduction under the Bothnia microcontinent was active at 1.89–1.88 Ga (Figure 508
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11.10E–F and cross-section in Figure 11.11). A low-grade 1870 ± 2 Ma metarhyolite (Lundqvist et al., 1998) from the Bothnian basin shows that magmatism lasted until 1.87 Ga. The 1.88–1.87 Ga volcanism in the Skellefte district (e.g., Billström and Weihed, 1996; Allen et al., 1996b) is thought to represent back-arc volcanism linked with the N-directed subduction. Subduction terminated at about 1.87–1.86 Ga, as indicated by the age of the second major deformation episode south of the Skellefte district (Rutland et al., 2001a,b). It is suggested that another subduction system developed under the Norrbotten microcontinent and the attached arc in the west (Figure 11.10F) – this is to explain the 1.89– 1.87 Ga calc-alkaline and alkaline magmatism north of the Skellefte district in northern Sweden; these represent subduction-related and extensional-type magmatism, respectively (Bergman et al., 2001). The 1.89–1.87 Ga period was the major collisional stage associated with voluminous continental growth in the central Fennoscandian Shield. Molasse-type sediments accumulated in the Keitele and Lapland areas, as did some graywackes in the Bothnian basin (Claesson et al., 1993; Lehtonen et. al., 1998; Lahtinen et al., 2002). The microcontinent accretionary assemblage amalgamated to form a continent that occupies the major part of the present Fennoscandian Shield (Figure 11.10). The accretionary stage ended at ~1.87 Ga and was followed by an Andean- or Himalayan-type evolution (Figure 11.12).
5.6. Attempted orogenic collapse and related magmatism at 1.89–1.87 Ga An attempt of collapse of the Lapland–Savo orogen occurred simultaneously with the collision in the south (Fennian orogeny). The Karelia–Norrbotten/Keitele continent–continent collision during the Lapland–Savo orogeny thickened the crust and maximum shortening
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occurred in areas where the crust was originally thinnest, e.g., in the proposed suture zones (Figure 11.10). Thick-skinned stacking of the lithospheric plates at the Karelia–Keitele boundary, associated with inversion of extensional faults, caused thickening of both crust and subcontinental lithospheric mantle. A N–S compression at the southern margin during the Fennian orogeny inhibited large-scale extension, and later-stage transtensional shearing, following thrusting (see above), dominated at the Karelia–Keitele boundary. Both dry and wet magmatism characterized the Karelia–Keitele boundary zone. Mantle-derived magmas caused melting of arc-affinity mafic rocks (Savo belt), producing calc-alkaline granitoids associated with a slightly later episode of alkaline granitoids (Chapter 10). The locus of mantle-derived magmatism was first on the thickened suture zone and subsequently shifted to the northeast and southwest. In the Karelian craton, this magmatism is seen as an abundance of granitoid rocks and dikes that are 1.87–1.86 Ga in age (Huhma, 1986) and in the decreasing age of the thermal peak towards the northeast (Pajunen and Poutiainen, 1999). The attempted orogenic collapse in the southwest resulted in a domain of extension and associated dry magmatism at 1.87 Ga (Figure 11.12A; Nironen et al., 2000b). Alkalic-calcic to calc-alkaline 1.89–1.86 Ga magmatism is found along the Karelian– Norrbotten boundary zone (Väänänen, 1998; Bergman et al., 2001). The calc-alkaline and alkaline perthite monzonite suites west of the boundary zone are considered to represent a shift from subduction-related to extensional magmatism (Bergman et al., 2001) in a backarc region related to E-directed subduction (Figures 11.10F and 11.12A). Magmatic activity at 1.89–1.86 Ga in the colder plate (Karelian craton) was limited to a with of about 50 km and 100 km at the Karelia–Norrbotten boundary and the Karelia–Keitele boundary, respectively. The CHAPTER
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more extensive occurrence of magmatism at the Karelia–Keitele boundary correlates with the existence of abnormally thick crust (Figure 11.4A), especially thick high-velocity lower crust (Korsman et al., 1999). Only a thin high-velocity lower crustal layer is found at the Karelia–Norrbotten boundary, where the overall crustal thickness increases from 41 km in the north to a maximum of 48 km along the Fennolora profile (Guggisberg, 1986; Figures 11.4 and 11.5).
5.7. The end of the Fennian orogeny at 1.87–1.85 Ga: orogenic collapse North–south trending compressional structures have been observed in the southern part of the Bergslagen area (Stålhös, 1976), in the Häme belt (<1.85 Ga; Nironen, 1999), at the western edge of the Keitele area (<1.87 Ga; Mäkitie, 1999), and in the Finnish Lapland (<1.88 Ga; Lehtonen et al., 1998). It is not clear whether these structures are related to single or multiple events, but they imply E–W compression after the Fennian orogeny. In the model these structures are attributed to convergence at the western margin of the Fennoscandian Shield between 1.87 Ga and 1.84 Ga (Figure 11.9 A through C). Subduction in the west migrated southward along the western edge of the amalgamated continent at 1.87 Ga (see Figures 11.10F and 11.12A). Subduction reversal, following the Fennian orogeny, took place at the southern margin (Figure 11.12A, B). The change in plate motions was associated with a large-scale extensional stage at 1.86–1.84 Ga in the hinterland, causing the formation of extensional basins (Figure 11.12B and crosssection d-d’ in Figure 11.13). The effects of the extensional stage varied depending on the nature of earlier structures and the temperature of the crust. Areas involved in the Lapland–Savo orogeny were already rather stable. Thinning caused by extension was concentrated in the Berg-
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d-d’ 1.86–1.85 Ga Bergslagen mc
e-e’ 1.82 Ga Amazonia
Knaften ac
f-f’ 1.81–1.80 Ga Unknown mc
Bergslagen mc HSZ
Karelian craton
Keitele mc
Karelian craton
Keitele mc
Bothnia mc
50 km
Knaften ac
Vertical exaggeration 2:1 500 km
Mostly Archean crust
Mafic magmatism
Archean and Proterozoic crust
Felsic to intermediate magmatism
2.1–2.0 Ga crust
Sedimentary rocks
2.0–1.95 Ga crust
Oceanic crust
Island arc
Remnants of oceanic crust
Umeå allochton
Extension
Magmatic underplate Fig. 11.13. Crustal cross-sections pertaining to profiles d-d’, e-e’, and f-f’ in Figure 11.12.
slagen–Häme and Bothnian areas (Figure 11.9B) as these were still hot after the Fennian orogeny. Thus the extension was related to a large-scale orogenic collapse of hot crust away from a cooler core. Lithospheric extension, leading to astenospheric upwelling and thinning of the lithosphere, would rapidly increase the temperature in the lower and middle crust and thus provide a heat source for migmatization and granite formation. Metavolcanic rocks with ages of 510
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1871 ± 7 Ma (Dobbe et al., 1995) and 1867 ± 9 Ma (Welin, 1987) as well as mature shallow-water quartzites with a depositional age of 1.87–1.86 Ga (Claesson et al., 1993; Lahtinen et al., 2002) are related to this stage. The molasse-type sediments in northern Finland and Sweden (Bergman et al., 2001; Chapter 4) could be correlated with this stage. Regional extension and crustal thinning in southern Finland is proposed based on preliminary data (Rod Holcombe, Nick Oliver, and
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Matti Pajunen, pers. comm., 2001). Indirect evidence of the extension is the well-preserved stacking structure of migmatite complexes in the BABEL profiles 1, 6, and C (Figure 11.8; Korja and Heikkinen, 2000). The age of this event is poorly constrained, but <1.87 Ga is proposed based on syntectonic rocks in southern Finland (Väisänen and Hölttä, 1999).
5.8. Svecobaltic orogeny: Andean-type active margin and continent–continent collision at 1.84–1.80 Ga The Karelia–Keitele and Keitele–Tavastia bounda ries were stabilized at 1.85–1.81 Ga and, consequently, magmatism of this age is lacking from these areas. Neither are 1.85–1.82 Ga magmatic ages common at the Karelia–Norrbotten boundary, where the situation is more complex due to strong magmatic and thermal overprinting at 1.82–1.77 Ga (Bergman et al., 2001). Collisional granites and migmatites were formed in the Häme belt and the Bergslagen area at 1.83–1.80 Ga (e.g., Welin, 1992; Korsman et al., 1999). The available data imply that Andean-type magmatism was abundant in the southwestern part of the Fennoscandian Shield at 1.85–1.79 Ga (see below). Thus simultaneous subduction toward north–northwest in the southwest and collision toward the northwest in the southeast (Figure 11.12C through E) took place during the Svecobaltic orogeny. In the southwestern part of the Fennoscandian Shield, volcanism related to the Andean-type magmatic stage includes 1836 ± 9 Ma metatuffites in the center of the Bergslagen area (Dobbe et al., 1995) and plutonic and volcanic rocks in the Oskarshamn–Jönköping belt (OJB in Figure 11.2). The latter is situated south of the Bergslagen area and is characterized by supracrustal rocks intruded by 1.83 Ga calc-alkaline I-type syntectonic granitoids (Mansfeld, 1996). A possible continuation of the belt is found in western Lithuania (Skridlaite and Motuza, 2001). Along the CHAPTER
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southwestern margin of the Bergslagen area, a NW-trending belt of 1.85–1.84 Ga mafic rocks and granitoids has been found; these are interpreted to have formed in a continental arc setting (Figure 11.9; Andersson and Wikström, 2001; Åhäll and Larson, 2000). Mafic rocks of more continental affinity are also found (Andersson, 1997). We propose that an Andean-type active margin was retreating southwestward and caused cyclic periods of subduction-type and marginal basin-type magmatism (Figure 11.12C). Regional magnetic (Figure 11.3) and gravity patterns that trend northwest can be followed from central and southern Sweden to Lithuania, Latvia, and Estonia. The anomalies are truncated by the NNE-trending discontinuities parallel to the western edge of Sarmatia (Skridlaite and Motuza, 2001; Mansfeld, 2001). The northwesterly trend is parallel to the proposed fossil subduction zone (Figure 11.5) that is seen as a dipping mantle reflector and layer boundary in reflection seismic profiles in the Baltic Sea area (Abramovitz et al., 1997; Ostrovsky, 1998). In the east, however, the Fennoscandia–Sarmatia suture has been proposed to dip southeast to south–southeast (Stephenson et al., 1996). The age of volcanism and metamorphism decreases toward the northwest from the Fennoscandia–Sarmatia junction (Claesson et al., 2001). The westernmost NNE-trending belt records metamorphism that began before 1.8 Ga (Skridlaite and Motuza, 2001). The inferred subduction direction, southeast to south–southeast, at the Fennoscandia–Sarmatia junction differs from the NE-directed subduction inferred for the southwestern part of the Fennoscandian Shield from seismic reflection, lithologic, and geochronologic data (Abramovitz et al., 1997; Balling, 2000; Åhäll and Larson, 2000; Beunk and Page, 2001; Skridlaite and Motuza, 2001). These two domains with different inferred subduction directions and structural grain are separated by a 30- to 50-km-wide zone (Skridlaite and
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Motuza, 2001) in which a considerable change in crustal thickness, from 42–44 km to 50 km, is observed (EUROBRIDGE Seismic Working Group, 2001). These features probably indicate two subduction zones with opposite polarities and separated by a transform fault (Figure 11.12B). In the early stage of the Svecobaltic orogeny (1.84–1.82 Ga), compressional structures in southern Finland and central Sweden were related to the Fennoscandia–Sarmatia collision (Figure 11.12C). Later the area was divided into two segments by a crustal-scale shear zone (Figure 11.12D), now depicted as a magnetic lineament (Figure 11.6). Oblique collision of Fennoscandia and Sarmatia in a NNW direction produced transpressional structures in southern Finland (Ehlers et al., 1993; Korsman et al., 1999; Väisänen et al., 1999). The Fennoscandia– Sarmatia collision started with a N-directed thrusting followed by thrusting toward the northwest and folding during transpression. The continuous shortening of thickened crust led to the development of E–W shear zones and reactivation of thrust surfaces striking south-southwest. Due to both SW and SE contraction, the transform fault developed into a crustal-scale shear system striking west–northwest (Figure 11.12D). Dextral displacement (Högdahl and Sjöström, 2001) transported the Bergslagen microcontinent to its present position at ~1.82 Ga (Figure 11.12D). The crustal-scale shear system is responsible for the current NW-verging curved shape of the suprcacrustal belts in southwestern Finland (Ehlers et al., 1993). Compression east of the crustal-scale shear system seems to have been active until ~1.80 Ga. The continent–continent collision (Figure 11.12C, D) led to inversion and imbrication of the 1.86–1.84 Ga extensional basins (Figure 11.12B). The basin inversion caused thickening of the hot and extended crust and resulted in migmatization and granite magmatism. The migmatization and melt-induced deformation 512
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during and after thrusting and thickening was continuous between 1.84 and 1.80 Ga. Stålhös (1976) regarded the E-trending fold structures younger than the N-trending fold structures in the Bergslagen area. He interpreted the former as a result of a N–S compression that was simultaneous with the peak of regional metamorphism. Northwesterly thrusting and folding in southern Finland (see Ehlers et al., 1993; Korsman et al., 1999; Väisänen and Hölttä, 1999) was also associated with migmatization and the emplacement of 1.84–1.81 Ga granites (Huhma, 1986; Patchett and Kouvo, 1986; Suominen, 1991). Local domains of granulite-facies metamorphism and migmatization (1.83–1.81 Ga) are found in southwestern Finland (Väisänen et al., 2000). The cyclic nature of events caused repeated compressional peaks and ramping of earlier thrust planes; this led to a complicated stacking structure in southern Finland. The final emplacement of inverted migmatite basins over the Tavastia island arc and farther to the north probably occurred at a late stage, between 1.82 Ga and 1.81 Ga. In the model, the areas west of the crustalscale shear system show alternating compressional and neutral tectonics because of an Andean-type retreating NNE-directed subduction zone at 1.83–1.81 Ga (Figure 11.12C through E). Beunk and Page (2001) have interpreted metamorphism and deformation (1825–1800 Ma) in a back-arc environment north of the Oskarshamn–Jönköping belt (OJB) as having been caused by oblique accretion of a magmatic arc onto the continental margin. According to Gorbatschev (2001), a WNWto E-trending, partly gneissic 1.81–1.77 Ga granitoid belt occurs in the southeast corner of Sweden. The evidence of 1.81–1.77 Ga granitoids implies that a second (continental) arc accreted to the Bergslagen microcontinent from the southwest. This late collision (Figure 11.12E and cross-section f-f ’ in Figure 11.13) is seen as later ‘cold’ transpressive shortening in a N–S direction at 1.80–1.78 Ga
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north–northeast of the Oskarsham–Jönköping belt (Beunk and Page, 2001). It should be noted that, at 1.82–1.80 Ga, all the shortening southwest of the crustal-scale shear system (Figure 11.12D) resulted from the NE-directed collisional events. During the Svecobaltic orogeny, crustal shortening and high-temperature shearing seem to have been concentrated in the areas extended at 1.86–1.84 Ga (see above). A notable feature is folding in Lapland with E-striking axial planes. This event deformed the post-1.88 Ga molasse-type sediments deposited discordantly on top of already metamorphosed and deformed rocks (Lehtonen et al., 1998). Late (post-1.87 Ga) thrusting of the Lapland granulite belt overprinted the NNE-trending folds (Korja et al., 1996; Nironen and Mänttäri, 2003). It is proposed that the final thrusting and emplacement of the Lapland granulite belt atop the Karelian craton took place between 1.84 Ga and 1.81 Ga simultaneously with the Svecobaltic orogeny. The emplacement resulted from the final docking of Fennoscandia to Laurentia (Figure 11.12C). The amalgation reactivated the Belomorides (Archean), where titanite-rutile thermochronometry shows metamorphic ages of 1.87–1.82 Ga in the central parts of the region (Bibikova et al., 2001).
5.9. The Nordic orogeny: continent– continent collision at 1.82–1.79 Ga The core of the presently exposed part of the Nordic orogen is the ~1400-km-long Transs candinavian igneous belt (TIB; Patchett et al., 1987; Gorbatschev and Bogdanova, 1993), which is a collective term for a linear N-trending batholithic belt partly covered by Caledonian rocks. The TIB 1 (1.81–1.77 Ga) magmatism (Larson and Berglund, 1992) represents magmatism in a convergent margin environment (e.g., Wilson, 1980; Gorbatschev and Bogdanova, 1993; Andersson, 1997; Åhäll and Larson, 2000). Andersson (1991) and GorCHAPTER
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batschev and Bogdanova (1993) also included the Revsund granitoids in central Sweden in the TIB1. We favor a model of two different subduction directions followed by two continent–continent collisions (Figure 11.12B through E). The NW-trending TIB1 belt in southern Sweden is related to the NE-directed subduction of the Svecobaltic orogeny (see above) and only the N–S belt in the north is included in the Nordic orogen. The TIB1 rocks are related to the waning phase of subduction and subsequent orogenic collapse. A NW–SE to E–W collision is proposed at 1.82–1.80 Ga (Figure 11.12D and crosssection e-e’ in Figure 11.13). The proposed continent–continent collision, possibly Fennoscandia–Amazonia (see Chapter 15), mainly affected the northern and central parts of Fennoscandia. Allen et al. (1996b) described a second major deformation phase in the Skellefte district (SD) causing N-trending, open to tight folds. Late, ~1.80 Ga, N–S compressional deformation zones are found in northern Sweden (Bergman Weihed, 1997; Bergman et al., 2001). Granites and high-temperature, low- to intermediate-pressure metamorphism were associated with the deformation zones. Deformation zones (D5 of Lehtonen et al., 1998), probably coeval, trending approximately north-south, are found in the western part of Finnish Lapland, close to the proposed suture zone. Migmatization and related thermal metamorphism at decreasing pressure have also been documented in the Central Lapland granitoid complex (CLGC; Perttunen et al., 1996). Younger metamorphism and associated S-type granites at 1.82–1.80 Ga (Claesson and Lundqvist, 1995; Billström and Weihed, 1996) are correlated with this stage. To explain the different characteristics of the upper and lower crust in the BABEL profiles 2 (not shown) and 3&4 (Figure 11.8), a basin inversion followed by large-scale, thickskin thrusting is proposed (Figure 11.12B). A migmatite complex, the Umeå allochthon (Figure 11.9B and cross-sections e-e’ and f-f ’
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in Figure 11.13), comprising the ‘pre-1.9 Ga sequence’ (Rutland et al., 2001b) and overlying younger rocks in the Bothnian basin, was thrust toward the southeast at 1.83–1.81 Ga (Figure 11.12D,E and cross-section e-e’ in Figure 11.13). This thrust slice continues to western Finland. The crustal-scale thrusting could also explain the tectonic discordance (Rutland et al., 2001a) between the Skellefte district and the migmatite area south of it. The bottom of the subhorizontal conductive layer (Rasmussen et al., 1987; Figure 11.6C) could also be interpreted as a basal thrust plane. The thickening of pre-heated crust led to local migmatization and formation of minimum-melt Härnö-type granites from non-migmatized ‘younger’ sediments. The I- to A-type 1.81–1.79 Ga TIB1 and Revsund magmatism (see below), related to the waning phase of subduction and subsequent extension, also added heat to the system.
5.10. End of the Nordic orogeny and orogenic collapse at 1.79–1.77 Ga The TIB1 granitoids were intruded during extension in either a convergent margin or intracratonic setting at 1.80–1.78 Ga (e.g., Andersson, 1997; Åhäll and Larson, 2000). The ‘postorogenic’ 1.80–1.78 Ga Revsund granitoids are a rather homogenous group of predominantly A- to I-type granitoids and form a linear belt of batholiths in the central part of the Nordic orogen (e.g., Claesson and Lundqvist, 1995; Andersson, 1997). Farther to the north, two types of 1.80–1.78 Ga granitoids are found. The granite–syenitoid–gabbroid association (Bergman et al., 2001; Edefors-type by Öhlander and Skiöld, 1994) has a wide compositional range and is characterized by high alkali and Zr contents, whereas the granite–pegmatite association (Bergman et al., 2001; Lina-type by Öhlander and Skiöld, 1994) shows a limited compositional variation and is dominated by minimum-melt granites (Öhlander and Skiöld, 1994). 514
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1.80–1.77 Ga granites are common also in Finnish Lapland, especially in the Central Lapland granitoid complex (e.g., Haapala et al., 1987; Korsman et al., 1997; Chapter 10). Generally, the Central Lapland granitoid complex, composed of granites and migmatites, is poorly documented. New mapping results (Jukka Väänänen, pers. comm., 2001) indicate that gabbros, monzodiorites, and granodiorites are also found in the central part of the complex. A porphyritic granite that cuts migmatites is dated at 1770 ± 8 Ma (Lauerma, 1982) and another granite at 1843 ± 23 Ma Ga (Huhma, 1986), but no age data are yet available for the other types of granitoid intrusions in the Central Lapland granitoid complex. The I-type character and Nd isotope composition of the granites indicate Archean crust as the main source (Huhma, 1986; Haapala et al., 1987), but the spatial relationships of migmatites and granites suggest also local derivation from metasediments. The close relationship between granites and migmatites indicates a ~1.80 Ga age for the latest stage of metamorphism and migmatization, possibly related to decompression (Perttunen et al., 1996). We suggest that the onset of crustal melting and 1.80–1.77 Ga granite formation is related to extension (Figure 11.12F), but that a mantle-derived heat component is also needed to account for melting of ‘dry’ Archean lower crust. Bergman et al. (2001) proposed that the granite–syenitoid–gabbroid association rocks may be more widespread than previously known and that they may form a large subsurface mafic intrusion. A large gravity anomaly within the Central Lapland granite complex (Ruotoistenmäki et al., 1997) and distribution of mafic rocks in the Central Lapland granitoid complex area could also be interpreted in the same manner. It is thus suggested that mantle-derived heat and extensional magmatism was closely related to the formation of the 1.80–1.77 Ga granites. 1.80–1.75 Ga pegmatites and resetting of
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the U-Pb system at 1.78–1.75 Ga in the Belomorian mobile belt (Bibikova et al., 2001) and at ~1.78–1.76 Ga in Finnish Lapland (Evins, 2001; Corfu and Evins, 2002) mark the end of the thermal activity in the northeastern part of the Fennoscandian Shield. Pegmatites in western and southwestern Finland define a tectonomagmatic episode at 1.80–1.79 Ga (Alviola et al., 2001) and similarly, at 1.80–1.77 Ga, also in Sweden (Romer and Smeds, 1994, 1997). 1.80 Ga zircons from mantle and lower crustal xenoliths in eastern Finland verify large-scale mantle activity at this time (Hölttä et al., 2000; Peltonen and Mänttäri, 2001). In southern Finland, exhumation, either by erosion or extension, during 1.81–1.80 Ga has been proposed (Korsman et al., 1999; Väisänen et al., 2000). The termination of ductile deformation at 1.80–1.78 Ga (Beunk and Page, 2001) and the intrusion of the youngest TIB1 granites at 1.77 Ga (see Åhäll and Larson, 2000) mark the end of tectonomagmatic activity between 1.78 Ga and 1.77 Ga also in southern Sweden. Postcollisional metamorphism at 1.79 Ga (e.g., Claesson et al., 2001) was followed by cooling at 1.77–1.73 Ga (Bogdanova et al., 2001) in the southeastern part of the Fennoscandian crustal segment (Figure 11.1). The intrusion of 1.79–1.73 Ga anorthosite–rapakivi magmas in the Ukrainian Shield (Amelin et al., 1994) also indicates cratonization in the Sarmatian crustal segment at 1.78 Ga. We conclude that, after the Nordic orogeny, a large-scale orogenic collapse took place within the Fennoscandian Shield at 1.79–1.77 Ga (Figure 11.12F). The shield records a neutral to tensional tectonic regime from 1.77 Ga to 1.73 Ga (cf. Åhäll and Larson, 2000) before the onset of the Gothian evolution.
6. Gothian evolution at 1.73–1.55 Ga
margin of the Fennoscandian Shield (Figures 11.1, 11.2, and 11.9B); this region was further reworked during the Sveconorwegian–Grenvillian orogeny (Gaál and Gorbatschev, 1987). Different deformation episodes, the areal extent of which is often unclear and interpretation controversial, have been suggested within the time span of 1.73–1.55 Ga (e.g., Åhäll and Larson, 2000; Andersson and Wikström, 2001; Åhäll and Larson, 2001). We have preferred not to separate this complicated period into different orogenic events and, consequently, we use the combined term Gothian evolution (cf. Åhäll and Larson, 2000). We only review the latest studies and do not propose a tectonic model for the Gothian evolution that started in the late Paleoproterozoic and continued to early Mesoproterozoic; although beyond the actual scope of this study, the Mesoproterozoic part is also included in the discussion. Several growth stages at 1.73 Ga (initial), 1.69–1.65 Ga, 1.62–1.58 Ga, and 1.56–1.55 Ga have been proposed for the Gothian evolution (Åhäll and Larson, 2000) and the last three were followed by rapakivi magmatism at a distal inboard setting (Åhäll et al., 2000). These three magmatic stages and the intervening deformation have been modeled by westward-stepping subduction (see also Berthelsen, 1980) and periodic arc accretion (Åhäll and Larson, 2000; Åhäll et al., 2000). Åhäll et al. (2000) also proposed that recurring subduction controlled episodic mantle melting and subsequent bimodal rapakivi magmatism at faraway positions (see Chapter 12). Docking of Sarmatia and Fennoscandia occurred at ~1.84 Ga and the East European craton was thus formed. The ~1.7 Ga shearing, the 1.67 Ga late reactivation of deformation zones in central Sweden (Högdahl, 2000), and the 1.55–1.45 Ga E–W shear zones in the subsurface of the Baltic countries (Bogdanova et al., 2001) indicate crustal-scale movements within the craton during the Gothian evolution.
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7. Discussion The extent to which the concept of plate tectonics can be applied to the Archean is still unclear, but a broad consensus exists regarding plate tectonic processes operating in the Proterozoic. Our model is based on plate tectonic-type interactions affecting the crustal growth and assembly of the Fennoscandian Shield. The model is, of course, a working hypothesis. Many of the proposed stages are difficult to verify, because they are interpreted from reflection seismic data or are based on isotopic-geochemical modeling of non-exposed crustal components. Some stages can be tested and either approved or rejected, others will probably remain hypothetical or will be reinterpreted in future studies.
7.1. Comparison with modern analogues The distance of the volcanic front from the trench in arc systems varies normally from 80 km to 250 km. This distance, together with the width of the volcanic arc (normally up to 150 km), is negatively correlated with the angle of subduction (e.g., Gill, 1981; Tatsumi and Eggins, 1995). In the model (Figures 11.11 and 11.13), we have used a shallow angle (~30˚) of subduction for illustration purposes, but it should be noted that usually the angle is more than 30˚ (e.g., Gill, 1981). Shallow subduction produces wider volcanic arcs with distinct trench-side and back-arc volcanic chains forming when the underlying subducting slab reaches the depths of ~110 km and ~170 km, respectively. In steeply dipping subduction zones the volcanic chains overlap and the volcanic arc is narrower (Tatsumi and Eggins, 1995). In the model, a complex accretion of microcontinents is suggested to explain the assembly history between 1.92 Ga and 1.87 Ga (Figure 11.10). This situation is similar to that postulated for the Banda Sea (Lee and Lawver, 1995), as proposed earlier by Ward 516
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(1987) and Nironen (1997). The Banda Sea is one example of the occurrence of microcontinents and many coeval subduction zones with subduction directions at high angles to each other or even in opposite directions. The current tectonic setting of the southernmost Andes, the Scotia plate, and the northern part of the Antarctic Peninsula is very complex (Diraison et al., 2000) and it has similarities to some stages in our model. Two intervening subduction zones at high angles are the key to the 1.87–1.81 Ga evolution of the Fennoscandian Shield (Figure 11.12). They account for the observed magmatic and structural grains. Subduction zones at high angles are found around the Sagami Trough and Japan Trench offshore Japan (Tatsumi and Eggins, 1995) as well as within the Molucca Sea and Celebes Sea in Indonesia (see Figure 16 in Lee and Lawver, 1995). The curved Banda arc displays an extremely complicated subduction system that is caused either by two separate lithospheric slabs or by a single, rapidly eastward-retreating slab (see discussion in Milsom, 2001). One critical aspect to be taken into account when the dimensions of the structures in modern tectonic environments are compared with those in the Proterozoic is that shortening during collision has changed the geometry of Proterozoic structures. Although 50% to 80% shortening is possible in mountain building (Le Pichon et al., 1982; Coward, 1994), shortening has rarely been calculated. Kilpeläinen (1998) found that upright folding had caused over 50% shortening of already thickened crust in the Tampere belt (Keitele) accretionary prism and that the total amount of shortening must have been even higher. Maximum shortening and relative thickening is confined to mobile zones that are thinnest before collision. These are plastic than rather rigid crustal indentors that either remain largely intact or are thickened by crustal-scale thrusting. Thrusting involving both the Paleoproterozoic cover and the Archean basement is needed to
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account for the great crustal thickness at the Archean–Proterozoic boundary zone. Local shortening and displacement may also occur during shearing along strike-slip zones, as proposed in Figure 11.12D.
7.2. Comparison with earlier studies and models We agree with Kohonen (1995) in that the final breakup of the Karelian craton took place at ~2.06 Ga. This is based on observations such as basin formation and abundant volcanism at ~2.1 Ga at the present western margin, the ~2.0 Ga ancient oceanic lithosphere within the Kittilä allochthon, and the profound change in carbon isotope ratios at 2.06 Ga. Peltonen et al. (1996) favored the later age of 1.95 Ga for the continental breakup on the basis of the absence of rift sedimentation and the presence of deep-water sediments above the Jormua ophiolite. We suggest that the 1.95 Ga ophiolites are related to a marginal basin rather than breakup. Paleomagnetic data indicate that the Karelian and Kola cratons were probably at least 1000 km apart at 2.4 Ga (Fedotova et al., 1999; Mertanen and Pesonen, 2000). This can be interpreted in two ways: either the Kola and Karelian cratons were separate cratons at 2.4 Ga or they were distant parts of a supercontinent that broke up at 2.1 Ga. In the latter case, the difference in the paleomagnetic poles could be explained by the relative rotation of the Kola nucleus before the accretion of the Kola and Karelian cratons. Our model for the evolution of the Lapland–Kola orogen has many similarities with the model proposed by Berthelsen and Marker (1986a) and Marker (1990). It explains both the opposite vergences found along the Lapland–Kola orogen and the felsic nature of granulites that indicates a continental backarc rather than a trench setting. An opposite subduction polarity with a suture zone located under the Lapland granulite belt (Barbey et al., CHAPTER
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1984; Krill, 1985; Daly et al., 2001) is also possible. If the later hypothesis is correct, the ~1.91 Ga metamorphism occurred at the base of an active continental margin. Sedimentary rocks would then have been transported along a shallowly dipping slab beneath the more than 30-km-thick continental crust to account for the highest calculated pressures in the granulites. The thermal peak would be more difficult to explain, but could be linked with arc-forearc magmatism. Subduction of a hot mid-ocean ridge could also have contributed heat or, actually, produced the thermal peak. Most of the tectonic models for the central Fennoscandian Shield have focused on the nature of the Archean–Proterozoic boundary. These models can be divided into three different concepts: (a) continent–arc/continent collision (favored), (b) back-arc/retro-arc basin related to NE-directed subduction occurring farther in the southwest, and (c) a strike-slip model with all the Proterozic parts considered exotic. Structural evidence of thrusting (Koistinen, 1981; Kohonen, 1995; Tuisku and Laajoki, 1990) favors the first two models, but an earlier separate thrusting event is needed in the strike-slip model. Refraction seismic studies (Luosto and Heikkinen, 2001) indicate a Sto SW-dipping mantle reflector underneath the Archean–Proterozoic boundary zone, favoring a continent–arc/continent collision model. The 1.95 Ga Jormua ophiolite is interpreted as a fragment of marginal basin crust and there is no evidence for a back-arc origin of the Jormua ophiolite (Peltonen et al., 1996). Thus, we consider that there is a substantial amount of data in favor of the continent–arc/continent collision model. We suggest that the two microcontinents Norrbotten and Keitele, both having a juvenile island arc attached, collided with the Karelian craton. In the early stages, the collision was continent–arc collision that subsequently developed into continent–continent collision (Lapland–Savo orogen). Traditionally, the Archean–Proterozoic boundary and the western edge of the Karelian
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craton have been placed along a NW–SE line continuing from Finland to Sweden. Although an Archean–Proterozoic boundary in this direction is observed in Sweden (Öhlander et al., 1993), we propose that in the north, the western margin of the Karelian craton is parallel to the N–S megashear (BBZ) (Berthelsen and Marker, 1986b). This is based on the reinterpretation of BABEL reflection seismic and potential-field data (Figure 11.4B). If our model for the N–S suture is wrong, it would indicate more complicated northeastward transportation of the Kittilä allochthon. Ward et al. (1989) and Sorjonen-Ward et al. (1997) proposed that the N-verging structures and N- to NE-trending folds in central Lapland were formed by interference between deformations caused by the coeval Svecofennian and Lapland–Kola orogenies. This contradicts our model that supposes the early SW-verging thrusting to have been associated with the Lapland–Kola orogeny and the N–S compression and N–S to NE–SW structures having been associated with different collisional stages. The reinterpretation of BABEL reflection seismic (Figure 11.8) and potential-field data combined with the observed lithologic and structural grain have led us to propose a complex history to preserve the paleosubduction zone noticed by the BABEL Working Group (1990). The fossil arc margin with an attached accretionary prism is located in the middle and lower crust, whereas the uppermost 15 km to 25 km are interpreted to image the crustalscale Umeå allochthon (Figures 11.8, 11.9B, and 11.12D, F, and cross-sections e-e’ and f-f’ in Figure 11.13). Critical to the preservation of the convergent margin geometry is docking of rigid microcontinents and development of a transform fault (Figure 11.10D through F). The transform fault further divided the collisional collage into two separate units and thus explains the different tectonic evolution of Finland and Sweden (Figure 11.7D through F). The differences are displayed in the migmatite 518
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belts between TB and HB and within BA and BB (Figure 11.2). The migmatites between TB and HB show subvertical structures, indicating higher rates of shortening compared with the more shallowly dipping structures in the belt between BA and BB in central Sweden (Figure 11.8). We propose that the 1.88–1.87 Ga volcanism in the Skellefte district and south of it resulted from N-directed subduction under the Bothnia microcontinent (Figure 11.10E). The Skellefte district volcanism represents back-arc volcanism or extensional arc volcanism, as proposed by Allen et al. (1996b). Simultaneous eastward subduction under the Norrbotten nucleus at ~1.88 Ga is suggested to explain the N-trending lithologic grain and magmatism in the Norrbotten area (Figures 11.10F and 11.12A). An attempt of orogenic collapse at the Archean–Proterozoic boundary of the Lapland– Savo orogen caused voluminous magmatism at 1.89–1.87 Ga. Full collapse was inhibited by coeval compression from the south (Fennian orogeny). The net result was the stabilization of a thick crust, a characteristic feature of the central Fennoscandian Shield. One of the key stages in the model is the large-scale orogenic collapse of hot crust far from the old core at the end of the Fennian orogeny around 1.86–1.84 Ga. The regional extent of the extensional collapse is still poorly constrained and there are controversial interpretations on the nature of the subhorizontal structures (Väisänen and Hölttä, 1999; Korsman et al., 1999). Extension could explain the heat source needed for the early migmatization, granite magmatism, and continued migmatization and granite formation during crustal thickening of the 1.84–1.80 Ga Svecobaltic orogeny. Korsman et al. (1999) favored magmatic underplating and extensional collapse as a cause for this second high-temperature metamorphism. According to Väisänen et al. (2000; see also Väisänen and Hölttä, 1999), thickening and
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shortening of the upper crust and the extension and mantle upwelling in the lower crust occurred at the same time. They proposed that the post-collisional magmatism (e.g., Eklund et al., 1998) caused voluminous intraplating that increased crustal temperatures and led to the observed metamorphism and crustal anatexis in southern Finland. Postcollisional magmatism is volumetrically insignificant at the present level of erosion; it is evidenced as small plutons that crosscut the surrounding metamorphic rocks. The proposed genesis of these rocks involved small percentage melting of fluid-rich metasomatized parts of the mantle (see discussion in Väisänen et al., 2000). It is unlikely that such a melting could have formed a significant volume of hightemperature magmas at the lower–upper crust interface. We include these postcollisional rocks into the earliest stage of the collapse of the Svecobaltic orogen. On the basis of paleomagnetic data, Elming et al. (2001) proposed that the Ukrainian Shield was not in its present position relative to Fennoscandia between 2.0–1.72 Ga. Bogdanova et al. (2001) proposed that the final docking of Fennoscandia and the combined Sarmatia and Volgo–Uralian crustal segments occurred between 1.71 Ga and 1.66 Ga along zones trending north and north–northeast. A pre-docking common history, where Sarmatia and Bergslagen were juxtaposed and erosion of Sarmatia produced sedimentary material for the Bergslagen microcontinent at 1.9 Ga, has also been proposed (Lahtinen et al., 2002). These interpretations of 1.71–1.66 Ga docking conflict with our model as we propose that the Sarmatia–Fennoscandia collision caused a NNW transpression at 1.84–1.80 Ga. The ~1.7 Ga shear zones that trend north and north–northeast can be interpreted to be linked with the Gothian evolution. Earlier interpretations from the Baltic countries favor a dominant NE to NNE trend for all belts west of Sarmatia, although with a turn to the northwest in Estonia (Gorbatschev CHAPTER
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and Bogdanova, 1993). A more recent interpretation is that NW-trending magnetic and gravity patterns can be followed from Sweden to Lithuania, Latvia, and Estonia (Svecobaltia in Figures 11.3, 11.6, and 11.9B) and that they are truncated by the discontinuities trending north–northeast along the western edge of Sarmatia (Skridlaite and Motuza, 2001). We propose an oblique collision and a transform fault between Sarmatia and southwestern Fennoscandia to explain the different inferred subduction polarities and the different structural grain of these areas. The oceanic transform fault further developed into a continental one (Figure 11.12B through D and cross-section f-f ’ in Figure 11.13) that separated collisional events in the east from a subduction system in the west. In central Sweden, north of the transform fault, the 1.86–1.85 Ga granitoids are largely undeformed, whereas 1.85–1.84 Ga granitoids south of it are penetratively deformed and metamorphosed to amphibolite facies (Högdahl, 2000; Högdahl and Sjöström, 2001). Our interpretation of a large-scale displacement along this shear system implies that the granitoids south of the shear system were strongly deformed during the Svecobaltic orogeny and that they were later juxtaposed with well-preserved ‘northern’ granitoids at 1.82 Ga (Figure 11.12D; Högdahl and Sjöström, 2001). One of the key items in our model is two subduction zones at high angles and operational between 1.86 and 1.81 Ga. E-directed subduction has long been proposed as a source of the TIB1 magmatism (see discussion in Åhäll and Larson, 2000), but the NW-oriented magmatic grain in southern Sweden and in the Baltic countries indicates subduction toward the northeast. A collision starting in the north at ~1.82 Ga (Nordic orogeny) is proposed to explain the 1.82–1.80 Ga metamorphism and crustal anatexis in central and northern Sweden. This stage was also responsible for the basin inversion leading to the transportation of the Umeå
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allocthon towards the east (Figure 11.12D). Extension and mafic magmatism (mantle upwelling) that led to large-scale melting of the Archean lower crust is proposed as the cause for the abundant magmatism and decompression-type migmatization in northern Fennoscandia at 1.80–1.77 Ga. Simultaneously with the Nordic orogeny, another collision along the southwestern margin took place, as an unknown microcontinent (cf. Abramovitz et al., 1997) was added to the Svecobaltic orogen (Figure 11.12E). We suggest that a large-scale orogenic collapse occurred within the Fennoscandian Shield at 1.79–1.77 Ga. The 1.79–1.72 Ga anorthosites and rapakivi granites farther east in the Ukrainian Shield (Amelin et al., 1994) could indicate stabilization of both Fennoscandia and Sarmatia. This stable stage was followed by the Gothian, possibly Andean-type, evolution at the southwestern margin of Fennoscandia between 1.73 Ga and 1.55 Ga (see Åhäll and Larson, 2000). Crustal growth by accretion is generally accepted as the cause for the Paleoproterozoic evolution of the Fennoscandian Shield (e.g., Park et al., 1984; Gaál and Gorbatschev, 1987; Baker et al., 1988; Gorbatschev and Bogdanova, 1993; Lahtinen, 1994; Nironen, 1997; Åhäll and Larson, 2000). Although more complex, our model of the microcontinental accretionary stage at 1.92–1.87 Ga (Figure 11.10) can be regarded as a continuum of these earlier models. Rutland et al. (2001a, b) proposed a different model where a large and long-lived 2.1–1.9 Ga back-arc basin existed in central Fennoscandia. In particular, they considered a back-arc basin model to be more plausible than an accretionary model for generation of thick mature crust by 1.89 Ga. Rutland et al. (2001a, b) also suggested that a synformal structure can explain both the paleosubduction-type pattern in the BABEL data (Figure 11.8) and the conductivity anomalies of the area.
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8. Concluding remarks Gorbatschev and Bogdanova (1993) emphasized that “orogeny in Fennoscandia cannot be described in terms of a series of repetitive cyclic orogenies succeeding each other” and they proposed that the term “Svecofennian” should be used for a semi-continuous crustal formation period. Although we agree about the complexity of the “Svecofennian” evolution, we disagree in that it cannot be separated into succeeding orogenies. We believe that new data, especially from crustal-scale studies, justify an actualistic plate tectonic concept. We propose that there were several collisions from different directions and that these cannot be related to one long-lived Andean-type margin that would produce a lithologic grain only in one orientation. The orogens record both collisional and extensional tectonics as in modern orogens. The time frame of the Paleoproterozoic orogenies was similar to that of the Phanerozoic orogenies. We propose, for the time period between 1.92 Ga and 1.79 Ga, five orogenies, with partly overlapping ages and different structural grains. The associated tectonic evolutionary model varies from less to more tentative, depending on the available information. From a uniformitarian standpoint, our tectonic model is feasible and does not disagree with modern tectonic processes. This does not validate the model but it makes it possible; our primary intention has been to discuss possible geological processes that operated during the Paleoproterozoic time in the Fennoscandian Shield, and to produce a model for further testing.
Acknowledgments We are indebted to geoscientists, past and present, who have been working on the Fennoscandian Shield. We are especially grateful to Per Weihed for doing the tough job of critically reviewing the first manuscript and giving
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valuable comments on the Swedish part of the shield. Moreover, Hugh O’Brien and Yrjö Kähkönen made valuable suggestions that improved the manuscript. The editorial comments improved the clarity and legibility of the text. Juha Korhonen and Pekka Heikkinen kindly provided geophysical images.
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M., 2001. Svecofennian rare-element granitic pegmatites of the Ostrobothnia region, western Finland; their metamorphic environment and time of intrusion. In: H. Mäkitie (Ed.), Svecofennian granitic pegmatites (1.86-1.79 Ga) and quartz monzonite (1.87 Ga), and their metamorphic environment in the Seinäjoki region, western Finland. Geol. Surv. Finland, Spec. Pap. 30, 9–29. Amelin, Yu.V., Heaman, L.M., Skobelev, V.M., 1994. Geochronological constraints on the emplacement history of an anorthositerapakivi granite suite: U–Pb zircon and baddeleyite study of the Korosten Complex, Ukraine. Contrib. Mineral. Petrol. 116, 411–419. Andersson, U.B., 1991. Granitoid episodes and mafic–felsic magma interaction of the Svecofennian of the Fennoscandian Shield, with main emphasis on the ~1.8 Ga plutonics. In: I. Haapala, K.C. Condie (Eds.), Precambrian granitoids, petrogenesis, geochemistry and metallogeny. Special Issue. Precambrian Res. 51, 127–149. Andersson, U.B., 1997. Petrogenesis of some Proterozoic granitoid suites and associated basic rocks in Sweden (geochemistry and isotope geology). Sveriges Geol. Unders., Rapp. Meddel. 91, 1–216. Andersson, U.B., Wikström, A., 2001. Growth-related 1.85–1.55 Ga magmatism in the Baltic Shield; a review addressing the tectonic characteristics of Svecofennian, TIB 1related and Gothian events – A discussion. GFF 123, 55–61. BABEL Working Group, 1990. Evidence for Early Proterozoic plate tectonics from seismic reflection profiles in the Baltic Shield. Nature 348, 34–38. Baker, J.H., Wellingwerf, R.H., Oen, I.S., 1988. Structure, stratigraphy and ore-forming processes in Bergslagen: implications for the development of the Svecofennian of the Baltic Shield. Geologie en Mijnbouw 67, 121–138. Balling, N., 2000. Deep seismic reflection evidence for ancient subduction and collision zones within the continental lithosphere of northwestern Europe. Tectonophysics 329, 269–300.
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South Finland. Contrib. Mineral. Petrol. 92, 1–12. Patchett, P.J., Todt, W., Gorbatschev, R., 1987. Origin of continental crust of 1.9–1.7 Ga age: Nd isotopes in the Svecofennian orogenic terrains of Sweden. In: G. Gaál, R. Gorbatschev (Eds.), Precambrian geology and evolution of the central Baltic Shield. Special Issue. Precambrian Res. 35, 145–160. Peltonen, P., Mänttäri, I., 2001. An ion microprobe U-Pb-Th study of zircon xenocrysts from the Lahtojoki kimberlite pipe, eastern Finland. Bull. Geol. Soc. Finland 73, 47–58. Peltonen, P., Kontinen, A., Huhma, H., 1996. Petrology and geochemistry of metabasalts from the 1.95 Ga Jormua Ophiolite, northeastern Finland, J. Petrol. 37, 1359–1383. Perttunen, V., Hanski, E., Väänänen, J., Eilu, P., Lappalainen, M., 1996. Rovaniemen karttaalueen kallioperä. Summary: Pre-Quaternary rocks of the Rovaniemi map-sheet area. Geological map of Finland 1:100 000. Explanation to the maps of Pre-Quaternary rocks, Sheet 3612 Rovaniemi. Geol. Surv. Finland, Espoo. 1–63. Rämö, O.T., Vaasjoki, M., Mänttäri, I., Elliott, B.A., Nironen, M., 2001. Petrogenesis of the post-kinematic magmatism of the Central Finland Granitoid Complex I : radiogenic isotope constraints and implications for crustal evolution. J. Petrol. 42, 1971–1993. Rankama, K., Welin, E., 1972. Joint meeting of the Precambrian stratigraphy groups of Denmark, Finland, Norway and Sweden in Turku Finland, March 1972. Geol. Newsletter 4, 265–267. Rasmussen, T.M., Roberts, R.G., Pedersen, L.B., 1987. Magnetotellurics along the Fennoscandian Long Range Profile. Geophys. J. Royal Astron. Soc. 89, 799–820. Romer, R.L., Smeds, S.-A., 1994. Implications of U–Pb ages of columbite–tantalites from granitic pegmatites for the Palaeoproterozoic accretion of 1.90–1.85 Ga magmatic arcs to the Baltic Shield. Precambrian Res. 67, 141–158. Romer, R.L., Smeds, S.-A., 1997. U–Pb columbite chronology of post-kinematic Palaeoprote-
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rozoic pegmatites in Sweden. Precambrian Res. 82, 85–99. Ruotoistenmäki, T., 1996. A schematic model of the plate tectonic evolution of Finnish bedrock. Geol. Surv. Finland, Rep. Invest. 133, 1–23. Ruotoistenmäki, T., Aaro, S., Elo, S., Gellein, J., Gustavsson, N., Henkel, H., Hult, K., Kauniskangas, E., Kero, L., Kihle, O., Lehtonen, M., Lerssi, J., Sindre, A., Skilbrei, J., Tervo, T., Thorning, L., 1997. Aeromagnetic anomaly map of northern and central Fennoscandia: total intensity referred to DGRF65. Scale 1 : 2 000 000. Espoo : Trondheim : Uppsala: Geological Survey of Finland : Geological Survey of Norway : Geological Survey of Sweden. Rutland, R.W.R., Kero, L., Nilsson, G., Stølen, L.K., 2001a. Nature of a major tectonic discontinuity in the Svecofennian province of northern Sweden. Precambrian Res. 112, 211–237. Rutland, R.W.R., Skiöld, T., Page, R.W., 2001b. Age of deformation episodes in the Palaeoproterozoic domain of northern Sweden, and evidence for a pre-1.9 Ga crustal layer. Precambrian Res. 112, 239–259. Simonen, A., 1953. Stratigraphy and sedimentation of the Svecofennidic, early Archean supracrustal rocks in southwestern Finland. Bull. Comm. géol. Finlande 160, 1–64. Skridlaite, G., Motuza, G., 2001. Precambrian domains in Lithunia: evidence of terrane tectonics. Tectonophysics 339, 113–133. Sorjonen-Ward, P., Nironen, M., Luukkonen, E., 1997. Greenstone associations in Finland. In: M.J. De Wit, L.D. Ashwal (Eds.), Greenstone belts. Oxford monographs on geology and geophysics 35, 677–698. Sorjonen-Ward, P., Claoué-Long, J., Huhma, H., 1994. Shrimp isotope studies of granulite zircons and their relevance to Early Proterozoic tectonics in Northern Fennoscandia. In: M.A. Lanphere, G.B. Dalrymple, B.D. Turrin (Eds.), Abstracts of the eight International Conference on Geochronology, Cosmochronology, and Isotope Geology. U.S. Geol. Surv., Circular 1107, 299. Stålhös, G., 1976. Aspects of the regional tecton-
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ics of eastern central Sweden. Geol. Fören. Stockholm Förhandl. 98, 146–154. Stålhös, G., 1981. A tectonic model for the Svecokarelian folding in east central Sweden. Geol. Fören. Stockholm Förhandl. 103, 33–46. Stephenson, R.A., Bogdanova, S., Julin, C., 1996. Palaeo zoic rifting in a 2.0 Ga Andean-type magmatic belt in the East European Craton: structural and rheological implications. In: Seventh International Symposium on Deep Seismic Profiling of the Continents, 15–20 September 1996, Asilomar, California, 46. Suominen, V., 1991. The chronostratigraphy of southwestern Finland, with special reference to Postjotnian and Subjotnian diabases. Geol. Surv. Finland, Bull. 356, 1–100. Tatsumi, Y., Eggins, S., 1995. Subduction zone magmatism. Blackwell Science. 1–211. Tuisku, P., Laajoki, K., 1990. Metamorphic and structural evolution of the Early Proterozoic Puolankajärvi Formation, Finland – II. The pressure-temperature-deformation-composition path. J. Metam. Geol. 8, 375–391. Tuisku, P., Huhma, H., 1998. SIMS dating of zircons: metamorphic and igneous events of the Lapland granulite belt are 1.9 Ga old, provenance is Paleoproterozoic and Archaean (2.0–2.9 Ga) and the tectonic juxtaposition about 1.9–1.88 Ga old. In: N. Philippov (Comp.), SVEKALAPKO : EUROPROBE project workshop, Repino, Russia, 26.-29.11.1998: abstracts. St. Petersburg: Ministry of Natural Resources of Russian Federation: State company “Mineral”, 64–65. Väänänen, J., 1998. Kolarin ja Kurtakon karttaalueiden kallioperä. Summary: Pre-Quaternary rocks of the Kolari and Kurtakko map-sheet areas. Geological map of Finland 1:100 000. Explanation to the maps of Pre-Quaternary rocks, sheets 2713 and 2731. 1–87. Vaasjoki, M., Huhma, H., 1999. Lead and neodymium isotopic results from metabasalts of the Haveri Formation, southern Finland: evidence for Palaeoproterozoic enriched mantle. In: Y. Kähkönen, K. Lindqvist
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(Eds.), Studies Related to the Global Geoscience Transects / SVEKA Project in Finland. Bull. Geol. Soc. Finland 71, 143–153. Väisänen, M., Hölttä, P., 1999. Structural and metamorphic evolution of the Turku migmatite complex, southwestern Finland. In: Y. Kähkönen, K. Lindqvist (Eds.), Studies Related to the Global Geoscience Transects / SVEKA Project in Finland Bull. Geol. Soc. Finland 71, 177–218. Väisänen, M., Mänttäri, I., Kriegsman, L.M., Hölttä, P., 2000. Tectonic setting of postcollisional magmatism in the Palaeoproterozoic Svecofennian Orogen, SW Finland. Lithos 54, 63–81. Valbracht, P.J., Oen, I.S., Beunk, F.F., 1994. Sm-Nd isotope systematics of 1.9–1.8-Ga granites from western Bergslagen, Sweden: inferences on a 2.1–2.0-Ga crustal precursor. Chem. Geol. 112, 21–37. Vuollo, J., 1994. Palaeoproterozoic basic igneous events in Eastern Fennoscandian Shield between 2.45 and 1.97 Ga, studied by means of mafic dyke swarms and ophiolites in Finland. Acta Univ. Ouluensis, Ser. A., 250, 1–47. Vuollo, J., Piirainen, T., Huhma, H., 1992. Two Early Proterozoic tholeiitic diabase dyke swarms in the Koli–Kaltimo area, eastern Finland – their geological significance. Geol. Surv. Finland, Bull. 363, 1–32. Walther, C., Flüh, E.R., 1993. The POLAR Profile revisited: combined P- and S-wave interpretation. In: R. Gorbatschev (Ed.), The Baltic Shield. Special volume. Precambrian Res. 64, 153–168. Ward, P., 1987. Early Proterozoic deposition and deformation at the Karelian craton margin in southeastern Finland. In: G. Gaál, R. Gorbatschev (Eds.), Precambrian geology and evolution of the central Baltic Shield. Special Issue. Precambrian Res. 35, 71–93. Ward, P., Härkönen, I., Nurmi, P.A., Pankka, H.S.,
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1989. Structural studies in the Lapland greenstone belt, northern Finland and their application to gold mineralization. In: S. Autio (Ed.), Geological Survey of Finland, Current Research 1988. Geol. Surv. Finland, Spec. Pap. 10, 71–77. Wasström, A., 1990. Knaftenområdet – en primitiv tidig Proterozoisk vulkanisk öbåge söder om Skelleftefältet, norra Sverige. M.Sc. Thesis, University of Åbo Akademi, Turku, Finland. 1–128. (in Swedish) Wasström, A., 1993. The Knaften granitoids of Västerbotten County, northern Sweden. In: T. Lundqvist (Ed.), Radiometric dating results. Sveriges Geol. Unders., Serie C823, 60–64. Wasström, A., 1996. U–Pb zircon dating of a quartz-feldspar porphyritic dyke in the Knaften area, Västerbotten County, northern Sweden. In: T. Lundqvist (Ed.), Radiometric dating results 2. Sveriges Geol. Unders., Serie C828, 34–40. Wegmann, C.E., 1928. Über die Tektonik der jüngeren Faltung in Ostfinnland. Fennia 50, No. 16, 1–22. Weihed, P., Bergman, J., Bergström, U., 1992. Metallogeny and tectonic evolution of the Early Proterozoic Skellefte district, northern Sweden. In: G. Gáal, K.J. Schulz (Eds.), Precambrian metallogeny related to plate tectonics. Special Issue. Precambrian Res. 58, 143–167. Welin, E., 1987. The depositional evolution of the Svecofennian supracrustal sequence in Finland and Sweden. In: G. Gaál, R. Gorbatschev (Eds.), Precambrian geology and evolution of the central Baltic Shield. Special Issue. Precambrian Res. 35, 95–113. Welin, E., 1992. Isotopic results of the Proterozoic crustal evolution of south-central Sweden; review and conclusion. Geol. Fören. Stockholm Förhandl. 114, 299–312. Wilson, M.R., 1980. Granite types in Sweden. Geol. Fören. Stockholm Förhandl. 102, 167–176.
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O.T. Rämö, I. Haapala 533
Cover page: Wiborgite, Wiborg batholith, southeastern Finland. Photo: Archives of the Geological Survey of Finland.
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Rämö, O.T., Haapala, I., 2005. Rapakivi granites. 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. 533–562. © 2005 Elsevier B.V. All rights reserved.
The classic Finnish rapakivi granite intrusions are found as four large batholiths (Wiborg, Åland, Laitila, Vehmaa) and several smaller plutons that sharply crosscut the Paleoproterozoic metamorphic bedrock of southern Finland. They belong to a mid-Proterozoic (1.67–1.47 Ga) anorogenic magmatic province that extends to central Sweden in the west, the Baltic countries in the south, and Russian Karelia in the east. The rapakivi granite intrusions are composed of a series of granitic rocks ranging from primitive fayalite-hornblende granite through hornblende granite, biotite-hornblende granite, and biotite granite to alkali-feldspar granite, the latter characterized by accessory magmatic topaz and associated tin mineralization. Geochemically, the Finnish rapakivi granites are aluminous (metaluminous to marginally peraluminous) A-type granites (high Fe/Mg, K/Na, Ga/Al, Zr, F, and LREE except Eu) and they do not include significant volumes of sodic silicic rocks. They show a bimodal magmatic association with temporally and spatially associated leucocratic gabbroic bodies (e.g., leucogabbronorite, anorthosite) and tholeiitic diabase dike swarms. Intermediate magmatic rocks (jotunite, ferrodiorite) are found only locally. Evidence for volcanic or subvolcanic lithologic units is present but sparse. The major rapakivi intrusions are associated with thinned crust and seismically complex, bulging upper mantle. This is particularly well established for the 1.65–1.62 Ga Wiborg batholith of southeastern Finland. The Nd isotope composition of the Finnish rapakivi granites (slightly negative initial εNd values) complies with the evolution path of the surrounding Paleoproterozoic crust, whereas the composition of the associated mafic rocks is, on average, somewhat more radiogenic than that of the granites. The genesis of the rapakivi granite melts is related to magmatic underplating and resultant anatexis of deep crust in an extensional tectonic regime that did not favor wholesale mixing of the mantle- and crust-derived melts. The cause of the mid-Proterozoic thermal perturbations is debated – possible scenarios include deep mantle plumes and distant convergent processes. The rapakivi texture (characterized by ovoidal plagioclase-mantled alkali feldspar megacrysts) is considered magmatic in origin and may relate to abrupt decrease of pressure during the emplacement of the intrusions and to hybridization processes.
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1. Introduction The Paleoproterozoic Svecofennian bedrock of southern Finland formed during a multiphase convergence and collision of lithospheric plates ~1900 Ma ago. The next event to strongly reshape the Finnish bedrock was the emplacement of the rapakivi granites ~1600 Ma ago, when the Svecofennian mountain chain had already been eroded down to its roots. At this time changes occurred in the Earth’s mantle within Fennoscandia, which led to rearrangements within the bedrock and the crystallization of rapakivi granites and accompanying mafic and intermediate rocks in the upper and middle parts of the continental crust. Sauna and rapakivi are the only Finnish words known in their original form in all civilized languages. The common people in Finland have applied the word rapakivi to describe the way certain rock outcrops weather into an easily crumbling rock or gravel (Figure 12.1). In literature, rapakivi is mentioned for the first time in 1694 in Urban Hjärne’s guide for mineral recognition and ore prospecting “En kort Anledning till Åtskillige Malm och Bergarters, Mineraliers etc. efterspöriande och angifvande” (in Swedish). The first scientific study on the Finnish rapakivi granites is the dissertation of J. Moliis, “Om Finska Sielffrätsten”, from 1798, which reported the regional distribution of the Finnish rapakivi granites and their mineral composition. J.J. Sederholm brought the term rapakivi granite into international geological literature in 1891 with his study “Ueber die finnländischen Rapakiwigesteine”. From then on, southern Finland has been regarded as the world’s type area for rapakivi granites. Sederholm’s work in solving geological problems involving rapakivi granites has been successfully carried on e.g., by Walter Wahl, Pentti Eskola, Victor Hackman, Th.G. Sahama, Antti Savolahti, Atso Vorma, and Matti Vaasjoki. This chapter is a state-of-the-art review of 536
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the locus classicus rapakivi granites as we now know them after more than 100 years of continuous research. We will also point out issues that still are inadequately deciphered and need to be focused on in the future.
2. What is rapakivi granite? The characteristic feature of the rapakivi granites is the presence of large (diameter 2–5 cm) rounded alkali feldspar crystals or ovoids that are surrounded by a grayish or geenish mantle of plagioclase feldspar. This is the rapakivi texture. In his study on the rapakivi granites of Fennoscandia, Vorma (1976) summarized a more detailed definition of the rapakivi texture: (1) The alkali feldspar phenocrysts are ovoidal; (2) Most (not all) ovoids are mantled by a rim of oligoclase–andesine; (3) Both alkali feldspar and quartz have crystallized in two generations, the early quartz generation as drop-like high quartz. However, all granites classified as rapakivi granites do not contain alkali feldspar ovoids. If the rapakivi texture is well developed, the rapakivi granite is called wiborgite (Figure 12.2A). If most or the majority of ovoids do not posses plagioclase mantles, the rock is a pyterlite (Figure 12.2B). In addition to wiborgite and pyterlite, many porphyritic and evengrained granites (Figure 12.2C) are considered rapakivi granites, and can be identified by the presence of drop-like quartz. A typical feature of the Finnish rapakivi granites is also their massive texture – they do not exhibit a penetrative fabric caused by orogenic movements. Granitic rocks of ~1700–1500 Ma age that have crystallized close to Earth’s surface, sharply cross-cut the bedrock, and generally exhibit rapakivi texture have in Finland been regarded as rapakivi granites sensu stricto (Vorma, 1976). Granites of any age containGRANITES
thus proposed a new definition for rapakivi granites that takes into consideration their special attributes but does not pose an age limit: Rapakivi granites are A-type granites characterized by the presence, at least in the larger batholiths, of granite varieties showing the rapakivi texture.
3. Distribution, mode of occurrence, and age
Fig. 12.1. Weathered rapakivi granite (moro in Finnish) in the Wiborg rapakivi batholith at Kymi. Photo: Ilmari Haapala.
ing alkali feldspar ovoids with plagioclase mantles have been called rapakivi granites sensu lato. Rapakivi granites were studied intensively during the 1990’s in different parts of the world. It has been established that rapakivi granites are found on every continent and have formed at least during the last ~3 billion years and that they almost exclusively represent an anorogenic evolution phase of the continental crust in any one area (Haapala and Rämö, 1999, and references therein). In addition to their characteristic mode of occurrence, they exhibit a bimodal rock association consisting of both felsic and mafic rocks, and show the petrographic and mineralogic features typical of A-type granites. Haapala and Rämö (1992)
On a global scale, rapakivi granites are particularly abundant within a zone that extends from the Ukraine to Fennoscandia and onwards via Greenland accross the North American continent to California (Rämö and Haapala, 1995). Many rapakivi granites are also known in South America, with associated rich tin occurrences (particularly in Brazil). Other rapakivi granites are known at least in Sudan, Tanzania, Botswana, the Ural Mountains, India, China, Siberia, Australia, and the Queen Maud Land in Antarctica. Rapakivi granites are found typically within or close to Paleo- or Mesoproterozoic continental crust formed 2000–1500 Ma ago and they are clearly younger than the surrounding crust. For example, the Finnish rapakivi granites are 1650–1540 Ma old and are surrounded by the ~1900 Ma Svecofennian orogenic crust of southern Finland. Figure 12.3 shows that there are four large rapakivi batholiths in southern Finland (Wiborg, Åland, Laitila, Vehmaa) and a number of smaller bodies (Suomenniemi, Ahvenisto, Onas, Bodom, Obbnäs, Peipohja, Mynämäki, Eurajoki, Reposaari, Siipyy, Fjälskär, Kökarsfjärden). In addition to these, on the northern shore of Lake Ladoga in Russian Karelia lies the Salmi rapakivi granite area, of which a considerable portion was part of Finland before World War II. Associated with the rapakivi plutons are quartz-feldspar porphyry dikes, which compositionally correspond to the rapakivi
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granites, and often cross-cut, in addition to the surrounding bedrock, the rapakivi granites themselves. Mafic rocks temporally and spatially associated with the rapakivi granites include gabbro–anorthosite plutons within or at the fringes of rapakivi granite areas and diabase dikes that form several dike swarms cross-cutting the Svecofennian bedrock (and, locally, the rapakivi granites). Several studies (e.g., Rämö, 1991; Eklund, 1993; Salonsaari, 1995) have shown that the felsic and mafic rocks of the rapakivi areas form also composite dikes and hybrid rocks. This means in effect that felsic and mafic melts coexisted in the crust during the emplacement of the rapakivi granites. In addition to felsic and mafic rocks, subordinate intermediate monzodioritic rocks (jotunite, ferrodiorite) are found within the rapakivi areas. Geophysical studies indicate that the Finnish rapakivi granites form thin, maximally 10-km-thick plate-like intrusive masses in the upper part of the continental crust (e.g., Korja and Elo, 1990). According to the present opinion, rapakivi granite melts formed in the lower parts of the continental crust and intruded into higher crustal levels (by cauldron subsidence and magmatic stoping) in an extensional tectonic setting. Deep shear zones and faults, formed already before the formation of the rapakivi plutons, guided the way of the rapakivi melts and the associated mafic
magmas through the continental crust. The Finnish rapakivi granites and the associated mafic and intermediate rocks are more than 150 Ma younger than the surrounding Svecofennian rocks. At present, there are abundant U-Pb data from various rapakivi areas (e.g., Suominen, 1991; Vaasjoki et al., 1991). They demonstrate that rapakivi granites may be divided into two age groups (Figure 12.4). The rapakivi plutons in southeastern Finland (Wiborg, Suomenniemi, Ahvenisto, Onas, Obbnäs, Bodom) are 1650–1620 Ma old. From the Häme diabase dike swarm west–northwest of the Wiborg batholith (Figure 12.3) there also exists a somewhat higher age of 1665 Ma (Vaasjoki and Sakko, 1989). The southwest Finnish plutons (Åland, Laitila, Vehmaa, Peipohja, Mynämäki, Reposaari, Siipyy, Eurajoki, Fjälskär, Kökarsfjärden) are about 50 Ma younger, with ages ranging from 1590 to 1540 Ma.
4. Lithologic association Typical of the rock association within the rapakivi areas is its bimodality – the vast majority of the rock types are either felsic (> 66 wt.% SiO2) or mafic (< 52 wt.% SiO2) in composition. Intermediate rocks are rare within the rapakivi areas. In this context, we discuss the felsic, mafic, and intermediate plutonic rocks as well as the dikes and volcanic
Fig. 12.2. (facing page) Different rock types of Finnish rapakivi granites, associated dikes, and mafic and intermediate rocks. (A) Wiborgite within the Wiborg batholith. Summa, Hamina; (B) Pyterlite within Obbnäs intrusion west of Helsinki; (C) Weakly porphyritic variety of the so-called Väkkärä granite within the Eurajoki stock (cf. Figure 12.6) in western Finland; (D) A large plagioclase crystal (diameter 10 cm) in olivine leucogabbronorite at Suopelto in the northeastern part of the Ahvenisto rapakivi area; (E) A pillow-like structure formed by a dark monzodiorite and granitic material at Pärnäjärvi, Jaala, in the southeastern part of the Ahvenisto rapakivi area; (F) An alkali feldspar diabase dike cross-cutting Svecofennian rocks on Kirkkovuori, Mäntyharju, northwest of the Suomenniemi rapakivi area; (G) A close-up on the Kirkkovuori alkali feldspar diabase; (H) A quartz-feldspar porphyry dike with abundant angular alkali feldspar phenocrysts at Taivaannaapuri, Heinola, southwest of the Ahvenisto rapakivi area. Photos: A, B, and D through H by O. Tapani Rämö, C by Ilmari Haapala. C H A P T E R 1 2 • R A PA K I V I G R A N I T E S •
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Finland
Russia
Ahvenisto 1643–1632 Ma
62° Siipyy
Salmi 1547–1530 Ma
Reposaari
Suomenniemi 1640–1635 Ma
Bothnian Sea Eurajoki
Peipohja Laitila 1570–1540 Ma
Åland 1576–1568 Ma
Lake Ladoga
Kymi Vehmaa 1570 Ma Bodom
Föglö swarm
60°
Onas Obbnäs 24°
Fjällskär
Wiborg 1646–1615 Ma Suursaari
Gulf of Finland
30°
100 km
Archean crust
Paleoproterozoic crust
Gabbro, anorthosite
Monzodiorite
Rapakivi granite
Diabase dikes
Phanerozoic sedimentary rocks Fig. 12.3. A generalized bedrock map of southeastern Fennoscandia. The ages of the most significant rapakivi areas (cf.Vaasjoki, 1996) are given. Modified from Rämö and Haapala (1995).
rocks separately.
4.1. Felsic plutonic rocks The felsic plutonic rocks of the rapakivi areas are granites sensu stricto and they are distinguished from other granitic rocks in Finland by, e.g., their mode of occurrence and magmatic association. Compared to the Paleoproterozoic calc-alkaline granites typical of the Svecofennian continental crust, the rapakivi granites are compositionally more homogeneous, contain more alkali feldspar, and their mafic silicates exhibit a higher Fe/Mg ratio (Simonen and Vorma, 1969; Rieder et al., 1996; Elliott, 2001). The felsic plutonic rocks within the Finnish rapakivi areas have crystallized in many stages and and several different intrusive phases cross-cut the older ones within the rapakivi plutons (Figures 12.5 through 12.7). The early and main intrusive phases are represented by granites in which the principal mafic 540
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silicates are hornblende (hastingsite), iron-rich biotite (annite, siderophyllite), and fayalite (sometimes also ferroaugite, orthoferrosilite, and pigeonite); later phases are characterized by granites with biotite as the sole major mafic mineral. Typical wiborgite contains biotite and hornblende as the main mafic silicates, whereas pyterlite contains only biotite. The presence of fayalite together with quartz and magnetite in the more mafic rapakivi granite indicates reducing conditions. The most common mineral in rapakivi granites is alkali feldspar, which usually forms ~50% of the rock and is found as euhedral or subhedral, often ovoidal crystals. In some dark, early rapakivi phases (e.g., the tirilite in the Wiborg batholith and the Tarkki granite of the Eurajoki stock), plagioclase is more abundant than alkali feldspar and is euhedral in respect to it. A typical feature of the rapakivi granites is that the mafic silicates (biotite, hornblende) fill the interstices between feldspars and quartz, GRANITES
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SW Finland
SE Finland
Number of age determinations
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1550
1600 Age (Ma)
1650
Felsic plutonic rocks
Felsic dikes
Mafic plutonic rocks
Mafic dikes
1700
Intermediate plutonic rocks Fig. 12.4. The distribution of zircon and baddeleyite U-Pb ages determined from Finnish rapakivi granites and associated mafic and intermediate plutonic rocks as well as felsic and mafic dike rocks. The diagram is based on 71 age determinations (Suominen, 1991;Vaasjoki, 1996; Alviola et al., 1999).
which have crystallized earlier. In granites forming the early and main intrusive phases, accessory minerals are fluorite, zircon, apatite, ilmenite, magnetite, anatase, and allanite; monazite is found instead of allanite in the varieties dominated by biotite. Plagioclasemantled alkali feldspar ovoids are common in the three largest rapakivi batholiths (Wiborg, Åland, Laitila), but are relatively rare in the others. The latest crystallizing phases of rapakivi granites are leucocratic topaz-bearing alkali feldspar granites (Figure 12.2C) with frequent associated greisen-type tin-beryllium-tungsten-zinc-lead mineralization (Haapala, 1977b, 1997; Edén, 1991; Lukkari, 2002; Figures 12.6 and 12.7). These granites cross-cut the other rapakivi varieties. They can be of even-grained
or porphyritic texture, but do not contain mantled alkali feldspar ovoids. The major minerals are quartz, microcline perthite, and albite. The dark mica (generally less than 5%) is lithium-bearing siderophyllite. Accessory minerals include topaz, fluorite, monazite, bastnäsite, xenotime, ilmenite, cassiterite, columbite, and thorite. Zircon, apatite, and magnetite, typical of the early and main rapakivi phases, are rare. The topaz-bearing granites contain often, especially in the upper parts on the intrusions, miarolitic cavities, which indicates that magmas of the late granite phases were saturated with water. At times the upper contact of the topaz-bearing granites exhibits a pegmatite zone (stockscheider), which also demonstrates the concentration of volatile matter in the up-
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Ahvenisto Suomenniemi FINLAND
Lappeenranta Jaala-Iitti 61°
Kouvola RUSSIA
Hamina Kymi 60°30’
Kotka
Gulf of Finland
Onas
Someri 25 km Suursaari
26°
v
Svecofennian bedrock
Porphyry aplite
Monzodiorite
Wiborgite
Pyterlite
Even-grained granite Hornblende or biotitehornblende granite Biotite granite
Diabase dikes Quartz-feldspar porphyry dikes Volcanic formation
Porphyritic granite
Topaz-bearing granite
Coarse-grained granite
Gabbro, anorthosite
Dark wiborgite
v
Fig. 12.5. Petrological map of the Finnish part of the Wiborg rapakivi batholith and its satellitic intrusions (Ahvenisto, Suomenniemi, Onas). The Jaala−Iitti complex that was investigated in detail by Salonsaari (1995) is shown in purple on the northwestern flank of the Wiborg batholith. Dashed line denotes the margin of the Wiborg batholith in the Gulf of Finland. Based on Vorma (1980), Rämö (1991), and Salonsaari (1995).
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Eurajoki
Kymi
61°15’
H
60° 30’
H
Sn, Be
H
Sn Sn
Sn, Be
Be
Sn, W, Be
Be, W Pb, Zn, Cu
A
Sn, W
2 km
1 km
B
Topaz-bearing granite
Even-grained topaz-bearing granite
Biotite granite
Porphyritic topaz-bearing granite
Faylite-biotite-hornblende granite
Marginal pegmatite (stockscheider)
Quartz porphyry dike (H = hybrid)
Quartz porphyry dike
Greisen and quartz veins
Greisen and quartz veins
Rapakivi granite (Laitila batholith)
Rapakivi granite (Wiborg batholith)
Crystalline basement (Paleoproterozoic) Fig. 12.6. Petrological maps of the (A) Eurajoki and (B) Kymi stocks (cf. Figure 12.3), which contain tin, tungsten, beryllium, zinc, and copper mineralization. According to Haapala (1977a, b).
per fringe of the magma reservoir. An excellent example of this is the Kymi stock (Figure 12.6B), where the zoned granite cupola is rimmed by a pegmatite zone containing topaz, tourmaline, beryl, arsenopyrite, molybdenite, and columbite (Haapala and Ojanperä, 1972; Kaartamo, 1996; Haapala and Lukkari, 2005). Ore formation associated with the topaz-bearing granites is linked to greisen veins and irregular greisen lenses that contain topaz, chlorite, cassiterite, wolframite, sulfides, and beryllium minerals (beryl, genthelvite, bertrandite; Haapala, 1977a,b). The bedrock map of the Wiborg rapakivi area is presented in Figure 12.5. Most of the Finnish part of the pluton consists of wiborgite.
Simonen and Vorma (1969) made a distinction between normal wiborgite and dark wiborgite, the latter containing plagioclase megacrysts in addition to wiborgitic ovoids. Both wiborgite and dark wiborgite contain hornblende in addition to biotite, but dark wiborgite contains more of it as well as fayalite. According to Simonen and Vorma (1969) ~80% of the Wiborg batholith in Finland consists of wiborgite or dark wiborgite. The rest is formed by pyterlite, porphyritic rapakivi granite (biotite-dominated granite with angular alkali feldspar phenocrysts), dark even-grained rapakivi granite (fayalite-hornblende granite, e.g., tirilite), even-grained biotite granite and porphyry aplite (leucocratic biotite granite
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Ahvenisto Fennoscandia Wiborg batholith
30 km
Lake Juolasvesi
Lake Enonvesi
Lake Vuohijärvi
5 km
Ahvenisto complex
Hornblende-biotite granite
Quartz-feldspar porphyry dike Greisen veins
Biotite granite
Monzodiorite
Topaz-bearing granite
Anorthosite
Porphyritic aplite and aplite
Olivine-bearing gabbroids
Hornblende granite
Leucogabbronorite Uralite gabbro Olivine diabase Svecofennian rocks Granitoids Schists
Fig. 12.7. Petrological map of the Ahvenisto rapakivi granite–gabbro–anorthosite complex. Location relative to Fennoscandia and the Wiborg rapakivi granite batholith is indicated. After Alviola et al. (1999).
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with rare large alkali feldspar crystals). The second largest rapakivi area in Finland lies in the Åland archipelago, and comprises likewise many different types of granites, which also represent granites crystallized from rapakivi melts during different phases. Bergman (1981) has recognized the following plutonic rocks from the Åland batholith: (1) quartz-feldspar porphyry (the oldest), (2) porphyritic hornblende rapakivi, (3) wiborgite, (4) pyterlite, (5) even-grained rapakivi granite, and (6) aplitic rapakivi granite (the youngest). Most of the Åland batholith consists of wiborgite and pyterlite. The third largest rapakivi area in Finland, the Laitila batholith consists principally of a pyterlite, where alkali feldspar ovoids are well developed, but generally without a plagioclase mantle (Vorma, 1976). In addition to pyterlite, several hornblende, hornblende-biotite, and biotite granites are found. Northwest of the Laitila batholith lies the Eurajoki stock, which contains a relatively low-silica fayalite-bearing hornblende granite (Tarkki granite) and a leucocratic, topaz-bearing alkali feldspar granite (Väkkärä granite, Figure 12.2C; Haapala, 1977a and 1997). A characteristic feature of the other rapakivi granite plutons in Finland is that they usually consist of at least a few types of granites, ranging from hornblende- to biotite-bearing varieties (often also topaz-bearing granites). Detailed descriptions, in addition to those of the Laitila (Vorma, 1976) and Eurajoki (Haapala, 1977a) plutons, have been published from the Vehmaa (Kanerva, 1928; Lindberg and Bergman, 1993), Ahvenisto (Savolahti, 1956; Johanson, 1984; Edén, 1991; Alviola et al., 1999), Suomenniemi (Rämö, 1991), and Jaala-Iitti (Salonsaari, 1995) areas.
4.2. Mafic plutonic rocks The majority of the mafic plutonic rocks are leucocratic gabbros (leucogabbronorites, olivine leucogabbronorites, leucogabbros, leuco-
norites, leucotroctolites) with plagioclase, pyroxenes, and locally olivine as the major minerals. There is also some anorthosite with plagioclase in excess of 90%. The mafic plutonic rocks of the rapakivi association are found either as inclusions within the rapakivi plutons or as small intrusions on their margins. U-Pb ages do not show any measurable differences (in excess of a few million years) between the felsic and mafic plutonic rocks of the rapakivi areas (Suominen, 1991; Vaasjoki et al., 1991; Alviola et al., 1999). The most significant occurrence of mafic plutonic rocks is the Ahvenisto gabbro–anorthosite complex at Mäntyharju northwest of the Wiborg rapakivi area (Figure 12.7). The complex has the shape of a horse-shoe and surrounds the Ahvenisto rapakivi granite pluton. Most of the gabbro–anorthosite complex consists of coarse-grained leucogabbronorite, which contains scattered large (up to 10 cm in diameter) plagioclase crystals. In the northwestern part of the complex, there are olivine-bearing mafic plutonic rocks, olivine leucogabbronorite and leucotroctolite, also containing large plagioclase crystals and representing an earlier phase in the evolution of the mafic magma (Figure 12.2D). Anorthosite within the Ahvenisto complex is found as lens-shaped, maximally 200-m-long bodies in the more mafic rock types of the complex. The contact between the anorthositic bodies and the surrounding mafic rocks is sharp. The most typical minerals in the rocks of the Ahvenisto gabbro–anorthosite complex are plagioclase (An45–70), orthopyroxene, olivine, augite, titanian magnetite, and ilmenite. In addition to Mäntyharju, rapakivi-associated gabbros and anorthosites are found in Ylämaa in the central part of Wiborg batholith (Figure 12.5). Several mafic inliers, their dimensions ranging from a few hundreds of meters to a few kilometers, occur within the rapakivi, their plagioclase being iridescent andesine–labradorite. Small gabbro–anorthosite
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bodies are also known at Väärälampi on the northern flank of the Suomenniemi pluton, at Kolinummi southeast of the Laitila batholith, and in the western and southwestern parts of the Åland batholith (Figure 12.3). In areas where the continental crust has been eroded to a deeper section than in Finland, a much larger proportion of gabbros and anorthosites is present. This is the case for example in Labrador, where roughly equal amounts of felsic and mafic plutonic rocks are found. Significantly, recent seismic studies indicate that a considerable gabbro–anorthosite pluton probably lies beneath the Wiborg batholith at a depth of 10–15 km (Korja, 1995). Recent magnetic and gravimetric surveys (Elo et al., 1996) suggest that the Ahvenisto gabbro–anorthosite complex is a small outcropping part of this pluton.
4.3. Intermediate plutonic rocks Although the magmatic association of the rapakivi granites is clearly bimodal, some intermediate rocks are also present. These have been described in detail from the Ahvenisto gabbro–anorthosite complex (Johanson, 1984; Alviola et al., 1999). The intermediate rocks of the Ahvenisto complex are monzodiorites and quartz monzodiorites, which contain distinctly more alkali feldspar than the mafic rocks of the complex. They are found as dikes within the complex and at its margins (Figure 12.7). The intermediate dikes cross-cut the gabbros and anorthosites, whereas rapakivi granites generally cut the monzodiorites and quartz monzodiorites; locally also mingling of monzodiorites and granite has occurred (Figure 12.2e). The monzodioritic rocks of the Ahvenisto complex are medium-grained, often porphyritic and darker than the gabbros and anorthosites. Their major minerals are plagioclase (An35–50), hypersthene, and hornblende. Accessory minerals are alkali feldspar, quartz, iron-titanium oxides, apatite, zircon, baddeleyite, and, in some cases, Fe-rich olivine. 546
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In addition to Ahvenisto, intermediate plutonic rocks associated with rapakivi granites have also been reported from the Åland batholith (Eklund et al., 1994).
4.4. Dikes and volcanic rocks The present level of erosion in southern Finland represents a section of the Mesoproterozoic bedrock at an estimated depth of a few kilometers and dike rocks related to the rapakivi granites are found rather widely in the bedrock of southern Finland. Volcanic rocks, however, are rare. Within the Wiborg batholith itself, dike rocks are few (Figure 12.5). The best known are the quartz porphyry dikes around Hamina in the central part of the batholith (Simonen, 1987). Rapakivi-associated volcanism is probably represented by the Ruoholampi (and possibly also Taalikkala) roof pendant near Lappeenranta, in which amygdaloidal porphyritic diabase has been observed (Vorma, 1975). In the Ruoholampi roof pendant there also is a felsite porphyry interpreted as a volcanic rock. The best evidence for volcanic activity assosiated with the Wiborg batholith is found on the Island of Suursaari (Hogland), which lies on the southern margin of the pluton (Figure 12.3). The bedrock of Suursaari was studied before World War II by Ramsay (1890), Kranck (1929), and Wahl (1938). The lowermost formation on the island is the Paleoproterozoic Svecofennian bedrock and its erosion surface. The latter is covered by a quartzose conglomerate with no evidence for later deformation. Volcanic activity in the area started with an eruption that formed a thin layer of ash and ejecta on the conglomerate. After this, basaltic and andesitic lavas flowed into the valleys and tephra with tuffaceous breccia structure was deposit atop of them. The tuffaceous breccias and lavas of the first eruption phase are overlain by locally over 100-m-thick quartz-feldspar porphyries. The dikes around the Suomenniemi GRANITES
5m
Quartz-feldspar porphyry
Mixed rock
Diabase
Svecofennian gneissic granite
Diabase with alkali feldspar and quartz xenocrysts Fig. 12.8. Schematic diagram of a zoned diabase–quartz-feldspar porphyry dike cross-cutting Paleoproterozoic Svecofennian rocks at Korpijärvi, Mäntyharju, northwest of the Suomenniemi rapakivi pluton. Silicic magma has first intruded a crack in the Svecofennian bedrock and has been followed by diabase magma in a manner that the melts have partly intermingled (Rämö, 1991). Dashed lines indicate the probable continuation of the different zones.
rapakivi pluton are exceptionally varied in character (Rämö, 1991). Northwest of the pluton is a NW-trending tholeiitic diabase dike swarm that extends over 80 km away from the Suomenniemi pluton (Figure 12.5). A number of dikes have passed through a silicic magma chamber and have taken with them abundant alkali feldspar ovoids, some with a plagioclase mantle (Figure 12.2F, 12.2G). In the same swarm with the diabases, but only to a distance of about 12 km from the pluton, there are also quartz-feldspar porphyry dikes. The diabase
and quartz-feldspar porphyry magmas were melts at the same time, as both have entered the same fissures with resultant local mingling structures (Figure 12.8). The Suomenniemi pluton is also cut by a dike swarm that consists of peralkaline alkali feldspar syenite and alkali feldspar quartz syenite (Rämö, 1991). The dikes are generally a few meters wide, NW-striking, and of variable dip. The major minerals are mesoperthite and aegirine-augite. Alkali amphibole, titanite, titanian andradite, zircon, and quartz are some of the accessory
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minerals. Between the Suomenniemi and Wiborg rapakivi batholiths is the ~1645 Ma Lovasjärvi mafic intrusion that is slightly older than the granitic batholiths. It is a 5.4-km-long and 800–1500-m-wide vertical sheet-like intrusion and consists of diabase, olivine diabase, and melatroctolite (Alviola, 1981; Siivola, 1987). This mafic intrusion represents an early, internally differentiated tholeiitic magma chamber related to the diabase magmatism of the area. Northwest of the Wiborg rapakivi batholith is the Häme diabase dike swarm (Figure 12.3) with two sets of diabase dikes deviating from each other in direction, age, and composition. The older of them strikes west-northeast and is ~1665 Ma old (Vaasjoki and Sakko, 1989), ~100 km long, and consists of medium-grained diabase with abundant olivine (Laitakari, 1987). Plagioclase is not present as phenocrysts or large fragments in these dikes (Laitakari and Leino, 1989). The younger set of dikes within the swarm is ~1645 Ma old (Laitakari, 1969). It strikes towards northwest and is ~150 km long. In the vicinity of the Ahvenisto rapakivi granite complex, also quartzfeldspar porphyry magmas have invaded the same system of cracks. In the Häme swarm, the largest diabase dikes lie in the eastern end of the swarm close to the Ahvenisto complex. The dikes become narrower towards the west and, e.g., at Kuru in the central part of the swarm, the widest of them measures no more than 10 m (Laitakari, 1987). The dikes of the younger set of the Häme swarm are mainly of olivine tholeiites, but they contain clearly less olivine than those of the older set; some dikes are quartz tholeiitic. The pyroxene is principally augite, but orthopyroxene has been met with occasionally. Plagioclase is found as phenocrysts, sometimes even as ~20-cm-long megacrysts that are concentrated in the central parts of the dikes. In the thinnest, only a few centimeters wide dikes basaltic glass has been preserved 548
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(Lindqvist and Laitakari, 1980). The dike-like Jaala–Iitti complex at the northwestern margin of the Wiborg batholith was studied in detail by Salonsaari (1995). The complex is arcuate, 22 km long, the width varying from 0.1 to 1.5 km, and it sharply cross-cuts the granite of the Wiborg pluton and the surrounding Svecofennian bedrock. By its composition the complex is hybrid rock and has, according to Salonsaari (1995), been formed through mingling of basaltic and rhyolitic melts in a magma chamber deeper in the crust at temperatures between 950 and 750 oC. The Jaala–Iitti complex contains indications of both end members (mafic basaltic inclusions, granites) and of the hybrid rock types (monzogranites, quartz monzonites) resulting from magma mixing and mingling. In addition to the dike rocks associated with the Wiborg batholith, both quartz-feldspar porphyry and diabase dikes have been found around the small rapakivi plutons (Onas, Bodom, Obbnäs) in the vicinity of Helsinki (e.g., Törnroos, 1984; Kosunen, 1999). Their strike is almost parallel to that of the more northerly Häme swarm. In the rapakivi areas of southwestern Finland, dike rocks are present between Åland and Turku (Ehlers and Ehlers, 1977; Suominen, 1987, 1991), in the southwestern part of the Åland rapakivi batholith (Eklund, 1993), in the Laitila batholith (Haapala, 1977a), and near the town of Pori north of Laitila (Pihlaja, 1987). The dikes in the Åland and Turku areas comprise pyroxene diabase, amphibole diabase, and quartz-feldspar porphyry. Of the pyroxene diabases, the best known is the 35-km-long Föglö swarm (Figure 12.3) between the Åland and Kökarsfjärden rapakivi plutons. It consists of several diabase dikes emplaced in an en echelon manner. The Föglö swarm is not known to cross-cut the rapakivi granites, but the association is clear as its isotopic ages (~1570–1540 Ma) correspond to the age of the Åland rapakivi granites. Many of the Föglö dikes contain large plagioclase fragments, even several tens of centimeters GRANITES
in diameter. In association with the Eurajoki rapakivi pluton (Figures 12.3 and 12.6) are felsic, locally topaz-bearing quartz porphyry dikes and more mafic, dark porphyry dikes. The dark dikes contain plagioclase (An55–60), quartz, alkali feldspar, and ilmenite mega crysts (Haapala, 1977a). The plagioclase and quartz megacrysts are usually strongly resorbed, whereas the alkali feldspar megacrysts are more euhedral. The ground mass of the dikes consists of plagioclase, biotite, hornblende, alkali feldspar, quartz, and apatite. The dark porphyry dikes are the result of mingling of mafic and felsic magmas typical of the rapakivi association (Haapala, 1997). It is quite probable that the present level of erosion of the Finnish bedrock represents at the Åland batholith a shallower section than around the plutons in southeastern Finland (cf. Bergman, 1986). Subvolcanic complexes, the evolution of which is constrained by the bimodal magmatism typical of rapakivi associations and local hybridization processes (e.g., Eklund, 1993; Eklund et al., 1994), have been described in conjunction with the Åland batholith. Also an ignimbrite-like lithologic unit has been described from the area (Eklund et al., 1996).
5. Chemical composition In addition to their singular texture and mode of occurrence, the rapakivi granites also exhibit a singular chemical composition that sets them apart from many other granites. The rapakivi granites contain more Si, K, F, Rb, Ga, Zr, Hf, Th, U, Zn, and REE but less Ca, Mg, Al, P, and Sr than granitic rocks in general (Rämö and Haapala, 1995). In the late topaz-bearing intrusive phases of the rapakivi plutons, Sn, F, Ga, Rb, Sn, and Nb are especially abundant and contents of Ba, Sr, and Zr are very low. In terms of aluminum saturation, the rapakivi granites range from metaluminous
(early intrusive phases) to marginally peraluminous (late intrusive phases). Typical features of the rapakivi granites are also high Fe/Mg, K/Na, and Rb/Sr ratios. These geochemical traits are evident in Table 12.1 that shows the area-weighted average chemical compositions of the Wiborg and Laitila batholiths in comparison to the granite average of Turekian and Wedepohl (1961). The area-weighted means of the two rapakivi batholiths can be considered as fair estimates of the composition of their initial magmas (Rämö and Haapala, 1995). The two batholiths are close to each other compositionally, save for the Wiborg batholith being slightly more mafic (higher Mg, Fe), calcic, and sodic. Comparison to the average granite of Turekian and Wedepohl (1961) shows that, at the roughly similar SiO2 values of ~70.5 wt.%, the rapakivi batholiths have higher FeO* (~3.7 vs. 2.8 wt.%), K2O (~5.6 vs. 4.1 wt.%), and F (0.24 vs. 0.07 wt.%) as well as lower Al2O3 (13.8 vs. 14.6 wt.%), MgO (0.32 vs. 0.91 wt.%), and Na2O (2.8 vs. 3.7 wt.%). Figure 12.9 depicts the chemical composition of the Finnish rapakivi granites on the discrimination diagrams from Pearce et al. (1984) and Whalen et al. (1987). The Finnish rapakivi granites plot within the field of the within-plate granites (Figure 12.9A) and have the characteristics of the A-type granites (Figure 12.9B). Chondrite-normalized REE abundances of the main rapakivi granite types (e.g., wiborgite, pyterlite) indicate strong overall enrichment in the LREE and a small Eu minimum; the late-stage topaz-bearing intrusive phases have flat chondrite-normalized patterns and a very pronounced Eu minimum (Figure 12.10). In the classification scheme of Frost et al. (2001), the Finnish rapakivi granites mostly comply with the ferroan alkalicalcic granitoids (Figure 12.11).
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Table 12.1. Area-weighted mean compositions of the Wiborg (Finnish part) and Laitila rapakivi batholiths compared to average granite composition.
A/CNK c
0.95 0.91 1.89 2.23
1.05 0.93 2.16 2.50
1.02 0.76 1.11 0.52
Rb (ppm) Sr Ba Zr
252 113 1062 367
265 106 1175 344
140 270 630 158
VAG 10
12
•
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1000 Y+Nb (ppm)
10000
100
B
10
A-type
FG
OGT 1 100
1000 Zr+Nb+Ce+Y (ppm)
10000
Fig. 12.9. The geochemical composition of the Finnish rapakivi granites (data from Rämö and Haapala, 1995) on (A) a (Y+Nb) vs. Rb diagram (Pearce et al., 1984) and (B) a (Zr+Nb+Ce+Y) vs. (K2O+Na2O)/CaO diagram (Whalen et al., 1987). Abbreviations for the fields: FG−extensively fractionated felsic granites; OGT−unfractionated M-, I-, and S-type granites; syn-COLG−granites associated with collisions of tectonic plates; WPG−within plate granites; ORG−ocean ridge granites;VAG−volcanic arc granites.
The rapakivi texture, especially the alkali feldspar ovoids mantled with plagioclase, has been subject to many studies and interpretations, but no generally accepted model for its origin has emerged so far. Models worth consideration should be able to explain both the roundish form of the feldspar crystals and the mechanism of the mantle formation. Most models propose that the rapakivi texture was formed during crystallization of the rapakivi melt. CHAPTER
ORG
1 100
6. Origin of the rapakivi texture
•
WPG
100
a Based on the areal percentage and mean composition of the granite types present in each batholith (Rämö and Haapala, 1995). b After Turekian and Wedepohl (1961). c Molecular Al2O3/(CaO+Na2O+K2O). d Weight ratio, Fe* denotes total iron.
550
synCOLG
Rb (ppm)
Area-weighted meansa Granite averageb Wiborg Laitila 70.41 70.68 70.71 0.45 0.43 0.38 13.66 13.98 14.55 3.93 3.54 2.82 0.05 0.04 0.06 0.37 0.28 0.91 1.96 1.59 2.13 2.95 2.61 3.65 5.57 5.63 4.05 0.13 0.08 0.17 0.22 0.26 0.07 99.70 99.12 99.50 0.09 0.11 0.03 99.61 99.01 99.47
Fe*/(Fe*+Mg) d K/Na Rb/Sr
A
1000
(K2O+Na2O)/CaO
SiO2 (wt.%) TiO2 Al2O3 FeOTOTAL MnO MgO CaO Na2O K2O P2O5 F Total –O=F2 Total
10000
Magmatic origin is supported by the fact that the alkali feldspar crystals are at times broken and the plagioclase mantle is either thinner or totally lacking on the broken surface. This demonstrates that the plagioclase mantles have grown on surfaces of alkali feldspars existing in a partly crystallized granite magma. The main theories on the origin of the
GRANITES
1.0
1000 Finnish rapakivi granites
A
0.9
1
Selected granite types from the Wiborg batholith: Wiborgite Pyterlite Topaz-bearing granite
0.2 La Ce Pr Nd Sm EuGd Tb Dy Ho Er Tm Yb Lu
Fig. 12.10. The rare earth element (REE) contents of the Finnish rapakivi granites (gray field – data from Rämö and Haapala, 1995) relative to the chondritic composition. Compositions of three selected rapakivi granites from the Wiborg batholith are also shown (data from Haapala et al., 2005).
rapakivi texture can be grouped in the following manner: (1) Exsolution in a solid state, with the exsolving plagioclase migrating to the rims of grains (Elders, 1968; Dempster et al., 1994). (2) Crystallization from a viscous granite melt with few nuclei for crystal growth. Because of the slow rate of diffusion, (a) the alkali feldspar crystals grow roundish and (b) a potassium-depleted and calcium-enriched boundary layer is created around the alkali feldspar crystals, which causes the crystallization of the plagioclase shell (Sederholm, 1928; Savolahti, 1962). (3) A rapid decrease of pressure coupled with a slow decrease in temperature may stabilize plagioclase at the expense of alkali feldspar (and quartz); alkali feldspar and quartz crystals are resorbed and rounded, and plagioclase starts to crystallize on the surfaces of the alkali feldspar while the orientation of the crystal structure is retained (Nekvasil, 1991). (4) Mingling and mixing of the crystal-
0.8
ferroan
0.7 magnesian 0.6 0.5
Finnish rapakivi granites
0.4 50 55 12 10 B 8 6 alkalic 4 2 cic 0 -cal kali l a -2 ali -alk -4 calc -6 -8 50 55
60
65
70
75
80
65 70 SiO2 (wt.%)
75
80
A-type
Na2O+K2O–CaO
Sample/Chondrite
10
FeO*/(FeO*+MgO)
A-type 100
60
Fig. 12.11. The composition of the Finnish rapakivi granites (Rämö and Haapala, 1995) in (A) FeO*/(FeO* + MgO) vs. SiO2 and (B) (Na2O + K2O – CaO) vs. SiO2 classification diagrams of Frost et al. (2001). Composition range of ~500 A-type granites (as referred to in Frost et al., 2001) is also shown.
lizing granitic melt (crystal mush) with mafic melt can lead to the resorption of alkali feldspar and quartz crystals and the crystallization of plagioclase on surfaces of alkali feldspar (Hibbard, 1981; Stimac and Wark, 1992). Of these hypotheses, the exsolution theory (1) cannot be applied as a general model, as plagioclase mantles are in some cases found around granite inclusions, but there is no mantle on the alkali feldspar crystals within the inclusions (Figure 12.12). Neither is the slowness of diffusion in a viscous magma (2) generally applicable, because, for example, the high fluorine content typical of rapakivi magma very effectively decreases viscosity and increases rates of diffusion. Models (3)
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Fig. 12.12. Plagioclase-mantled granite autolith in wiborgite. Erratic boulder from the Wiborg rapakivi granite area. Photo: O. Tapani Rämö.
and (4) are based on petrological observations and experiments and are both possible mechanisms for the origin of rapakivi texture. They require that pressure, temperature, water fugacity, and chemical composition within the rapakivi magma change in a manner that renders alkali feldspar unstable (resorption, rounding) but plagioclase stable, and thus allow plagioclase to crystallize on the alkali feldspar megacrysts. Salonsaari (1995) showed that rounded mantled alkali feldspar ovoids were formed in the Jaala–Iitti complex (Figure 12.5) as a result of magma mingling and mixing (hybridization). In this process, alkali feldspar crystals from the granitic crystal mush were enclosed by more mafic magma, with which they reacted and became resorbed on their edges. The mantles of the ovoids usually con552
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sist of a micrographic plagioclase-quartz intergrowth (thus not a normal rapakivi texture), sometimes of plagioclase, and were formed at a temperature of 850 oC to 750 oC and a pressure under 2 kb. According to Salonsaari there is, however, no evidence of large-scale magma mingling within the Wiborg rapakivi pluton, although 80% of its area is covered by rapakivi granites containing mantled ovoids. Thus he did not regard hybridization as the principal formation mechanism for rapakivi texture, but concluded that a temperature rise caused by mafic magmas resulted in the formation of rapakivi texture in a granitic crystal mush. It may also be envisaged that hot mafic magmas retard the decrease of temperature in a rising, partly crystalline magma and thus allow rapakivi texture to develop according to the model of Nekvasil (1991). By comparing mineral chemical, pressure, and temperature determinations of rapakivi granites and experimental data, Eklund and Shebanov (1999) concluded that subisothermal pressure decrease in ascending magma is a viable mechanism for the formation of the rapakivi texture.
7. Origin of the rapakivi magma The rapakivi granites in Finland are about 350–150 Ma younger than the surrounding Paleoproterozoic Svecofennian rocks. The large age difference justifies the opinion that the rapakivi granites are anorogenic in respect to the Svecofennian orogeny. This conclusion is supported by the discordant mode of occurrence, bimodal magmatic association, and the mineralogical and geochemical (A-type) characteristics of the rapakivi granites. Studies carried out at the Institute of Seismology, University of Helsinki (e.g., Luosto et al., 1990; Korja et al., 1993; Luosto, 1997) demonstrate that the rapakivi plutons are found in areas where continental crust is thinner than in the surroundings, i.e., the mantle bulges upwards in a dome-like fashGRANITES
A
Rapakivi granite
B
C
6.0 6.1
6.5 6.45 6.6 6.7 7.0 7.2
E
Wiborg batholith
E
6.2
10
D
D
km
30
C 50
HVL
6.1 6.2 6.4 6.45 6.65 6.8 7.0
8.0 8.2
M1
8.4
M2
7.3 8.5
B 70 A Gulf of Finland
0
50
100
150
200
250 km
Fig. 12.13. Vertical section based on seismic studies (Luosto et al., 1990) through the continental crust and upper mantle in the western part of the Wiborg rapakivi batholith and north of it. The dashed lines indicate changes in the velocities of seismic longitudinal waves, HVL is a high velocity layer (most likely gabbro and anorthosite), and M1 and M2 indicate Moho discontinuities at crustmantle boundary and within the uppermost mantle. Modified from Rämö and Haapala (1996).
ion (Figures 12.13, 12.14). In the areas of the domes there is in the uppermost parts of the mantle a zone, where the velocities of seismic longitudinal waves are lower than in mantle peridotite in general, but clearly higher than in the lower parts of continental crust. Apparently this zone consists of a mixture of mantle peridotite, mantle-derived mafic rock, and partially melted crustal rocks. The lower part of the crust has thinned in the area of the domes (Figure 12.13). The swarms of subparallel diabase dikes around rapakivi complexes are indicative of an extensional tectonic regime at the time of rapakivi magmatism (Haapala, 1988). Similarly, an extensional environment is demonstrated by some normal faults and shallow-dip listric faults (Korja and Heikkinen, 1995). Isotope geological and geochemical studies (e.g., Rämö, 1991, 2001) have demonstrated that the parent magmas of the Finnish rapakivi granites formed through partial melt-
ing of the deep parts of the Paloproterozoic Svecofennian crust (Figure 12.15), in the case of the Suomenniemi pluton possibly by partial melting of intermediate-felsic rock material. In contrast, a large part of the material of the Salmi rapakivi batholith lying in the border zone between the Archean and Proterozoic domains (cf. Figure 12.3) represents remelted Archean crust (Figure 12.15; Rämö, 1991; Neymark et al., 1994). The diabase dikes and other mafic rocks (gabbros and anorthosites) have a Nd isotope composition that partly overlaps that of the granites, but is slightly more radiogenic (higher ⑀Nd values) on average (Figure 12.15). This suggests that they were crystallized from mantle melts that were variably contaminated by crustal material (Rämö, 1991). The results of the studies presented above can be collected to a model depicting the origin of the rapakivi magmatism as follows (Figure 12.16):
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43
55
Finland
43
59
1.55–1.53
RAGUNDA
57
61
AHVENISTO
1.53–1.47 41
Baltic Sea
45
Sweden 47 North 16° Atlantic
Lake Ladoga Gulf of Finland
43
1.67– 4 Estonia 1.62 9
43 41
400 km
45
Russia
47
0
100
200 km
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Rapakivi granite complexes Rapakivi-age mafic intrusions and dikes
RIGA
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SW Scandinavian domain Phanerozoic
47
WIBORG
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57 51
1.59– 1.54
A
41 43
45
68°
68°
Archean Paleoproterozoic
SALMI
ÅLAND
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16°
SUOMENNIEMI
59
Bothnian Sea
TIB
41
63
57
49
46
53
55 53
NORDINGRÅ
51
47 45
51
45 49
Latvia
57 59 Moho depth contours
Fig. 12.14. Map showing the 1.67–1.47 Ga rapakivi granite complexes and contours of crustal thickness of the south-central part of the Fennoscandian Shield. The complexes delineate four age clusters from west to east: 1.53–1.47, 1.59–1.54, 1.67–1.62, and 1.56–1.53 Ga; the latter three are indicated with blue lines (cf. Haapala et al., 2005). The inset shows the area relative to the major crustal domains of the shield. TIB is the 1.85–1.7 Ga Transscandinavian igneous belt. The green area marked by “A” in the inset denotes 1.69–1.55 Ga magmatic arcs in southwesternmost Sweden (Ahl et al., 1997; Åhäll et al., 2000). Note that the area south of the Gulf of Finland is covered by Phanerozoic sedimentary rocks (not shown in principal map). Compiled mainly after Koistinen (1994), Rämö et al. (1996), Andersson (1997), Korja et al. (2001), Korsman et al. (1999), and Persson (1999).
(1) Mafic magma created by partial melting of the upper mantle caused magmatic underplating and thus led to partial melting of the lower continental crust to form granitic rapakivi magma. (2) Mafic magmas that rose along deep fissures in the crust formed diabase dikes, gabbro–anorthosite plutons, and layers of basaltic 554
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lavas; granitic magmas formed quartz-feldspar porphyry dikes, rapakivi granite plutons, and rhyolitic supracrustal rocks. Magma mingling and mixing occurred locally at varying depths. The emplacement of all the large plutons and many of the smaller stocks occurred as a result of several sequential magma pulses. The lower part of the continental crust thinned in the proGRANITES
1.9 Ga Svecofennian crust
Depleted Mantle
+5
Finnish rapakivi granites
CHUR
Evo zoic lution o crus f Pa leop t
0
εnd rote ro-
Evolution of the source of the Finnish granites
–5
Evo
luti
on
Russian Karelian rapakivi granites (mixed source)
of A
rch e
Mafic rocks associated with the Finnish rapakivi granites 2.0
–10
an c
rus
t –15
Age (Ga)
1.5
Fig. 12.15. An ⑀Nd vs. age diagram illustrating the Nd isotope evolution of the rapakivi granites, Svecofennian crust, Archean crust, and upper mantle depleted in the LREE from 2.1 to 1.1 Ga ago. Data points of the Finnish rapakivi granites (Rämö, 1991) are marked by capital “F”, those of the associated mafic rocks by blue dots. The diagram shows that the Finnish rapakivi granites probably represent remelted Svecofennian crust. In contrast, the material of the rapakivi granites in Russian Karelia (Rämö, 1991; Neymark et al., 1994; Amelin et al., 1997: cf. Figure 12.3) is a mixture of two sources (Paleoproterozoic and Archean crust). CHUR indicates the evolution of undifferentiated Earth (DePaolo and Wasserburg, 1976).The evolution path of the depleted mantle is from DePaolo (1981), that of the Svecofennian crust from Huhma (1986) and Patchett and Kouvo (1986), and that of the Archean crust from Rämö et al. (1996).
cess because of the shift of material – rapakivi granite melts were extracted from this domain – and because of plastic stretching resulting
from extension. The bimodal rapakivi granite magmatism occurred at different times in different places
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Present level of erosion Upper crust Middle crust Crust
Lower crust Mantle peridotite + gabbro
mantle
Mantle peridotite Basic magma or pluton
Quartz-feldspar porphyry dike
Diabase dike
Composite dike
Silicic magma or pluton (rapakivi granite)
Silicic/basic volcanic rock
Fig. 12.16. A two-stage model of the formation of rapakivi granites according to Haapala (1989) and Rämö and Haapala (1996). For detailed explanation, see the text.
(Figure 12.14): the Wiborg rapakivi batholith with its satellitic plutons and diabase dikes was emplaced at 1670–1620 Ma, the rapakivi intrusions in southwestern Finland (the Åland, Laitila, and Vehmaa batholiths and associated intrusions) at ~1590–1540 Ma, and the Salmi rapakivi complex in Russian Karelia at 1550–1530 Ma. This shows that diapiric mantle upwelling and resultant crustal melting occurred below and within the cratonized Paleoproterozoic crust of southeastern Fennoscandia during more than 100 Ma without a regular age-locality pattern.
8. Tectonic scenarios There is nowadays a wide consensus that the heating effect of mafic magmatism of mantle origin formed the rapakivi magmas via melting of the deep crust and that the rapakivi granites crystallized from these magmas (Haapala and Rämö, 1999, and references therein). However, the cause for the partial melting of the mantle is yet unsolved. Overall, the Fennoscandian rapakivi granite batholiths (Figure 556
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12.14) form four age-associated clusters. The batholiths are found as relative thin (~5–10 km) sheet-like bodies in the upper part of the crust, and the Moho contours depicted in Figure 12.14 indicate particularly steep ovoid thinnings in the crust associated with the rapakivi intrusions (Elo and Korja, 1993; Luosto, 1997). The crust hosting the rapakivi intrusions is also characterized by listric seismic reflectors and thinned lower crust (Korja et al., 2001). These observations fit well the magmatic underplating model and they can also explain the extensional tectonic setting and the bimodal nature of the magmatism. The ultimate cause of the underplating is controversial, however. Plausible mechanisms include active or passive rifting, extensional collapse of orogen, and deep mantle plumes (Haapala and Rämö, 1999, and references therein). Furthermore, petrologically unstable domains in the lithospheric mantle (related to earlier or contemporaneous distant subduction zones) could have controlled the loci of magmatism (Haapala and Rämö, 1992; Rämö and Haapala, 1995). The Fennoscandian rapakivi granites have recently been related to intermitGRANITES
tent subduction events on the southwestern flank of the Fennoscandian Shield (Åhäll et al., 2000). Roughly NS-trending magmatic arcs related to this postulated process are located within a thin sliver of crust on the eastern flank of the southwestern Scandinavian domain (Figure 12.14 inset). This “inboard model” is, however, unable to account for the non-linear age distribution of the Fennsocandian rapakivi granite complexes.
9. Future challenges Regarding the petrogenesis of the Finnish rapakivi granites, and the mid-Proterozoic rapakivi granites in general, several unanswered questions remain. Because of the comprehensive field, geochronologic, petrologic, and geophysical data available, the classic Fennoscandian rapakivis will provide a healthy ground for further investigations. Major issues that need to be thoroughly evaluated include: (1) Source characteristics (mantle vs. crust, mafic vs. felsic crustal sources); (2) Redox budget of the silicic magmas (reduced vs. oxidized rapakivi granites; cf. Frost and Frost, 1997; Kosunen et al., 2002; Rämö et al., 2002); (3) Tectonic framework (mantle plume vs. inboard models); and (4) Role of associated rift basins and subsequent continental flood basalt magmatism (Rämö et al., 2001, 2004; see also Chapter 13). Better handles on these issues will not only improve our knowledge about granite petrogenesis in general, but will also contribute to continental reconstructions, the Laurentia–Baltica connection in particular.
Acknowledgments We are indebted to the late Matti Vaasjoki for translating an early version of the manuscript from Finnish to English and to Brent A. Elliott (Alabama) and Stephen Frindt (Namibia) for assisting with the illustrations. Editorial comments from Martti Lehtinen are highly appreciated.
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kallioperä. Summary: Pre-Quaternary rocks of the Signilskär, Mariehamn, and Geta map sheet areas. Geological map of Finland 1 : 100 000, Explanation to the maps of PreQuaternary rocks. Geol. Surv. Finland, Espoo. 1–72. Bergman, L., 1986. Structure and mechanism of intrusion of postorogenic granites in the archipelago of southwestern Finland. Acta Acad. Aboensis, Ser. B, Nr. 46. 1–74. Dempster, T.J., Jenkin, G.R.T., Rogers, G., 1994. The origin of rapakivi texture. J. Petrol. 35, 963–981. DePaolo, D.J., 1981. Neodymium isotopes in the Colorado Front range and crust-mantle evolution in the Proterozoic. Nature 291, 193–196. DePaolo, D.J., Wasserburg, G.J., 1976. Nd isotopic variations and petrogenetic models. Geophys. Res. Lett. 3, 249–252. Edén, P., 1991. A specialized topaz-bearing rapakivi granite and associated mineralized greisen in the Ahvenisto complex, SE Finland. Bull. Geol. Soc. Finland 63, 25–40. Ehlers, C., Ehlers, M., 1977. Shearing and multiple intrusion in the diabases of Åland archipelago, SW Finland. Geol. Surv. Finland, Bull. 289, 1–31. Eklund, O., 1993. Coeval contrasting magmatism and magma mixing in Proterozoic post- and anorogenic granites, Åland, SW Finland. Åbo Akademis tryckeri, Åbo, 1–57. Eklund, O., Shebanov, A.D., 1999. The origin of rapakivi texture by sub-isothermal decompression. In: I. Haapala, O.T. Rämö (Eds.), Rapakivi granites and related rocks. Precambrian Res. 95, 129–146. Eklund, O., Fröjdö, S., Lindberg, B., 1994. Magma mixing, the petrogenetic link between anorthositic suites and rapakivi granites, Åland, SW Finland. Mineral. Petrol. 50, 3–19. Eklund, O., Shebanov, A., Fröjdö, S., Yli-Kyyny, K., Andersson, U.B., 1996. A low-foliated ignimbrite related to the Åland rapakivi granite in SW Finland. Terra Nova 8, 548–557. Elders, W.A., 1968. Mantled feldspars from the granites of Wisconsin. J. Geol. 76, 37–49. Elliott, B.A., 2001. Crystallization conditions of the
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Wiborg rapakivi batholith, SE Finland: an evaluation of amphibole and biotite mineral chemistry. Mineral. Petrol. 72, 305–324. Elo, S., Korja, A., 1993. Geophysical interpretation of the crustal and upper mantle structure in the Wiborg rapakivi granite area, southeastern Finland. In: R. Gorbatschev (Ed.), Baltic Shield, Special Volume, Precambrian Res. 64, 273–288. Elo, S., Rastas, J., Rämö, O.T., 1996. Gravity and aeromagnetic anomaly patterns of gabbro-anorthosites associated with rapakivi granites of southern Finland. In: I. Haapala, O.T. Rämö, P. Kosunen (Eds.), The Seventh International Symposium on Rapakivi Granites and Related Rocks, University of Helsinki, July 24–26, 1996, Abstract Volume. Helsinki University Press. 22–23. Frost, C.D., Frost, B.R., 1997. Reduced rapakivitype granites: The tholeiite connection. Geology 25, 647–650. Frost, B.R., Barnes, C.G., Collins, W.J., Arculus, R.J., Ellis, D.J., Frost, C.D., 2001. A geochemical classification for granitic rocks. J. Petrol. 42, 2033–2048. Haapala, I., 1977a. Petrography and geochemistry of the Eurajoki stock, a rapakivi-granite complex with greisen-type mineralization in southwestern Finland. Geol. Surv. Finland, Bull. 286, 1–128. Haapala, I., 1977b. The controls of tin and related mineralizations in the rapakivi-granite areas of south-eastern Fennoscandia. Geol. Fören. Stockholm Förhandl. 99, 130–142. Haapala, I., 1988. Metallogeny of the Proterozoic rapakivi granites of Finland. In: R.P. Taylor, D.F. Strong (Eds.), Recent advances in the geology of granite-related mineral deposits. Can. Inst. Mining Metall., Spec. Vol. 39, 124–132. Haapala, I., 1989. Suomen rapakivigraniiteista (English summary: Rapakivi granites of Finland). Acad. Scient. Fennica, Year Book 1988–1989, 135–140. Haapala, I., 1997. Magmatic and postmagmatic processes in tin-mineralized granites: topaz-bearing leucogranite in the Eurajoki rapakivi granite stock, Finland. J. Petrol. 38, 1645–1659.
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Haapala, I., Lukkari, S., 2005. Petrological and geochemical evolution of the Kymi stock, a topaz granite cupola within the Wiborg rapakivi batholith, Finland. In: O.T. Rämö (Ed.), Granitic Systems−Ilmari Haapala Volume. Lithos 80, 347−362. Haapala I., Ojanperä P., 1972. Genthelvite-bearing greisens in southern Finland. Geol. Surv. Finland, Bull. 259, 1–22. Haapala, I., Rämö, O.T., 1992. Tectonic setting and origin of the Proterozoic rapakivi granites of southeastern Fennoscandia. Trans. R. Soc. Edinburgh, Earth Sci. 83, 165–171. Haapala, I., Rämö, O.T., 1999. Rapakivi granites and related rocks: An introduction. In: I. Haapala, O.T. Rämö (Eds.), Rapakivi granites and related rocks. Precambrian Res. 95, 1–7. Haapala, I., Rämö, O.T., Frindt, S., 2005. Comparison of Proterozoic and Phanerozoic rift-related basaltic-granitic magmatism. In: O.T. Rämö (Ed.), Granitic Systems−Ilmari Haapala Volume. Lithos 80, 1−32. Hibbard, M.J., 1981. The magma mixing origin of mantled feldspars. Contrib. Mineral. Petrol. 76, 158–170. 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. Johanson, B.S., 1984. Ahvenisto gabro-anortositkomplex - en petrografisk och mineralogisk undersökning. M.Sc. Thesis, University of Helsinki, Finland. 1–85. (in Swedish) Kaartamo, K., 1996. Kymin stokin reunapegmatiittimuodostuman (Stockscheider) rakenteesta ja mineralogiasta. M.Sc. Thesis, University of Helsinki, Finland. 1–82. (in Finnish) Kanerva, I., 1928. Über das Rapakiwigebiet von Vehmaa im südwestlichen Finnland. Fennia 50 (N:o 40), 1–25. Koistinen, T. (Ed.), 1994. Precambrian basement of the Gulf of Finland and surrounding area, 1:1 mill. Geol. Surv. Finland, Espoo. Korja, A., 1995. Upper crust of the Baltic Profile, Finland. Department of Geophysics, University of Oulu, Report No. 19. 1–19. Korja, A., Elo, S., 1990. Crustal and upper mantle structure of the Wiborg batholith, SE Fin-
land. Abstracts of the second symposium on the Baltic shield held in Lund, Sweden, June 5–7, 1990, 57. Korja, A., Heikkinen, P.J., 1995. Proterozoic extensional tectonics of the central Fennoscandian Shield: Results from the Babel and Bothnian Echoes from the Lithosphere experiment. Tectonics 14, 504–517. Korja, A., Korja, T., Luosto, U., Heikkinen, P., 1993. Seismic and geoelectric evidence for collisional and extensional events in the Fennoscandian Shield – implications for Precambrian crustal evolution. Tectonophysics 219, 129–152. Korja, A., Heikkinen, P., Aaro, S., 2001. Crustal structure of the northern Baltic Sea palaeorift. Tectonophysics 331, 341–358. Korsman, K., Korja, T., Pajunen, M., Virransalo, P., GGT/SVE-KA Working Group, 1999. The GGT/SVEKA Transect: structure and evolution of the crust in the Paleoproterozoic Svecofennian Orogen in Finland. Int. Geol. Rev. 41, 287–333. Kosunen, P., 1999. The rapakivi granite plutons of Bodom and Obbnäs, southern Finland: petrography and geochemistry. Bull. Geol. Soc. Finland 71, 275–304. Kosunen, P., Rämö, O.T., Vaasjoki, M., 2002. The Bo dom and Obbnäs rapakivi granites, southern Finland: distinct composition imply a Paleoproterozoic terrane boundary. In: R. Lahtinen, A. Korja, K. Arhe, O. Eklund, S.-E. Hjelt, L.J. Pesonen (Eds.), Lithosphere 2002 – Second Symposium on the Structure, Composition and Evolution of the Lithosphere in Finland. Espoo, Otaniemi, November 12-13, 2002, Programme and Extended Abstracts. Institute of Seismology, University of Helsinki, Report S-42, 49-53. Kranck, E.H., 1929. Hoglands berggrund. Fältiakttagelser utförda sommaren 1927. Geol. Fören. Stockholm Förhandl. 51, 173–198. (in Swedish) Laitakari, I., 1969. On the set of olivine diabase dikes in Häme, Finland. Bull. Comm. géol. Finlande 241, 1–65. Laitakari, I., 1987. Hämeen subjotuninen diabaasijuoniparvi. The Subjotnian diabase dyke
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swarm of Häme. In: K. Aro, I. Laitakari (Eds.), Suomen diabaasit ja muut mafiset juonikivilajit. Diabases and other mafic dyke rocks in Finland. Geol. Surv. Finland, Rep. Invest. 76, 99–116. (in Finnish with English abstract and figure and table captions) Laitakari, I., Leino, H., 1989. A new model for the emplacement of the Häme diabase dyke swarm, central Finland. In: S. Autio (Ed.), Geological Survey of Finland. Current Research 1988. Geol. Surv. Finland, Spec. Pap. 10, 7–8. Lindberg, B., Bergman, L., 1993. Vehmaan karttaalueen kallioperä. Summary: Pre-Quaternary rocks of the Vehmaa map-sheet area. Geological map of Finland 1:100 000, Explanation to the maps of Pre-Quaternary rocks, Sheet 1042. Geol. Surv. Finland, Espoo, 1–56. Lindqvist, K., Laitakari, I., 1980. Glass and amygdules in Precambrian diabases from Orivesi, southern Finland. Bull. Geol. Soc. Finland 52, 221–229. Lukkari, S., 2002. Petrography and geochemistry of the topaz-bearing granite stocks in Artjärvi and Sääksjärvi, western margin of the Wiborg rapakivi granite batholith. Bull. Geol. Soc. Finland 74, 115–132. Luosto, U., 1997. Structure of the earth’s crust in Fennoscandia as revealed from refraction and wide-angle reflection studies. In: L.J. Pesonen (Ed.), The lithosphere in Finland – a geophysical perspective. Geophysica 33, 3–16. Luosto, U., Tiira, T., Korhonen, H., Azbel, I., Burmin, V., Buyanov, A., Kosminskaya, I., Ionkis, V., Sharov, N., 1990. Crust and upper mantle structure along the DSS Baltic profile in SE Finland. Geophys. J. Intern. 101, 89–110. Nekvasil, H., 1991. Ascent of felsic magma and formation of rapakivi. Am. Mineral. 76, 1279–1290. Neymark, L.A., Amelin, Yu.V., Larin, A.M., 1994. Pb-Nd-Sr isotopic constraints on the origin on the 1.54-1.56 Ga Salmi rapakivi granite-anorthosite batholith (Karelia, Russia). Mineral. Petrol. 50, 173–194.
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Patchett, P.J., Kouvo, O., 1986. Origin of continental crust of 1.9–1.7 Ga age. Nd isotopes and U-Pb zircon ages in the Svecokarelian terrain of South Finland. Contrib. Mineral. Petrol. 92, 1–12. Pearce, J.A., Harris, N.B.W., Tindle, A.G., 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. J. Petrol. 24, 956–983. Persson, A.I., 1999. Absolute (U–Pb) and relative age determinations of intrusive rocks in the Ragunda rapakivi complex, central Sweden. In: I. Haapala, O.T. Rämö (Eds.), Rapakivi granites and related rocks. Precambrian Res. 95, 109–127. Pihlaja, P., 1987. Porin seudun subjotuniset diabaasit. The Subjotnian diabases of the Pori region, southwestern Finland. K. Aro, I. Laitakari (Eds.), Suomen diabaasit ja muut mafiset juonikivilajit. Diabases and other mafic dyke rocks in Finland. Geol. Surv. Finland, Rep. Invest. 76, 133–150. (in Finnish with English abstract and figure and table captions) 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. Rämö, O.T., 2001. Isotopic composition of pyterlite in Vyborg (Viipuri), Wiborg batholith, Russia. Bull. Geol. Soc. Finland 73, 111–115. Rämö, O.T., Haapala, I., 1995. One hundred years of rapakivi granite. Mineral. Petrol. 52, 129–185. Rämö, O.T., Haapala, I., 1996. Rapakivi granite magmatism: A global review with emphasis on petrogenesis. In: D. Demaiffe (Ed.), Petrology and geochemistry of magmatic suites of rocks in the continental and oceanic crusts. A volume dedicated to Professor Jean Michot. Université Libre de Bruxelles, Royal Museum for Central Africa. 177–200. Rämö, O.T., Huhma, H., Kirs, J., 1996. Radiogenic isotopes of the Estonian and Latvian rapakivi granite suites: new data from the concealed Precambrian of the East European Craton. Precambrian Res. 79, 209–226.
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Rämö, O.T., Korja, A., Haapala, I., Eklund, O., Fröjdö, S., Vaasjoki, M., 2000. Evolution of the Fennoscandian lithosphere in the MidProterozoic: the rapakivi magmatism. In: L.J. Pesonen, A. Korja, S.-E. Hjelt (Eds.), Lithosphere 2000: a Symposium on the Structure, Composition and Evolution of the Lithosphere in Finland. Espoo, Otaniemi, October 4-5, 2000, Programme and Extended Abstracts. Institute of Seismology, University of Helsinki, Report S-41. 129–136. Rämö, O.T., Mänttäri, I., Vaasjoki, M., Upton, B.G.J., Sviridenko, L., 2001. Age and significance of Mesoproterozoic CFB magmatism, Lake Ladoga region, NW Russia. Geol. Soc. Am. Abstr. Progr. 33, A-139. Rämö, O.T., Dall’Agnol, R., Macambira, M.J.B., Leite, A.A.S., de Oliveira, D.C., 2002. 1.88 Ga oxidized A-type granites of the Rio Maria region, eastern Amazonian craton, Brazil: Positively anorogenic! J. Geol. 110, 603–610. Rämö, O.T., Mänttäri, I., Kohonen, J., Upton, B.G.J., Vaasjoki, M., Luttinen, A.V., Lindqvist, V., Lobaev, V., Cuney, M., Sviridenko, L.P., 2004. The Lake Ladoga basin; preliminary insights into geochronology, igneous evolution, and tectonic significance. In: C. Ehlers, O. Eklund, A. Korja, A. Kruuna, R. Lahtinen, L.J. Pesonen (Eds.), Lithosphere 2004 − Third Symposium on the Structure, Composition and Evolution of the Lithosphere in Finland. Institute of Seismology, University of Helsinki, Report S-45, 105−106. Ramsay, W., 1890. Om Hoglands geologiska byggnad. Geol. Fören. Stockholm Förhandl. 12, 471–490. Rieder, M., Haapala, I., Povondra, P., 1996. Mineralogy of dark mica from the Wiborg rapakivi batholith, southeastern Finland. Eur. J. Mineral. 8, 593–605. Salonsaari, P.T., 1995. Hybridization in the subvolcanic Jaala-Iitti complex and its petrogenetic relation to rapakivi granites and associated mafic rocks of southeastern Finland. Bull. Geol. Soc. Finland 67 (1b), 1–104.
Savolahti, A., 1956. The Ahvenisto massif in Finland. The age of the surrounding gabbroanorthosite complex and the crystallization of rapakivi. Bull. Comm. géol. Finlande 174, 1–96. Savolahti, A., 1962. The rapakivi problem and the rules of idiomorphism in minerals. Bull. Comm. géol. Finlande 204, 33–112. Sederholm, J.J., 1891. Ueber die finnländischen Rapakiwigesteine. Tscherm. Mineral. Petrograph. Mitth. 12, 1–31. Sederholm, J.J., 1928. On orbicular granites, spotted and nodular granites etc. and on the rapakivi texture. Bull. Comm. géol. Finlande 83, 1–105. Siivola, J., 1987. Lovasjärven mafinen intruusio. The mafic intrusion of Lovasjärvi. In: K. Aro, I. Laitakari (Eds.), Suomen diabaasit ja muut mafiset juonikivilajit. Diabases and other mafic dyke rocks in Finland. Geol. Surv. Finland, Rep. Invest. 76, 121–128. (in Finnish with English abstract and figure and table captions) Simonen, A., 1987. Kaakkois-Suomen rapakivimassiivin kartta-alueiden kallioperä. Summary: Pre-Quaternary rocks of the rapakivi massif in southeastern Finland. Geological map of Finland 1:100 000, Explanation to the maps of Pre-Quaternary rocks, Sheets 3023+3014, 3024, 3041, 3042, 3044, 3113, 3131, and 3133. Geol. Surv. Finland, 1–49. Simonen, A., Vorma, A., 1969. Amphibole and biotite from rapakivi. Bull. Comm. géol. Finlande 238, 1–28. Stimac, J.A., Wark, D.A., 1992. Plagioclase mantles on sanidine in silicic lavas, Clear Lake, California. Implications for the origin of the rapakivi texture. Geol. Soc. Am. Bull. 104, 728–744. Suominen, V., 1987. Lounais-Suomen mafiset juonikivet. Mafic dyke rocks in southwestern Finland. In: K. Aro, I. Laitakari (Eds.), Suomen diabaasit ja muut mafiset juonikivilajit. Diabases and other mafic dyke rocks in Finland. Geol. Surv. Finland, Rep. Invest. 76, 151–172. (in Finnish with English abstract and figure and table captions) Suominen, V., 1991. The chronostratigraphy of
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southwestern Finland with special reference to Postjotnian and Subjotnian diabases. Geol. Surv. Finland, Bull. 356, 1–100. Törnroos, R., 1984. Petrography, mineral chemistry and petrochemistry of granite porphyry dykes from Sibbo, southern Finland. Geol. Surv. Finland, Bull. 326, 1–43. Turekian, K.K., Wedepohl, K.H., 1961. Distribution of elements in some major units of the Earth’s crust. Geol. Soc. Am. Bull. 72, 175–192. Vaasjoki, M., 1996. Explanation to the geochronological map of southern Finland: The development of the continental crust with special reference to the Svecofennian orogeny. Geol. Surv. Finland, Rep. Invest. 135, 1–30. Vaasjoki, M., Sakko, M., 1989. The radiometric age of the Virmaila diabase dyke: Evidence for 20 Ma of continental rifting in Padasjoki, southern Finland. In: S. Autio (Ed.), Geological Survey of Finland. Current Research 1988. Geol. Surv. Finland, Spec. Pap. 10, 43–44. Vaasjoki, M., Rämö, O.T., Sakko, M., 1991. New U-Pb ages from the Wiborg rapakivi area: constraints on the temporal evolution of the
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rapakivi granite–anorthosite–diabase dyke association of southeastern Finland. In: I. Haapala, K.C. Condie (Eds.), Precambrian granitoids. Petrogenesis, geochemistry and metallogeny. Special Issue, Precambrian Res. 51, 227–243. Vorma, A., 1975. On two roof pendants in the Wiborg rapakivi massif, southeastern Finland. Geol. Surv. Finland, Bull. 272, 1–86. Vorma, A., 1976. On the petrochemistry of rapakivi granites with special reference to the Laitila massif, southwestern Finland. Geol. Surv. Finland, Bull. 285, 1–98. Vorma, A., 1980. The Wiborg rapakivi massif. In: K. Hytönen (Ed.), Precambrian bedrock of southern and eastern Finland, Guide to excursions 001 A+C, 26th International Geological Congress Paris 1980. Geol. Surv. Finland. 6–8. Wahl, W., 1938. Några iakttagelser från Wiborgsrapakiviområdets södra gränsgebiet. Geol. Fören. Stockholm Förhandl. 60, 88–96. (in Swedish) Whalen, J.B., Currie, K.L., Chappell, B.W., 1987. A-type granites; Geochemical characteristics, discrimination, and petrogenesis. Contrib. Mineral. Petrol. 95, 407–419.
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Chapter 13
SEDIMENTARY ROCKS, DIABASES, AND LATE CRATONIC EVOLUTION
J. Kohonen, O.T. Rämö 563
Cover page: Mesoproterozoic fluvial sandstone with dendrite. Panelia, Kiukainen, southwestern Finland. Photo: Erkki Halme.
Kohonen, J., Rämö, O.T., 2005. Sedimentary rocks, diabases, and late cratonic evolution. 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. 563–604. © 2005 Elsevier B.V. All rights reserved.
The late evolution of the shield area is envisioned as an overall dynamic process, in part linked to the Sveconorwegian and Caledonian orogenies, and is divided into six phases, each with a characteristic rock record and tectonic significance: (1) The intracratonic rift basin stage (~1600–1300 Ma). After the emplacement of the locus classicus ~1.67–1.54 Ga rapakivi granites and related rocks, thick fluvial deposits began to fill the developing intracratonic rift basins. These rock sequences (e.g., the Satakunta and Muhos Fms.) are preserved in tectonic depressions and graben structures delineated by reactivated fracture zones. (2) Crustal extension episodes and the Sveconorwegian orogeny (~1300–900 Ma). ~1265 and 1100–1000 Ma basaltic dikes record minor juvenile additions to the shield. These ages correspond to the initial extension and final closing stages, respectively, of the Sveconorwegian orogeny in the southwest. (3) The Neoproterozoic exhumation stage (~900–600 Ma). In the late Precambrian, the crystalline basement became exposed over large areas in northern Europe. Very little is known, however, about the timing and causes of this uplift and exhumation. (4) The stage of platform sedimentation (~600–420 Ma). In response to the opening of the Iapetus Ocean, a shallow marine continental margin was created. Siliciclastic sediments (mainly Cambrian) and platform carbonates (mainly Ordovician) are preserved in the Bothnian Bay area, while minor cover remnants (e.g., the Lauhanvuori and Hailuoto Fms. and the Dividal Group) and other scattered indications (e.g., clastic dikes) are found in the mainland. The Finnish kimberlites were emplaced during the onset of this platformal stage. (5) The Caledonian foreland stage (~420–350 Ma). Minor rock units belonging to the Scandinavian Caledonides are present in the far northwest part of the country. Recent fission track studies indicate that mainly Devonian Caledonian foreland sediments once also covered other parts of Finland. However, none of these strata have survived. The Sokli and Iivaara alkaline complexes were formed at this stage. (6) Final exhumation and birth of the shield (~350–5 Ma). The final exhumation of the shield is not well constrained; several distinct phases can, however, be distinguished. In the late Paleozoic and early Mesozoic (at ~200 Ma), denudation and isostatic uplift of the Caledonides resulted in almost total peneplanation. The mainly Paleozoic sedimentary cover of Finland was probably removed only during the Paleogene to Neogene uplift of western Scandinavia. CHAPTER
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1. Introduction In Finland, the emplacement of the classic 1.65–1.54 Ga rapakivi granites was the latest major crustal increment to the Fennoscandian Shield. In the later Mesoproterozoic as well as the Neoproterozoic and Paleozoic eras, both the supracrustal record and evidence for igneous activity are sparse, and Mesozoic and Cenozoic record is virtually absent (Figure 13.1). In this chapter, the unmetamorphic sedimentary cover, igneous rocks younger than 1500 Ma, and the Finnish Caledonides are reviewed. First we describe these rock units and their geologic settings, and thereafter propose a model for the tectonic and basin evolution for Mesoproterozoic through Cenozoic. The Meso- to Neoproterozoic and Paleozoic supracrustal sequences in Finland are local remnants resting atop the Archean–Paleoproterozoic crystalline basement. All these supracrustal units consist of unmetamorphosed, subhorizontally layered strata, which define a prominent angular unconformity towards their regionally metamorphosed and multiply folded basement. The most important Mesoproterozoic magmatic event after 1500 Ma was the formation of diabase sills and dikes in southwestern Finland (Figure 13.2). These form the eastern part of a ~1270 Ma mafic igneous province that also comprises much of central Sweden. In the Lake Ladoga region of Russian Karelia, however, a clearly older system of diabase sills is recognized, whereas in northern Finland scattered 1100–1000 Ma mafic dikes are found. In the far northwest part of the country, the allochthonous units of the Finnish Caledonides include ~435 Ma gabbroic and ultramafic rocks. A separate chapter in the present volume is devoted to the ~600 Ma kimberlites and Devonian alkaline intrusions in northeastern Finland. In most previously published reviews and map compilations, the Mesoproterozoic sedimentary rocks have been regarded as part of 566
the Fennoscandian shield per se, whereas the late Neoproterozoic (“Vendian”) sedimentary units have been associated with the Paleozoic platform cover sequence (e.g., Laitakari et al., 1996; Koistinen et al., 2001). For this reason, the Mesoproterozoic units are normally included in the “bedrock”. In order to make a difference between the “bedrock” and the depositional basement of the Mesoproterozoic sediments, “crystalline basement” is here employed to describe the latter, while “cover sequence” is used for all unmetamorphosed supracrustal rocks regardless of age. In Finland, the general chronostratigraphic division of the Proterozoic cover sequence has followed the traditional (e.g., van Eysinga, 1975) terminology scheme employing the concepts of “Riphean” (1600–650 Ma) and “Ven dian” (650–570 Ma). The term “Eocambrian” has been used especially by microfossil specialists in a sense more or less synonymous with “Vendian”. In addition, such traditional terms as “Jotnian”, “Subjotnian”, and “Postjotnian” are still widely adhered to. Originally, the term “Jotnian” was introduced by J.J. Sederholm (1897) for unmetamorphic sandstones and associated igneous rocks in the Precambrian of Fennoscandia. Currently the term is poorly defined and only used informally for certain ~1500–1300 Ma sedimentary units. “Subjotnian” and “Postjotnian” refer to diabases regarded, respectively, as older and younger than these sediments (e.g., Laitakari et al., 1996). To avoid confusion, the IUGS (2000) recommendations have been employed in this paper. Accordingly, the Mesoproterozoic (1600–1000 Ma) and Neoproterozoic (1000–540 Ma) eras have the following subdivisions: • MP1 (1600–1400 Ma), MP2 (1400–1200 Ma), MP3 (1200–1000 Ma), and • NP1 (1000–850 Ma), NP2 (850–650 Ma), NP3 (650–540 Ma).
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
SVECOFENNIAN OROGENY
FINLAND
Rapakivi granites/ Mafic intrusions Mafic dikes
Kimberlites
Intrusions Supracrustal rocks 3200 Ma
2500 Ma
1600 Ma
ARCHEAN
1000 Ma
540 Ma
PROTEROZOIC Paleoproterozoic
Alkaline intrusions
PHANEROZOIC
Mesoproterozoic Neoproterozoic
Mafic dikes
FENNOSCANDIA IN GENERAL
GOTHIAN OROGENY SVECOFENNIAN OROGENY
SVECONORWEGIAN OROGENY
CALEDONIAN OROGENY
THIS CHAPTER
Fig. 13.1. Geological time scale with a sketch diagram of igneous activity and the presence of supracrustal rocks. Major orogenic events are also indicated.
As implicit in the IUGS (2000) scheme, the term “Tertiary” has been discarded, “Paleogene” and “Neogene” being used instead. All the other chronostratigraphic terms are informal and are marked by italics. At present there exists no comprehensive lithostratigraphic code for Mesoproterozoic and younger cover of Finland, the employed litostratigraphic nomenclature having been adapted mainly from Kousa and Lundqvist (2000).
2. Mesoproterozoic sedimentary sequences 2.1. Regional setting The best known Mesoproterozoic sedimentary rock units in Finland are the sandstones of the Satakunta Formation and the Muhos Formation. Both have submarine extensions beneath the Gulf of Bothnia (Winterhalter et al., 1981; cf. Figure 13.2). A sea-bottom occurrence area is also known from the Åland Sea southwest of the Åland Islands. Mesoproterozoic (“Jotnian”) sedimenCHAPTER
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tary sequences including arkose, siltstone, shale, and conglomerate have been described from some ten localities in the Fennoscandian Shield. All these occupy shallow basins, tectonic depressions or grabens bordered by fractures or fault zones many of which trend northwest. The largest occurrence is in central Sweden where the Dala (Dalarna) sandstones cover an area of about 50 km by 150 km. These dominantly continental sediments (e.g., Lundqvist, 1979, and references therein; AlDahan, 1985) are preserved in the core of an open, N–S-oriented syncline. The Nordingrå (Welin and Lundqvist, 1984) and Gävle sandstone areas are seemingly tied to the Satakunta sandstone in Finland (Figure 13.3). Other Swedish occurrences described as “Jotnian” are in the Svartälven, Mälaren, and Almesåkra farther south. In Russia, the Mesoproterozoic supracrustal units cluster in the Lake Ladoga–Salmi and White Sea–Tersk regions and along the northernmost edge of the Kola Peninsula (Figure 13.3). Many of these cover sequences are spatially associated with the rapakivi granites and the 1670–1550 Ma (“Subjotnian”) mafic ROCKS,
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NORWAY Kautokeino
Laanila
Sokli CENTRAL LAPLAND Siurunmaa Savukoski Naruskajärvi Pelkosenniemi
Akanvaara Salla
Kuusamo Iivaara Saarijärvi Hailuoto Bothnian Bay
RUSSIA
Oulu MUHOS
SWEDEN
KAINUU
oth
nia
Kuhmo
Gu
lf o
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FINLAND
Nordingrå
Vaasa
Lappajärvi
Söderfjärden
Lauhanvuori
Karstula
Karikkoselkä
Suvasvesi Iso-Naakkima Paasselkä Virtasalmi
Bothnian Sea Pori
Sääksjärvi
Salmi
Finngrundet Gävle
Västra banken
SATAKUNTA
Lake Ladoga
Lumparn Åland
Sea
inland
Gulf of F
0
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Uljaste
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
dikes and also with the ~1265 Ma (“Postjotnian”) diabases. In Finland, the latter define the minimum deposition age of the Satakunta sandstones. In the Lake Ladoga region, a sill in the upper part of the local sedimentary cover sequence (Amantov et al., 1996) has a U-Pb age of ~1460 Ma (Rämö et al., 2001). This indicates that, at least in that region, basin formation commenced relatively rapidly after – or maybe even during – the emplacement of the Ladoga–Salmi rapakivi granite plutons at ~1560–1530 Ma.
2.2. The Satakunta Formation and its submarine extensions The Mesoproterozoic bedrock in western Finland is dominated by 1580–1540 Ma rapakivi granites and associated mafic rocks (Vaasjoki, 1977; Suominen, 1991), the Satakunta sandstone, and olivine tholeiitic diabase dikes (Figure 13.4). In Satakunta, “Subjotnian” diabase dikes are considered to be older than the rapakivi granites of the region (~1650 Ma; Pihlaja, 1987). The Satakunta sandstone covers a north-
west elongated, fault-bounded area about 15 km by 100 km in size. The only direct evidence regarding its minimum thickness comes from a drill hole south of the town of Pori; this penetrated 591 m of sandstone without reaching its base. According to gravimetric surveys (Elo, 1976, 1982), the maximum thickness of the Satakunta Formation may be more than 1500 m. Along its southwestern edge, the sandstone forms a step-like structure (Laurén, 1970), gradually thickening towards the northeast. It has a subvertical, faulted contact against the Svecofennian basement in the northeastern. A few inliers of Svecofennian rocks in the fringe area probably represent fault blocks and irregular basement along the edge of an ancient rift valley. The Satakunta Formation exhibits grain sizes varying from conglomeratic to siltstone and mudstone, but various kinds of sandstone form the bulk of these strata. In outcrop, coarse-grained, purplish sandstones prevail, but the deep drill core consist dominantly of medium-grained sandstone, thin mudstone interlayers, and coarse sandstone in its lowermost and uppermost parts (Kohonen et al.,
Alkaline intrusions ~360 Ma
Mafic dikes ~1.1–1.0 Ga
Sedimentary rocks (PZ)
Mafic dikes ~1.46–1.22 Ga
Allochthonous Caledonian rocks
Mafic dikes ~1.67–1.47 Ga
Sedimentary rocks (NP3)
Impact site
Mafic intrusions ~1.46–1.22 Ga
Sandstone dike province
Sedimentary rocks (MP1–MP2)
Kimberlite province
Rapakivi granites ~1.65–1.54 Ga Crystalline basement (Paleoproterozoic) Crystalline basement (Archean) Fig. 13.2. Simplified geological map of Finland highlighting the Mesoproterozoic to Phanerozoic rock units. Map data mainly according to the Geological Map of the Fennoscandian Shield (Koistinen et al., 2001; Pesonen et al., 2000). The boxes refer to Figures 13.4 and 13.14. Localities mentioned in the text are indicated with black dots. CHAPTER
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Kola Peninsula Tersk
fro n
t
SWEDEN
led on
ian
White Sea
oth
nia
Ca
Bothnian Bay
RUSSIA
lf o
Gu
Nordingrå
fB
NORWAY Dala
FINLAND
front
Bothnian Sea
ui
st
lin
Lake Ladoga
Mälaren
ESTONIA Almesåkra
Sveco
e
norw
nq
Gävle Åland
egian
To r
Svartälven
Baltic Sea
0
200 km
Mafic dikes (NP3 to Cambrian)
Mafic dikes (~1.67–1.47 Ga)
Mafic dikes (~1.1–1.0 Ga)
Sedimentary cover (NP3 and PZ)
Mafic dikes (~1.46–1.22 Ga)
Sedimentary cover (MP1–MP2)
Fig. 13.3. Distribution of Mesoproterozoic to Cambrian sedimentary rocks and mafic dike swarms in the central Fennoscandian Shield to the east of Sveconorwegian and Caledonian tectonic fronts. Dikes and other rock units simplified from Koistinen et al. (2001) and Mertanen et al. (1996a).
1993). Basal quartz-pebble conglomerates have been observed along the southwestern contact of the formation. Lithologically, the sandstones are, in order of abundance, arkoses, subarkoses, and quartz 570
arenites. Petrographical studies (Simonen and Kouvo, 1955; Marttila, 1969) indicate that the detritus had been derived from the Svecofennian basement, not from the adjacent rapakivi granites. Also, an isotope study of detrital zir-
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
Reposaari
Pori Sääksjärvi
Vammala Luvia
Harjavalta
Kokemäki BOTHNIAN SEA Eurajoki Rauma
Säkylä Pyhäjärvi
Laitila
20 km
Uusikaupunki
Sandstone
Crystalline basement; supracrustal/igneous
Diabase
Impact structure
Rapakivi granite Diabase dikes
Paleocurrent direction
Fig. 13.4. Map of the Satakunta sandstone area in southwestern Finland. The paleocurrent directions are according to Kohonen et al. (1993). For location, see Figure 13.2.
cons and monazite (Vaasjoki and Sakko, 1987) excluded rapakivi granites as a major source and indicated a Svecofennian provenance. Typical sedimentary structures in the Satakunta sandstone include planar and trough CHAPTER
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cross-bedding, ripple marks, mud cracks, clay galls, and raindrop imprints (Figure 13.5). All the reported primary features and facies associations are in accord with a fluvial depositional environment (e.g., Kohonen et al., ROCKS,
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B
C A Fig. 13.5. Primary structures in the Satakunta sandstone. (A) Ripple marks at Knapernummi, Luvia; (B) Cross-bedding at Lammaistenkoski power station Harjavalta; (C) Rain drop imprints. Metsäkulma, Pori; (D) Desiccation cracks in sandstone. Power station Harjavalta. The length of index plate is 16 cm. Photos: Hannu Kujala. D
1993). However, the presented fluvial models are based mainly on outcrops representing the uppermost part of the preserved rock sequence – in contrast, a major part of the “Jotnian” Dala sandstone in Sweden has been interpreted as eolian in origin (Pulvertaft, 1985). Observed paleocurrent patterns and grain size distribution indicate that during the deposition of the presently exposed, uppermost sediments, the distal part of the basin was situated in the northwest (Kohonen et al., 1993). 572
Nevertheless, the filling of the Satakunta basin may have had a long and complex history. The major trend of both of the mafic dike swarms (~1650 Ma and ~1265 Ma) is subparallel to the axis of the Satakunta basin (see Figures 13.2 and 13.4), but there is no conclusive evidence in regard to the age relationship between the older dikes and the sandstone. It would thus appear that the upper part of the sandstone was deposited ~1400–1300 Ma ago (e.g., Simonen,1960; Neuvonen, 1973; Pesonen
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
et al., 1989), but it may still be possible that the basin history extends all the way back to initial rifting at ~1650 Ma. The final preservation and the present margins of the sandstone were plausibly controlled by local subsidence related to the emplacement of basaltic magma at ~1265 Ma. According to Winterhalter et al. (1981), the Satakunta sandstone continues into the Bothnian Sea and links together the Mesoproterozoic sandstone areas of the Satakunta graben, Gävle Valley, and the Nordingrå area, the latter two in Sweden (Figure 13.3). The areal extent of the submarine sandstones has been interpreted from seismic profiles. The interpretation has been confirmed by the presence of red arkosic sandstone (Winterhalter, 1972). The thickness of the sandstone formation probably exceeds 1000 m (cf. Axberg, 1980), but a thickness of no less than 3 km to 4 km has been suggested for the Mesoproterozoic sedimentary rocks in the northern part of the Bothnian Sea (Korja et al., 2001). The Åland Sea depression contains a separate submarine occurrence of the rocks of the sedimentary cover (Winterhalter et al., 1981; Söderberg, 1993). A major part of that sequence probably consists of Mesoproterozoic arkosic sandstone with a maximum thickness ~1200 m (Söderberg, 1993).
2.3. The Muhos Formation and its submarine extensions The Muhos Formation occupies a fault-bounded, SE-trending basin south of Oulu, extending ~50 km inland from the coast (Figure 13.2). Its rocks are very poorly exposed and observations have mainly been made in drill cores. The age estimate for the Muhos Formation is poor and the assumption of a Mesoproterozoic age is based on a rather old K-Ar dating (~1300 Ma for the diagenesis of shale; cf. Simonen, 1960) and microfossil data (Tynni, 1978; Tynni and Uutela, 1984). A thickness of 895 m was observed in diamond drilling (cf. Kousa and CHAPTER
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Lundqvist, 2000). The typical rocks of the Muhos Formation are interlayered reddish siltstones and mudstones. Interbeds of arkosic sandstone are present throughout the sequence, but otherwise sandstones become more abundant and coarse-grained only towards the base of the formation. In the type section at Muhos township (a 527m drill core), the lowermost 20 m consist of conglomerate interbedded with arkosic sandstone. Lithostrathigraphic sections and more detailed rock descriptions can be found in Simonen and Kouvo (1955). Cross-bedding is common in the sandstones of the Muhos Formation, and a floodplain environment of deposition was advocated by the pioneer researchers (Simonen and Kouvo, 1955). While there are no recent detailed sedimentological investigations, the overall nature of the sequence appears to be compatible with an alluvial plain or a floodplain depositional model. A special rock type exposed in the only outcrop representing the Muhos Formation is the Kieksi conglomerate at the northern edge of the area. The conglomerate consists of subangular pebbles of granite, gneiss, and schist in a sandy matrix with some carbonate cement (Kesola, 1985). The stratigraphic position of the conglomerate is not obvious, but it has been correlated with the basal units of the type section (Brenner, 1944). Red arkosic sandstones, mudstones, and conglomerates (cf. Veltheim, 1969), tentatively correlated with the Muhos Formation, are also observed in drillings at the Hailuoto Island near the town of Oulu (Figure 13.2). Here, the Muhos Formation is overlain by the Neoproterozoic (NP3) Hailuoto Formation (see below). According to seismic data, a major part of the bottom of the Bothnian Bay is occupied by sedimentary rocks (Winterhalter et al., 1981). The distribution of the Mesoproterozoic sediments in the Bothnian Bay area has been summarized by Winterhalter (2000).
ROCKS,
DIABASES, AND...
•
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2.4. Minor occurrences The Lappajärvi impact crater (Figure 13.2) is surrounded by an angular graben representing the terrace zone of the structure (Pesonen et al., 2000). The age of the impact is ~73 Ma (Mänttäri and Koivisto, 2001). In places, remnants of a pre-impact sedimentary cover have been preserved together with down-faulted blocks of crystalline basement (Pipping and Lehtinen, 1992; Vaarma and Pipping, 1997). The presence of these sedimentary rocks was disclosed by the Lappajärvi drilling program. The thickness of the observed siltstone–sandstone sequence (the Pokela Formation) is around 18 m, these sedimentary rocks being underlain by strongly weathered Paleoproterozoic schist. According to microfossil studies (Uutela, 1990) the depositional age plausibly corresponds to that of the Muhos Formation. The abundance and distribution of glacial sandstone boulders resembling the Satakunta sandstone have inspired many authors to speculate on the presence of unexposed Mesoproterozoic rock occurrences (cf. Salonen, 1991). Such remnants may well be preserved, but none have been verified so far.
3. Mesoproterozoic igneous rocks 3.1. Introduction Late Mesoproterozoic igneous rocks in the Finnish part of the Fennoscandian Shield are scarce and comprise ~1265 Ma to 1000 Ma mafic dikes and sills in the southwestern and northern parts of the country. These hypabyssal rocks, the older ~1265 Ma dikes and sills in particular, have counterparts in Sweden and Norway. Comparable, but older, lithologic units are also found in the Lake Ladoga region of Russian Karelia. Here we describe the mode of occurrence, petrography, and geochemistry of the ~1265 Ma mafic hypabyssal rocks of southwestern Finland and the 1100–1000 574
Ma mafic dikes in the northern part of the country. We also compare the “Postjotnian” ~1265 Ma dikes to the corresponding dikes of central Sweden and to the ~1460 Ma mafic sill exposed in the Lake Ladoga basin.
3.2. The ~1265 Ma magmatism Regional setting The ~1265 Ma (“Postjotnian”) hypabyssal rocks of southwestern Finland have been described in several reconnaissance-type works (e.g., Aro and Laitakari, 1987; Amantov et al., 1996) and unpublished theses (e.g., Inkinen, 1963; Hämäläinen, 1985). These mafic rocks are found in large amounts in the Satakunta region of southwestern Finland (Figure 13.2) where they comprise extensive, vertical to subhorizontal dikes that cut all the other rock units of that region, i.e., the Svecofennian granitoids and metamorphic rocks, the rapakivi granites, and the Satakunta sandstone. Apart from Satakunta, mafic hypabyssal rocks of this age group are found in the Åland region in the far southwest part of the country (Hausen, 1964; Bergman, 1979) and in the Vaasa region farther to the north (Ervamaa, 1962; Aro, 1987); most of the latter are hidden beneath the sea. The crystallization age of the mafic dikes is well-constrained. Suominen (1991, cf. also Vaasjoki, 1996) reported U-Pb zircon/baddeleyite ages from eight samples in four different localities (Figure 13.6), with upper intercept ages averaging 1265 ± 7 Ma (2σ). Thus the dikes probably represent a relatively short igneous episode; such conditions are, in general, typical of Phanerozoic continental flood basalt associations and Precambrian dike swarms that are considered to represent the feeder dikes for flood basalts (Basaltic Volcanism Study Project, 1981). The age established for the “Postjotnian” dikes of southwestern Finland has also been assumed as the crystallization age of corresponding, mostly lopolithic intrusions in central Sweden (Figure 13.6; Patchett
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
+2.0, +2.7 +2.7
+1.6
+2.2 1268± 13 Ma
+3.3
FINLAND
+3.7
Valamo diabase: εNd (at 1460 Ma) –9
+3.0 +3.3
SALMI
1264±12 Ma
Bothnian Sea
+3.3 +2.3 1258± 13 Ma
+1.6
Lake Ladoga WIBORG
+2.2
ÅLAND 60°
+3.5 1265± 6 Ma
RUSSIA 100 km
Gulf of Finland
SWEDEN
30°
ESTONIA Caledonides
Paleoproterozoic rocks
Mesoproterozoic diabase
Archean rocks
Mesoproterozoic cover rocks
+3.3
Rapakivi granite complexes
1265±6 Ma
εNd (at 1265 Ma) value U-Pb age
Fig. 13.6. Map showing the distribution of the Mesoproterozoic diabase dikes, cratonic sedimentary cover, and rapakivi granite intrusions in southern Finland and surrounding areas. Initial εNd values of the diabase dikes (Rämö, 1990; Patchett et al., 1994; this work) are indicated, as are the U-Pb zircon/ baddeleyite ages for the Finnish dikes (Suominen, 1991).
et al., 1994). In addition, a lamproite dike with a comparable age (~1250 Ma; U-Pb on perovskite) has been described from the Kuhmo area in eastern Finland (Peltonen et al., 2000).
Petrography and geochemistry The “Postjotnian” dikes and sills of southwestern Finland are typically subophitic, equigranular, medium- to coarse-grained, homogeneous rocks with calcic plagioclase, intermediate olivine, augite, and Fe-Ti oxides as the main minerals. The individual subhorizontal dikes can be several tens of meters thick, some of them containing megaophitic parts (Figure 13.7). The latter are geochronologically valuable as they often contain ample CHAPTER
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amounts magmatic zircon for U-Pb isotope analysis (e.g., Suominen, 1991). Geochemically, the ~1265 Ma rocks are transitional tholeiites and are relatively evolved. Their SiO2 and MgO contents range from 45 to 48 wt.% and 3.0 to 7.7 wt.%, respectively, and Mg numbers are intermediate to low (~60 to 40; Hämäläinen, 1987; Rämö, 1990; Upton et al., 1998). The trace element patterns of the dikes suggest OIB affinity, save for prominent negative Th and Nb spikes in mantle-normalized diagrams (Figure 13.8). Nd isotope data are available for four localities in Satakunta (Rämö, 1990). For this work, we also analyzed two additional dikes, one from the Åland Islands (Märket) in the far ROCKS,
DIABASES, AND...
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575
1000
~1265 Ma diabases, southwestern Finland
MORB EMORB OIB IAT
Mantle normalized
100
10
Rb Ba Th K Nb La Ce Sr Nd P Sm Zr Ti Y
Fig. 13.8. Multielement diagram showing the composition of the ~1265 Ma diabases from the Satakunta region, southwestern Finland (data from Rämö, 1990). The normalizing values and compositions of mid-ocean ridge basalt (MORB), enriched mid-ocean ridge basalt (EMORB), oceanic-island basalt (OIB), and island arc tholeiite (IAT) are from Sun (1980).
Source characteristics and magmatic evolution Fig. 13.7. ~1265 Ma tholeiitic diabase from the Island of Säppi, offshore Pori, southwestern Finland. Medium- to coarse-grained diabase shows to the left, megaophitic diabase to the right. Photo: Pekka Pihlaja.
southwest part of the country and the other from the Vaasa archipelago (Korsnäs) somewhat farther north (Figures 13.2 and 13.6); these data are shown in Table 13.1. All of the six analyzed dikes have relatively radiogenic Nd isotope compositions. The Satakunta samples have εNd (at 1265 Ma) values between +1.3 and +3.3 and those of the Märket and Korsnäs diabases are +3.5 and +2.2, respectively. The depleted mantle model ages (DePaolo, 1981) range from 1.71 to 1.44 Ga. Whole-rock and feldspar Pb isotope data on the Satakunta dikes indicate relatively unradiogenic compositions, intermediate between the growth curves for crustal and mantle lead (Rämö, 1990). 576
The “Postjotnian” diabases of southwestern Finland were crystallized from relatively evolved basaltic magmas that obviously had undergone a long fractionation history before reaching their present level in the crust. Thus they may not provide a direct indication of the composition of the Mesoproterozoic subcontinental mantle in the region. The initial εNd values are, however, clearly positive and indicate that the magmas were originally derived from at least mildly LREE-depleted sources. Nevertheless, the dikes are less radiogenic (εNd ~ +2 to +3) than the contemporaneous depleted mantle of the DePaolo (1981) model (εNd +5; Figure 13.9). Whether the lower εNd values of the dikes are due to crustal contamination associated with low-pressure fractional crystallization (cf. Rämö, 1990) or reflect interaction of lithospheric and asthenospheric mantle domains (cf. Patchett et al., 1994), is yet to be determined. In this respect, it is interesting to examine
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
Table 13.1. Nd isotope data for two Mesoproterozoic (1265 Ma) diabases, southwestern Finland. Sample
Location
Sm
(Map sheet; Grid coordinates) (ppm) Märket, Eckerö, Åland archipelago A562 0043 07; 6688.20–0562.80
17.03
Nd
147
(ppm)
144
73.07
0.1409
Sm/ a
Nd
143
Nd/
144
εNdc
b
Nd
0.512354 ± 7
TDMd (Ma)
+ 3.5
1436
Norrgrynnan, Korsnäs,Vaasa archipelago A732 1242 03; 6978.47–1500.72 8.07 32.57 0.1498 0.512362 ± 9 + 2.2 1615 a Estimated error for 147Sm/144Nd is less than 0.5%. b 143 Nd/144Nd normalized to 146Nd/144Nd = 0.7219. Within-run error expressed as 2σ in the least significant digits. c Initial εNd values, calculated using 143Nd/144Nd = 0.512638 and 147Sm/144Nd = 0.1966. Maximum error is ± 0.40 ε-units. d Depleted mantle model ages according to the model of DePaolo (1981). Description of analytical methods: Analyses were performed by O.T. Rämö at the Unit for Isotope Geology, Geological Survey of Finland. Rock powders (~200 mg) were dissolved in Teflon beakers at 180 oC in a mixture of HNO3 and HF, dissolved in HCl, and spiked with a 149Sm-150Nd tracer. Light REE were separated using standard cation exchange chromatography and Sm and Nd were purified by and on quartz columns (Richard et al., 1976). The total procedural blank was <300 pg for Nd. Isotope ratios of Sm and Nd were measured on a VG Sector 54 mass spectrometer. Repeated analyses of La Jolla Nd standard gave 143Nd/144Nd of 0.511845 ± 0.000009 (mean and external 2σ error of 14 measurements); external error in the reported 143Nd/144Nd is thus better than 0.002 %.
the isotope composition of the samples analyzed from the entire ~1.26 Ga dike suite of southwestern Finland and north-central Sweden (Figure 13.6). The dikes in central Sweden are quite radiogenic (εNd +1.6 to +3.7) and this is also true for the dikes of southwestern Finland (εNd +1.6 to +3.5). As the diabases of north-central Sweden are on the radiogenic side, this rules out crustal contamination as a major factor in the petrogenesis of the dikes, because this particular section of the Swedish bedrock is presumably underlain by a major Archean source component in the lower crust (Andersson et al., 2002; see also Chapter 12). Conversely, the ~1460 Ma diabase sills intercalated with the sedimentary rocks of the Lake Ladoga basin (Figure 13.6) show a clear Archean source component with initial εNd values of –9.2 to –8.6 and depleted mantle model ages of 2.58 to 2.55 Ga (Figure 13.9; Upton et al., 1998; Rämö et al., 2001, 2004). The Lake Ladoga basin was developed just south of the Archean Karelian craton (Figure 13.6) and the sills probably tapped a Neoarchean enriched lithospheric mantle source (Rämö et al., 2001).
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3.3. The 1100–1000 Ma magmatism Regional setting The bedrock of northernmost Finland includes several unmetamorphosed basaltic dikes that cut all the Paleoproterozoic and Archean rock types of the region. In the Salla–Pelkosenniemi region, a 100-km-long swarm of WNW-trending olivine tholeiitic dikes is found (Figure 13.2). These dikes are up to 100 m wide and show distinct symmetric differentiation patterns (Figures 13.10, 13.11). Zircon and titanite fractions from three samples of the Salla swarm have yielded an upper intercept age of 1122 ± 5 Ma (Lauerma, 1995). Other, presumably slightly younger swarms of basaltic dikes are found in the far northern part of the country and in adjacent Norway (Figure 13.2). In the Laanila–Ristijärvi region, these dikes run in a north–northeast, show an en-echelon pattern, are up to 200 m wide, and can be followed for more than 100 km along strike (Pihlaja, 1987). Typically, these rocks are medium-grained and ophitic with calcic plagioclase, augite, and serpentinized olivine as the main minerals. In aeromagnetic maps, the dikes are characterized by negative anomalies. Nd isotopic ROCKS,
DIABASES, AND...
•
577
+7
Kautokeino diabase
+6
TLE
DEPLETED MAN
+5
Laanila diabase
Ristijärvi diabase
+4 +3 +2 +1
CHUR
0 εNd
Key to the ~1265 Ma diabases: E1 E4 E7 A691c A562 A732
Southwestern Finland (Rämö, 1990; this paper)
R72b R120a
Central Sweden (Andersson, 1997)
–1 –2 –3 –4 –5 –6
84076 P1131 Central Sweden P1150 (Claesson, 1987) 73108 &73109
Valamo diabase, Russian Karelia
J-18 J-58
–7 –8
Central Sweden (Patchett et al., 1994)
–9 –10
1500
1400
1300
1200 Age (Ma)
1100
1000
900
Fig. 13.9. The initial Nd isotope compositions (εNd values) of the ~1460 Ma Valamo diabase in the Lake Ladoga region, Russian Karelia (Rämö et al., 2001), the ~1265 Ma “Postjotnian” diabases of southwestern Finland (Rämö, 1990; this work) and central Sweden (Andersson, 1997; Claesson, 1987; Patchett et al., 1994), and the 1100–1000 Ma diabase dikes in northern Finland (Mertanen et al., 1996b). CHUR is the Chondritic Uniform Reservoir (DePaolo and Wasserburg, 1976), depleted mantle evolution is according to DePaolo (1981).
578
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
0
10
20
30
40
50
60
70
80
Olivine basalt
Porphyritic pyroxene diabase
Banded pyroxene diabase
Quartz-bearing diabase
90
100 m
Fig. 13.10. Sketch cross-section of the Salla diabase dike showing pronounced internal differentiation (after Väänänen, 1965).
studies (Mertanen et al., 1996b) have yielded whole-rock–mineral ages of 1042 ± 50 Ma and 1013 ± 32 Ma for dike samples from Laanila and Ristijärvi, respectively. Mertanen et al. (1996b) also reported a 1066 ± 34 Ma age for a similarly oriented swarm ~150 km west of the Laanila–Ristijärvi dikes in Kautokeino, northern Norway (Figure 13.2).
Geochemistry and source characteristics Geochemical data for the 1100–1000 Ma mafic dikes are relatively few. According to Pihlaja (1987), the 1042 ± 50 Ma Laanila dike is an olivine tholeiite with 46 to 49 wt.% SiO2 and 5.4 to 6.8 wt.% MgO. This dike is thus quite similar to the 1265 Ma Satakunta diabases in southwestern Finland. The Sm-Nd whole-rock–mineral data of Mertanen et al. (1996b) indicate markedly positive initial εNd values (+4.8 to +5.8) for the Laanila–Ristijärvi and Kautokeino dikes. The initial values cluster about the model depleted mantle curve (Figure 13.9). Given that the diabase magmas transected a Paleoproterozoic–Archean crustal domain with a distictly unradiogenic Nd isotope composition (Huhma, 1986; Figure 13.6), the mafic magmas probably experienced little, if any, crustal contamination on their way from the mantle to upper crustal levels. CHAPTER
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4. Neoproterozoic and early Paleozoic sedimentary sequences 4.1. Regional setting Within the mainland area of Finland, only minor remnants of Neoproterozoic/early Paleozoic deposits are known, but a major occurrence of preserved Paleozoic cover is found beneath the Gulf of Bothnia (e.g., Winterhalter et al., 1981). In Estonia and northwestern Russia, the extensive Paleozoic cover sequence (Figure 13.3) is typically underlain by Vendian (NP3) clastic sediments. An obvious reference for correlations is the cover succession in northern Estonia (Figure 13.12). Here, the NP3 sequence is conformably overlain by Cambrian sedimentary rocks. Both the Vendian (NP3) and Cambrian strata comprise sandstone, siltstone, and mudstone, whereas the NP3–Cambrian boundary is lithologically not well marked. This, in part, explains the apparent difficulty of correlating the corresponding sedimentary remnants in Finland lithostratigraphically, and creates uncertainty in regard to their ages of deposition.
ROCKS,
DIABASES, AND...
•
579
Fig. 13.11. Photograph of banded pyroxene diabase in a marginal part of the Salla diabase (see Figure 13.10). Photo: Tuomo Manninen.
4.2. The Hailuoto Formation and its submarine extensions On the Hailuoto Island, the Mesoproterozoic Muhos Formation is conformably overlain by the Neoproterozoic (NP3) Hailuoto Formation (HlF). The available age estimates are based solely on microfossil studies and lithologic correlations, while Tynni and Donner (1980) pointed out that such comparisons do not justify any precise ages for these deposits. However, if the proposed estimates are even close to correct, the disconformity between the Muhos and Hailuoto Formations must represent a time gap of 500 to 600 Ma. The Hailuoto sedimentary sequence has been intersected by drillings at three localities near the western shore of the island. The boundary between the Mesoproterozoic and the Neoproterozoic is by no means distinctive. Tynni and Donner (1980) suggested that at the site of drillhole R2 (cf. also Veltheim, 1969) the uppermost 60 m represent the HlF, whereas the underlying 160 m belong to the Muhos Formation. The Hailuoto Formation consists of interbedded sandstone, siltstone, and mudstone (Veltheim, 1969; Tynni and Donner, 1980). The dominant rock type is a medium-grained, pale pink or light greenish subarkose. As the sedimentological features have not been reported systematically, the environment of 580
deposition remains uncertain, but the overall lithology and some reported details (e.g., cross-bedding) suggest fluvial plain as the most probable environment of deposition. According to map compilations based on seismic data (Winterhalter et al., 1981; Lundqvist et al., 1996; Koistinen et al., 2001), sedimentary deposits comparable to the Hailuoto Formation cover a large part of the bottom of the Bothnian Bay. In addition, it has been suggested that also remnants of Cambrian sedimentary rocks are preserved in that area (e.g., Winterhalter et al., 1981; Wannäs, 1989; Winterhalter, 2000).
4.3. The Lauhanvuori Formation The hill of Lauhanvuori in western Finland (Figure 13.2) is a geomorphological peculiarity inasmuch as these unmetamorphosed sedimentary rocks form a topographic high. The reasons for the preservation of Lauhanvuori sequence are not fully understood, but exceptionally weak Quarternary glacial erosion (Söderman et al., 1983) was presumably one important control. The main occurrence of Lauhanvuori sandstone covers an area of ~15 km by 5 km (Simonen and Kouvo, 1955) with some outliers farther to the north. The age of that formation is not finally resolved but, on a petrographical basis, Simonen and Kouvo (1955) correlated Lauhanvuori with the Cambrian “dikes” of clastic sandstone in southwestern Finland. The only direct age indication comes from fossil traces of crawling annelids (Figure 13.13) recognized in local sandstone boulders by Tynni and Hokkanen (1982). These authors therefore concluded that the Lauhanvuori Formation is younger than 700 Ma. In recent regional map compilations (e.g., Korsman et al., 1997), the sandstone is regarded as Neoproterozoic (NP3), but a Cambrian or even younger age of deposition cannot be excluded. The maximum thickness of the Lauhanvuori Formation has been estimated as ‘some
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
N
S
Quaternary
Gulf of Finland 50 m
Ordovician
Sea level
Cambrian -100
NP3 (Vendian)
-200
Crystalline basement (Paleoproterozoic)
100 m Limestone
Sandstone
Siltstone to mudstone
Fig. 13.12. Cross-section showing the uniform thickness of the cover units in northern Estonia to the north of Uljaste (for location see Figure 13.2). Simplified from Puura et al. (1996).
ten meters’ (Simonen and Kouvo, 1955), while a drilling at Tiiliharju, near the northern margin of the main occurrence area, penetrated 11 m of sandstone (Söderman et al., 1983). That sandstone is underlain by a deeply weathered Paleoproterozoic granite. The age of the local kaolinitic weathering is not known, but as feldspar clasts in the sandstone have been strongly altered, it appears to be coeval or even younger than the deposition of the sandstone. The composition of the sandstone is quartz-arenitic. However, part of the matrix plausibly represents alteration products of clastic feldspar, and in places the original composition has probably been nearly subarkosic. The median grain size typically ranges between 0.2 and 0.5 mm, but conglomeratic interbeds with rounded quartz pebbles are also common (Simonen and Kouvo, 1955). The type section of the Lauhanvuori Formation is defined by the Tiiliharju drillcore. It begins with a very coarse-grained, conglomeratic sandstone. The maximum size of the clasts is around 20 mm and the average CHAPTER
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grain size several millimeters. The clasts are dominantly quartz with some fragments of strongly altered feldspar. In addition, pebbles of purplish siltstone (cf. Söderman et al., 1983) are present. Upwards, the remaining 10 m of the drillcore consists of medium- to coarse-grained subarkosic to quartz-arenitic sandstone with scattered, well-rounded quartz clasts up to five mm in size. The color of the sandstone varies from light brownish yellow to pale pink in the uppermost part. The whole sequence is small-scale cross-bedded. Asymmetric ripple marks and lensoid siltstone interbeds have been reported from boulders (Tynni and Hokkanen, 1982). A fluvial depositional environment appears most probable (cf. Sauramo, 1916).
4.4. The bottom of the Bothnian and Åland seas The geology of the Bothnian Sea is known in considerable detail (Winterhalter et al., 1981, and references therein). The Mesoproterozoic ROCKS,
DIABASES, AND...
•
581
Fig. 13.13. Creeping tracks of worm-like annelids in the Lauhanvuori sandstone. The length of the sample is ~30 cm. Photo: Helena Halme.
sandstone is here unconformably overlain by Cambrian to Ordovician strata. However, in places (e.g., Finngrundet and Västra Banken), the Paleozoic rocks lie directly on the crystalline basement (Thorslund and Axberg, 1979). The youngest preserved rocks in the Bothnian Sea area are Upper Ordovician calcilutitic limestones (the ‘Baltic limestone’ of Thorslund, 1960). The maximum thicknesses of the Cambrian mudstone–sandstone sequence and the carbonate-dominated Ordovician sequence are around 200 m and 350 m, respectively (Winterhalter et al., 1981). Based on seismic data, Söderberg (1993) suggested that Paleozoic sedimentary rocks are also present in the Åland Sea area. The total thickness of the inferred Cambrian and Ordovician strata was estimated at 120 m and 230 m, respectively.
4.5. The Dividal Group of northwestern Lapland A small remnant of Neoproterozoic (NP3) to Cambrian sedimentary cover has been preserved immediately beneath the basal thrust of the Finnish Caledonides. In most previous reviews, these autochtonous rocks have been classified as part of the Caledonides (e.g., Lehtovaara, 1988, 1995; Korsman et 582
al., 1997), but in the present text the Dividal Group is separated from the Caledonides. This is done to emphasize that its deposition was hardly related to the Caledonian orogeny but rather to the sedimentation of the other coeval autochtonous cover sequences in Fennoscandia. Nevertheless, the boundary towards the structurally overlying Jerta Nappe is not sharp and sedimentary units comparable to the Dividal Group are also present within that nappe (Lehtovaara, 1988, 1989). The name of the Dividal Group has been adopted from adjacent areas in Norway. In Finland, the sequence is exposed along the Caledonide front (Figures 13.2, 13.14). The maximum thickness of the basal conglomerate unit is 10 m and that of the whole group ~200 m. Based on fossil studies, the major part of the Dividal Group is considered Cambrian, but the deposition of lower parts may have taken place already during the latest Precambrian (cf. Lehtovaara, 1988, and references therein). In Finland, the crystalline basement of the Dividal Group is Archean in age and mostly consists of gneissose granites and granodiorites. The nonconformity between the sedimentary sequence and the unweathered basement is sharp, while the contact dips gently (2–4o) beneath the Caledonides in the northwest. The following rock descriptions are according to Lehtovaara (1982a, 1988, 1995). The lowermost Dividal rocks are conglomerates and sandstones. In the conglomerates, well-rounded quartz and quartzite pebbles and cobbles are characteristic, but there are also a few angular basement-rock fragments in the basal parts. The amount of sandy matrix varies and cross-bedded sandstone interbeds are locally present. The uppermost sandstones of the basal conglomerate unit contain interbeds of mudstone. With decreasing amounts of sandstone, the sediment grades into mudstone in the upper unit (the sandstone–shale sequence) of the group. In that unit, shales (originally mudstones and siltstones) predominate over sandstones. Ripple marks and current bedding
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
FINNISH CALEDONIDES
NORWAY
N
FINLAND
0
SWEDEN
10 km
Vaddas Nappe: Cumulates/sills
Nalganas Nappe
Nabar Nappe; lower/upper part
Dividal Group and Jerta Nappe
Fig. 13.14. Simplified sketch map of the Finnish Caledonides (after Lehtovaara, 1989). For location, see Figure 13.2.
are common primary features in the quartzarenitic sandstones. A near-shore to tidal-flat depositional environment in a regionally transgressive regime has been suggested for the Dividal rocks in Finland (Lehtovaara, 1988).
4.6. Minor occurrences The meteorite impact structure of IsoNaak kima (Figure 13.2) has preserved a 100-m-thick sequence of unmetamorphosed sedimentary rocks (Elo et al., 1993). That sequence was deposited on a deeply weathCHAPTER
13
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ered Paleoproterozoic crystalline basement. Its lowermost units consist of conglomerate and quartz arenite. Upwards, the type section continues with siltstone, sandstone, and, finally, red or purplish mudstone (shale). On the basis of microfossil studies (Elo et al., 1993), a Neoproterozoic (NP1 to NP2) depositional age has been suggested. The rather small impact structure of Saarijärvi, with a diameter of ~1.5 km, also includes remnants of a sedimentary cover. The basement of the 156-m-thick sequence consists of Archean gneisses. The succession comprises ROCKS,
DIABASES, AND...
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an interbedded system of siltstone, mudstone, and sandstone. The bed thicknesses of the different units vary substantially. A Neoproterozoic (NP3) to Cambrian age estimate is based on microfossil data (Tynni and Uutela, 1985; Öhman et al., 2000). The circular depression of Söderfjärden is currently also considered an impact structure (Lehtovaara 1984, 1992; Abels et al., 2000). Drilling and other investigations have shown that this structure is filled by a ~240-m-thick sequence of Cambrian sedimentary rocks (Lehtovaara, 1982b; Tynni, 1982a). Mudstone (shale) dominates the succession, but also sandstone, siltstone, and conglomerate are present. The crystalline basement consists of a Svecofennian migmatitic granodiorite (Lehtovaara, 1982b). The carbonate rock occurrence in Lumparn Bay in the Åland Islands has been known since the early 1900’s. It comprises sedimentary cover rocks with a maximum thickness of ~120 m (Lehtovaara, 1982b), preserved in a fault bounded depression. The greatest part of the deposit is beneath the sea (cf. Winterhalter, 1982). The lowermost sediments have been deposited atop of a weathered erosional surface of rapakivi granite. The sequence consists of Lower Cambrian siltstone and sandstone, Lower Ordovician glauconitic limestone, and Middle Ordovician limestone (Tynni, 1982b; Lehtovaara, 1982a, and references therein). The origin of the down-faulted block of Lumparn Bay is still controversial. The first to propose an impact origin was Merrill (1979). That idea was initially rejected (e.g., Lehtovaara, 1982a; Winterhalter, 1982), but the impact hypothesis has been re-established as shock features were recognized (Svensson, 1994; Abels et al., 1998, 2000, 2002). In southwestern Finland, clastic sandstone dikes have been observed in many localities (cf. Figure 13.2), where sandstone is found as fissure fillings in cracks of the unweathered crystalline basement. According to microfossil studies (Tynni, 1982b), most of these dikes 584
are Lower Cambrian, while some have been assumed to be Lower Ordovician in age. The sandstone dikes and their origins have been discussed in detail by Bergman (1982) and Donner (1996). In the impact crater of Karikkoselkä, sedimentary rocks are found as blocks and fragments within an allochthonous breccia. The major part of the observed siltstones, sandstones, and mudstones are interpreted as Cambrian or Precambrian, but also Ordovician microfossils have been reported (Uutela, 2001). The overall setup of the crater, age estimates, and detailed rock descriptions can be found in Arkonsuo (2000). A remarkable concentration of sandstone boulders is known from Karstula, some 150 km east–northeast of Lauhanvuori and a local source is probable (e.g., Sauramo, 1916; Simonen and Kouvo, 1955; Lehtovaara, 1982a). This sandstone resembles that of Lauhanvuori, but its age of deposition is unknown.
5. Allochthonous rocks of the Finnish Caledonides 5.1. Introduction and regional setting The Caledonide orogenic belt represents a continent–continent collision of Laurentia and Baltica at the closure of the Iapetus Ocean. Thrust tectonics characterize the Scandinavian Caledonides, and stacks of detached and tectonically transported nappes (thrust sheets) with a prevailing southeastern sense of movement are typical. The Caledonide frontal thrust marks the border of the Caledonides towards the Fennoscandian Shield and its autochtonous cover (the Dividal Group). In Scandinavia, the main Scandian phase (see Roberts and Gee, 1985) of Caledonian compressional deformation and metamorphism took place during the late Silurian and early Devonian. Particularly in northern Scandinavia, an earlier Finnmarkian phase with
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
deformation in the late Cambrian and the early Ordovician has been identified (e.g., Sturt et al., 1978; Ramsay et al., 1985). Most of the rocks belonging to the Lower and Middle Allochtons represent the tectonically telescoped and shortened ancient western continental margin of Baltica, whereas the Upper Allochton consists of rocks from the transition zone between the continental margin and the oceanic crust, and terranes derived from areas outboard Baltica (e.g., Roberts and Stephens, 2000). The Scandinavian Caledonides of Finland occupy the extreme northwest part of Finnish Lapland (Figures 13.2 and 13.14). The allochthonous rocks of the Finnish Caledonides have been classified (Lehtovaara, 1989, 1995) by applying the tectonostratigraphic scheme developed in Norway by Zwaan and Roberts (1978) and Zwaan (1988). In the following, the allochthons and their suddivision are depicted according to Lehtovaara (1995); the rock descriptions are mainly adopted from Lehtovaara (1988, 1989, 1995).
5.2. The Lower Allochthon (Jerta Nappe) The parautochtonous rocks of the Jerta Nappe, except for the dolomites, correspond lithologically to those of the autochtonous Dividal Group. Structurally, the Jerta Nappe is considered to be an imbrication fan below the major sole thrust (Lehtovaara, 1995). The maximum thickness of this nappe has been estimated at 500 m (Lehtovaara, 1988). Typical rock types are bluish quartzite and slate, which represent recrystallized sandstone and shale, respectively. Dolomite occurs as a solitary bed near the upper contact of the Jerta Nappe. It displays a remarkable amount of shearing, and due to this deformation, the bed thickness varies from zero to some tens of meters.
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5.3. The Middle Allochthon (Nalganas and Nabar Nappes) The sharp lower contact of the Middle Allochthon plausibly represents the Caledonian sole thrust in Finland. According to the correlation by Lehtovaara (1988), the main thrusting occurred during the Finnmarkian phase, but reactivation during the Scandian phase may also have taken place. All the tectonostratigraphic units of the Middle Allochthon in Finland correspond to the indigenous rocks of the Kalak Nappe complex in Norway (cf. Lehtovaara 1989, 1995; Zwaan, 1988). The Nalganas Nappe comprises most of the Finnish Caledonides (Figure 13.14) and the strongly foliated, greenish or grayish arkosic quartzite of the Nalganas Nappe is their most typical rock type (the ‘Feldschiefer’ of Hausen, 1942). The main minerals are quartz, alkali feldspar, and plagioclase. The tectonic fabric is striking and protomylonitic to mylonitic textures are characteristic of the quartzite. The lower contact of the Nabar Nappe is far less distinctive than that of the Nalganas Nappe and is in the field recognized mainly by the first appearance of amphibolite. The Nabar unit is lithologically heterogeneous, except for its upper parts. Rock types include sericite quartzite, muscovite gneiss, mica gneiss, amphibolite, and granite pegmatite. Owing to strong deformation and amphibolite facies metamorphism, the protolith of the gneisses is ambiguous in most cases, but a Proterozoic to Archean age has tentatively been suggested (Lehtovaara, 1988). The top third of the Nabar Nappe consists solely of a banded, garnetiferous gneiss.
5.4. The Upper Allochthon (Vaddas Nappe) The Vaddas Nappe consists mostly of ultramafic to mafic rocks of the Halti–Ridnitsohkka igneous complex (Sipilä, 1992). The lower thrust of this allochthon is marked by strongly ROCKS,
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sheared rocks of that igneous complex (Lehtovaara, 1995) and intense refolding of the banded gneiss of the Nabar Nappe (Sipilä, 1992). The igneous complex consists of two parts: layered dunite–troctolite–olivine gabbro cumulates (Halti cumulates) and an interlayered system of tholeiitic sills and sillimanite-garnet gneisses (Ridnitsohkka gabbro sills). These have been considered to represent two magmatic phases – first the intrusion of the tholeiitic sills and, somewhat later, the emplacement of the layered complex (Sipilä, 1992; Vaasjoki and Sipilä, 2001). Vaasjoki and Sipilä (2001) reported a U-Pb baddeleyite age of 434 ± 5 Ma for the tholeiitic sills of Ridnitsohkka. Accordingly, this tectonostratigraphic unit has been correlated with the Vaddas Nappe of the Upper Allochton representing the Scandian phase of orogenic evolution (Sipilä, 1992; Lehtovaara, 1995).
6. Paleosols and Cenozoic sedimentary remnants Deep chemical weathering probably played a major role in the Mesozoic–Cenozoic evolution of Finland. Preserved in situ weathering crusts are valuable indications of ancient land surfaces and paleoclimate. Nevertheless, the geological interpretation of paleosols is often difficult, because dating of the weathering products is a cumbersome task and the results are often far less than precise. Paleosols are known in various parts of the country and they are especially common in northern Finland. Only paleosols directly overlain by Quaternary glacial deposits are described here; the paleosols associated with sedimentary cover (e.g., Lauhanvuori) have been discussed above. In southern Finland, several localities with minor indications of weathering are known, but the only major occurrence is the Virtasalmi kaolin deposit. At that locality, the saprolite developed on Svecofennian crystal586
line rocks and the thickness of the weathering profile typically reaches 30–40 m (~100 m at maximum). K-Ar dating of authigenic illite has yielded an age of 1180 Ma, and a Mesoproterozoic age of the weathering is further supported by indications of saprolite beneath Neoproterozoic sediments of the nearby IsoNaakkima sequence (Sarapää, 1996; see also above). Significant occurrences of weathering crusts are scarce in central Finland. In addition to the Lauhanvuori area, saprolite patches are only preserved along the coast of the Gulf of Bothnia (Söderman et al., 1983). From Kainuu, eastern Finland, a few kaolinitic occurrences are known. The parent rock of the saprolite there is often a Paleoproterozoic feldspathic quartzite and the preserved weathering profiles are up to 35 m thick (Sarapää, 1996). The age of weathering is not known. Kaolin occurrences similar to those in Kainuu are also present in northern Finland (e.g., at Siurunmaa, Savukoski, and Kuusamo). In central Lapland, northern Finland, the weathering crust is semi-continuous over wide areas and saprolites have been developed on crystalline bedrock. These crusts are typically some meters thick, but also tens of meters thick weathering profiles have been reported (e.g., Hirvas, 1991; Aario and Peuraniemi, 2000). The weathering crust apparently envelops the recent topography; thus saprolite has been observed in lowlands, on hill slopes, and also atop of summits (Hirvas, 1991). This indicates that Quarternary glacial erosion has been relatively weak in central Lapland. According to Aario and Peuraniemi (2000), the saprolite compositions in Lapland point to tropical or subtropical weathering conditions. These authors tentatively suggest a Cretaceous to Tertiary age for the saprolite formation. In general, the ‘preglacial’ chemical weathering history is poorly known in Finland. In a recent review (Migòn and Lidmar-Bergström, 2001), the main stage of chemical weathering in Finland and northern Sweden was consid-
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
ered late Mesozoic. In Finland, however, an older weathering component is present (e.g., Sarapää, 1996) and saprolites younger than Mesozoic (Paleogene to Neogene) may be present as well. The only indications of Cenozoic sedimentation in Finland are from eastern Lapland. The following description is according to Hirvas and Tynni (1976) and Tynni (1982c). At Akanvaara (205 m above sea level; Figure 13.2), unweathered bedrock is overlain by a 0.8-m-thick layer of stiff, symmictic clay that contains kaolinite and montmorillonite and is compositionally different from Quaternary clays of Finland. It is covered by a few meters of Quaternary till. On the basis of microfossils, a marine depositional environment and a Paleogene age have been suggested (Tynni, 1982c; Fenner, 1988). Some 70 km northeast of Akanvaara, at Naruskajärvi (270 m above sea level), a deposit of diatomaceous earth with Neogene freshwater microfossils is present. The deposit is found as an interlayer in a gravel–sand–silt deposit overlain by glacial till. An in situ nature of the deposits (no major glacial transport and redeposition) has been favored for both occurrences. Thus the observed remnants bear very significant information concerning the late geological evolution of northern Finland.
7. Tectonic evolution from the Mesoproterozoic to the Cenozoic 7.1. Introduction In this section, we place the Mesoproterozoic to Phanerozoic geological features of Finland into a regional tectonic context. The significance of the preserved sedimentary record and the dated igneous events (cf. Figures 13.1 and 13.2) are considered in the light of the evolutionary history of Fennoscandia. Especially important are major tectonic events, such as CHAPTER
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extension and rifting leading to the formation of new continental margins and episodes of accretion along the western margin of the craton. The traditional models concerning the Neoproterozoic and Phanerozoic evolution of the Finnish part of the Fennoscandian Shield maintain that this area had remained tectonically stable and that only minor sedimentation had occurred. These ideas have presumably arisen from the lack of an extensive preserved sedimentary record. Nevertheless, it is evident that in particular the western part of the craton underwent major tectonic reworking at about 1100–900 Ma (during the Sveconorwegian orogeny) and again at about 450–350 Ma (during the Caledonian orogeny). The latter was preceded by the opening of the Iapetus Ocean at ~600 Ma and succeeded by the opening of the North Atlantic in late Mesozoic and early Cenozoic times. The scattered 1100–1000 Ma dikes in northern Finland, dikes of similar ages in Scandinavia, the ~600 Ma kimberlites (Chapter 14), and the Devonian alkaline complexes of the Kola province (Chapter 14) can all be considered magmatic consequences of these tectonic episodes and show their prominent effects also in the eastern part of the Fennoscandian Shield. The Mesoproterozoic–Phanerozoic burialexhumation history of the shield can be assessed using (1) the sedimentary record; (2) other indications of ancient land surfaces, such as paleosols and impact structures; (3) the erosion depths of Paleozoic intrusions; and (4) isotope and fission track studies. A fundamental question related to the remnants of the sedimentary record is whether these reflect original basin configurations (e.g., Van Balen and Heeremans, 1998) or the selective preservation of originally more extensive cover sequences due to later tectonic movements (e.g., Cederbom et al., 2000). That question can be assessed, at least to a certain extent, by analyzing the original variations of thickness and sedimentary facies and their ROCKS,
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overall mode of occurrence (i.e., the amounts of folding, the nature of bordering faults, etc.). The usefulness of paleosols in such reconstructions is limited by difficulties in dating, several possible successive stages of weathering, and possible confusion with epigenetic hydrothermal alteration products. Reliably dated impact structures, like Lappajärvi and Sääksjärvi in Finland, are valuable piercing points marking the ancient erosional surface. In addition, remnants of pre-impact and/or post-impact cover may be preserved within such structures (cf. Kohonen and Vaarma, 2001; Abels et al., 2002). During the last fifteen years, fission track analysis in the Fennoscandian Shield (e.g., Zeck et al., 1988; Tullborg et al., 1996; Larson et al., 1999a; Cederbom, 2001; Murrell and Andriessen, 2001) has raised questions concerning the role of the Sveconorwegian and Caledonian orogenies and, especially, the original extent of foreland basins. It has been suggested (Larson et al., 1999b) that the Caledonide foreland sediments covered large parts of Finland during the Late Paleozoic and Mesozoic. Murrell and Andriessen (2001) have further indicated that the final exhumation of the shield occurred not earlier than during the Cenozoic. These models are in obvious conflict with the traditional ideas of only minor Phanerozoic deposition and the absence of a major sedimentary cover.
7.2. The intracratonic rift basin stage (~1600–1300 Ma) While detailed modeling of the mid-Proterozoic tectonic evolution is beyond the scope of this chapter (cf. Chapter 12), a brief discussion is included in order to delineate the depositional history of the supracrustal rocks. The extensive geophysical evidence of crustal thinning and mafic underplating, the long duration of magmatism, and the substantial volume of crust-derived magmas all require a major thermal source in the subcontinental mantle 588
when the rapakivi association was formed. The existence of such a thermal head is compatible with a mantle plume model (cf. Haapala and Rämö, 1992; Rämö and Korja, 2000), but that model has been challenged by Åhäll et al. (2000) who related the Fennoscandian rapakivi magmatism to successive subduction events along the southwestern edge of the shield. This “inboard model” cannot, however, explain the non-linear age distribution of the rapakivi granite batholiths (cf. Rämö et al., 2000; Chapter 12). Whatever the ultimate tectonic explanation of the magmatism, the following features (cf. also Rämö and Korja, 2000) are characteristic of the rapakivi stage: • major tectonothermal activity with mantle upwelling, mafic underplating, and emplacement of rapakivi granites; • crustal thinning and formation of intracratonic basins. The alluvial deposits (e.g., the Satakunta sandstone) are well compatible with an overall intracratonic setting. An eolian origin of the Dala sandstone in central Sweden (Pulvertaft, 1985) would indicate that sedimentation may have occurred also outside of major rift valleys. For the Satakunta sandstone, a minimum age of 1265 Ma is well constrained, whereas recent results from the Lake Ladoga area (Figures 13.2 and 13.6) indicate that graben formation and sedimentation of arkosic sandstones occurred there shortly after or perhaps concomitantly with the emplacement of the 1560–1530 Ma Salmi rapakivi granite complex (Amelin et al., 1997; Rämö et al., 2001). In general, however, the depositional age of the Mesoproterozoic sequences is poorly constrained. It is possible that the preserved strata represent multiple depositional stages and the basin histories may thus be longer and more complex than previously envisaged. The original extent of the intracratonic basin system (MP1 to MP2) and the primary
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
thicknesses of the sediments are difficult to estimate reliably. However, comparisons with the assumed scales of Mesoproterozoic (“Riphean”) basins in the Russian platform (e.g., Kumpulainen and Nystuen, 1985; Bogdanova et al., 1996; Nikishin et al., 1996) suggest that deposition had been more widespread than what is indicated by the preserved basins. This is supported also by the presence of Mesoproterozoic sedimentary rocks within the Lappajärvi impact structure. Similarly, the abundance and scattered occurrence of sandstone glacial boulders is best explained by assuming that unexposed sandstone remnants are more common in Finland than currently indicated in geological maps.
7.3. Crustal extension episodes and the Sveconorwegian orogeny (~1300–900 Ma) The younger Mesoproterozoic magmatic event at ~1265 Ma and the mafic dike swarms of Salla (~1125 Ma) and Laanila (~1040 Ma) presumably reflect short-lived extensional episodes in northern Finland. Their time span approximately corresponds to that of the Sveconorwegian orogeny (~1.2–0.9 Ga; Johansson et al., 1991). The ~1265 Ma dikes could record the onset of the Sveconorwegian–Grenvillian orogeny (e.g., Gorbatschev et al., 1987; Rämö, 1990) or, from a supercontinental point of view, the breakup of an earlier Proterozoic supercontinent, which created the separate cratons of Baltica and Laurentia (e.g., Elming and Mattson, 2001; Pesonen et al., 2003). Major intracratonic tectonic movements and extensive rift systems have also been indicated for the period between ~1400 and 1100 Ma in the interior of the East European craton (e.g., Nikishin et al., 1996; Bogdanova et al., 1996). Both in Finland and in Sweden, the ~1265 Ma dikes are spatially connected with older Mesoproterozoic (MP1 or “Subjotnian”) intrusions and sedimentary deposits (MP1/ CHAPTER
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MP2). Their emplacement plausibly occurred along pre-existing (MP1) crustal ruptures. It is also possible that initial subsidence related to this stage of evolution triggered the deposition of the upper parts of the Satakunta and Muhos formations. The down-faulting related to the final events of this extensional stage was apparently the main control of the preservation of the Satakunta Formation (Laitakari, 1983; Kohonen et al., 1993), perhaps also that of the the Muhos Formation. The kaolinitic Virtasalmi saprolite (~1180 Ma; Sarapää, 1996) lying directly on the crystalline basement manifests local absence of sedimentary cover and conditions of long-lasting continental weathering in central Finland. Nevertheless, our model implies that, at the end of the Mesoproterozoic, large parts of Finland were covered by sedimentary rocks of intracontinental depositional origin (Figure 13.15). The influence of the Sveconorwegian orogeny on the foreland in the northeast is speculative, but it has been suggested that a km-scale pile of foreland sediments had been deposited in the vicinity of the Sveconorwegian front in southwestern Sweden at ~950 Ma (Tullborg et al., 1996; Larson et al., 1999). In Finland, the Iso-Naakkima sedimentary sequence overlies a saprolite resembling that in Virtasalmi. Its depositional age is very poorly constrained (~1000 to 650 Ma) and it is thus difficult to speculate in regard to connections with the sedimentation in the Sveconorwegian foreland.
7.4. The Neoproterozoic exhumation stage (~900–600 Ma) Except for the case of Iso-Naakkima (see above), evidence for Neoproterozoic (NP1 and NP2) geological evolution are virtually absent in Finland. Also, the overall early Neoproterozoic (NP1/NP2) geological history of the surrounding areas is insufficiently known (cf. Vidal and Moczydlowska, 1995). After that time, the development of an ROCKS,
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~1300 Ma
Sveconorwegian orogenic front
~900 Ma
Iapetus Ocean
~700 Ma
Tornquist Sea
~450 Ma
Caledonian orogenic front
Estimated margin of major lithospheric flexure ~360 Ma
590
~30 Ma
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
ancient North European shield area with little or no sedimentary cover is evident during the Neoproterozoic (NP2/NP3 or “pre-Vendian”). This is manifested by erosion of the pre-existing cover, continental conditions, and subaerial weathering profiles in many parts of the present East European craton (e.g., Korkutis, 1981; Puura et al., 1996). For example, to the south and southeast of Finland, in Estonia and Russia, late Neoprotezoic (NP3 or “Vendian”) deposition occurred directly atop of the metamorphic crystalline basement (e.g., Winterhalter et al., 1981; Amantov et al., 1988). These sediments are typically underlain by a kaolinitic paleosol. In places, especially in the west, the weathering crust is directly covered by Cambrian deposits (Puura et al., 1996). This indicates that the deposition of the latest Neoproterozoic (“Vendian”) sediments and subaereal weathering took place virtually simultaneously. The preserved “Vendian” successions in Estonia and Lake Ladoga–St. Petersburg area show no decreasing thicknesses towards the north. This suggests that these deposits were once present also in southern Finland (Puura et al., 1996, 1999). In Finland, the deposits referred to as late Neoproterozoic (NP3 or “Vendian”) are poorly dated. These include the Hailuoto and Lauhanvuori Formations, and possibly also the lowermost part of the Dividal Group, but their relationship to the overall basin configuration remains to be resolved.
In summary, although Finnish data are scarce, it is suggested that gradual uplift and erosion characterized most of the Neoproterozoic Era in Finland. As a prologue to the platformal stage, some local subsidence and deposition may have occurred at the end of this stage after a prolonged period of erosion and subaereal weathering.
7.5. The stage of platform sedimentation (~600–420 Ma) Observations from southern Finland, such as clastic dikes of Cambrian sandstone and the ~560 Ma (Muller et al., 1990) impact structure of Sääksjärvi, indicate that the present erosional level is very close to the major unconformity between the crystalline basement and the NP3/Cambrian strata. This is also supported by isotope studies. The Phanerozoic crystallization of galena (Vaasjoki, 1977; Bergman and Lindberg, 1979; Sundblad et al., 2002) and uraninite (Vaasjoki et al., 2002) in fractures within the crystalline basement in southern Finland may reflect the proximity of the ancient unconformity. These mineralizations were potentially related to increased permeability and hydrothermal activity along the unconformity just above the present erosional level during this particular stage of evolution. In the latest Precambrian and during the Cambrian, fluvial and shallow marine de-
Fig. 13.15. (facing page) Model illustrating schematic paleogeographic snapshots of Fennoscandia; the blue areas indicate the present distribution of Mesoproterozoic cover, the horizontally ruled areas represent active depositories, the dotted areas show the sedimentary cover. (A) Alluvial–fluvial–eolian deposition in intracratonic rift basins; the red arrow indicates the main direction of sediment transport (see Figure 13.4); (B) Sveconorwegian foreland basin; examples of assumed ~1270 Ma normal faults (jagged lines) controlling the preservation of earlier (MP1 to MP2) cover are also indicated; (C) Remnants of sedimentary cover ~700 Ma ago; (D) The Ordovician carbonate platform; the assumed continental margin normal fault system (jagged lines) is also outlined; (E) The Caledonian foreland basin; the estimate of major lithospheric flexure (the deepest part of the basin) according to Samuelsson and Middleton (1998); (F) Assumed main sediment transport system (red arrows) and reactivation of earlier faults (jagged lines) ~30 Ma ago; schematic contours of late Mesozoic and Paleogene uplift are also indicated (after Riis, 1996). CHAPTER
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position prevailed in the slowly submerging northwestern part of the East European craton. Tectonically, basin formation at that time was contemporaneous with the breakup of the Meso- to Neoproterozoic Rodinia supercontinent and the opening of the Iapetus Ocean at ~600 Ma. Extension along the Tornquist line in the southwest occurred approximately at the same time. In Finland, the age of kimberlites approximately corresponds the time of that supercontinent breakup, but possible causal connections have not yet been worked out. The Cambrian basins to the south and southeast of Finland, in Estonia, and the St. Petersburg area have been studied and modeled in detail (e.g., Puura et al., 1996; Mens and Pirrus, 1997; Artyushkov et al., 2000). Detailed description of the stratigraphy and inferred trangression and regression phases are beoynd the scope of this chapter, but a short summary of basinal stages is given. In Estonia, the lowermost Cambrian mudstone and siltstone formations show increasing thicknesses first towards the northeast and east and later, variably, towards the west and southwest. Towards the end of the Cambrian, sandstones and siltstones representing shallow marine depositional conditions become dominant throughout the area. The basin was apparently deepening towards the newly formed continental margin in the west. The original extents and geometries of the Cambrian basins in Finland are difficult to reconstruct in detail. Nevertheless, it has been suggested (cf. Puura et al., 1996) that, at the end of the Cambrian, southern Finland had been covered by a pile of fine clastic sediments with a thickness varying between 100 and 350 m. During the Ordovician and Silurian, an extensive carbonate platform existed at the western margin of the ancient continent (e.g., Bruton et al., 1985; Basset, 1985; Nestor and Einasto, 1997). Carbonate rocks and marly clastic sediments, 200 to 400 m in thickness, represent this stage in the Baltic countries (e.g., Puura and Vaher, 1997) and, before the 592
onset of the Caledonian collisional tectonism, the axis of maximum thickness of sediments appears to have extended along a line connecting the western coast of Estonia with the coast of Poland. That basin was probably open towards the Tornquist Ocean in the south and the Iapetus Ocean in west (Torsvik, 1998; Puura et al., 2000). Puura et al. (1996, 2000) suggested that these sediments covered large parts of the present Fennoscandian Shield. In Finland, this view is supported by recent observations of Ordovician rocks in the Karikkoselkä impact structure (see Figure 13.2). In addition, the Ordovician in Estonia shows neither a systematic decrease of bed thickness nor a change of sedimentary facies towards the north. The thickness estimates for the Cambrian and Ordovician strata in the Bothnian Sea roughly correspond to those in the Baltic countries and there is no systematic decrease in sedimentary unit thickness towards the margins of the present cover rock area in the Bothnian Sea (cf. Axberg, 1980). This, together with the current bowl-shaped nature of the structure, indicates that the Bothnian Sea is a tectonically preserved remnant of an originally extensive cover sequence rather than a primary Paleozoic basin (cf. van Balen and Heeremans, 1998). In summary, the present occurrences of Paleozoic cover in Finland represent sporadic erosional remnants of an extensive platform cover. The development of early Paleozoic marine basins may tentatively be connected with the development of passive-margin-type conditions, followed by initial foreland subsidence and formation of an early foredeep-type basin in response to continent collision in the west. However, it appears that the development of the Silurian Baltic basin was tectonically controlled not only by the formation of the Scandinavian Caledonides but also by the development of the North German–Polish Caledonides in the southwest (e.g., Poprawa et al., 1999).
• C H A P T E R 1 3 • S E D I M E N TA RY R O C K S , D I A B A S E S , A N D. . .
7.6. The Caledonian foreland stage (~420–350 Ma) and the final exhumation of the shield Sedimentary rocks younger than Ordovician are not known in Finland, but the influence of the Caledonian orogeny can hardly be ignored. Models presented for the Silurian and Devonian evolution in Finland can only be based on isotope methods, geological modeling, and comparisons with adjacent areas, especially in Sweden, Norway, and Estonia. The formation of a thrust and fold belt in Scandinavia at ~400 Ma and related crustal thickening formed a foreland basin on the eastern side of the Caledonian orogen. As the present crustal level in Scandinavia represents a deep section through the Scandinavian Caledonides, vast amounts of sediments must have been deposited during the Devonian. A corresponding heating of the basement due to burial has been recorded and modeled by, e.g., Zeck et al. (1988), Tullborg et al. (1996), Samuelsson and Middleton (1998), and Larson et al. (1999a). The width of the foreland basin and the original thickness of the sedimentary piles in different parts of the basin are subject to speculation (e.g., Larson et al., 1999b; Samuelsson and Middleton, 1999). The thick Devonian “Old Red” sandstones in the Baltic countries have been interpreted as part of the foreland basin (e.g., Plink-Björklund and Björklund, 1999). The Silurian–Devonian depositional hiatus plausibly records a rapidly changing base level configuration and a basin setup related to foreland basin initiation. The role of the Caledonides in the Paleozoic evolution of the Baltic region is reflected in the shape of that basin. The axis of the basin, as deduced from both the early Paleozoic carbonate deposition and the Devonian sandstone formations, apparently followed the overall northeastern trend of the Scandinavian Caledonides (Puura et al., 2000). Larson et al. (1999a) estimated CHAPTER
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that the detritus derived from the Scandinavian Caledonides had buried most of Finland, and that practically all of Finland was covered by 0.5- to 1.5-km-thick sedimentary strata (Figure 13.15). The final exhumation of the Finnish part of the Fennoscandian Shield occurred probably in late Mesozoic and the early Cenozoic. This estimate is supported by the fact that, at ~75 Ma, the Lappajärvi meteorite impact site was still buried beneath a sedimentary cover. A high amount of post-Devonian erosion in northern Finland is undisputably manifested by the exposed ~365 Ma alkaline intrusions (Iivaara, Sokli). The denudation of the Caledonide mountain range had reached a stage of a hilly landscape during the late Mesozoic (Riis, 1996). As a response to the opening of the North Atlantic, substantial tectonic uplift (1–2 km) took place during the Cenozoic in northwestern Scandinavia. According to the estimates given (e.g., Riis, 1996; Stuevold and Eldholm, 1996), uplift in northern Finland was ~500 m during the late Cretaceous and Paleogene, continuing into the Neogene. The present morphology of the eastern and western coasts of the Bothnian Sea – the level Finnish coast and the faulted uneven Swedish coast – has been interpreted by different amounts of Cenozoic uplift (van Balen and Heeremans, 1998). Paleogene to Neogene evolution has probably had a significant impact on the exhumation history and current topography of Finland, especially in its northern part.
7.7. Concluding remarks All the discussed magmatic events, the mafic dikes at ~1265 Ma, ~1100 Ma, and 1000 Ma, the kimberlites at ~600 Ma, and the alkaline intrusions at ~360 Ma, have at least a temporal relationship with the well-known major tectonic events on the Fennoscandian Shield. It also appears that the particulars of the admittedly scarce sedimentary record ROCKS,
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can be interpreted rather soundly by applying simple tectonic modeling of the basin types and depositional processes derived from the various stages of the tectonic histories of the areas adjoining Finland. The presently preserved unmetamorphic sedimentary rocks in Finland represent only small remnants of an originally much more extensive sedimentary cover. Thus the current map patterns of these occurrences reflect the result of selective preservation rather than original extent of the depositional basins. The occurrence areas of Mesoproterozoic sedimentary rocks may still reflect, at least in part, the location of one-time major rifts, but most of these successions were probably eroded prior to the formation of the late Neoproterozoic/Cambrian basins. At end of the Silurian, most of present Finland was covered by a thick pile of late Neoproterozoic to Paleozoic platform sediments. The destruction of this cover was either related to the development of the foreland bulge during the Devonian or a consequence of the uplift that followed the denudation of the Caledonide mountain range during the late Paleozoic and Mesozoic. The slow denudation during late Mesozoic to Neogene uplift in the west and northwest put a finishing touch on the Fennoscandian Shield as we know it today.
Acknowledgments It is our pleasure to acknowledge Eira Kuosmanen who provided invaluable assistance with map data processing. Tuomo Manninen, Pekka Pihlaja, Hannu Kujala, and others provided photographs and Anneli Lindh and Pirkko Kurki helped in drafting the figures. Discussions with Peter Sorjonen-Ward and Seppo Paulamäki were very useful. A particularly devoted review of the manuscript by Roland Gorbatschev, and comments by Satu Mertanen, Matti Vaasjoki, and Martti Lehtinen are sincerely appreciated. 594
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Norway. In: D.G. Gee, B.A. Sturt (Eds.), The Caledonide Orogen; Scandinavia and related areas; Vol. 1. John Wiley & Sons, Chichester, 163–184. Richard, P., Shimizu, N., Allègre, C.J., 1976. 143 Nd/146Nd, a natural tracer: an application to oceanic basalts. Earth Planet. Sci. Lett. 31, 269–278. Riis, F., 1996. Quantification of Cenozoic vertical movements of Scandinavia by correlation of morphological surfaces with offshore data. Global Planet. Change 12, 331–357. Roberts, D., Gee, D.G., 1985. An introduction to the structure of the Scandinavian Caledonides. In: D.G. Gee, B.A. Sturt (Eds.), The Caledonide Orogen; Scandinavia and related areas; Vol. 1. John Wiley & Sons, Chichester, 55–68. Roberts, D., Stephens, M.B., 2000. Caledonian Orogenic Belt. In: Th. Lundqvist, S. Autio (Eds.), Description to the bedrock map of central Fennoscandia (Mid-Norden). Geol. Surv. Finland, Spec. Pap. 28, 79–104. Salonen, V.P., 1991. Glacial dispersal of Jotnian sandstone fragments in southwestern Finland. In: S. Autio (Ed.), Geological Survey of Finland, Current Research 19891990. Geol. Surv. Finland, Spec. Pap. 12, 127–130. Samuelsson, J., Middleton, M.F., 1998. The Caledonian foreland basin in Scandinavia: constrained by thermal maturation of the Alum Shale. GFF 120, 307–314. Samuelsson, J., Middleton, M.F., 1999. The Caledonian foreland basin in Scandinavia: constrained by thermal maturation of the Alum Shale. Final reply. GFF 121, 342. Sarapää, O., 1996. Proterozoic primary kaolin deposits at Virtasalmi, southeastern Finland. Geol. Surv. Finland, Espoo, 1–152. Sauramo, M., 1916. Über das Vorkommen von Sandstein in Karstula, Finland. Fennia 13, 1–13. Sederholm, J.J., 1897. Om indelningen af de prekambriska formationerna i Sverige och Finland och om nomenklaturen för dessa äldsta bildningar. Geol. Fören. Stockholm Förhandl. 19, 20–53. Simonen, A., 1960. Pre-Quarternary rocks in
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May 31-June 5, 1994; Abstract volume, 1. Thorslund, P., 1960. Cambro-Silurian. In: N.H. Magnusson, F. Brotzen, O. Kulling, P. Thorslund (Eds.), Description to accompany the map of the pre-Quaternary rocks of Sweden. Sveriges Geol. Unders. Ba 16, 69–110. Thorslund, P., Axberg, S., 1979. Geology of the southern Bothnian Sea; Part I. Bull. Geol. Inst. Univ. Uppsala, New Series 8, 35–62. Torsvik, T.H., 1998. Palaeozoic palaeogeography: A North Atlantic viewpoint. GFF 120, 109–118. Tullborg, E.-L., Larson S.Å., Stiberg, J.-P., 1996. Subsidence and uplift of the present surface in the southeastern part of the Fennoscandian Shield. GFF 118, 126–128. Tynni, R., 1978. Microfossils of the Muhos formation (in Finnish with English abstract). Geol. Surv. Finland, Rep. Invest. 30, 1–18. Tynni, R., 1982a. New results of studies on the fossils in the Lower Cambrian sediment deposits of the Söderfjärden basin. Bull. Geol. Soc. Finland 54, 57–68. Tynni, R., 1982b. On Paleozoic microfossils in clastic dykes in the Åland Islands and in the core samples of Lumparn. Geol. Surv. Finland, Bull. 317, 35–114. Tynni, R., 1982c. The reflection of geological evolution in Tertiary and interglacial diatoms and silicoflagellates in Finnish Lapland. Geol. Surv. Finland, Bull. 320, 1–40. Tynni, R., Donner, J., 1980. A microfossil and sedimentation study of the Late Precambrian formation of Hailuoto, Finland. Geol. Surv. Finland, Bull. 311, 1–27. Tynni, R., Hokkanen, K., 1982. Annelidien ryömimisjälkiä Lauhanvuoren hiekkakivessä. Traces of crawling by annelids in Lauhanvuori sandstone. Geologi 34, 129–134. (in Finnish with English summary) Tynni, R., Uutela, A., 1984. Microfossils from the Precambrian Muhos formation in Western Finland. Geol. Surv. Finland, Bull. 330, 1–38. Tynni, R., Uutela, A., 1985. Myöhäisprekambrinen ajoitus Taivalkosken savikivelle mikrofossiilien perusteella. Late Precambrian shale formation of Taivalkoski in northern Fin-
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land. Geologi 37, 61–65. (in Finnish with English summary) Upton, B.G.J., Rämö, O.T., Vaasjoki, M., Sviridenko, L.P., Svetov, A., 1998. Constraints on petrogenesis of Mesoproterozoic CFBs from the SE Fennoscandian Shield: Trace element and Nd isotopic evidence. Abstracts of ICOG-9, 1998 Beijing. Chinese Sci. Bull. 43, 134. Uutela, A., 1990. Proterozoic microfossils from the sedimentary rocks of the Lappajärvi impact crater. Bull. Geol. Soc. Finland 62, 115–120. Uutela, A., 2001. Proterozoic and early Palaeozoic microfossils in the Karikkoselkä impact crater, central Finland. Bull. Geol. Soc. Finland 73, 75–85. Väänänen, P., 1965. Diabaasijuonesta Sallassa. M.Sc. Thesis, University of Helsinki, Finland. 1–62. (in Finnish) Vaarma, M., Pipping, F., 1997. Alajärven ja Evijärven kartta-alueiden kallioperä. Pre-Quaternary rocks of the Alajärvi and Evijärvi map-sheet areas. Geological map of Finland 1:100 000. Explanation to the maps of PreQuaternary rocks. Sheets 2313 and 2314. Geol. Surv. Finland, Espoo, 1–83. (in Finnish with English summary) Vaasjoki, M., 1977. Rapakivi granites and other postorogenic rocks in Finland: their age and the lead isotopic composition of certain associated galena mineralizations. Geol. Surv. Finland, Bull. 294, 1–64. Vaasjoki, M., 1996. Explanation to the geochronological map of southern Finland: The development of the continental crust with special reference to the Svecofennian orogeny. Geol. Surv. Finland, Rep. Invest. 135, 1–30. Vaasjoki, M., Sakko, M., 1987. Zirkoni-indikaatio Satakunnan hiekkakiven alkuperästä. Summary: A zircon indication on the provenance of the middle Proterozoic Satakunta sandstone, Finland. Geologi 39, 184–187. Vaasjoki, M., Sipilä, P., 2001. U-Pb isotopic determinations on baddeleyite and zircon from the Halti-Ridnitsohkka intrusion in Finnish Lapland: a further constraint on Caledonide evolution. In: M. Vaasjoki (Ed.), Radio-
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Pleistocene glaciation. Geol. Surv. Finland, Bull. 258, 1–66. Winterhalter, B., 1982. The bedrock geology of Lumparn Bay, Åland. Geol. Surv. Finland, Bull. 317, 115–130. Winterhalter, B., 2000. Sedimentary rocks underlying the Gulf of Bothnia. In: Th. Lundqvist, S. Autio (Eds.), Description to the bedrock map of central Fennoscandia (Mid-Norden). Geol. Surv. Finland, Spec. Pap. 28, 76–77, 79. Winterhalter, B., Flodén, T., Ignatius, H., Axberg, S., Niemistö, L., 1981. Geology of the Baltic Sea. In: A. Voipio (Ed.), The Baltic Sea. Elsevier, Amsterdam, 1–121.
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Zeck, H.P., Andriessen, P.A.M., Hansen, K., Jensen, P.K., Rasmussen, B.L., 1988. Paleozoic paleo-cover of the southern part of the Fennoscandian Shield; fission track constraints. Tectonophysics 149, 61–66. Zwaan, K.B., 1988. Nordreisa, berggrunnsgeolokisk kart - M 1:250 000. Norges Geol. Unders. Zwaan, K.B., Roberts, D., 1978. Tectonostratigraphic succession and development of the Finnmarkian nappe sequence, North Norway. Norges Geol. Unders., Bull. 48, 53–71.
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Chapter 14
KIMBERLITES, CARBONATITES, AND ALKALINE ROCKS
H.E. O’Brien, P. Peltonen, H. Vartiainen
Cover page: An example of hypabyssal kimberlite from Kaavi–Kuopio (Pipe 10, Ryönä) displaying the classic suite of lithospheric mantle-derived xenocrysts: olivine (two generations, now pseudomorphed by serpentine), red pyrope, gray Mg-ilmenite, and bright green chromian diopside. Photo: Helena Saarinen.
O’Brien, H.E., Peltonen, P., Vartiainen, H., 2005. Kimberlites, carbonatites, and alkaline 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. 605–644. © 2005 Elsevier B.V. All rights reserved.
Finland contains some of the classic examples of carbonatites and alkaline rocks in the world, and more recently kimberlites have also been found. The carbonatites span the age range from Archean, the Siilinjärvi carbonatite complex, to Proterozoic, the Laivajoki and Kortejärvi occurrences, to the Devonian Sokli complex, one of the largest in the world at just over 20 km2 in size. Study of carbonatites with such a range in ages gives a unique temporal perspective of the development of this type of carbonated mantle source and although there is a considerable range in levels of REE enrichment in these carbonatites, radiogenic and stable isotope compositions are little changed from the Archean to Devonian examples. This implies derivation from a well-mixed portion of the Earth’s upper mantle. Surprisingly, isotope data from the Iivaara ijolite complex, the southernmost expression of the huge Devonian Kola alkaline province, show highly enriched isotope signatures. This is believed to be the result of extreme crustal contamination of magmas originally with isotope compositions similar to the carbonatites. The 600 Ma Group I kimberlites in the Kaavi–Kuopio region are typical of those found elsewhere in the world, including large quantity of entrained lithospheric mantle material that allows deciphering of the stratigraphy of the underlying mantle. Xenolith and xenocryst studies show that this mantle has at least three distinct layers, an upper, extremely depleted layer composed mostly of harzburgites down to ~110 km, a middle layer composed of dominant lherzolite with subordinate harzburgite and wehrlite down to ~180 km, and a lower more fertile layer down to at least 240 km. The latter represents either refertilized Archean mantle or a recent Proterozoic underplate during continent collision at ~1.88 Ga. Despite this large entrained load of lithospheric mantle, these kimberlites are isotopically unaffected by this process and show derivation from a sublithospheric well-mixed mantle source. In contrast, the Group II kimberlite-lamproite hybrids of the Kuhmo region contain considerably more aged lithospheric component that is apparent in their extreme isotope compositions. These phlogopite-rich rocks in places also contain significant lithospheric mantle material. Preliminary data from mantle xenocrysts suggest that the lithospheric mantle stratigraphy here is less heterogeneous, with a depleted, high-Mg lherzolite- and harzburgite-bearing horizon extending from the very top of the garnet-bearing mantle at about 80 km depth, to a depth of 250 km or more. An areally diverse group of Proterozoic primitive mica-rich lamprophyres in eastern Finland probably imply a similar significant aged lithospheric component.
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1. Introduction Even though alkaline rocks represent a small fraction of all igneous rocks, in fact less than 1%, they cover such a wide range of chemical and mineralogical diversity that they account for nearly half of all igneous rock names. The petrogenesis of these rocks is particularly interesting, in part due to their great variability and in part because they are economically important, containing most of the global reserves of, for example, the rare earth elements (REE), Nb, and Ta and such minerals as apatite and diamond. Finland contains classic examples of many of these rock types, including one of the oldest carbonatites in the world at Siilinjärvi and the type locality for the nepheline-bearing rock ‘ijolite’ at Iivaara. Recently, diamondiferous kimberlites similar to those from southern African and Siberia have been discovered and have proven to be extremely valuable in a scientific sense because of the unique samples of lithospheric mantle underlying eastern Finland that they have transported to the surface. Alkaline rock research by Finnish geologists (Figure 14.1) has a long tradition, and dates back to 1857, when H.J. Holmberg collected the first samples from the ijolite intrusion at Iivaara in Kuusamo (Holmberg, 1857; Ramsay and Berghell, 1891). Other important targets for research have been the Lovozero (Lujaur Urt) and Khibina (Umptek) nepheline syenite/carbonatite intrusions discovered on the Kola Peninsula, Russia, by Wilhelm Ramsay in 1887 (Ramsay, 1889; Ramsay and Hackman, 1894), the Kola Peninsula Turjanniemi alkaline rock district investigated by Håkan Kranck (Kranck, 1928) and the Iivaara area in Kuusamo studied by Mauno Lehijärvi (Lehijärvi, 1960). Research on alkaline rocks in Finland underwent resurgence in the latter half of the 1960’s with the discovery of the Sokli carbonatite intrusion at Savukoski (Paarma, 1970) and the carbonate-phosphorus ore at Siilinjärvi (Puustinen, 1970, 1971). Alkaline rocks have been intruded into 608
the bedrock of Finland since at least the Late Archean, and this chapter will describe the occurrences from the oldest to the youngest, starting with one of the oldest carbonatites in the world.
2. Description of alkaline rock complexes of Finland 2.1. The Archean Siilinjärvi carbonatite The Siilinjärvi carbonatite complex is located in eastern Finland close to the city of Kuopio (Figure 14.1). It consists of a steeply dipping lenticular body roughly 16 km long with a maximum width of 1.5 km and a surface area of 14.7 km2 (Figure 14.2) intruded into granite gneiss. It was discovered in 1950 after samples of carbonatite were found by local mineral collectors; studies of these samples at the Geological Survey of Finland (GTK) ultimately led to the body. Exploration drilling began in 1958 and continued along with laboratory and pilot plant work until 1979 when an open pit mine for phosphorus ore was commissioned (Figure 14.3). Present production at the Siilinjärvi mine is about 9.2 Mt of ore per annum. The carbonatite within the Siilinjärvi complex occurs as a central tabular 600–700m-wide body of calcite- and dolomite-bearing phlogopite rocks running the length of the complex surrounded by a fenite margin. Although not strictly zoned, cross-cutting relationships and xenoliths suggest that, at least at the present level of exposure, some of the syenites formed early, followed by a relatively carbonate-poor ultramafic magmatic pulse that created the majority of the phlogopite rocks, finally culminating in a carbonate-dominated pulse. Crosscutting the surrounding bedrock, the fenite halo, and the central intrusive body is a 4-km-long, 20–30-m-wide, N-trending melasyenite dike (or series of dikes, Heikki Lukkarinen, pers. comm., 2003) within the
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Barents Sea
Kola alkaline province Lovozero Kovdor
Khibina
Sokli
NORWAY
Terskii Laivajoki
Iivaara
Verkhotina Zolotitsa
Kortejärvi
Kostamuksha
SWEDEN
Lentiira
RUSSIA Siilinjärvi
FINLAND
Kemozero
Kuopio Kaavi
Halpanen
Svecofennian mobile belt
Baltic Sea
0
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Archean orthogneiss, migmatite
Hardrock kimberlite occurrences
Archean greenstones, supracrustal rocks
Carbonatites and related rocks
Post-Archean, mainly Svecofennian rocks Caledonian orogenic belt
Southwest margin of the Archean craton
Phanerozoic sedimentary cover Fig. 14.1. Location map of kimberlites, carbonatites, and undersaturated alkaline rocks in Finland with the general geology and the Devonian Kola alkaline province outlined.
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27°44’
Pien-Varpanen
Suur-Varpanen
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Pahkalampi Mustinjärvi
Diabase Metasyenite dike Fenite Carbonatite Quartz diorite Mine
63°10’ Saarinen
Pitkälampi
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Jaakonlampi Kortteinen
Sulkavanjärvi Särkilampi
Kuusilampi
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Sulkavanjärvi
27°44’
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536
Juuruvesi
Fig. 14.2. Geological map of the Siilinjärvi carbonatite complex, eastern Finland. After Puustinen (1971) with slight modification from recent mapping compiled by Lukkarinen (2000).
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complex area (Figure 14.2) that appears to have a lamprophyric character and may be related to the same intrusive event as the carbonatite (Puustinen, 1971). White-green medium-grained pure carbonatite formed during the carbonate-dominated pulse is relatively rare and in general true carbonatite (>50 modal % carbonates) is a relatively minor rock type at Siilinjärvi. The vast majority of the central body is formed of phlogopite-rich rocks ranging from almost pure glimmerite (biotitite) via carbonate glimmerite to silicocarbonatites and finally to carbonatites. Blue-green richterite forms up to 30% of the rock in places. Figure 14.4 shows the distribution of these rocks types, as mapped in 1979, in the area of the present-day open pit mine (Mikkonen et al., 1980; Härmälä and Liferovich, 2001). The phlogopite-rich nature of the Siilinjärvi intrusion is apparent from this diagram. Figure 14.5 shows some examples of the variety of rock types from the carbonate-rich pulse, and also displays some of the less common minerals, including zircon (14.5). Even though all varieties of this magmatic pulse contain apatite, apatite is nonetheless concentrated in the carbonate-rich rocks. The overall mode of the carbonatite–glimmerite portion of the complex, as indicated by the average composition of the Siilinjärvi ore (Härmälä, 2001), is 65% phlogopite (including tetraferriphlogopite), 20% carbonates (with a 4:1 calcite:dolomite ratio), 5% richterite and 10% apatite (equivalent to 4% P2O5 in the whole rock). Other, relatively rare accessory minerals at Siilinjärvi include barite, strontianite, monazite, pyrochlore, baddeleyite, ilmenite, magnetite, pyrite, pyrrhotite, and chalcopyrite. Fenites surrounding the carbonatite–biotitite central core developed as a result of Na metasomatism of the surrounding granite gneiss country rocks. The main minerals in the fenites are microcline, amphibole and pyroxene but there exists a wide variety of syenite types including: pyroxene, amphibole, carbonate,
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Fig. 14.3. Aerial photograph of the Siilinjärvi mine in 1997. The present mine is similar in shape, but somewhat deeper. Present production is 11.1 Mt of which 9.2 Mt is ore grading 4.2 wt.% P2O5 with 170 Mt of reserves and 380 Mt of probable reserves.
quartz, aplitic, and quartz-aegirine syenites. Compositions of the fluids that are likely to have produced these fenites have been determined from fluid inclusions within magmatic zircon and apatite (Poutiainen, 1995). Zircon crystals, which are found predominately in the amphibole-rich parts of the intrusion, contain two types of fluid inclusions trapped prior to emplacement of the carbonatites. Type 1 fluid is a H2O-CO2 mixture with low salinity (1–4 wt.% NaCl equivalent), whereas type 2 is of moderate salinity (7–18 wt.% NaCl equivalent), alkali- and H2O-rich. Type 1 inclusions surround rounded, presumably older zircon cores while type 2 inclusions surround type 1. Apatite crystals contain only type 2 inclusions and this is consistent with the fact that apatite crystallized predominately after zircon, although rare minute apatite daughter crystals in some of the type 1 inclusions in zircon indicate initiation of apatite saturation at this stage. Consequently, the development of H2Oand alkali-rich late-stage fluids that formed the fenite halo was a direct consequence of the early crystallization of predominantly
carbonate + apatite (Poutiainen, 1995). Ascent and hydrofracturing by the evolving H2O-rich fluid may have facilitated the ascent of these ultramafic and carbonatite magmas along deep crustal shears, with attendant fenitization along the path. A concordant zircon U-Pb age of 2609 ± 6 Ma (Olavi Kouvo, pers. comm.) shows that Siilinjärvi is one of the oldest carbonatites in the world. Ion microprobe analysis of the older rounded zircon cores (described above) remains to be done in order to determine their ultimate origin, i.e., crustal or mantle.
2.2. Proterozoic Kortejärvi and Laivajoki intrusions The Kortejärvi and Laivajoki carbonatites in the Koillismaa area of northeastern Finland (Figure 14.1) were discovered in 1961 as a result of mineral exploration by Rautaruukki Oy. The carbonatite bodies are located within the Hirvaskoski shear zone between the Kuhmo and the Pudasjärvi Archean blocks with amphibolites as the main country rocks and
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Glimmerite Carbonate & apatite glimmerite Silicocarbonatite Diorite porphyry Apatite rock Fenite Carbonatite Diabase Aplite
2000
50 m
Carbonate
6950
6950
1800
50 m
Apatite
1800 <10%
10–15%
1800 0–10%
>15% Waste rock
50 m
Richterite
6950
1800 Black mica Dark brown mica
Reddish brown mica
10–20% 20–30% 30–40% >40%
50 m
Phlogopite
6950
7400
7300
7100
7000
50 m
7200
1900
1800 <5%
5–10%
>10%
Fig. 14.4. Top: Geological map of an exposed portion of the Siilinjärvi complex at an early stage of mine development. Below: Modal distribution maps for apatite, carbonate, and richterite and an approximate phlogopite color distribution map for the same exposure. According to grade estimates, the only waste rock areas in this exposure (shown in gray in the mineral distribution maps) were the dikes, the larger blocks of syenite and a few parts of the apatite-poor glimmerite. Note that the apatite-rich portions of the exposure match relatively well with those areas rich in carbonate and where reddish brown mica is dominant. The latter is, at least in part, tetraferriphlogopite (Puustinen, 1973). The richterite-rich zones appear to cut the main mineralogical trends, and may have resulted as reaction by late-stage fluids. It is also notable that at Siilinjärvi, the overall amount of carbonatite, i.e., rocks with >50 modal % carbonate (in yellow, top map), is relatively limited; the bulk of the intrusion is formed by glimmerite. Modified after Härmälä (1981) and Mikkonen et al. (1980).
612
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B
A
C
D E Fig. 14.5. Examples of some of the Siilinjärvi complex rock types. (A) Medium-grained calcite carbonatite with apatite (greenish) and one 3-cm-wide phlogopite book; (B) Zircon crystals in silicocarbonatite; exposed portion of larger crystal is 0.8 cm; (C) Calcite (pink) carbonatite with large (5 cm) books of phlogopite; (D) Abundant, several cm long crystals of richterite in calcite and dolomite carbonatite; (E) Typical example of mixing between carbonatite and glimmerite with large fractured crystals of apatite. Photos by the first author for this figure and all succeeding photographs unless stated otherwise. C H A P T E R 1 4 • K I M B E R L I T E S , C A R B O NAT I T E S , A N D A L K A L I N E RO C K S •
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Laivajoki
Kortejärvi
Fig. 14.6. Aeromagnetic map of the zone between the Kuhmo and Pudasjärvi blocks. The Kortejärvi and Laivajoki carbonatites are easily distinguished because of their strong magnetic signatures (dark areas), due to the magnetite they contain. The bright spots in the carbonatite anomalies result from remanent magnetization of magnetite. Background image by Meri-Liisa Airo using Geological Survey of Finland databases.
lesser mica and quartz-feldspar gneisses. On aeromagnetic maps the carbonatite intrusions exhibit highly elongate shapes (Figure 14.6) concordant to the nearly vertical lithological layering of the country rocks, indicative of strong stretching during Proterozoic activation of the Hirvaskoski zone. Based on aeromagnetic data and limited drill results, Nykänen et al. (1997) estimated that the Kortejärvi deposit is about 60 m thick and 2 km long while Laivajoki is about 20 m thick and 4 km long. At Kortejärvi the main rock type is bluish to white calcite carbonatite with calcite grains averaging 1.4 mm but ranging up to 7.5 mm, and accessory minerals including dolomite (up to 10%), magnetite with ilmenite exsolution lamellae, tetraferriphlogopite, olivine, serpentine, tremolite, actinolite, apatite, allanite, and monazite. Significant amounts of slightly coarser yellowish dolomite carbonatite and 614
dolomite-calcite carbonatite also occur, the former containing the most apatite-rich layers at Kortejärvi. Silicate dominated rock types at Kortejärvi include: (1) dark green to brownish green glimmerite, in layers up to 4 m thick, composed of phlogopite grains 0.4 mm to 1.8 mm in diameter with minor actinolite, edenite and accessory calcite, dolomite, apatite, and rare diopside, zircon, allanite, sulfide, and magnetite; (2) olivine-magnetite rock layers less than 1 m thick, in which olivine varies from fresh to completely altered to serpentine, iddingsite and bowlingite, magnetite invariably contains ilmenite exsolution lamellae and accessory minerals include dolomite, tetraferriphlogopite, richterite, and zircon. These rocks differ from typical foskerites from, for example Phalaborwa in South Africa, in that there is no apatite. Laivajoki is dominated by rocks called tremolite-rich carbonatite by Nykänen et al. (1997) that show large modal variations in carbonate and tremolite contents but nevertheless average greater than 50% carbonates. Other minerals include magnetite (ranging from accessory to a main mineral), ilmenite, phlogopite, relatively abundant zircon, allanite, and a few sulfides. Apatite is rare, as in the Kortejärvi magnetite-rich rocks. Rarely pyroxene remnants can be seen within the tremolite grains, perhaps suggesting this variety of carbonatite was originally pyroxene-rich. The main true carbonatite rock at Laivajoki is calcite carbonatite, but here the carbonatite is almost pure calcite with only occasional grains of dolomite, no olivine or serpentine, and the phlogopite does not show the reverse pleochroism indicative of tetraferriphlogopite. Accessory minerals include apatite, actinolite, magnetite, ilmenite, tremolite, zircon, allanite, and a few sulfides. The other major rock type at Laivajoki is serpentine-talc-dolomite rock with the first two minerals as alteration products of olivine and, in addition to dolomite, it contains small amounts of phlogopite, magnetite, ilmenite, sparse sulfide grains, and rare
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apatite. Thin layers of glimmerite similar to that at Kortejärvi also occur at this locality. A single U-Pb zircon age of 2020 Ma is thought to represent the primary age of these carbonatites (personal communication by Olavi Kouvo in Vartiainen and Woolley, 1974), but the zircon population is heterogeneous and several younger zircon generations exist (Karhu et al., 2001). An earlier K-Ar age of 1875 Ma reported by Kresten et al. (1977) from Kortejärvi is probably a metamorphic age that represents the last major event along the Hirvaskoski shear zone.
2.3. Proterozoic lamprophyre dikes A relatively unusual group of mafic to ultramafic alkali-rich rocks, which typically form dikes and commonly carry diopside, phlogopite, and/or amphibole phenocrysts, are lamprophyres, a term based on the Greek for ‘glistening porphyry’ coined by C.W. von Gümbel in 1874. The predominance of essential phlogopite or amphibole phenocrysts and the lack of early feldspar indicate that the magmas that formed these dikes were particularly fluid-rich. Lamprophyres can be divided into three main groups based on overall mineralogy: calc-alkaline, alkaline, and ultramafic and each of these groups can be further subdivided based on mineral modes and mineral chemistry. Table 14.1 compares the mineralogical characteristics of calc-alkaline (exemplified by minette) and ultramafic lamprophyres with those of kimberlites and lamproites, which are discussed in a later section. For the lamprophyres, the increase in alkaline and ultramafic character shown by the range from calc-alkaline to alkaline to ultramafic varieties corresponds to a decrease in SiO2 content and accompanying changes in mineralogy from phlogopite and amphibole with some feldspar, to more SiO2-poor minerals such as perovskite, and finally to meliliterich and/or monticellite-bearing varieties with or without significant carbonate (Table 14.1).
Further discussion of the ultramafic lamprophyres is left for the section on the Sokli dikes (see below). Lamprophyres are most often found as dikes, sills and plugs, but may also form breccia pipes and lava flows. Examples of breccia structures, some with classic pipe form, include the alnöite diatremes north of the Alnö complex in Sweden (von Eckermann,1948; Kresten, 1990), the Bulljah lamprophyre pipe in Western Australia (Hamilton and Rock, 1990) and the Buell Park diatreme in the Four Corners area of Arizona, United States (Roden and Smith, 1979). Lamprophyres also form lavas, as attested to by the remnant volcanic necks in the Navajo area of the southwestern United States and more recent volcanism in the Colima graben of northwestern Mexico (Luhr and Carmichael, 1981; Allan and Carmichael, 1984) where volcanoes are built from lamprophyre lava flows. Most relevant to this section on Finnish Proterozoic lamprophyres are the calc-alkaline varieties, which include minette (essential phlogopite phenocrysts and alkali feldspar > plagioclase in matrix), vogesite (calcic hornblende, alkali feldspar > plagioclase), kersantite (phlogopite, plagioclase > alkali feldspar) and spessartite (calcic amphibole, plagioclase > alkali feldspar). The term calc-alkaline stems from their affinities to arc-related magmas in that they have very similar trace-element signatures albeit at higher overall enrichment levels than typical primitive arc rocks (Rock et al., 1991). First described by Hackman (1914, 1933), and later by Huhma (1981) and Laukkanen (1983), the lamprophyres from the lakes district (Figure 14.1) are represented almost exclusively by minette and kersantite. According to Laukkanen (1983) the dikes are found in four different areas. The largest swarm of 25 dikes occurs at Haukivesi, where they are 10– 100 cm wide. Their age is ~1840 Ma (Neuvonen et al., 1981). Although several of the dikes plot into the camptonite chemical field, all of
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Table 14.1. Mineralogical characteristics of kimberlites, lamproites, and lamprophyres. Mantle Olivine Mica
Xenoliths Xenocrysts Macrocrysts Phenocrysts Macrocrysts Phenocrysts Groundmass
Kimberlites C C C C C, phlogopite
Orangeites C C C C C, phlogopite C, tetraferriphlogopite Rare, Mg-chromite to Ti-magnetite
Lamproite rare rare rare C C, phlogopite to Ti-phlogopite C, Ti-tetraferriphlogopite Rare, Mg-chromite to Ti-magnetite
--C, Al- & Ti-poor rare, Sr- & REE-rich
--C, Al- & Ti-poor rare, Sr- & REE-rich C, Sr- & REE-rich ---
--C C, Al- & Ti-rich C, Al- & Ti-rich --C, Sr- & REEpoor C, Sr- & C, Sr- & REEREE-poor poor -----
--C C C C C C very rare C
rare rare abundant ------very rare C ---
Spinels
Groundmass
Monticellite Diopside Perovskite
Groundmass Groundmass Groundmass
C, phlogopite kinoshitalite abundant, Mg-chromite to Mg-ulvöspinel C --C, Sr- & REE-poor
Apatite
Groundmass
C, Sr- & REE-poor
Primary Serpentine Calcite Sanidine
Groundmass
abundant
abundant, Sr- & REE-rich C
Groundmass Phenocrysts Groundmass K-richterite Phenocrysts Groundmass K-Ba-titanite Groundmass Zr-silicates Groundmass Mn-ilmenite Groundmass Leucite Phenocrysts
abundant --------very rare very rare rare ---
C --rare --rare C C C rare pseudomorphs
Minettes --rare --rare C, phlogopite
UML rare rare rare C C, phlogopite
C, Al-biotite
C, Al-biotite
C, Mg-chromite C, Mg-chromite to Ti-magnetite to Ti-magnetite
C to abundant ------------rare ---
C = common, --- = absent, = critical K2L matching characteristic, = important K2L matching characteristic, = matching characteristic with evolved K2L endmember. UML = ultramafic lamprophyres. Table is modified after Mitchell (1995b).
them have very similar mineral compositions, and it is likely that crystal accumulation has had sufficient effect to displace some of the minettes into the camptonite compositional field. Figure 14.7 shows photomicrographs of representative samples from this locality. Note the predominance of phlogopite in the samples, indicative of minette. At Nilsiä nine minette dikes with widths of 4–40 cm are found. From Kaavi seven dikes are known, one giving a U-Pb zircon + titanite age of ~1830 Ma (Huhma, 1981). These dikes are 10–70 cm wide and comprise camptonite and kersantite. 616
In addition, one minette dike over 40 cm wide is known from the Pielavesi area. Undoubtedly greater numbers of lamprophyres occur in the Finnish basement than have been reported, but they are not easily recognized because of their dark, fine-grained nature and the fact that their mica-rich character renders them substantially less robust to weathering than their typical host rocks.
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A B Fig. 14.7. Photomicrographs of minette samples from the Lakes District of eastern Finland. Phlogopite occurs as phenocrysts and is the major mafic component of the groundmass. Diopside grains with reaction rims and one rounded resorbed quartz xenocryst (gray) are visible. Other samples contain greater amounts of magmatic hornblende. (A) Plane polarized light, (B) Crossed polars; width of field is 6.5 mm.
2.4. Proterozoic Halpanen carbonatite The small Halpanen occurrence was recognized as a carbonatite by the exploration staff of Rautaruukki Oy and was first described by Puustinen (1986). It is located 12 km northeast of the city of Mikkeli in southeastern Finland, along a major, deep north–south fracture that also includes the Siilinjärvi carbonatite complex (Figure 14.1). According to geophysical data the approximately 8-m-wide, shallowly dipping (30–35˚) dike-like body is about 1.5 km long. It consists of fine-grained, massive or weakly banded calcite carbonatite and has accessory minerals of apatite, magnetite, pyrite, barite, monazite, and fluorite with the first two in places as phenocrysts >1 cm in diameter. At the eastern contact of the dike within the main quarry there is a roughly 10-cm-wide apatite-rich zone that varies from 20–80% calcite and 20–80% apatite. Fenite alteration around the intrusion is limited to about 1 m, and is shown mostly by elevated SrO and BaO in the surrounding quartz-feldspar rocks. Preliminary age data from monazite give a Paleoproterozoic age of 1800–1700 Ma (Puustinen and Karhu, 1999).
2.5. Proterozoic Group II kimberlites – olivine lamproites (K2L) Kimberlite pipes are formed from ultramafic, volatile-charged, incompatible element-rich magmas that represent a mixture of liquid, mantle peridotite, and eclogite detritus carried from depth, and typically megacryst suite minerals such as titanian pyrope, magnesian ilmenite, and subcalcic clinopyroxene. There are two end-member kimberlite types, based on examples from South Africa: Group I with abundant large, rounded grains (macrocrysts) of olivine, in a matrix of subhedral to euhderal olivine, monticellite, perovskite, spinel, mica, calcite, and serpentine, and Group II typically with abundant phlogopite ± olivine in a matrix of phlogopite, potassium richterite, and other diagnostic minerals (Table 14.1). Olivine lamproites show some similarities to Group II kimberlites but exist, for example, in Western Australia (Argyle diamond mine), Montana and Wyoming (U.S.A.), and southern Spain, whereas no rocks absolutely identical to Group II kimberlites have been found outside of southern Africa. In the area of Kuhmo, in eastern Finland
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tPh KR
D
B
A
Fig. 14.8. Photomicrographs of Kuhmo K2L dikes. (A) Primitive olivine macrocryst-rich phlogopite K2L dike rock that also contains abundant euhedral olivine phenocrysts (serpentinized) in a matrix of Ti-rich phlogopite, potassium richterite, Mn-rich ilmenite, Cr-rich spinel zoned to titanian magnetite, apatite, perovskite, calcite, and serpentine (European Diamonds PLC, Lentiira Prospect). (B) Slightly more evolved K2L rock from Kuhmo (Malmikaivos Oy, Prospect no. 16; Seitaperä). Grains surrounding former pool of late stage liquid include tetraferriphlogopite (tPh), diopside (D), and potassium richterite (KR). Olivine pseudomorphs rimmed by perovskite in a matrix of phlogopite, apatite, and calcite are also apparent. Both images in plane polarized light; width of field is 2.3 cm (A) and 2.55 mm (B).
(Figure 14.1), there is a series of dike rocks that show mineralogical similarities to both olivine lamproite and Group II kimberlite (Table 14.1). In hand specimen, the most distinctive feature of the Kuhmo potassic, ultramafic rocks is their phlogopite-rich nature. Phlogopite occurs rarely as macrocrysts, but is abundant as phenocrysts and microphenocrysts with relatively Ti-rich compositions similar to those of lamproite microphenocrysts. The more primitive Kuhmo potassic rocks may also contain large amounts of olivine macrocrysts (Figure 14.8A) and in some cases abundant xenocrysts and xenoliths of mantle peridotite. Additional groundmass minerals include potassium richterite, Mn-rich ilmenite, Cr-rich spinel zoned to titanian magnetite, apatite and perovskite in a calcite + serpentine matrix. More evolved versions of this rock type contain abundant olivine phenocrysts (rimmed by perovskite; Figure 14.8B) rather than macrocrysts, low-Al clinopyroxene, and phlogopite that is zoned to low-Ti tetraferri618
phlogopite, similar to mica zoning trends in Group II kimberlites (Mitchell, 1995a). A suite of dikes and small breccia pipes ranging from leucite lamproite to olivine lamproite to Group II kimberlite has been identified and studied in the Kostamuksha region of Russian Karelia (Proskuryakov et al., 1990; Zhuravlev et al., 1995), about 40 km northeast of the Kuhmo occurrences. These rocks have been termed K2L by Mahotkin et al. (1998), in reference to their intermediate mineralogies between Group II kimberlite and olivine lamproite and it would appear that such a terminology would also be appropriate for the Kuhmo rocks. Use of the K2L terminology is, however, only a convenience to keep more rock names from entering the already crowded field of potassic ultramafic rocks. Another viewpoint is that the rock types at Kuhmo are sufficiently different as to warrant a new name, given that each craton appears to contain a characteristic potassic ultramafic magma type derived predominantly from metasomatized
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B
zones within the subcontinental lithospheric mantle. The source mineralogy of these magmas is governed by the particular metasomatic history of the mantle source and melts derived from these domains will have characteristics unique to that mantle (the so-called metasomatized mantle melts of Mitchell, 1995a). In this regard, it is interesting to note the general similarity of the K2L magmas of Kuhmo and Kostamuksha to the diamondiferous micaceous kimberlites of the Arkhangelsk area of Russia and the fact that they occur within the same Karelia–Kola–Kuloi cratonic block. Limited age-dating on perovskite from the Kuhmo dikes give ages of around 1230 Ma (Figure 14.9A). This age is consistent with a Rb-Sr mineral isochron age of 1231 ± 8.9 Ma for the K2L rocks of Kostamuksha (Belyatsky et al., 1995), and is clearly distinct from the much younger Devonian (~360 Ma; Sablukov, 1984) Arkhangelsk kimberlites.
2.6. Neoproterozoic Group I kimberlites C
Fig. 14.9. (A) U-Pb age data measured by ion microprobe on perovskite from the Kuhmo K2L Seitaperä locality. The spectrum of apparent ages is large, with the oldest intercept of 1230 Ma being close to age determinations for K2L rocks at Kostamuksha, Russian Karelia (Belyatsky et al., 1995). (B) and (C) U-Pb age data measured by ion probe on perovskite from two Kaavi–Kuopio Group I kimberlites. Although the age data do not overlap within error, an average of ~600 Ma is taken as the time of intrusion for the Kaavi–Kuopio kimberlites because of the strong geochemical coherency of the entire group.
Twenty kimberlite occurrences have been discovered so far in the combined Kaavi and Kuopio clusters in eastern Finland (Fig 14.10). All have typical Group I mineralogies, major and trace element compositions and intrusion morphologies (O’Brien and Tyni, 1999). They range from hypabyssal kimberlites, to tuffisitic kimberlite breccias (TKB) formed in steepsided funnel- or carrot-shaped pipes. Their morphology ranges from dikes 500 m by 30 m in size (Fig 14.11) to nearly circular diatremes up to 4 ha in size. None of the Kaavi–Kuopio pipes appear to have crater-facies materials extant, due to erosion of the upper portions of the pipes. The hypabyssal facies rocks are hard and compact, with dark gray to black matrices enclosing coarser minerals, particularly olivine, and crustal and mantle xenoliths (see Section 5.1). Pipe 1 (Figure 14.12A) may represent the deepest exposure of root zone material because of its well-developed segregationary
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50 km
Fig. 14.10. Location map of the kimberlites discovered in the Kaavi (eastern) and Kuopio (western) clusters on an aeromagnetic background by Maija Kurimo using Geological Survey of Finland (GTK) databases.
texture, which in some samples, particularly near the edge of the intrusion, develops into globular segregations in which late crystallizing serpentine and calcite form irregular pools in a more uniform silicate matrix. A more typical hypabyssal kimberlite from Kaavi–Kuopio is shown in Figure 14.12B, and displays the classic suite of lithospheric mantle-derived xenocrysts: olivine (pseudomorphs), red pyrope, gray magnesian ilmenite, and bright green chromian diopside. The diatreme facies rocks are much less well indurated and span the color spectrum from green to gray to brown to dark red. Diagnostic textures of the diatremes facies rock types include rounded pelletal lapilli in which kernels of small crustal xenoliths or mineral grains, particulary olivine (pseudomorphs), 620
are surrounded by serpentine- and carbonaterich matrix material. Juvenile magmaclasts are also common. The content of crustal material incorporated into the diatreme during formation is large (Figure 14.12C), raising silica contents from original levels of ~30 wt.% to 44 wt.% SiO2 or more (O’Brien and Tyni, 1999). Phenocryst and macrocryst olivine compositions from the Kaavi–Kuopio kimberlites show either a restricted compositional range, Fo92–87, or are bimodal, with a similar Fo93–89 population and an additional Fo86–83 population representing megacryst-suite olivines. Monticellite, mostly from 10 µm to 50 µm in size, is common in Pipe 1, rare in Pipe 14, found only as inclusions in mica and titanite in Pipe 10 and absent from the remaining less
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3.
5.
2.
4.
9.
10.
6. 14.
23.
100 m Fig. 14.11. Shapes and relative sizes of Kaavi– Kuopio kimberlite intrusions. Dark green represents pipes that are wholly or predominantly hypabyssal kimberlite and light green represents those that are predominantly tuffisitic kimberlite breccia, presumably formed in a similar fashion to the classic kimberlites from South Africa. Updated from O’Brien and Tyni (1999).
pristine pipes. Ba-rich mica (kinoshitalite) occurs in the groundmass of virtually all of the Kaavi–Kuopio kimberlite samples. Its abundance ranges from very sparse as in Pipe 9, to as much as 10% of the matrix, as in Pipe 10. Relative to the kinoshitalite micas reported from the Iron Mt. kimberlite (Mitchell, 1995), these examples range to extremely Ba-rich compositions (up to 17.8 wt.% BaO), especially those from Pipe 1. They also contain a large amount of fluorine. Spinel is common in all of these kimberlites, although the most spectacular examples are the atoll structures from Pipes 1, 2, 3, and 5 (Figure 14.12D). In thin section the translucent dark red cores of titanian aluminous magnesian chromite (TIMAC) show typical kimberlite magmatic
trend 1 zoning (Mitchell, 1986). An amorphous serpentine-like mineral that proves to be an alteration product of titanium-bearing, nearly chrome-free pleonaste spinel typically surrounds these cores; the pleonaste has only rarely been found intact. The succeeding mantle of magnesian ulvöspinel (MUM) is nearly unzoned and is uniformly surrounded by a thin outer rim of magnetite. Abundant apatite occurs as acicular grains commonly in radiating stellate clusters in the groundmass and as larger more prismatic grains grown primarily within calcite segregations. They are relatively Si-rich (0.7–1.1 wt.% SiO2) and Sr-poor (<1 wt.% SrO), characteristic of Group I kimberlite apatite (Mitchell, 1995a). The majority of the perovskite occurs as euhedral to subhedral discrete grains (rarely as aggregates) that are 0.02 mm to 0.1 mm across. Although only limited data exist, the perovskites appear to be typical of Group I kimberlites (op. cit.) with ~1.3 wt.% FeO, 1.5 wt.% Nb2O5 and 0.1–0.3 wt.% SrO. Mantle xenocrysts from these kimberlites include: (1) Mg-ilmenite that shows complex zoning and resorption features suggesting extensive magma mixing (see O’Brien and Tyni, 1999); (2) Pyrope garnet derived from a range of sources including high Cr, Ca-depleted harzburgite to Ca-saturated lherzolite and Ca-rich wehrlite, to Ti-rich, megacryst-compositions, and at lower MgO and higher CaO contents, orange garnets derived from mantle eclogite (see Section 5.2 and Lehtonen et al. 2004); (3) Clinopyroxene comprising lherzolitic, lowCr megacrystic, and eclogitic subgroups; (4) Spinels from upper mantle spinel lherzolites and rare chromites plotting within the diamond inclusion field.
2.7. Devonian Sokli carbonatite complex The Sokli carbonatite complex (Figure 14.1) was discovered in Finnish Lapland in 1967 when Rautaruukki Oy was prospecting for iron ore in the area of Finland nearest to the
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B
E
Fig. 14.12. Images of representative Kaavi–Kuopio kimberlites. (A) Sawn surface of hypabyssal kimberlite (Malmikaivos Oy, Prospect no. 1; Koskenniemi). Rounded olivine macrocryst have been resorbed and abraded as peridotite xenoliths were disaggregated during rapid kimberlite magma ascent from mantle depths. Width of picture corresponds to 3 cm. (B) Hypabyssal kimberlite containing abundant indicator minerals (Malmikaivos Oy, Prospect no. 10; Ryönä). The rounded indicator minerals are chrome-rich red pyrope, green diopside, and steel-gray magnesian ilmenite. In addition, the sample contains two generations of altered olivine grains. Mantle xenocryst compositions demonstrate that some sampling occurred at depths greater than 150 km, where diamond is a stable form of carbon,
622
• K I M B E R L I T E S , C A R B O NAT I T E S , A N D A L K A L I N E RO C K S • C H A P T E R 1 4
Table 14.2. Petrographic features of the magmatic carbonates at Sokli. Magmatic phase
Rock type
Typical minerals
Texture
I
Phoscorite
Olivine, phlogopite, magnetite, calcite, apatite, U-Ta-pyrochlore, baddeleyite
Massive, coarsegrained
II
Calcitic carbonatite
Calcite, phlogopite, olivine, magnetite, U-pyrochlore, baddeleyite
Massive, mediumor coarsegrained
III*
Altered phoscorite
Tetraferriphlogopite, clinohumite, richterite, iddingsite, olivine, magnetite, apatite, sulfides, Th-pyrochlore, zircon
Massive, mediumor coarsegrained
IV
Calcitic and dolomitic carbonatite
Dolomite, calcite, tetraferriphlogopite, richterite, apatite, pyrrhotite, Th-pyrochlore, zircon
Orientated, banded, fineor mediumgrained
V
Dolomitic late dikes
Dolomite, barytocalcite, sulfides, hematite, ancylite
Massive, partly amygdaloidal, medium- or coarse-grained
* = pneumatolytic-hydrothermal phase
and the sample is diamond-bearing. Width of image corresponds to 3 cm. (C) Sawn slab of tuffisitic kimberlite (Malmikaivos Oy, Prospect no. 14; Kaatronlampi). The pale and green xenoliths derive from the crust and consist principally of quartz and feldspar. The angularity of the crustal xenoliths is a function of relatively short transport distances within the kimberlite compared to the peridotite detritus in (A) and (B) above. The prospect contains microdiamonds. Width of image corresponds to 3 cm. Photos (A) through (C) by Helena Saarinen. (D) Backscattered electron image (BEI) from sample in (a) showing atoll spinels. In these examples, as in virtually all such grains, the middle zone of pleonaste spinel between the chromite core (only visible in one grain) and the MUM mantle, has been replaced by serpentine. Width of field is 460 m. (E) Backscattered electron image (BEI) from sample in (C) showing final stages of kimberlite crystallization. Laths of Ba-rich phlogopite (kinoshitalite) are growing across boundaries of final stage matrix minerals calcite (angular patches) and serpentine (dark green). Width of field is 460 m. C H A P T E R 1 4 • K I M B E R L I T E S , C A R B O NAT I T E S , A N D A L K A L I N E RO C K S •
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C
D
U(ppm)
K(%)
Th(ppm)
E B
A
F 1 U ppm
Magmatic
2
core
3
Weathered cap
4 5 6 km
G
Th ppm
Metasomatic carbonatite Metasomatites Fenite
624
Phosphorus ore Carbonatite core and transgressive dikes
• K I M B E R L I T E S , C A R B O NAT I T E S , A N D A L K A L I N E RO C K S • C H A P T E R 1 4
Russian Kovdor alkaline complex. When the dimensions of Sokli became apparent (total area ~20 km2), and its ore potential was appreciated, a comprehensive study of the complex was undertaken (Vartiainen and Paarma, 1979). After over 20 years of investigation, the Sokli complex is one of the better-known carbonatite complexes of the world. The Sokli complex consists of a concentrically zoned, funnel-shaped plutonic body (Figure 14.13) as indicated by deep seismic soundings (Paarma et al., 1981). The youngest part of the intrusion, the plug-like central magmatic core, is 2.5 km wide at the surface and 1 km wide at 5 km depth. This core is surrounded by a nearly continuous ring of slightly older, metasomatically affected carbonatites, followed outwards by a zone of metasomatized ultramafic silicate rocks and an unusually wide fenite halo (up to 1 km wide). Petrographic features of the central magmatic core carbonatites at Sokli are summarized in Table 14.2 and examples are shown in Figure 14.14. Gradational contacts have been documented from both drill core and exposed rock outcrops. The magmatic carbonatites were intruded in five major phases (Table 14.2). The rocks of each phase can be differentiated using petrographic, mineralogical, and compositional criteria. The second and
fourth carbonatite phases are dominant. The most curious group of rocks is formed by the late stage 5 carbonatite dikes (Vartiainen and Vitikka, 1993), which with further mineralogical study will undoubtedly add to the list of minerals recognized from the complex. Metasomatic carbonatites formed as a result of carbonitization of pyroxenites and coarse-grained magnetite olivinites, the oldest rocks of the complex. The slightly carbonitized variants are massive, the more strongly replaced ones generally banded. Metasomatic silicate rocks formed from the ultramafic rocks of the original intrusion through complex replacement and substitution processes in a manner that at its extreme resulted in almost pure phlogopite rock. As intermediate products, there are host rock types containing variable amounts of amphibole, aegirine and phlogopite as major minerals. Fenites surrounding the carbonatite–biotitite central core developed by Na metasomatism of the surrounding granite gneiss, amphibolite, and hornblende schist. The Sokli fenite halo is developed up to 3 km from the carbonatite core, and is manifested by the development of alkali feldspar, pyroxene (aegirine and aegirine-augite), alkali amphibole (arfvedsonite and eckermannite) and phlogopite giving the rock a greenish cast (Vartiainen
Fig. 14.13. (facing page) (A) Schematic cross-section through the Sokli carbonatite complex showing the fenite zone (salmon), the metasomatic silicate rocks (green), the metasomatic carbonatites (peach), and the youngest magmatic carbonatite pulse (orange) with associated transgressive dikes. Phosphorus ore is shown in yellow. (B) Bouguer gravity anomaly map of the Sokli carbonatite and nearby Tulppio olivinite ultramafic body to the southwest (Vartiainen, 1980). Data from 118 gravity stations were used by Seppo Elo of GTK to produce this image. (C) Aeromagnetic map of Sokli showing the obvious ring structure. (D) through (G) Ternary color image, K, equivalent U and equivalent Th radiometric maps of the Sokli complex area. The ternary image and the Th anomaly outline the area of the complex extremely well while K (intense phlogopite metasomatism) is concentrated in the NE sector of the intrusion. The U anomaly covers the area of the latest magmatic pulse, and along with Th, shows a plume in the down-ice direction (SE) probably representing carbonatite material transported by the last glaciation. Images (C) through (G) by Meri-Liisa Airo. All geophysical data are from GTK databases except gravity, which was supplied by the Finnish Geodetic Institute.
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A
B
C
D
Fig. 14.14. Photomicrographs of Sokli carbonatites. (A) Calcite-tetraferriphlogopite-apatite rock from magmatic pulse, plane-polarized light; (B) Same as (A), polars crossed; (C) Calcite-apatite-olivine-magnetite-pyrochlore rock from magmatic pulse. Plane-polarized light; (D) Same as (C), polars crossed. Width of field in all four images is 6.85 mm.
and Woolley, 1976). Levels of fenitization increase towards the carbonatite intrusion and in the zone of most intense fenitization proximal to the intrusion, K metasomatism is dominant, forming extensive phlogopitite and phlogopite-alkali amphibole rocks (Figure 14.13E). All of the Sokli area is covered by weathered bedrock of variable thickness. The most important mineralogical changes associated with weathering are the partial dissolution of carbonates, local replacement of phlogopite vermiculite, intense alteration of olivine, partial replacement of magnetite by hematite (martitization), alteration of pyrochlore and the total removal of sulfides. As Sokli lies within the continental ice-divide and within a topographic low, its weathered crust and 626
the associated regolithic phosphorus ore have survived the erosional action of the continental ice sheet rather well. The phosphorus ore developed from the carbonatite and underwent complex weathering, leaching, recrystallization, and lithification processes (Nuutilainen, 1973) driven by the prevailing tropical climatic conditions (Finland lay on the equator ~400 Ma ago). The end product is a reddish brown layer, averaging 26 m in thickness, which varies from solid rock to a soil in which the carbonate has been totally dissolved. Apatite, magnetite, hydrated mica, and patchy pyrochlore occur as restite minerals. Recrystallized phases include francolite (carbonate-fluorapatite), goethite, and manganese oxides (Vartiainen et al., 1990). Owing to this process the P2O5 values have
• K I M B E R L I T E S , C A R B O NAT I T E S , A N D A L K A L I N E RO C K S • C H A P T E R 1 4
363 Ma, proving that it is a member of the Devonian Kola alkaline province (Kramm et al., 1993).
2.8. Devonian Sokli ultramafic lamprophyre dikes
A
B Fig. 14.15. (A) Sawn slab of Sokli ultramafic lamprophyre dike showing abundant obvious steel gray magnetite, dark gray olivine, very dark gray to black phlogopite, and reddish-brown alteration product of olivine. Sample height is 3 cm. (B) Photomicrograph taken from the corresponding thin-section of the same rock. Olivine is mostly fresh, but approximately 20% has been converted to alteration minerals, in this case predominantly iddingsite. The phlogopite shows systematic zoning from a cloudy core, to clear subhedral mantles, to very late discontinuous tetraferriphlogopite rims (see large grain middle right). Calcite was late to crystallize, but abundant. The remaining minerals include magnetite, apatite, and richterite. Plane-polarized light, width of field is 5.0 mm.
been elevated from the original 4–5% to up to 30%. Rb-Sr dating of five Sokli samples, using biotite, carbonate, apatite, and whole rocks, give ages for Sokli ranging from 368 Ma to
Associated with many ultramafic alkaline complexes are ultramafic, phlogopite-rich, olivine- and carbonate-bearing dikes that are termed ultramafic lamprophyre. They typically have low silica contents and experimental evidence suggests that they represent very small partial melts of carbonated peridotite source rocks that compositionally can be gradational to Group I kimberlite (Dalton and Presnall, 1998). The best-known Fennoscandian ultramafic lamprophyre complex is the 560 Ma Alnö intrusion and peripheral diatreme pipes in eastern Sweden (Kresten, 1990). In some examples, such as West Greenland (Mitchell et al., 1999) ultramafic lamprophyres have been derived from sufficiently deep sources to carry xenocrysts of lithospheric mantle material including chromian pyrope, chromian diopside, and diamond. Everywhere within the Sokli complex, and also up to several km away from it, ultramafic lamprophyre dikes are found. They are generally dark rocks varying from reddish to greenish gray, occur as swarms with different orientations and, according to drilling results, average about a half meter in thickness. Vartiainen et al. (1978) divided the dikes into four subgroups based on texture: porphyritic, xenolithic (= autolithic), massive, and micarich. The last two groups dominate, while the first two groups occur mainly within the fenite zone and the bedrock area outside of the carbonatite proper. A typical example of a Sokli ultramafic lamprophyre is shown in Figure 14.15. The mineral composition of the Sokli lamprophyres varies considerably. In the porphyritic group, the proportion of phenocrysts varies between 27% and 37%. Typically,
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olivine is strongly altered and is absent from the groundmass. The groundmass is very fine-grained (~0.1 mm) and crystallized in the order opaques, phlogopite/calcite, and richterite. Rock fragments in the xenolithic variety within the carbonatite complex are autoliths, i.e., they derive from the rocks of the complex itself and include phoscorite, carbonatite, and fenite. According to their chemical compositions (Table 14.3), the Sokli rocks are ultramafic lamprophyres in the classification of Rock et al. (1991) and their mineralogy is consistent with this categorization (Table 14.1). Minerals characteristic of kimberlite are lacking from the lamprophyres at Sokli and additionally they contain richterite, which does not occur in kimberlites. Ti-rich garnet and phenocrystic diopside, generally typical of ultramafic lamprophyres, are apparently absent. A heavy mineral study of the esker covering the Sokli complex established that this material does not contain critical minerals indicative of diamond-bearing kimberlites (Perttunen and Vartiainen, 1992).
2.9. Devonian Iivaara alkaline complex One of the many alkaline complexes associated with the Devonian Kola alkaline province, the 373–363 Ma (Doig, 1970; Kramm et al., 1993) Iivaara intrusion was one of the first in this province to be well studied (Ramsay and Berghell, 1891; Hackman, 1900) and became the type locality for the alkaline rock type ‘ijolite’ (see below). Oval in shape and covering 8.8 km2, the intrusion is relatively poorly exposed, and despite the existence of nine diamond drill holes the structure and mutual relationships of the various rock types is inadequately known. The intrusion at Iivaara can be divided into three main zones: an outer fenite zone, transitional rocks, and the main central mass of alkali rocks (Figure 14.16). Fenite formed as a result of the effusion of Na-rich fluids 628
derived from the intrusion and from the numerous dikes injected into the country rocks. Progressive metasomatic effects are seen as a change from typical granodiorite minerals to those having turbid feldspars and biotite reacting to aegirine to rocks with coarser grain size, nearly all biotite replaced by aegirine and the incipient development of albite. In the aegirine syenite stage, all quartz has been consumed, the feldspar is microperthitic, and biotite is absent. Calcite may be abundant, and wollastonite becomes apparent. Accessory minerals include titanite, apatite, cancrinite, and fluorite. The final stages of fenitization produces a cancrinite syenite in which nepheline appears. Near the summit of Iivaara, within a few dozen meters of the main intrusion, a zone of mixed rocks occur, in which the cancrinite syenitic fenite is brecciated by cancrinitenepheline-wollastonite rocks. In other parts of the transition zone, a dark, fine-grained pyroxene-amphibole-plagioclase rock exists, although the generally poor exposure makes the absolute position of this unit uncertain. Mafic minerals comprise >70% of the rock, and show large modal variations among hornblende, aegirine-augite, and biotite. Plagioclase is approximately An45 in composition. The entire central area of the Iivaara complex consists of nepheline-clinopyroxene rocks, urtite (nepheline > 70 modal %), ijolite (nepheline 30–70 modal %), and melteigite (nepheline <30 modal %). Nowhere in the central mass are there exposures of just one of these rock types. Modal variations are large on a small scale, but Lehijärvi (1960) suggested that in general ijolite is the dominant rock type, while melteigite is more abundant near the margins of the central mass on the northwestern slope of Iivaara, and urtite is concentrated near the summit. Cross-cutting dikes of modally and texturally different ijolite series rocks are ubiquitous, and it is clear that many pulses of magma were injected to form the central alkaline mass, negating any
• K I M B E R L I T E S , C A R B O NAT I T E S , A N D A L K A L I N E RO C K S • C H A P T E R 1 4
B
A
Iivaara Pieni-Näätälampi
Area enlarged at left
Iso-Näätälampi
Ahvenjärvi Ahvenvaara
Penikkavaara
Ijolite–urtite Melteigite–ijolite–urtite Microijolite and melteigite Fenite and hybrid rock
1 km
248 54 -123 -358 nT
Fig. 14.16. Geological map and aeromagnetic image of the Iivaara complex. (A) The geological map of the intrusion is very schematic as exposure is poor, and the rock types are highly mixed. (B) Aeromagnetic map over the Iivaara complex showing the obvious nearly circular structure of the complex. Aeromagnetic image by Meri-Liisa Airo using Geological Survey of Finland databases.
simple models of intrusion and subsequent differentiation. Mineralogically the rocks of the central mass are relatively straightforward (Figure 14.17). All contain aegirine-augite, which is typically zoned with higher Na and Fe contents toward their rims. Although nepheline ranges from gray to faintly reddish, it appears to have a very consistent composition throughout the intrusion with ~16 wt.% Na2O and ~6 wt.% K2O. Cancrinite is a common accessory mineral and is predominantly CO2-rich although some sulphatic cancrinite also occurs. Schorlomite (originally called iivaarite from this locality) commonly occurs in the ijolites and melteigites and crystals in late stage veins may show spectacular growth zoning in thin section (Lehijärvi, 1960; Figure 14.17D). Fluor-
hydroxylapatite represents about 3 modal % of the ijolites whereas titanite is concentrated in the melteigites. Other minerals that occur sparsely in the central alkaline mass, and/or in the surrounding fenites include hornblende, biotite, natrolite, sodalite, analcime, and pectolite.
3. Geochemistry of kimberlites, carbonatites, and alkaline rocks Although the range in composition of the undersaturated alkaline rocks and carbonatites in Finland is large (Table 14.3), several important characteristics and inferences about their sources tie this extremely varied group of rocks together: (1) They all represent low melt
C H A P T E R 1 4 • K I M B E R L I T E S , C A R B O NAT I T E S , A N D A L K A L I N E RO C K S •
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B
A
D C Fig. 14.17. Photomicrographs of Iivaara rocks from drill core samples. (A) Ijolite displaying aegirineaugite and fresh nepheline with inclusions of apatite and aegirine-augite. Plane polarized light, width of field is 6.5 mm. (B) Well-developed sector zoning is visible in brown-gray aegirine-augite grain at bottom right. Same view as (A), polars crossed. (C) Melteigite composed almost entirely of aegirineaugite with magnetite and melanite garnet filling the angular spaces between aegirine-augite grains. Well-developed sector zoning is displayed by one of the melanite grains (lower right). Plane polarized light, width of field is 6.5 mm. (D) Same as (C), crossed polars.
fraction, small volume melts formed preferentially from the melting of minor, more easily fused mantle minerals, e.g., phlogopite; (2) They are all formed from volatile-rich magmas. Carbonatites represent one end-member of this spectrum and in their extreme comprise pure carbonate magmas. Group I kimberlites represent another end-member of extremely volatile-rich magmas, with perhaps up to 25 wt.% volatiles of CO2 – H2O mixtures (Price et al., 2000). The lamprophyres may also contain a significant C-bearing volatile component 630
(e.g., ultramafic lamprophyres), but the more intermediate lamprophyre types, such as the minettes, typically have only a small CO2 component, and are instead water-dominated (Rock et al., 1991); (3) As a consequence of the high volatile content and the relatively depolymerized nature of the resulting melts, all of the magmas discussed here have very low viscosities. This enables such small volume melts to traverse thick lithosphere without undergoing “thermal death” and in many cases their ascent to the surface may
• K I M B E R L I T E S , C A R B O NAT I T E S , A N D A L K A L I N E RO C K S • C H A P T E R 1 4
Table 14.3. Representative analyses of carbonatites and undersaturated alkaline rocks in Finland. Rock Type
Halpanen
Sokli
Iivaara
Sokli
Niinilampi
Kuhmo Group II
Glimmerite Silicocar- Carbona- Carbonabonatite tite tite SiO2 40.76 14.97 1.59 0.62 TiO2 0.30 0.12 0.02 <0.01 Al2O3 10.13 1.70 0.10 0.01 Fe2O3 5.00 0.59 0.70 0.39 FeO 3.13 0.09 MnO 0.29 0.01 0.09 0.15 MgO 24.58 5.84 2.62 0.34 CaO 0.92 36.65 50.95 53.48 BaO 0.07 0.63 SrO 0.47 1.07 Na2O 0.30 0.81 1.42 <0.05 K2O 10.20 3.61 1.20 <0.01 P2O5 0.75 24.80 0.35 0.54 F 0.67 1.48 Cl 0.04 S 0.19 CO2 1.30 41.00 42.51 H2O 2.29 0.21 99.95 100.42 100.26 O=F;Cl;S 0.28 0.62 0.05 Total 99.67 99.80 100.27 100.21
Phoscorite 16.1 1.3 2.3 19.9 9.0 0.3 15.2 16.2 0.06 0.17
Ijolite
UML
CAL
46.15 0.38 15.70 6.59
27.00 2.30 2.00 14.10
43.40 1.98 12.71 7.55
Kimberlite 42.95 3.24 4.01 9.93
Kaavi– Kuopio Group I Kimberlite 32.88 2.33 5.27 11.99
0.18 5.52 14.16
0.37 20.70 13.20
0.34 8.64 9.67
7.24 2.61 0.77
1.20 2.20 2.50 0.55
3.06 4.66 3.44
0.14 21.68 5.44 0.14 0.06 0.40 5.73 0.12
0.26 24.08 10.29 0.13 0.09 0.09 0.75 0.71
0.66 9.50 3.20 99.48 0.40 99.08
0.27 2.56 0.95 99.23 0.07 99.16
0.25 1.35 4.71 99.90 0.06 99.96
0.03 0.06 1.32 9.49 99.74 0.02 99.72
Hackman O’Brien & (1914) Tyni (1999)
op. cit.
Ref.
Siilinjärvi
Puustinen (1971)
op.cit.
op. cit.
Puustinen & Karhu (1999)
2.9 7.4
3.8 8.1 0.93 102.73 0.99 101.74
100.23
Vartiainen (1980)
Lehijärvi (1960)
be facilitated by migration along pre-existing veins and fractures; (4) The three proceeding characteristics necessarily produce a fourth. All of these magmas inherit incompatible element-rich compositions, in some cases dominantly from their source regions, and in others dominantly by reacting with veins in the lithospheric mantle as they traverse ancient metasomatized lithospheric mantle. The incompatible element-rich nature of these magmas is clearly shown in a chondritenormalized REE plot (Figure 14.18). All of the magmas discussed here show moderate to extreme enrichments in the light rare earth elements (LREE) relative to typical basalts. In detail the Kuhmo K2L dike rock is extremely LREE enriched, displays a sigmoidal REE pattern and has relatively low Yb and Lu, which are all characteristics of the Arkhangelsk east-
Vartiainen et al. (1978 )
ern kimberlites and the Karpinskiy lamproite (Mahotkin, 2000). At slightly lower overall enrichment levels, and with concave rather than sigmoidal patterns, the Kaavi–Kuopio kimberlites REE patterns are fairly typical of Group I kimberlites in terms of concentration levels, but do not have the commonly seen linear Group I kimberlite REE patterns (Mitchell, 1986). The Finnish carbonatites are LREE enriched, with nearly linear patterns, and have uniformly higher Yb and Lu concentrations than in the silicate dominated rocks. The latter observation is probably due to a lesser role for garnet in the carbonatite melting reactions. The Lakes District lamprophyres show surprisingly strong LREE enrichments to levels well above what is typical for calcalkaline and alkaline lamprophyres (Rock et al., 1991), even up to the levels of the Kuhmo
C H A P T E R 1 4 • K I M B E R L I T E S , C A R B O NAT I T E S , A N D A L K A L I N E RO C K S •
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104
Sokli UML Siilinjärvi carbonatite Kuhmo K2L Lakes District lamprophyre Sokli carbonatite Iivaara ijolite Laivajoki–Kortejärvi carbonatites Kaavi–Kuopio kimberlites
Sample/Chondrite
103
102
101
100 La
Ce
Pr
Nd
Pm
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Fig. 14.18. Chondrite-normalized REE abundances for selected kimberlites, carbonatites, and undersaturated alkaline rocks. All of the rocks plotted show strong LREE enrichment and overall high levels of incompatible elements except the Iivaara ijolite, which is only moderately LREE enriched. Data sources: Iivaara–Laajoki and Makkonen (1994), Laivajoki and Kortejärvi–Nykänen et al. (1997), Sokli and Siilinjärvi carbonatites–Hornig-Kjarsgaard (1998), Sokli ultramafic lamprophyres–Vartianen et al. (1978), Lakes District lamprophyre–Laukkanen (1983), Kuhmo and Kaavi–Kuopio kimberlites–O’Brien and Tyni (1999). Normalization values from Boynton (1984).
K2L dike rock. Even more extreme values are registered in the more evolved minettes, but as crystal fractionation has played a role in this enrichment, they have not been plotted in Figure 14.18. In contrast, the Iivaara ijolite, related as it is to incompatible element-rich nephelinite magmas, shows only moderate LREE enrichment. A Nb-Zr diagram (Figure 14.19) helps to distinguish two principal types of trace element enrichment patterns characteristic of this suite of rocks, the kimberlite trend and the lamproite trend. Lamproites typically have Zr well in excess of Nb, and interestingly, the Lakes District lamprophyre plots in the direction of 632
lamproite enrichment. As would be expected, the Kaavi–Kuopio kimberlites show a Group I kimberlite enrichment, although more in the direction of the Aries (Western Australia) and Koidu (western Africa) kimberlite mantle sources than that which characterizes South African kimberlites (Taylor et al., 1994). Also plotted is the Zr-Nb correlation line for pristine hypabyssal kimberlite samples from the Jericho kimberlite in the Northwest Territories of Canada (Price et al., 2000). Note that the Kaavi–Kuopio kimberlite analyses plot almost directly on this line, which is believed to represent dilution of the highest Nb-Zr contents by the addition of mantle debris represented by an
• K I M B E R L I T E S , C A R B O NAT I T E S , A N D A L K A L I N E RO C K S • C H A P T E R 1 4
600
500
Alkaline rocks Kaavi–Kuopio kimberlites Kuhmo K2L Sokli UML Kortejärvi–Laivajoki carbonatite Lakes District lamprophyre
Aries 400
Koidu
Nb (ppm)
S. Africa Group I kimberlite Jericho
Olivine lamproites
200
W. Kimberley leucite lamproites
100
S. Africa Group II kimberlite 0 0
200
400
600
800
1000 Zr (ppm)
1200
1400
1600
1800
Fig. 14.19. Zr vs. Nb covariation plot for selected samples. Note that most of the rocks plot in the kimberlitic enrichment trend, with only the lamprophyres from the Lakes District showing substantially higher Zr/Nb ratios. The Kaavi–Kuopio kimberlites plot almost exactly along the line of pristine kimberlite melts from Jericho (Price et al., 2000). Data sources as in Figure 14.18.
increase in the content of olivine macrocrysts. A similar interpretation can be made for the Kaavi–Kuopio kimberlites with the highest Nb and Zr contents occurring in Pipe 1, the most massive hypabyssal kimberlite example from Finland (Figure 14.19). A similar kimberlitic enrichment trend, although at slightly higher Zr values, is displayed by the K2L rocks from Kuhmo (Figure 14.19).
4. Isotope composition of kimberlites, carbonatites, and alkaline rocks Mantle sources for kimberlites, carbonatites, and alkaline rocks can vary from dominantly
subcontinental lithospheric mantle, which in Finland is represented by the Archean Karelian craton, to those derived from the underlying asthenosphere. After much debate and the accumulation of a large body of data, it is still unclear what is the main mantle source for many of these magmas. However, radiogenic and stable isotopes provide a means to determine whether a particular magma is mostly derived from asthenospheric sources or contains a major contribution from old, aged, veined lithospheric mantle. Plotted in Figure 14.20 are the Sr and Nd isotope compositions of the rocks discussed here from Finland for which there are data, along with selected reference groups. The data for the Group I kimberlites from Kaavi–Kuo-
C H A P T E R 1 4 • K I M B E R L I T E S , C A R B O NAT I T E S , A N D A L K A L I N E RO C K S •
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Sokli +5
Other KAP carbonatites S. African Group I Kimberlite
Yo Ca ung rbo nat it
es
Arkhangelsk
0
Kaavi– Kuopio Group I
S. African transitional Arkhangelsk
εNd
-5
Tres Rancho (Brazil)
Kostamuksha Kuhmo K2L
-10
S. African Group II (Kimberlite)
Iivaara
-15
-20
0.702
0.704
0.706
0.708
0.710
87
Sr/86Sr
Fig. 14.20. Plot of initial 87Sr/86Sr vs. εNd for selected kimberlites, carbonatites, and undersaturated alkaline rocks.Young carbonatite box includes data from all carbonatites <200 Ma. Data from Sokli and other Kola alkaline province (KAP) carbonatites plot within or very near this box. Preliminary Nd isotope data from Siilinjärvi and Laivajoki–Kortejärvi also plot within the same range (Karhu et al., 2001) but Sr isotope data are still lacking. Data sources include Smith (1983), Harmer (1999), Mahotkin et al. (2000), Kramm (1993, 1994), and Belyatsky et al. (1995).
pio have been recalculated from O’Brien and Tyni (1999) to the correct age of 600 Ma. Recalculating using these older ages has significantly reduced the spread in the Sr-Nd isotope field for the kimberlites. In terms of Nd and Sr isotopes, the Finnish kimberlites are typical of Group I kimberlite compositions worldwide. This worldwide uniformity (Smith, 1983) strongly suggests that the source of Group I kimberlite is either well-mixed mantle, i.e., asthenosphere, or lithospheric mantle that has been converted physically by heating and chemically by melt infiltration into 634
asthenosphere-like mantle. Either way, Group I kimberlites do not directly provide information on the isolated aged roots of cratons they transect except through the xenoliths and xenocrysts they contain (see below). Group II kimberlites and related olivine lamproites, however, have Sr-Nd isotopic compositions that reflect long-term storage of Rb- and LREE-enriched mantle rocks separate from the asthenosphere (Figure 10.20). In addition to the compositional and mineralogical similarities of the K2L rocks from Kuhmo and Kostamuksha, and some of the diamondifer-
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ous rocks from Arkhangelsk, there is also a similarity in terms of isotope composition (Figure 14.20). However, as mentioned above, the Kuhmo and Kostamuksha dikes and pipes are considerably older, ~1230 Ma (Figure 14.9) vs. 365 Ma (Sablukov, 1984). The enriched isotope signatures of the Kuhmo and Arkhangelsk rocks indicate that the Karelian subcontinental lithospheric mantle (SCLM) contains veins of mica and amphibole, along with trace minerals that over time have developed extreme isotope compositions. In fact it is apparent that a strongly enriched Karelian craton SCLM already existed by 3.1 Ga (Peltonen et al., 2003). Sublithospheric magmas were either contaminated by material from these veins to produce the isotope signatures by mixing (e.g., O’Brien et al., 1995) or these magmas represent direct melts from metasomatized Karelian craton SCLM (e.g., Foley et al., 1999). Most carbonatites have similar, although not exactly the same, Sr-Nd isotope compositions as Group I kimberlites (Figure 14.20), and the straightforward interpretation is that they both originate in the well-mixed asthenospheric mantle. However, experiments on melting of carbonated peridotite indicate that most carbonatites should be produced at the solidus inflection at a depth of around 100 km, well within the lithospheric mantle (Wyllie and Lee, 1999). To resolve this dilemma it has been proposed that the asthenospheric isotope signatures result from multiple episodes of invasion and freezing of carbonatite melts in the lithospheric mantle, rapidly building zones of carbonated wehrlite. Melting of these modified zones can then produce carbonatites at 100 km depth with asthenospheric isotope compositions (Harmer, 1999). However, this process cannot take 100’s of million of years otherwise the isotope signatures would indicate an aged enrichment. The three Finnish carbonatites for which there are isotope data, despite their wide range in ages, all plot within the isotope space of
present day carbonatites (Figure 14.20). Sr and Nd isotope compositions of the Siilinjärvi carbonatite are near bulk earth both for Nd (Karhu et al., 2001) and Sr (εSr ~0 based on 87Sr/86Srmeas. = 0.701423 ± 7 for a calcite separate with very low Rb and a 2.6 Ga age; Hugh E. O’Brien, unpublished data). Laivajoki and Kortejärvi have a more depleted Nd isotope composition with initial εNd values of +1 to +3 (Sr data are lacking). Finally, Sokli has Sr and Nd isotope compositions suggesting a source with the lowest Rb/Sr and most LREE-depleted of the three carbonatites with initial εSr of –15 and εNd of +6 to +7 (Kramm, 1993). These values are very close to those of other carbonatites from the Kola alkaline province (Kramm, 1993) and the Kaavi–Kuopio Group I kimberlites described earlier (Figure 14.20). Because of the low abundance of C in the mantle, the volume of the source rock/melt required to generate carbonatite melt may be 1000–10,000 times higher than the volume of the carbonatite itself. Therefore, carbonatites (and carbonate-bearing kimberlites) can be expected to give a good estimate of the average C isotope composition of their ultimate mantle source. The Kaavi–Kuopio kimberlites contain 10–15% calcite, generally present as finegrained disseminations in the groundmass. Isotope compositions of C from four separate pipes give δ13C values ranging from –2.2‰ to –4.6‰ (PDB), with an average value of –3.5‰ (Peltonen et al., 2000). Surprisingly, carbonatite intrusions from the Karelian domain have δ13C values in the same range, irrespective of the age of the intrusion. Carbonates from the Archean Siilinjärvi carbonatite complex range from –3.1‰ to –4.5‰ (n = 8; Karhu et al., 2001), the Proterozoic Laivajoki and Kortejärvi carbonatites from –3.6‰ to –4.9‰ (n = 8; Nykänen et al., 1997) and the Devonian Sokli carbonatite from –2.7‰ to –4.1 ‰ (n = 10; op. cit.). Accordingly, no significant differences in δ13C signature can be shown to exist among the various carbonatite intrusions nor between carbonatites and the kimberlite pipes.
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Worldwide carbonatite complexes show only slightly more variation in their δ13C values with 91% of all carbonatites having δ13C values between –8‰ and –2 ‰ (Deines, 1989). The most likely explanation for this uniformity over several billions of years of kimberlite and carbonatite formation is that they are derived from a uniform asthenospheric C reservoir. This explanation is bolstered by the similarity in Nd and Sr isotopes described above between Group I kimberlites and the carbonatites. The isotope composition of the Iivaara rocks (Figure 14.20) is very different from the rest of the rocks in the Kola alkaline province showing strongly negative εNd from –9 to –19 and elevated initial Sr from 87Sr/86Sr = 0.70427 to 0.705752 (Figure 14.20). Kramm (1994) believed this is a result of extreme reaction of the ijolite magmas with country rocks and provides good evidence that the data lie on a mixing curve between Kola carbonatite compositions and fenites that formed from the surrounding Archean basement.
isotope dating has documented the close correlation between the ages of the stabilization of SCLM and formation of the overlying crust. In many shield areas (Kaapvaal, Siberia, Wyoming, Tanzania) the crust and mantle have remained coupled for billions of years (Pearson, 1999). Detailed study of the SCLM is also crucial for our understanding of crustal processes. This is because major modifications in the SCLM, e.g., by thermal erosion of the base of the SCLM by plume activity, or tectonic processes, e.g., rifting may cause uplifting, magmatic activity or formation of world-class mineral deposits in the crustal part of the lithosphere. Until recently, lack of mantle samples from the Fennoscandian lithospheric mantle has prevented the study of many fundamental aspects of this section of the lithosphere. Discovery of diamondiferous kimberlite pipes in eastern Finland has substantially improved the situation by providing us with mantle samples (xenoliths) from depths between 100–230 km (Kukkonen and Peltonen, 1999).
5. The kimberlite mantle samples
5.1. Mantle xenoliths
Most Archean cratons are underlain by anomalously thick (typically ~200 km), cold mantle keels generally distinguished by fast and anisotropic seismic velocities, relative to the underlying asthenospheric mantle (e.g., Polet and Anderson, 1995). Petrological studies of orogenic lherzolite massifs (exposed lithospheric mantle sections within continental shear zones) and mantle xenolith suites recovered from kimberlites and lamproites have implied that these keels consist of mantle peridotite depleted in basaltic constituents such as Ca, Al, and Fe (Boyd and Mertzman, 1987). Recent studies have also demonstrated a secular evolution in the composition of the subcontinental lithospheric mantle (SCLM) peridotites, which become less depleted from Archaean through Proterozoic to Phanerozoic time (Griffin et al., 1999). Progress in Re-Os
Three types of mantle xenoliths have been recovered from the Kaavi–Kuopio kimberlites: garnet-spinel peridotite, garnet peridotite, and bimineralic eclogite, some of which are highly diamondiferous. Garnet-spinel peridotites, most of which originate from depths of ~100–130 km, are all highly depleted, fine-grained granuloblastic and equigranular harzburgites. They have a cryptic and modal metasomatic overprint with the development of minor hydrous phases. The garnet peridotites are compositionally and texturally distinct from the garnet-spinel peridotites: They are of deeper origin (170–230 km) consisting of coarse-grained harzburgites, lherzolites and wehrlites, all being texturally similar to lithospheric mantle peridotites of other cratons such as Kaapvaal and Siberia. However, their Nd and Sr isotope composition is not typical
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Fig. 14.21. Diagram illustrating the possible configuration of the major lithospheric domains within the central Fennoscandian Shield. The stabilization age of the upper SCLM is believed to broadly coincide with the formation ages of the overlying crustal parts of the lithosphere. The middle part of the SCLM has a distinct composition relative to the upper SCLM. Xenolith and xenocryst data show that it is relatively coarse-grained and is composed of lherzolite, harzburgite, and minor wehrlite. The lowermost lherzolite layer may represent asthenospheric melt-infiltrated Archean peridotite, but even younger additions, related, for example, to the 1.88 Ga Svecofennian accretion event, the 1.6 Ga rapakivi granite event, and even the 0.6 Ga kimberlite magmatism, remain a possibility. Location of the Kaavi–Kuopio kimberlites and the approximate locus of origin of the studied mantle xenoliths (white stars) are indicated.
for Archean continental mantle, being more akin to off-craton lithospheric peridotites (Peltonen et al., 1999). The eclogite xenoliths are bimineralic Group I eclogites whose chemical and isotope compositions suggest that they represent mafic mantle cumulates rather than recycled ancient oceanic crust. Thermobarometry indicates that they have been derived from depths comparable to the deepest garnet peridotites, and they are inferred to occur as highly diamondiferous thin layers or pods within coarse-grained garnet peridotites (Peltonen et al., 2002). Probably the most important contribution of the mantle xenolith study is that SCLM within the central Fennoscandian Shield is compositionally and texturally heterogeneous.
This is also clearly apparent in the compositions of xenocrystic garnet (see next section). The sketch illustrated in Figure 14.21 attempts to combine the xenolith data with what is currently known about the geodynamic evolution of the craton margin. The uppermost harzburgitic part of the SCLM probably represents lithospheric mantle that was isolated from the convecting mantle at the time of the formation of the overlying crust. The present geometry of the boundary between the Karelian and Svecofennian upper SCLM was determined by the initial rifting of the craton ~2.0 Ga and by subsequent accretion of the Svecofennian lithosphere onto the craton margin. This lithospheric boundary is likely to be almost vertical, because the deepest Archean peridotite
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Fig. 14.22. Gem-quality xenocrysts from the Lahtojoki kimberlite in Kaavi pipe cluster, eastern Finland. The diamond octahedron is unresorbed, plane-faced and 1.2 mm in diameter (classified commercially as D-F,VS). Unresorbed diamond crystals are rare (about 10% of the diamonds) in the matrix material of the Kaavi kimberlites but common where still retained within xenoliths. The color of the garnets is indicative of their chemical composition: titanian pyrope (red), chromian pyrope (violet) and eclogitic garnet (brown). The dissolution of garnet and chromian diopside (green) to smooth grains is the result of boiling in caustic solution at the VTT mineral processing plant. The minerals were hand-picked under a stereomicroscope by Kari A. Kinnunen of GTK from the final heavy mineral concentrate of crushed kimberlite ore. Photo: Kari A. Kinnunen.
xenoliths originate from depths of ~180 km (Figure 14.21). The lower lithospheric mantle is distinct in texture and composition from that of the upper SCLM (Figure 14.21). As implied by the xenolith petrography and thermobarometry, the lower SCLM is composed of relatively fertile, coarse-grained garnet peridotite with some diamondiferous eclogite layers. No direct age determinations are available, but U-Pb dating of kimberlitic and lower crustal zircons indicate that the formation of the lower lithospheric mantle and mafic lower crust and the emplacement of postorogenic 1.8 Ga granites can all be related (Hölttä et 638
al., 2000; Peltonen and Mänttäri, 2001). This can be explained by a model in which a plume impinged on the base of the 3.5–1.9 Ga old lithosphere during the postorogenic stage resulting in thermal erosion of the deepest part of the pre-existing lithospheric mantle and its replacement – or strong refertilization – by asthenospheric melts originating from the 1.8 Ga plume head. The xenolith data thus imply that the SCLM beneath the Karelian craton margin is at least 230 km thick (Kukkonen and Peltonen, 1999; Kukkonen et al., 2003). This is, however, a minimum estimate as the deepest xenoliths originating from this level are still typical coarse-grained lithospheric peridotites with no evidence for being in the vicinity of the lithosphere–asthenosphere boundary. Although no “sheared” xenoliths are present in our sample suite, it is reasonable to assume – by inference with xenolith studies from other cratons – that the coarse garnet peridotites are underlain by a layer of “sheared” peridotites with mylonitic textures. Traditionally, such sheared lithospheric peridotites have been considered to be of asthenospheric origin (Nixon and Boyd 1973), but recent Re-Os isotope results have indicated that in most cases they give ancient formation ages, and thus actually represent the lowermost parts of the ancient cratonic roots which have been infiltrated by asthenospheric melts (Pearson, 1999). Therefore, we can infer that the true lithosphere–asthenosphere boundary within the Karelian craton margin lies some tens of km beneath the maximum depth of 230 km indicated by xenolith thermobarometry.
5.2. Mantle xenocrysts Mantle-derived xenocrysts (Figure 14.22) represent a more complete sampling of the mantle components entrained by kimberlite by virtue of the very large number of single grains relative to xenoliths found in kimberlite. Consequently xenocrysts data should provide
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a more complete mantle section. The main difficulty with using xenocrysts has been the inability to assess pressures and temperatures of single grains and consequently they could not be correctly placed within the mantle stratigraphy. New methods, including singlecrystal chromian diopside pressure-temperature calculations (Nimis and Taylor, 2000) and temperature determinations from chromian pyropes using Ni contents (Ryan et al., 1996) projected to the relevant geotherm, allow the mantle stratigraphy to be compiled. Combined major and trace element compositions of kimberlite pyrope xenocrysts reveal that there are three distinct layers in the lithospheric mantle at the Kaavi–Kuopio kimberlite localities (Lehtonen et al., 2004; Figure 14.21): (1) A low temperature, 700–850 °C, or 70–110 km (using the geotherm of Kukkonen and Peltonen, 1999) mantle layer containing Ca-rich but Ti, Y and Zr depleted wehrlite-composition pyropes. Garnets of this composition are in equilibrium with chromite and clinopyroxene and have been described in the same shallow mantle stratigraphic position from the Slave Province, northwestern Canada (Kopylova et al., 2000; Carbno and Canil, 2002). The low-temperature fine-grain ed harzburgites described above contain similar pyropes and are almost certainly the source of these xenocrysts; (2) A variably depleted lherzolite- and harzburgite-bearing horizon from 900–1200 °C, or 130 km to 180 km; (3) A fertile lherzolitic deeper layer (no harzburgitic pyropes known so far) from 180 km to 240 km, possibly representing Proterozoic underplating or a melt-enriched version of layer 2. The majority of the xenoliths come from this horizon. Xenocrystic pyrope data from the Kuhmo area show that the lithospheric mantle stratigraphy here is less heterogeneous, with a variably depleted lherzolite- and harzburgitebearing horizon extending from the very top of the garnet-bearing mantle at about 80 km depth, to 250 km or more (O’Brien et al.,
2003). Present data appear to show a thin melt-metasomatized layer at least a few tens of km thick at the very base of the Kuhmo area mantle lithosphere, possibly corresponding to the deep layer at Kaavi–Kuopio. Given the location of the Kaavi–Kuopio kimberlites at the edge of the craton the more complicated mantle stratigraphy at this location is quite reasonable. The mantle underlying the Kuhmo area apparently suffered less reactivation during this major collision event and post-collision processes and thus may represent a more pristine nucleus of the Karelian craton.
5.3. Diamonds Diamonds represent mantle samples from depths exceeding ~150 km and as such provide information on the physical and chemical properties of the deepest parts of the lithospheric mantle, and in the case of super deep diamonds also of the convective mantle. In the Kaavi–Kuopio kimberlites, diamonds occur mainly as xenocrysts in the kimberlite matrix but also as a rock-forming mineral in mantle eclogite xenoliths. Several of the Kaavi–Kuopio kimberlite pipes have yielded reasonable diamond grades (14–41 ct /100 t; Tyni, 1997). However, most of the crystals separated from the kimberlite matrix are resorbed. A photographic interpretation by Kinnunen (2001) indicated the proportion of resorbed to unresorbed crystals is about 9:1 in the small sample that was available. The primary crystal morphologies exhibited by these stones include octahedrons (Figure 14.22), twinned crystals, and aggregates of coalesced single stones (Kinnunen, 2001). Diamond crystals found within eclogite xenoliths have been studied in detail. One of the diamonds protruding from the outer surface of one xenolith is pseudohemimorphic. Its unresorbed part resided inside the xenolith while the exposed crystal faces have been resorbed. This implies that resorption took place by reaction between diamond and
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kimberlite melt. These eclogite xenoliths can be highly diamondiferous suggesting that some diamonds are present within very highgrade pods or thin seams in the lithospheric mantle. This indicates a genetic link between the eclogite host and its diamonds. However, there is no evidence to support a particularly carbon-rich parental magma for the eclogites. More probably, eclogite seams have acted as pathways for a carbon-bearing fluid phase. The high Fe3+ contents of garnet and clinopyroxene in diamondiferous eclogites compared to those in barren ones, suggest formation of diamond in situ through simultaneous reduction of CO2, CO, and CH4 in the fluid and oxidation of the Fe2+ in the adjacent silicates. Anomalous birefringence patterns in these diamonds imply intense plastic deformation and suggest that these eclogite seams also acted as sublithospheric shear zones at the time of diamond genesis (Peltonen et al., 2002, 2003).
Acknowledgments The authors are grateful to Alan Woolley for helping clarify several terminology problems and for doing a thorough yet rapid review of the text. Discussions with Juha Karhu, Raimo Lahtinen, Hannu Huhma, Matti Vaasjoki, and Kari A. Kinnunen, helped clarify some of the ideas presented here. Seppo Elo, Meri-Liisa Airo, and Maija Kurimo produced images based on geophysical data provided by the Geological Survey of Finland and this is greatly appreciated. Olli Härmälä, Kauko Puustinen, and Heikki Lukkarinen provided details on Siilinjärvi geology that were quite useful. Pirkko Kurki did a masterful job drafting Figures 14.2 and 14.4 and we thank her for that.
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Finnish) Lehijärvi, M., 1960. The alkaline district of Iivaara, Kuusamo, Finland. Bull. Comm. géol. Finlande 185, 1–62. Lehtonen, M., O’Brien, H.E., Peltonen, P., Johanson, B., Pakkanen, L., 2004. Layered mantle at the edge of the Karelian craton: P-T of mantle xenocrysts and xenoliths from eastern Finland kimberlites. Lithos 77, 593–608. Luhr, J.F., Carmichael, I.S.E., 1981. The Colima volcanic complex, Mexico: Part II. LateQuaternary Cinder Cones. Contrib. Mineral. Petrol. 76, 127–147. Lukkarinen, H., 2000. Geological Map of Finland, Pre-Quaternary rocks, Sheet–3331–Siilinjärvi. [1:100 000]. Geol. Surv. Finland, Espoo. Mahotkin, I.L., 1998. Petrology of Group 2 Kimberlite-Olivine lamproite (K2L) series from the Kostomuksha area, Karelia area, N.W. Russia. 7th Int. Kimberlite Conf., Ext. Abstracts, 529–531. Mahotkin, I.L., Gibson, S.A., Thompson, R.N., Zhuravlev, D.Z., Zherdev, P.U., 2000. Late Devonian diamondiferous kimberlites and alkaline picrite (Proto-kimberlite?) magmatism in the Arkhangelsk Region, NW Russia. J. Petrol. 41, 201–227. Mikkonen, A., Kauppinen, H., Kallioinen, J., 1980. Siilinjärven kaivos. Summary: Kemira Oy Siilinjärvi mine. Vuoriteollisuus–Bergshanteringen 38 (1), 10–15. Mitchell, R.H., 1986. Kimberlites – Mineralogy, Geochemistry, and Petrology. Plenum, New York, 1–442. Mitchell, R.H., 1995a. Kimberlites, Orangeites, and Related Rocks. Plenum, New York, 1–410. Mitchell, R.H., 1995b. The role of petrography and lithogeochemistry in exploration for diamondiferous rocks. J. Geochem. Explor. 53, 339–350. Mitchell, R.H., Scott Smith, B.H., Larsen, L.M., 1999. Mineralogy of Ultramafic Dikes from the Sarfartoq, Sisimiut and Maniitsoq Areas, West Greenland. In: J.J. Gurney et al. (Eds.), Proc. 7th Int. Kimberlite Conf., Vol. 2, 574–583. Neuvonen, K.J., Korsman, K., Kouvo, O., Paavola, J., 1981. Paleomagnetism and age relations of the rocks in the Main Sulphide Ore Belt
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in central Finland. Bull. Geol. Soc. Finland 53, 109–133. Nimis, P., Taylor, W.R., 2000. Single clinopyroxene thermobarometry for garnet peridotites. Part I. Calibration and testing of a Cr-inCpx barometer and an enstatite-in-Cpx thermometer. Contrib. Miner. Petrol. 139, 541–554. Nixon, P.H., Boyd F.R., 1973. Petrogenesis of the granular and sheared ultrabasic nodule suite in kimberlites. In: P.H. Nixon (Ed.), Lesotho Kimberlites, Lesotho National Development Corporation, 48–56. Nuutilainen, J., 1973. Soklin karbonatiittimassiivin geokemiallisista tutkimuksista. Geologi 25, 13–17. (in Finnish) Nykänen, J., Laajoki, K., Karhu, J., 1997. Geology and geochemistry of the early Proterozoic Kortejärvi and Laivajoki carbonatites, central Fennoscandian Shield, Finland. Bull. Geol. Soc. Finland 69, 5–30. O’Brien, H.E., Irving, A.J., McCallum, I.S., Thirlwall, M.F., 1995. Strontium, neodymium, and lead isotope evidence for the interaction of post-subduction asthenospheric potassic mafic magmas of the Highwood Mountains, Montana, USA, with ancient Wyoming craton lithospheric mantle. Geochim. Cosmochim. Acta 59, 4539–4556. O‘Brien, H.E., Tyni, M., 1999. Mineralogy and geochemistry of kimberlites and related rocks from Finland. In: J.J. Gurney et al. (Eds.), Proc. 7th Int. Kimberlite Conf., Vol. 2, 625–636. O’Brien, H.E., Lehtonen, M., Spencer, R.G.S., Birnie, A., 2003. Lithospheric Mantle in Eastern Finland: a 250 km 3D transect. 8th Int. Kimberlite Conf., Ext. Abstracts. Paarma, H., 1970. A new find of carbonatite in North Finland, the Sokli plug in Savukoski. Lithos 3, 129–133. Paarma, H.,Vartiainen, H., Litvinenko, V.I., Muzylev, V.V., 1981. Geological interpretations of seismic soundings at the Sokli carbonatite comples, northern Finland. Papers issued to the 10th General Meeting of the Finnish-Soviet Joint Geological Working Group, held in Rovaniemi, 7-11 September 1981, 213–224. Pearson, D.G., 1999. The age of continental roots. In: R. van der Hilst, W.F. McDonough (Eds.), Composition, deep structure and
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evolution of continents. Harvard University, Cambridge, Mass., 15–17 October 1997. Special Issue, Lithos 48, 171–194. Peltonen, P., Mänttäri, I., 2001. An ion microprobe U-Th-Pb study of zircon xenocrysts from the Lahtojoki Kimberlite pipe, eastern Finland, Bull. Geol. Soc. Finland 73, 47–58. Peltonen, P., Huhma, H., Tyni, M., Shimizu, N., 1999. Garnet peridotite xenoliths from kimberlites of Finland: Nature of the continental mantle at an Archaean Craton – Proterozoic mobile belt transition. In: J.J. Gurney et al. (Eds.), Proc. 7th Int. Kimberlite Conf., Vol. 2, 664–676. Peltonen, P., O’Brien, H., Karhu, J., Kukkonen, I.T., 2000. Kimberlites, carbonatites and their mantle sample: constraints for the origin and temporal evolution of the lithospheric mantle in Fennoscandia. In: L.J. Pesonen et al. (Eds.), Lithosphere 2000 - A symposium on the structure, composition and evolution of the lithosphere in Finland. Programme and extended abstracts, Espoo, Finland, October 4–5, 2000. Institute of Seismology, University of Helsinki. Report S-41, 63–69. Peltonen, P., Kinnunen, K.A., Huhma, H., 2002. Petrology of two diamondiferous eclogite xenoliths from the Lahtojoki kimberlite pipe, eastern Finland. Lithos 63, 151–164. Peltonen, P., Kinnunen, K.A., Woodland, A.B., Seitz, H.-M., 2003. Origin of eclogites and diamonds in the Fennoscandian continental mantle. Eur. Geophys. Soc., Geophys. Res. Abstracts 5, 06728. Perttunen, M., Vartiainen, H., 1992. Glaciofluvial transport of clasts and heavy minerals from the Sokli carbonatite complex, Finnish Lapland. Geol. Surv. Finland, Bull. 366, 1–21. Polet, J., Anderson, D.L., 1995. Depth extent of cratons as inferred from tomographic studies. Geology 23, 205–208. Poutiainen, M., 1995. Fluids in the Siilinjärvi carbonatite complex, Eastern Finland: Fluid inclusion evidence for the formation conditions of zircon and apatite. Bull. Geol. Soc. Finland 67, 3–18. Price, S.E., Russell, J.K., Kopylova, M.G., 2000. Primitive magma from the Jericho pipe, N.W.T., Canada: Constraints on primary kimberlite melt chemistry. J. Petrol. 41,
789–808. Proskuryakov, V.V., Uvadiev, L.I., Voinova, O.A., 1990. Lamproites of the Karelia-Kola region. Trans. USSR Acad. Sci. 314, 940–943. Puustinen, K., 1970. The carbonatite of Siilinjärvi in the Precambrian of Eastern Finland. A preliminary report. Lithos 3, 89–92. Puustinen, K., 1971. Geology of the Siilinjärvi carbonatite complex, Eastern Finland. Bull. Comm. géol. Finlande 249, 1–43. Puustinen, K., 1973. Tetraferriphlogopite from the Siilinjärvi carbonatite complex, Finland. Bull. Geol. Soc. Finland 45, 35–42. Puustinen, K., 1986. Halpasen karbonatiitti Mikkelin mlk:ssa. Summary: Halpanen, a new carbonatite occurrence in Finland. Geologi 38, 1–5. Puustinen, K., Karhu, J.A., 1999. Halpanen calcite carbonatite dike, southeastern Finland. In: S. Autio (Ed.), Geological Survey of Finland, Current Research 1997–1998. Geol. Surv. Finland, Spec. Pap. 27, 39–41. Ramsay, W., 1889. Geologische Beobachtungen auf der Halbinseln Kola. Fennia 3 (7), 1–52. Ramsay, W., Berghell, H., 1891. Das Gestein vom Iiwaara in Finnland. Geol. Fören. Stockholm Förhandl. 13, 300–312. Ramsay, W., Hackman, V., 1894. Das Nephelinsyenitgebiet auf der Halbinseln Kola. I. Fennia 11 (2), 1–225. Rock, N.M.S., Bowes, D.R., Wright, A.E., 1991. Lamprophyres. Blackie and Son Ltd., Glasgow. 1–275. Roden, M.F., Smith D., 1979. Field geology, chemistry and petrology of Buell minette diatreme, Apache County, Arizona. In: F.R. Boyd, H.O.A. Meyer (Eds.), Proc. 2nd Int. Kimberlite Conf., Vol. 1, 364–381. Ryan, C.G., Griffin, W.L., Pearson, N.J., 1996. Garnet geotherms: Pressure-temperature data from Cr-pyrope garnet xenocrysts in volcanic rocks. J. Geophys. Res. 101, 5611–5625. Sablukov, S.M., 1984. The question of origin stages and age of the Onega Peninsula pipes. Doklady Akademii Nauk SSSR 277, 168–170. (in Russian) Smith, C.B., 1983. Pb, Sr and Nd isotopic evidence for sources of southern African Cretaceous kimberlites. Nature 304, 51–54. Taylor, W.R., Tompkins, L.A., Haggerty, S.E., 1994.
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Vartiainen, H., Vitikka, A., 1993. The late dikes of the Sokli massif and their tectonic monitoring. Geochimia 8, 1241–1244. (in Russian) Vartiainen, H., Kresten, P., Kafkas, Y., 1978. Alkaline lamprophyres from the Sokli complex, northern Finland. Bull. Geol. Soc. Finland 50, 59–68. Vartiainen, H., Melnikov, I., Sullimov, B., 1990. The francolite ore deposits of Kovdor and Sokli. Proceedings of the Finnish-Soviet Symposium held in Helsinki, Finland. November 14-15. 1990. Research Report TKK-IGE A13, 7–14. Wyllie, P.J., Lee, W.-J., 1999. Kimberlites, carbonatites, peridotites and silicate-carbonate liquid immiscibility explained in parts of the system CaO-(Na2O+K2O)-(MgO+FeO) -(SiO2+Al2O3)-CO2. Proc. 7th Int. Kimberlite Conf., Vol. 2, 923–932. Zhuravlev, V.A., Shulga, T.F., Ushkov, V.V., 1995. Diamond-bearing lamproites of the Kostomuksha region of Karelia. Proc. 2nd Int. Symp. “Mineral Resources of Russia: Diamonds and Gold”, St. Petersburg, VSEGEI, 37–40.
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Chapter 15
DRIFT HISTORY OF THE SHIELD
S. Mertanen, L. J. Pesonen
Cover page: Russian Resurs 03 satellite photo mosaic including the Fennoscandian Shield. The mosaic was generated from a series of images taken over a period of three years (1996 to 1998). Source data were geometrically corrected and provided by Metria Satellus Kiruna. The mosaic, color balance, and final image composite was generated by WorldSat International Inc. Published with arrangements with Metria Kiruna, Sweden. © WorldSat International and Metria Satellus.
Mertanen, S., Pesonen, L.J., 2005. Drift History of the Shield. 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. 645–668. © 2005 Elsevier B.V. All rights reserved.
Paleomagnetic data and isotope ages from the Fennoscandian Shield provide direct evidence of the drift history of the shield and its relative position with respect to other cratons. During most of its geological history, the Fennoscandian Shield was located at low to moderate paleolatitudes, except during the Neoarchean. At 2.45 Ga, the shield formed a unity with Laurentia; this is indicated by the 2.45 Ga mafic dike swarms, which are parallel when placed according to the continental reconstruction presented here. The two continents were separated after 2.45 Ga probably in a breakup of a larger Archean landmass or a supercontinent. They collided again at ~1.9–1.8 Ga, during the Svecofennian orogeny in the Fennoscandian Shield and the Hudsonian orogeny in Laurentia. At that time, the Fennoscandian Shield was located at intermediate to low latitudes and probably formed part of the supercontinent Hudsonland. At ~1.25 Ga, both the Fennoscandian Shield and Laurentia experienced dike and sill magmatism and were joined in a rather similar configuration as at 1.83 Ga. This suggests that the North European and North American cratons may have formed a unity during the entire time period from 1.83 Ga to 1.25 Ga. At the same time, also the Amazonian craton had a close connection with the Fennoscandian Shield. Although geological evidence points to the existence of a laterally contiguous Laurentia–Fennoscandia–Amazonia landmass at ~1.83–1.50 Ga, the present paleomagnetic data cannot verify this. After ~1.2 Ga, the Fennoscandian Shield and Laurentia rifted apart and collided again at ~1.05 Ga to form the supercontinent Rodinia together with other Precambrian cratons.
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1. Introduction The geological evolution of the Fennoscandian Shield reflects an important part of the overall global history of the Earth’s lithospere and has a vital role in studying the evolution of supercontinents during the Precambrian. The importance of supercontinents for the Earth’s geological evolution has led to studies of global continental assemblies and their existence beyond the ~1050–780 Ma Rodinia assembly (e.g., McMenamin and McMenamin, 1990; Dalziel, 1991), i.e., in the Archean–Mesoproterozoic times (e.g., Piper, 1976; Rogers, 1996; Pesonen et al., 2003). Paleomagnetism combined with high resolution age data is the only method that provides direct knowledge on ancient latitudes and orientations of continents. It can therefore be used for testing geologically inferred supercontinent assemblies and also for proposing new continental assemblies to be verified by geological data. The Fennoscandian Shield includes wellexposed rock formations ranging in age from Archean to Paleozoic and a considerable amount of paleomagnetic data have been obtained from formations dated with U-Pb or 40 Ar- 39Ar methods. In Finland, paleomagnetic studies have been carried out to study the dispersal and aggregation of different blocks of the shield to constrain the drift movement of the craton, and to study the existence of supercontinents by defining the position of the shield in paleocontinental reconstructions in relation to other continents. In the following, updated paleomagnetic data are shown for the period from the Neoarchean to the early Neoproterozoic, covering the major geological events in Finland. Paleomagnetic studies in Finland have been carried out in keen cooperation with Scandinavian and Russian paleomagnetists. One expression of such cooperation are the Nordic Paleomagnetic Workshops that have been held every fourth year since 1986 in 648
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different Nordic countries (Pesonen, 1987, 1999; Abrahamsen et al., 2001). One of the main aims of these workshops has been to compile and qualitatively classify all the paleomagnetic data published in Fennoscandia. This has resulted in the Fennoscandian Paleomagnetic Database (Pesonen and Torsvik, 1990; Pesonen et al., 1991; Torsvik, 1998) that is continuously updated and forms part of the global paleomagnetic database (GPDB; http//www.dragon.ngu). So far there are ~800 data entries from the Scandinavian countries and western Russia. In the two latest paleomagnetic workshops the main aim has been the comparison of paleomagnetic data of the Fennoscandian Shield with data from other continents for making continental reconstructions (Buchan et al., 2000, 2001; Pesonen et al., 2003). In this chapter we review the main results of these studies.
2. Remanent magnetization in the Fennoscandian Shield The bedrock of Finland has experienced several geological events during the Precambrian. These events are reflected in the record of the Earth’s magnetic field, which is preserved as remanent magnetization in the rocks. In addition to primary remanence acquired during the cooling of a rock unit, the same rock may carry secondary remanence components formed in later geological processes. These remanent acquisition processes may relate to various mechanisms. If the rock is subjected to heating due to tectonic events, burial metamorphism or by contact metamorphism, it may lead to the acquisition of partial thermoremanent magnetization, which may be later obscured together with a pre-existing primary remanent magnetization (Figure 15.1A). Another important modifier of remanent magnetization are hydrothermal fluids that can either destroy pre-existing remanent magnetizations or form new magnetic minerals carrying stable
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Fig. 15.1. (A) Schematic illustration of the acquisition of two remanence components. Component A represents primary remanent magnetization acquired during the original cooling of a rock unit. The secondary component B is formed during a later event. In the Fennoscandian Shield, the latter could be, e.g., the Svecofennian orogeny that has an effect on the pre-existing Archean crust. Both components may be found within the same rock specimen. (B) The occurrence of components A and B is seen in the stereoplot as a slight movement of the direction of the natural remanent magnetization (NRM) during the course of demagnetization. Owing to their deviating unblocking temperatures, the components can be separated by multicomponent analysis. (C) Calculated paleopoles from components A and B. Only primary poles (such as A) are used in calculating the position of the Fennoscandian Shield at various times.
chemical or thermo-chemical magnetizations. If the temperature during metamorphism does not rise above the Curie point of the minerals carrying the primary remanence, it is possible to separate both the primary and secondary remanences by multicomponent analyzing methods (Figure 15.1B). Paleomagnetic poles are calculated from these remanence directions (Figure 15.1C) and, in ideal cases, they can be isotopically or paleomagnetically dated CHAPTER
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and linked to known geological events. One of the most conspicuous events in the Fennoscandian Shield was the Svecofennian orogeny at ~1.9–1.8 Ga that typically modified the remanent magnetization of older rocks. This is seen in many paleomagnetic studies of Finland and also in studies carried out in northern Sweden and western Russia (see Mertanen, 1995). In some situations, the rocks have been totally overprinted by this orogenic event so H I S T O RY
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that no primary remanences remain. In many cases, however, the Svecofennian orogeny has only partially overprinted the rocks so that both the primary and secondary magnetization components may be successfully separated. Good examples from rocks in Finland are the 2.45 Ga layered mafic intrusions and associated mafic dikes that carry a primary remanence and also a well defined Svecofennian secondary remanence formed at ~1.88 Ga (Mertanen et al., 1999). In the Varpaisjärvi area, central Finland, 2.1 Ga mafic dikes and the Archean granulite facies basement also show multicomponent remanent magnetizations (Neuvonen et al., 1997; Mertanen et al., 2004a). Most of the dikes are totally overprinted by a secondary remanence, while the Archean basement still carries the primary remanence in addition to the secondary one. In both rock types the secondary remanence was formed during the Svecofennian orogeny. Also younger rock formations yield secondary magnetizations. For example, the 1.63 Ga Sipoo quartz porphyry and diabase dikes in southern Finland carry a secondary remanence that may be connected to global scale tectonic events (Mertanen and Pesonen, 1995). In these dikes a secondary remanence, considered ~1.3 Ga old, was isolated in addition to the primary remanence, and may reflect crustal extension during the Postjotnian time or at the onset of the Sveconorwegian orogeny slightly later. In summary, the Fennoscandian Shield offers good prospects for isolation of remanent magnetizations of different ages and therefore provides good opportunities for making continental reconstructions. Although the difficulty in paleomagnetic studies has been the multicomponent nature of the remanence that may hinder the isolation of primary remanences, in many cases these primary remanences have been obtained. Primary remanences are the only type of remanence that are used in drift calculations and continental reconstructions, because they can be accurately dated with UPb or 40Ar- 39Ar methods. 650
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3. Fennoscandian drift history in the Precambrian In previous compilations of paleomagnetic data and drift history calculations (Pesonen et al., 1989; Elming et al., 1993), the Apparent Polar Wander Path (APWP) method was used. The difficulty with the APWP approach is that there often exist large age gaps between successive individual or ‘Grand Mean’ poles along the APWP. In recent studies, a new method using ‘key paleomagnetic poles’ has been adopted (Buchan and Halls, 1990; Buchan et al., 2000, 2001). The main paleomagnetic criteria for a pole to be a key pole is that it represents a statistically well-defined primary magnetization as verified by positive field test(s) and that it is obtained from an accurately dated rock unit (see Buchan et al., 2000, 2001). Key poles are used for establishing craton drift histories and, in order to create continental reconstructions, they are directly compared with poles of similar age from different cratons. There are a few important restrictions and ambiguities in paleomagnetic techniques used for producing such reconstructions. Only paleolatitude and orientation of a craton (with respect to its present geographic orientation) can be determined. The paleolongitude is not determinable due to the axial symmetry of the geocentric dipole field hypothesis (AGDF). The polarity of the AGDF forms another ambiguity. This becomes a problem if there are large age gaps between successive poles, in which case either normal or reversed poles can be accepted. Due to this polarity ambiguity, a continent can be plotted in correct latitudes in either hemisphere and in antipodal orientations (Buchan et al., 2000; Pesonen et al., 2003). Paleomagnetic data from the Fennoscandian Shield covers the time span from the Neoarchean to the Late Paleozoic. Figure 15.2 shows the latitudinal drift of the Fennoscandian Shield from ~2.55 Ga to 1.05 Ga. The model has been made by using the best defined
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Time (Ga) Fig. 15.2. The drift history of the Fennoscandian Shield from ~2.55 Ga to 1.05 Ga. The shield area is shown in yellow and the approximate growth of the shield in red. From 1.7 Ga onward, the shield is shown larger owing to the amalgamation of the Ukrainian Shield to the southern part of the Fennoscandian Shield. Only the paleolatitude and the orientation of the shield (with respect to its present orientation) can be defined paleomagnetically. The paleolongitude is not determinable. The horizontal axis denotes time.
paleomagnetic poles from isotopically dated rock units. However, most of the paleopoles used here are not considered key paleomagnetic poles with reference to the stringent criteria by Buchan et al. (2000, 2001). This is mainly due to lack of conclusive field tests that would show the primary nature of remanence and occasional large errors in the paleopoles or in the datings. Therefore, part of the paleopositions of the shield are less reliable and must be verified with further studies.
3.1. Neoarchean Paleomagnetic studies on Archean basement rocks have been carried out in several locations in Finland. In the case of the basement gneisses, the remanence intensities are generally too weak or the remanence directions too unstable to yield meaningful results. Furthermore, in many cases the remanence directions in the basement rocks show widespread Svecofennian overprinting acquired at ~1.9–1.8 CHAPTER
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Ga (Mertanen et al., 1989, 1999; Halls and Vuollo, 1999), thus emphasizing the penetrative effect of the Svecofennian orogeny deep into the Archean terranes. However, in areas of Archean granulitefacies metamorphism, high remanence intensities with stable directions occur. The calculated position of the Fennoscandian Shield at ~2.6–2.5 Ga (Figure 15.2) is based on studies on igneous enderbites and mafic granulites in the Varpaisjärvi area of the Iisalmi complex of the Karelian province (Neuvonen et al., 1981, 1997; Mertanen et al., 2004a) where, accodingly to U-Pb datings on monazite and zircon, the granulite facies metamorphism took place at 2.63 Ga (Hölttä et al., 2000). Sm-Nd garnet–whole-rock ages that record the cooling of the granulites at ~2.6–2.5 Ga (Hölttä et al., 2000) are considered to be in accord with the magnetization ages of the rocks. Due to their dry nature, these rocks have been less vulnerable to remagnetization than the lower grade metamorphic rocks. Similar paleomagnetic H I S T O RY
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directions (Mertanen, 2000) have been obtained also from igneous charnockites from the Lieksa area in eastern Finland where the granulite-facies metamorphism is assumed to have taken place contemporaneously with the metamorphism in the Varpaisjärvi area (Halla, 2002). In both areas, the Archean remanence has high coercivity and high unblocking temperatures supporting a primary origin of the remanence. Moreover, in the Varpaisjärvi and Lieksa areas the remanence directions are antipodal implying a geomagnetic field reversal during the cooling of these rocks. Naturally, when dealing with Archean rocks involved in later tectonomagmatic events, one has to take into account the tectonic movements after the acquisition of the primary Archaean remanences. The similarity of remanence directions both in the Varpaisjärvi and Lieksa regions support the interpretation that no large scale movements have taken place between these areas, although the steep inclinations in both formations allow minor rotations and displacements within error limits. Other studies on Archean rocks in the Karelian province have been carried out in the Vodlozero block (Krasnova and Gooskova, 1990) and in the Shilos greenstone structures (Arestova and Gooskova, 1998; Arestova et al., 2000a) in western Russia. However, because of the difficulty in obtaining the magnetization ages for the rocks, these results must be confirmed with further studies involving isotope datings and paleomagnetic field tests (Mertanen et al., 2004b).
3.2. Continental rifting at 2.4 Ga Continental rifting of the stabilized Archean craton was initiated at ~2.5–2.4 Ga. The rifting produced several layered mafic intrusions and associated mafic dike swarms both in the Karelian and Kola provinces (cf. Chapter 3 and 5). Numerous paleomagnetic studies have been carried out on these formations in both terranes. Figure 15.2 shows two positions of the Fennoscandian Shield (at 2.45 Ga and 652
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2.40 Ga) based on studies on dike rocks at the Lake Pääjärvi area in the Karelian province in Russia (Mertanen et al., 1999). The ages of the Lake Pääjärvi dikes range between 2470–2350 Ma (Sm-Nd or U-Pb on zircon or baddeleyite; Vuollo et al., 1996; Mertanen et al., 1999) and they are genetically related to the coeval Olanga layered mafic intrusions. The position of the Fennoscandian Shield at 2.45 Ga is based on the D’ pole that was isolated in a well preserved dike and contact host rock at the Lake Pääjärvi area (Mertanen et al., 1999) and is considered to be the primary remanence with a magnetization age of 2.45 Ga. The position of the Fennoscandian Shield at 2.40 Ga is also based on studies from the Lake Pääjärvi area and is calculated from the paleopole D, which is very often seen in the 2.4 Ga rocks of the Fennoscandian Shield. This D pole has been obtained from the Koillismaa layered mafic intrusions in Finland (Mertanen et al., 1989), from the Burakovka intrusion and associated dikes in southwestern Russia (Krasnova and Gooskova, 1990; Khramov et al., 1997; Fedotova et al., 1999; Mertanen et al., 2004b), and from the gabbro–norites and tholeiitic dikes (Krasnova and Gooskova, 1995) east of Lake Pääjärvi in Russia. Pole D’ is regarded as an original, primary remanent magnetization acquired during the early stages of rifting of the Archean crust and subsequent emplacement of the dikes and layered mafic intrusions at ~2.45 Ga. Pole D represents a slightly later magnetization, acquired at ~2.40 Ga during prolonged cooling of the layered mafic intrusions. It is also likely that the thermal effect of the voluminous plume-related (Hanski et al., 2001b) magmatic activity heated the Archean crust close to the layered mafic intrusions, thus creating the 2.4 Ga magnetization seen in many nearby Archean rocks (Mertanen et al., 1999, 2004b). If the assigned magnetization ages are correct, the Karelian province drifted 20° towards higher paleolatitudes and rotated some 45° clockwise during the time interval
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of 2.45–2.40 Ga. One of the targets of paleomagnetic investigations on early Paleoproterozoic formations has been to study the suggested dispersion and amalgamation of the Kola and Karelian provinces. Balagansky et al. (2001), Daly et al. (2001), and Glebovitsky et al. (2001) have suggested that the 2.5–2.4 Ga rifting developed to an ocean opening in the Kola Peninsula. On the other hand, it has also been suggested that the rifting probably did not lead to the formations of an ocean, but rather to an intracontinental rift basin (e.g., Melezhik et al., 1997). In recent years several paleomagnetic studies on the ~2.45 Ga layered mafic intrusions and related dikes have been carried out in the Kola Peninsula in Russia. Studies on the 2.50–2.45 Ga Mt. Generalskaya intrusion (Arestova et al., 1999, 2002), the 2.50–2.47 Ga Fedorova Tundra intrusion (Arestova et al., 1997), the 2.45–2.40 Ga Imandra intrusion (Arestova et al., 2000b), and the Main Range and Voche–Lambina intrusions forming parts of the 2.50–2.48 Ga Monchegorsk intrusion (Khramov et al., 1997; Fedotova et al., 1999) have revealed a primary ~2.45 Ga remanence (for isotope ages of the intrusions, see Hanski et al., 2001a, and references therein). However, the data from different intrusions are still controversial mainly because of probable errors in the remanence directions, owing to several superimposed magnetizations and possible uncorrected movements of the intrusions. In general, the suggested primary poles are in close agreement with the poles from the Karelian province. The results still have to be verified by further studies from well-preserved formations especially on the Kola Peninsula. In a recent paleomagnetic study on the 2.30 Ga (Hölttä et al., 2000) Tulisaari dolerite dike and some other dikes of the Varpaisjärvi area in the Iisalmi complex in central Finland, a presumably primary ~2.3 Ga remanent magnetization was obtained (Mertanen et al., 2003a). The direction of this remanence is similar to the 2.4 Ga D direction. If the ages CHAPTER
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of both remanent magnetizations are correct, it implies that during 2.4–2.3 Ga, the Fennoscandian Shield remained at a low paleolatitude (~30°) and without significant rotations. The result has important implications for the paleoclimate studies of the early Paleoproterozoic (Section 4.1).
3.3. Jatulian rifting and magmatism at 2.2–2.0 Ga Rifting of the Archean craton continued during the Jatulian time at ~2.2–2.0 Ga leading to the emplacement of extensive mafic dike swarms accompanied by mafic volcanic rocks and sills. However, despite several paleomagnetic studies on the Jatulian formations in Finland, northern Sweden, and Russia, the data are controversial. Most studies show either unstable remanences or strong Svecofennian overprinting. At present there are two different groups of directions that are suggested to represent the primary Jatulian remanence. One of these was obtained from diabase dikes and sedimentary rocks in Russia and is close to the 2.4 Ga direction (Damm et al., 1997; Khramov et al., 1997; Fedotova et al., 1999). The other one was obtained from 2.1 Ga dolerite dikes in the Varpaisjärvi area in central Finland (e.g., Neuvonen et al., 1981, 1997). However, as the latter remanence direction has also been obtained for younger formations in the Fennoscandian Shield, the age of this remanence is still questionable (Mertanen et al., 2004a). Hence, the position of the Fennoscandian Shield during Jatulian times is still unknown and further paleomagnetic studies are needed.
3.4. Onset of the Svecofennian orogeny at 2.0–1.9 Ga A geologically significant event in Finland was the Svecofennian orogeny, the first stages of which were probably initiated already at ~1.95 Ga (e.g., Nironen, 1997). However, from H I S T O RY
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Finland there are no paleomagnetic data of that age. In southwestern Russia, Pisarevsky and Sokolov (1999) studied the 1974 ± 27 Ma (Sm-Nd age; Puchtel et al., 1995) Konchozero sill located west of Lake Onega and obtained a paleomagnetic direction that they regarded as a primary 1.97 Ga remanence. This remanence has about the same direction as the 2.4 Ga D remanence which raises further doubt about the age of the D magnetization. If the D component is not a secondary magnetization, another explanation for the similarity of the two directions is that the Fennoscandian Shield roughly occupied the same paleolatitude and orientation at 2.4 Ga and 1.97 Ga, although having apparently drifted between those times. The oldest paleomagnetic data related to the Svecofennian orogeny in Finland and from the 1.93 Ga Tsuomasvarri gabbro–diorite intrusion in the hanging wall of the Lapland– Kola orogen (Mertanen and Pesonen, 1994). The Tsuomasvarri formation is related to the calc-alkaline magmatism in the Inari terrane dated 1.94–1.91 Ga (Kesola, 1991; Tuisku and Huhma, 1998; Daly et al., 2001). In the Tsuomasvarri intrusion, the characteristic paleomagnetic direction regarded as primary differs from the 2.4–2.3 Ga directions and places the Fennoscandian Shield near the equator, at ~20° paleolatitude. The collision of the Kola and Karelia provinces within the Lapland–Kola orogen took place probably slightly later, at ~1.9 Ga (Tuisku and Huhma, 1998). As it is likely that in the western Lapland–Kola orogen the Lapland–Kola ocean was closed at ~1.91 Ga (Balagansky et al., 2001; Daly et al., 2001); the nearly equatorial position of the Lapland–Kola orogen at 1.93–1.91 Ga (shown in Figure 15.2) may be applied to the Karelian province as well within paleomagnetic error limits.
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3.5. Svecofennian orogeny at 1.9–1.8 Ga There are several studies on isotopically dated 1.88 Ga synorogenic gabbroic intrusions in which a stable single component, primary thermoremanent magnetization, has been obtained (Pesonen and Stigzelius, 1972; Neuvonen et al., 1981; Elming, 1985, 1994; Mertanen and Pesonen, 1992). Some of the studied intrusions (e.g., the Kiuruvesi and Pohjanmaa gabbros) are located close to the Raahe–Ladoga belt, part of them further north (e.g., Vittangi gabbro in northern Sweden, Jalokoski gabbro in northern Finland). Due to prolonged cooling of these batholithic intrusions within the thermally elevated crust it is possible that the magnetization ages of the gabbros are younger (by some 10–20 Ma) than the isotope ages of ~1.88 Ga (see also Buchan et al., 2000). A younger remanence direction related to the late stages of the Svecofennian orogeny has been obtained from the 1837 Ma Haukivesi lamprophyric dikes (Huhma, 1981; Neuvonen et al., 1981) in central Finland. Recently, Pisarevsky and Sokolov (2001) obtained a stable remanence direction in the 1790–1770 Ma old Vazhinka River sedimentary rocks in northwestern Russia. Based both on polarities of remanence that define a magnetostratigraphic pattern and mineralogical evidence, they suggested that the remanence is primary and could be used as one of the key poles for the Fennoscandian Shield. In the 1770 Ma Rybreka sill that intrudes the Vazhinka sedimentary layers (A. Khramov, pers. comm., 2001), the remanence is close to the remanence direction of the Vazhinka formation and thus supports the pole’s ~1770–1780 Ma age. Figure 15.2 shows the positions of the Fennoscandian Shield calculated from the Svecofennian formations in the Karelian province during 1.88–1.77 Ga. The three poles (1.88 Ga, 1.84 Ga, and 1.77 Ga) indicate a slight paleolatitudinal movement of the Fenno-
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scandian Shield from the lower paleolatitude of ~15° at the early stages of the orogeny at 1.88 Ga to a higher paleolatitude of ~25° at 1.84 Ga, and again to a lower paleolatitude at ~1.77 Ga. Paleomagnetic data thus show that during the Svecofennian island arc accretion there was some latitudinal movement of the Shield but no significant rotation. Furthermore compared to the 2.45–2.40 Ga rifting episode, which was a period of high drift and rotation, the drift of the Fennoscandian Shield during the Svecofennian orogenic events was fairly slow.
3.6. Subjotnian magmatic interval at 1.65–1.5 Ga During the Subjotnian time at ~1.65–1.50 Ga, the rapakivi granites and associated anorthosites and dike swarms were emplaced in Finland and several paleomagnetic studies have been carried out on the rapakivi granites and quartz porphyry and diabase dikes. Paleopoles from the Subjotnian formations have been grouped according to their age into three different groups that all represent key paleopoles of the shield (Buchan et al., 2000). The oldest Subjotnian paleopole is obtained from the ~1630 Ma quartz porphyry dikes related to the Wiborg (Neuvonen, 1986) and Onas (Mertanen and Pesonen, 1995) rapakivi granites in southern Finland. A mean paleopole with an assigned age of 1570 Ma is obtained from the diabase and quartz porphyry dikes related to the Åland rapakivi granite (Pesonen and Neuvonen, 1981) and a slightly younger pole (1540 Ma; Pesonen and Neuvonen, 1981) characterizes dikes also associated with the Åland rapakivi batholith. Recent studies on ~1.5 Ga dike and intrusive rocks in central Sweden (Moakhar and Elming, 2000) support the earlier results on rocks of similar age in southern Sweden (Piper, 1979, 1980). Figure 15.2 shows that during the Subjotnian time, although characterized by anorogenic magmatism, the Fennoscandian Shield experienced CHAPTER
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relatively rapid paleolatitudinal movement.
3.7. Postjotnian time at ~1.26 Ga The Subjotnian time was followed by a paucity of magmatic activity until the Postjotnian when the Jotnian sedimentary units aged ~1500–1370 Ma were intruded by Postjotnian dolerite sills and dikes, collectively called the Central Scandinavian dolerite group (CSDG), at ~1270–1250 Ma (Suominen, 1991). One of the best defined paleomagnetic poles of Finland have been obtained from the Postjotnian Vaasa, Satakunta, and Märket dolerites in southwestern Finland (Neuvonen, 1965, 1966, 1973; Neuvonen and Grundström, 1969). The mean paleopole from these formations has been defined as a key pole for the Fennoscandian Shield (Buchan et al., 2000, 2001). Recent studies by Elming and Mattson (2001) from dolerite dikes in Sweden and in Finland confirm the similarity of paleomagnetic directions of the different dike groups (e.g., Bylund and Pesonen, 1987). Figure 15.2 shows the paleoposition of the Fennoscandian Shield at ~1.25 Ga when the shield had drifted from low northern latitudes to a paleolatitude of ~-15°.
3.8. Dike magmatism at 1.1–1.0 Ga The 1120 Ma Salla and 1042–1013 Ma (SmNd) Laanila–Ristijärvi diabase dikes (Mertanen et al., 1996) intruded the shield in northern Finland 100–200 Ma after the Postjotnian magmatism (see also Figure 13.2). A lot of paleomagnetic sampling has been done from the NW-trending Salla dikes (Mertanen et al., 1992), but owing to a strong Present Earth’s field component, no reliable results have been obtained so far. The Salla dikes trend almost perpendicular to the Laanila–Ristijärvi diabase dike swarm (Pesonen et al., 1986; Mertanen et al., 1996) that has the same general NE-trend as that of the Kautokeino and Karasjok diabase dikes in Finnmark, Norway. The dikes in H I S T O RY
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Kautokeino have a Sm-Nd age of 1066 ± 34 Ma (Mertanen et al., 1996). The remanences in the Laanila–Ristijärvi and Finnmark dikes are hard and stable and, according to baked contact tests, the remanences are primary. The ages of the Laanila–Kautokeino and Finnmark dikes correspond to the age of Sveconorwegian magmatism and emplacement of extensive dolerite dike swarms at ~1200–900 Ma in the Southwest Scandinavian domain in the southwestern part of the shield. Recent paleomagnetic studies on the 1041–980 Ma dolerite dikes of the Protogine zone in southern Sweden (Pisarevsky and Bylund, 1998) are in close agreement with the results from the 1200–975 Ma Bamble intrusions in southern Norway (Stearn, 1979; Stearn and Piper, 1984) and with the poles from the Laanila–Ristijärvi dikes. Correspondence of the poles from northern and southern parts of the shield implies that at that time the shield formed a unity with no significant block movements. Figure 15.2 shows the paleoposition of the Fennoscandian Shield at ~1.1 Ga based on data from the Bamble intrusions in southern Norway and at ~1.05 Ga based on the data from the Laanila–Ristijärvi dikes. At that time the Fennoscandian Shield was located at the paleolatitude of ~-30°. Compared to the position and orientation of the Fennoscandian Shield at 1.25 Ga, it had drifted southwards ~15° and rotated almost 90° clockwise. This rotation, suggested by Patchett and Bylund (1977), Poorter (1975), and Pesonen and Neuvonen (1981) is hence confirmed by the new data.
4. Position of the Fennoscandian Shield in the continental assemblies of the Precambrian Based on geological evidence, it has been suggested that supercontinents such as Ur, Arctica (Rogers, 1996) or Kenorland (Williams et al., 1991), and Vaalbara (Cheney, 1996) existed in the Archean. Reliable Archean paleomagnetic 656
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data do exist from Africa (Congo and Kalahari cratons), Australia (Pilbara), Laurentia (Superior), and the Fennoscandian Shield (Karelia), but owing to unmatching ages of the studied formations, no global-scale reconstruction is so far available. In the Nordic Paleomagnetic Workshop in 1999, continental reconstructions were made for the Proterozoic time (Mertanen and Pesonen, 2000; Pesonen et al., 2000, 2001, 2003; Elming et al., 2001; Pesonen and Mertanen, 2002). The data were compiled from all existing paleomagnetic resources involving most of the world’s cratons. The reconstructions are based on the most reliable paleomagnetic data that were first classified and filtered by using the Briden-Duff grading scheme (see Pesonen et al., 1989). Continental reconstructions were established for twelve time slots covering the time from 2.45 Ga to 1.00 Ga. The results were presented by Pesonen et al. (2003). In Figure 15.3, four of the reconstructions, 2.45 Ga, 1.83 Ga, 1.65 Ga, and 1.25 Ga, are shown.
4.1. Early Paleoproterozoic Figure 15.3A shows the 2.45 Ga reconstruction based on paleopoles from the Pääjärvi dikes of the Karelian craton of the Fennoscandian Shield (Mertanen et al., 1999), the Matachewan dikes of the Superior craton of Laurentia (Bates and Halls, 1990), the Ongeluk andesites (between 2489 ± 33 Ma and 2394 ± 26 Ma; Bekker et al., 2001) of the Kaapvaal craton of Africa (Evans et al., 1997), and the Widgiemooltha dikes of the Yilgarn craton of Australia (Evans, 1968). In this reconstruction, the Karelian and Matachewan dikes become parallel, suggesting that at ~2.45 Ga they formed a single dike swarm following the coeval mantle plume activity and rifting of the united Laurentia–Fennoscandia landmass that may have been part of the possible Neoarchean supercontinent Kenorland (see Pesonen et al., 2003). At about the same time with the rifting episode, during ~2.45–2.3 Ga, the Karelian,
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Fig. 15.3. Continental reconstructions at (A) 2.45 Ga, (B) 1.83 Ga, (C) 1.65 Ga, and (D) 1.25 Ga. B stands for Baltica (the Fennoscandian Shield), L for Laurentia, K for Kalahari, Au for Australia, Am for the Amazonian craton (the present South American continent is outlined with a dashed line), and C-Sf for Congo/SãoFrancisco craton. In (A) the Karelian (in B), Superior (in L), and Yilgarn (in Au) cratons are shown in green and the 2.45 Ga dike swarms in red. In (B) the 1.9-1.8 Ga Svecofennian (in B), Hudsonian (in L), and Ventuari–Tapajos (in Am) orogenic belts are shown in red. In (C) the ~1.7–1.5 Ga Gothian (in B), Labradorian (in L), and Rio Negro–Juruena belts (in Am) are shown in orange. In (D) the ~1.25 Ga Postjotnian (in B) and Mackenzie and Sudbury dike swarms (in L) are shown in red. The arrows indicate subsequent movement of the continents. In (B) and (C) the two possible positions of the Amazonian craton are shown as models (1) and (2). See text.
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Superior, and Kaapvaal cratons experienced glaciations (Mar mo and Ojakangas, 1984; Marmo, 1993; Evans et al., 1997; Williams and Schmidt, 1997; Schmidt and Williams, 1999; Ojakangas et al., 2001). The present paleomagnetic data from the Fennoscandian Shield thus support the growing evidence of early Paleoproterozoic glaciations at nearly equatorial latitudes. Different models have been presented to explain the enigma of low latitude glaciations during the Proterozoic, ranging from the ‘Snowball Earth’ hypothesis (Kirschvink, 1992; Hoffman et al., 1998) to the hypothesis of large Earth’s obliquity (Williams, 1993). The paleomagnetic data from the Yilgarn craton still need further studies, but according to the present data, at ~2.4 Ga the Yilgarn craton was located at a high paleolatitude and was thus not joined to other cratons.
4.2. Middle Paleoproterozoic Figure 15.3B shows the reconstruction at 1.83 Ga where data from the 1830 Ma Haukivesi lamprophyric dikes were used for the Fennoscandian Shield and data from the 1827 Ma Sparrow dikes for Laurentia (McGlynn et al., 1974). The relative position of the Fennoscandian Shield and Laurentia differs from that at 2.45 Ga (Figure 15.3A), implying that between 2.45 Ga and 1.83 Ga, the two continents were separated and eventually collided sometime before 1.83 Ga. Separation probably took place at ~2.2–2.0 Ga when both continents experienced extensive rifting and dike activity. Although contemporaneous paleomagnetic data from the Fennoscandian Shield still remains to be studied, the well-defined paleomagnetic data from Laurentia (Buchan et al., 2000) indicate considerable drift and rotation of Laurentia, which may have been due to a breakup of a supercontinent. The collision between the Fennoscandian Shield and Laurentia is manifested by the occurrence of the coeval 1.9–1.8 Ga Svecofennian 658
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and Hudsonian orogenies. The reconstruction at 1.83 Ga places the Kola Peninsula adjacent to the present day East Greenland. The reconstructions of the Fennoscandia and Laurentia differ slightly at 1.88 and 1.83 Ga when using the mean 1.88 Ga pole from the Svecofennian intrusions (Jalokoski, Kiuruvesi, Pohjanmaa, and Vittangi) for Fennoscandia and the pole from 1877 ± 7 Ma Molson dikes for Laurentia (Halls and Heaman, 2000). According to that reconstruction, the Kola Peninsula was adjacent to the present southeast Greenland. Therefore, paleomagnetic data imply that the final amalgamation of the Fennoscandian Shield and Laurentia took place no earlier than ~1.83 Ga (see Pesonen et al., 2003). Reliable 1.83 Ga paleomagnetic data also exist from the Amazonian craton (Onstott et al., 1984), where the successive subparallel 1.9–1.8 Ga Ventuari–Tapajos and 1.8–1.55 Ga Rio Negro–Juruena provinces represent belts of orogenic activity along the present western–southwestern margin of the craton (Tassinari et al., 2000). The tectonic setting of the Amazonian craton resembles that of the Fennoscandian Shield where the Svecofennian 1.9–1.8 Ga orogenic belt is succeeded by the 1.8–1.7 TIB and ~1.7–1.5 Gothian belts in the present southwest. Figure 15.3B shows two possible positions for the Amazonian craton depending on the choice of polarity. According to the first alternative (model 1 in Figure 15.3B), the Amazonian craton and the Fennoscandian Shield are both located at the latitude of about 20º, but so that the 1.9–1.8 Ga orogenic belts are oriented in antipodal directions towards the Archean inlands. The use of this polarity option implies that the Fennoscandian Shield and the Amazonian craton were departing from each other. In the second alternative (model 2 in Figure 15.3B) where the Amazonian craton is rotated 180º and shifted to the opposite latitude, the1.9–1.8 Ga orogenic belts continue almost linearly. According to geological indications (e.g., Geraldes et al., 2001), this model could be
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the more probable one, suggesting that the Fennoscandian Shield, Laurentia, and the Amazonian craton formed an amalgamated continent, possibly part of a large continent, Hudsonland (see Pesonen et al., 2003). However, based on present paleomagnetic data, the Fennoscandian Shield and the Amazonian craton are separated by a latitude of about 20º which cannot be explained purely by error limits of the poles (3º for Fennoscandia and 6º for Amazonia). Therefore, in spite of similarities in geological evolution, more paleomagnetic data are still needed to verify the relative position of the Amazonian craton with respect to the Fennoscandian Shield and Laurentia. After the Svecofennian orogeny the Ukrainian Shield was probably amalgamated to the Fennoscandian Shield in the southeast. Recent paleomagnetic data from the 1.77 Ga anorthosites and older rock formations from the Ukrainian Shield (Elming et al., 1999, 2000) suggest that the amalgamation of the Fennoscandian Shield and the Ukrainian craton took place sometime after ~1.77 Ga. When using the 1770 Ma remanence of the Fennoscandian Shield (Pisarevsky and Sokolov, 2001) and the 1750 Ma remanence (B component) from the Molson dikes of Laurentia (Halls and Heaman, 2000), a different reconstruction is produced as compared to that of 1.83 Ga. The difference is mainly due to the varying orientation and paleolatitude of Laurentia, as the Fennoscandian Shield maintained roughly the same orientation and latitude during 1.88–1.77 Ga. The 1.77 Ga data thus imply that the two cratons were separated after the Svecofennian orogenic events. On the other hand, as will be shown later, the Fennoscandian Shield and Laurentia had a very similar relative position at 1.25 Ga and 1.83 Ga (see also Buchan et al., 2000). Consequently, either the two cratons were separated between 1.83–1.75 Ga and then collided again at 1.25 Ga in a fairly similar configuration, or the ~1.75 Ga paleomagnetic data, especially those from Laurentia, are suspect. CHAPTER
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4.3. Late Paleoproterozoic Paleomagnetic data from the ~1.63 Ga diabase and quartz porphyry dikes of the Fennoscandian Shield (see section 3.6), from the 1.64 Ga mafic dikes of the Amazonian craton (Hargraves, 1968; Onstott et al., 1984), and from the ~1.64 Ma sedimentary rocks from the McArthur Basin of Australia (Idnurm et al., 1995) allow a reconstruction of these cratons. Tentatively, the less well-defined data from the 1.70–1.65 Ga quartzites in Laurentia (Chandler and Morey, 1992) have also been studied in order to see its relative position with respect to these continents (Pesonen et al., 2003). Compared to the paleoposition at 1.83 Ga, both the Fennoscandian Shield and the Amazonian craton had drifted to the equator, while Australia was located at a paleolatitude of ~30°. Laurentia and the Fennoscandian Shield were latitudinally ~15º apart, but considering the large errors in the data from Laurentia (concerning both the age and paleomagnetic pole) the two continents may have been joined as well. If the first polarity option of Figure 15.3B is applied to the reconstruction at 1.63 Ga (Figure 15.3C, model 1), the Fennoscandian Shield had rotated counterclockwise by ~40° and the Amazonian craton by ~10°, but both had preserved approximately the same relative orientation as at 1.83 Ga. Based on the similarities in paleolatitude and relative orientation, it was previously suggested that before ~1.65 Ga the Fennoscandian Shield collided with the Amazonian craton resulting in the formation of the Gothian orogenic belt in the southwestern Fennoscandian Shield and the Rio Negro–Juruena belt in the western Amazonian craton (Mertanen and Pesonen, 2000; Pesonen et al., 2000, 2001; Elming et al., 2001; Pesonen and Mertanen, 2002). However, according to the present interpretation (Pesonen et al., 2003), this model is not favored any more. According to Geraldes et al. (2001), the Rio Negro–Juruena province H I S T O RY
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was formed in two different events at 1.8–1.7 Ga and 1.56–1.4 Ga with a paucity in tectonic and magmatic activity between the periods that each involved accretion and formation of juvenile magmas. Because the Amazonian craton and the Fennoscandian Shield show similar long-lasting destruction of oceanic crust at the craton margins, Geraldes et al. (2001) suggested that the two cratons formed a laterally contiguous continental margin at 1.6–1.5 Ga. The joint assembly of the cratons is also supported by the occurrence of roughly coeval rapakivi granites on both continents (Rämö and Haapala, 1995; Bettencourt et al., 1999; Tassinari et al., 2000). However, the paleomagnetic reconstruction places both the Amazonian craton and the Fennoscandian Shield at the same equatorial paleolatitude, thus making it impossible that the cratons were linearly contiguous. The disagreement of paleomagnetic data with geological observations is not solved by polarity ambiguity (option 2 Figure 15.3C) or by error limits of the poles, as both cratons are at the same paleolatitude. Therefore, the present paleomagnetic data imply that the cratons were not joined at 1.65 Ga, although both cratons experienced almost coeval juvenation at craton margins. An alternative is that the pole ages from the Amazonian craton are erratic. The data come from dikes dated by the Rb-Sr or K-Ar methods and, consequently, the pole ages may represent magnetizations that postdate the primary ages of the dikes.
4.4. Middle Mesoproterozoic Figure 15.3D shows the reconstruction at 1.25 Ga. Data are available from the Fennoscandian Shield, Laurentia, Amazonia, and Congo/São Francisco cratons. Paleomagnetic studies from the 1236 ± 24 Ma late Kibaran formations in the Congo craton (Meert et al., 1994) indicate that the united Congo/São Francisco craton was at equatorial latitudes at ~1.25 Ga, but, as will be shown below, drifted towards higher 660
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latitudes later on. Data for the Amazonian craton come from the Nova Floresta gabbros and basalts in western Brazil (Tohver et al., 2002). For Laurentia, paleomagnetic data have been obtained from the 1267 Ma Mackenzie dikes (Buchan and Halls, 1990) and from the 1235 ± 7 Ma Sudbury dikes (Palmer et al., 1977). For the Fennoscandian Shield, the data come from the 1.26 Ga Postjotnian dikes (Section 3.7). Based on these data, the Fennoscandian Shield and Laurentia formed a unity at 1.25 Ga when a global scale tensional regime was manifested by the emplacement of coeval mafic dike swarms in both continents. This reconstruction is supported by studies on anisotropy of magnetic susceptibility (AMS) (Elming and Mattson, 2001) that show a similar magma flow direction for coeval mafic sills and basalts of Greenland and the Central Scandinavian dolerite group, when the Fennoscandian Shield and Laurentia are put in this relative orientation (see Pesonen et al., 2003). As discussed before, at 1.83 Ga the Fennoscandian Shield and Laurentia had approximately the same relative orientation as at 1.25 Ga and both continents had drifted ~30° towards the equator and rotated counterclockwise by ~90°. Despite some differences, the close resemblance of the 1.83 Ga and 1.25 Ga reconstructions may indicate that during this time interval the two cratons drifted together (see Buchan et al., 2000). The differences may be explained by errors of pole positions in both reconstructions as they allow for variations of ± 10° in both latitudes and rotations. Evidence supporting their having drifted together comes from the coeval ~1.6 Ga Gothian and Labradorian belts in Fennoscandia and Laurentia, respectively, that become continuous in this reconstruction (see Figure 15.3C).
4.5. Late Mesoproterozoic Continental reconstructions during the Late Mesoproterozoic–Neoproterozoic were dis-
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cussed in Pesonen et al. (2003). Here, a short summary of the present knowledge of the paleomagnetic evolution of the Fennoscandian Shield during ~1.1–1.05 Ga is given. At ~1.2 Ga, Fennoscandia and Laurentia started to drift apart, towards opposite paleolatitudes. The Fennoscandian Shield drifted towards southern paleolatitudes which led to the collision of the southwestern part of the Fennoscandian Shield with the Congo/São Francisco craton (see Figure 14.3D) at ~1.1 Ga. The collision resulted in the Sveconorwegian orogeny in the Fennoscandian Shield. Although reliable paleomagnetic data from that time from the Amazonian craton is missing, geological data imply that the Amazonian craton collided with Laurentia, resulting in the Grenvillian orogeny. A hypothetical Congo Sea was opened between the Congo/ São Francisco–Fennoscandia and Amazonia–Laurentia landmasses. At 1.1 Ga, eastern Gondwanaland, predominantly comprising Australia, East Antarctica, and India was located north of Laurentia. Later on, at ~1.05 Ga, Laurentia drifted to lower paleolatitudes, while the Fennoscandian Shield maintained roughly the same position as at 1.1 Ga. The drift of Laurentia led to a collision between southeastern Greenland and northern Fennoscandian Shield producing ‘late’ Grenvillian type events (see Pesonen et al., 2003). Finally, the Fennoscandian Shield and Laurentia amalgamated with the Congo/ São Francisco, Amazonia, Gondwanaland, Siberia, South China, and Kalahari to form the supercontinent Rodinia at ~1.05 Ga.
5. Conclusions Paleomagnetic evidence for the evolution of the Fennoscandian Shield shows that during several time periods, the Fennoscandian Shield has formed a part of a larger continental assembly or a supercontinent that has had significant effects on its geological evolution. According to the present knowledge, Laurentia, which has CHAPTER
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experienced similar orogenies, extensional regimes, and magmatism, is the continent that has most often been joined with the Fennoscandian Shield during the Precambrian. Paleomagnetic data imply that Laurentia collided with the present northern part of the Fennoscandian Shield, thus giving rise to parts of the Svecofennian and Hudsonian orogenies at ~1.9–1.8 Ga. In addition to similar orogenic belts, the shared geological history of the Fennoscandian Shield and Laurentia is also seen in the cratons’ simultaneous riftings. At ~2.45 Ga and 1.25 Ga, extensional tectonism in both cratons caused the emplacement of major mafic dike swarms. At both times, the dike swarms become aligned (2.45 Ga) or continuous (1.25 Ga) if the cratons are placed relative to each other as suggested by the reconstructions presented here. In many cases, the 1.8–1.25 Ga paleomagnetic data from the Amazonian craton also match those from the Fennoscandian Shield. However, the present paleomagnetic data cannot reliably demonstrate that the Amazonian craton was attached to the Fennoscandian Shield, although it has been close to the joined Fennoscandia–Laurentia during most of the Precambrian.
Acknowledgments We wish to thank all those people – Niels Abrahamsen, Göran Bylund, Sten-Åke Elming, Manoel D’Agrella-Filho, Håkan Mattson, Joe Meert, Sergey Pisarevsky, and Philip Schmidt – who participated in the Precambrian subgroup of the Fourth Nordic Paleomagnetic Workshop in Aarhus, Denmark in 1999, for their invaluable contribution to the preparation of the continental reconstructions. The paper was improved by the kind comments on the manuscript of Pesonen at al. (2003) by K.-I. Åhäll and C. Klootwijk.
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Programme and Extended Abstracts, Espoo, Finland, November 12-13, 2002. Institute of Seismology, University of Helsinki, Report S-42, 103-109. Pesonen, L.J., Neuvonen, K.J., 1981. Palaeomagnetism of the Baltic Shield - implications for Precambrian tectonics. In: A. Kröner (Ed.), Precambrian Plate Tectonics. Elsevier, Amsterdam, 623–648. Pesonen, L.J., Stigzelius, E., 1972. On petrophysical and paleomagnetic investigations of the gabbros of the Pohjanmaa region, MiddleWest Finland. Geol. Surv. Finland, Bull. 260, 1–27. Pesonen, L.J., Torsvik, T.H., 1990. The Fennoscandian palaeomagnetic database: compilation of palaeomagnetic directions and poles from the Northern Segment of the EGT. In: R. Freeman, S. Mueller (Eds.), Proceedings of the Sixth Workshop on the European Geotraverse (EGT) Project: data compilations and synoptic interpretation. Einsiedeln, Switzerland, 29.11.-5.12.1989, 389–399. Pesonen, L.J., Huhma, H., Neuvonen, K.J., 1986. Palaeomagnetic and Sm-Nd isotopic data of the Late Precambrian Laanila diabase dike swarm, northeastern Finland. 17e Nordiska Geologmötet, Helsingfors Universitet, Finland, Abstracts, 149. Pesonen, L.J., Torsvik, T.H., Elming, S.-Å., Bylund, G., 1989. Crustal evolution of Fennoscandia – palaeomagnetic constraints. Tectonophysics 162, 27–49. Pesonen, L.J., Bylund, G., Elming, S.-Å., Torsvik, T.H., Mertanen, S., 1991. Catalogue of palaeomagnetic directions and poles from Fennoscandia: Archaean to Tertiary. Tectonophysics 195, 151–207. Pesonen, L.J., Elming, S.-Å., Pisarevsky, S., Mertanen, S., D´Agrella-Filho, M., Meert, J., Schmidt, P., Bylund, G., 2000. A pre-Rodinia Supercontinent? - A paleomagnetic survey. 25th General Assembly, EGS Symposium, Nice, Geophys. Res. Abstracts 2. [CD-ROM] Pesonen, L.J., Mertanen, S., Elming, S.-Å., 2001. Reconstructions of continents during the Proterozoic - a way towards Rodinia. In: K.N. Sircombe, Z.X. Li (Eds.), From basins CHAPTER
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to mountains: Rodinia at the turn of the century. Geol. Soc. Australia, Abstracts 65, 82–84. Pesonen, L.J., Elming, S.-Å., Mertanen, S., Pisarevsky, S.A., D´Agrella-Filho, M.S., Meert, J.G., Schmidt, P.W., Abrahamsen, N., Bylund, G., 2003. Palaeomagnetic configuration of continents during the Proterozoic. Tectonophysics. (accepted for publication) Piper, J.D.A., 1976. Palaeomagnetic evidence for a Proterozoic supercontinent. Phil. Trans. Roy. Soc. London, A 280, 469–490. Piper, J.D.A., 1979. Palaeomagnetism of the Ragunda intrusion and dolerite dykes, central Sweden. Geol. Fören. Stockholm, Förhandl. 101, 139–148. Piper, J.D.A., 1980. Palaeomagnetic study of the Swedish rapakivi suite: Proterozoic tectonics of the Baltic Shield. Earth Planet. Sci. Lett. 46, 443–461. Pisarevsky, S.A., Bylund, G., 1998. Palaeomagnetism of a key section of the Protogine Zone, southern Sweden. Geoph. J. Int. 133, 185–200. Pisarevsky, S.A., Sokolov, S.J., 1999. Palaeomagnetism of the Palaeoproterozoic ultramafic intrusion near Lake Konchozero, Southern Karelia, Russia. Precambrian Res. 93, 201–213. Pisarevsky, S.A., Sokolov, S.J., 2001. The magnetostratigraphy and a 1780 Ma palaeomagnetic pole from the red sandstones of Vazhinka River section, Karelia, Russia. Geoph. J. Int. 146, 531–538. Poorter, R.P.E., 1975. Palaeomagnetism of Precambrian rocks from southeast Norway and south Sweden. Phys. Earth Planet. Int. 10, 74–87. Puchtel, I.S., Bogatikov, O.A., Kulikov, V.V., Kulikova, V.V., Zhuravlev, D.Z., 1995. The role of crust and mantle sources on the petrogenesis of continental magmatism: isotopic and geochemical data from Early Proterozoic picrites of Onega plateau, Baltic shield. Petrologia 3, N4, 397–419. (in Russian) Rämö, O.T., Haapala, I., 1995. One hundred years of Rapakivi Granite. Mineral. Petrol. 52, 129–185. Rogers, J.J.W., 1996. A history of continents in the past three billion years. J. Geol. 104, H I S T O RY
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91–107. Schmidt, P.W., Williams, G.E., 1999. Paleomagnetism of the Paleoproterozoic hematitic breccia and paleosol at Ville-Marie, Québec: further evidence for the low paleolatitude of Huronian glaciation. Earth Planet. Sci. Lett. 172, 273–285. Stearn, J.E.F., 1979. Palaeomagnetism of the Sveconorwegian rocks of the Fennoscandian Shield. Ph. D. Thesis, Sub-Department of Geophysics, University of Liverpool, England. Stearn, J.E.F., Piper, J.D.A., 1984. Palaeomagnetism of the Sveconorwegian mobile belt of the Fennoscandian Shield. Precambrian Res. 23, 201–246. Suominen, V., 1991. The chronostratigraphy of southwestern Finland with special reference to Postjotnian and Subjotnian diabases. Geol. Surv. Finland, Bull. 356, 1–100. Tassinari, C.C.G., Bettencourt, J.S., Geraldes, M.C., Macambira, M.J.B., Lafon, J.M., 2000. The Amazonian Craton. In: U.G. Cordani, E.J. Milani, A. Thomaz-Filho, D.A. Campos (Eds.), Tectonic Evolution of South America. 31st Internat. Geol. Congr., Rio de Janeiro, Brazil, pp. 41–95. Tohver, E., Van der Pluijm, B.A., Van der Voo, R., 2002. Refining Rodinia: new paleomagnetic results from Amazonia and paleogeographic implications for the Grenville orogeny. EGS XXVII General Assembly, Nice, France, April 2002 [CD-ROM] Torsvik, T.H., 1998. Fennoscandian Data Base Software for Windows 98 [CD-ROM]
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Tuisku, P., Huhma, H., 1998. SIMS dating of zircons: metamorphic and igneous events of the Lapland granulite belt are 1.9 Ga old, provenance is Palaeoproterozoic and Archaean (2.0-2.9 Ga) and the tectonic juxtaposition about 1.9-1.88 Ga old. In: N. Philippov (Comp.), SVEKALAPKO, An Europrobe project, 3rd Workshop, Repino, Russia, 26.-29.11.1998. Abstracts, 64–65. Vuollo, J., Kamo, S.L., Huhma, H., Mertanen, S., Halls, H.C., Stepanov, V.S., 1996. Geochemistry and U-Pb and Sm-Nd isotope studies of a 2.45 Ga mafic dyke swarm in the eastern Fennoscandian Shield and correlation to North Atlantic Area. In: F.C. Gower (Comp.), Proterozoic evolution in the North Atlantic realm. Goose Bay, Labrador, Canada, July 29th-August 2nd, 1996, Program and Abstracts, St. John’s: Geological Survey. Department of Mines and Energy, 186–187. Williams, G.E., 1993. History of the Earth’s obliquity. Earth-Sci. Rev. 34, 1–45. Williams, G.E., Schmidt, P.W., 1997. Paleomagnetism of the Paleoproterozoic Gowganda and Lorrain formations, Ontario: low paleolatitude for Huronian glaciation. Earth Planet. Sci. Lett. 153, 157–169. Williams, H., Hoffman, P.F., Lewry, J.F., Monger, J.W.H., Rivers, T., 1991. Anatomy of North America: thematic portrayals of the continent. Tectonophysics 187, 117–134. , Bull. 260, 1–27.
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Chapter 16
PALEOPROTEROZOIC CARBON ISOTOPE EXCURSION
J.A. Karhu
Cover page: Paleoproterozoic dolomite with columnar stromatolite structure Columnacollenia rautamaa, Peräpohja belt, Peuranpalo, Tervola. Photo: Jukka Lehtinen.
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Karhu, J.A., 2005. Paleoproterozoic carbon isotope excursion. 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. 669–680. © 2005 Elsevier B.V. All rights reserved.
The isotope composition of carbon in Paleoproterozoic carbonate sediments indicates that the isotope composition of marine dissolved carbon underwent a large positive excursion between ~2.2 Ga and 2.1 Ga. First indications for the excursion were obtained from Fennoscandian supracrustal successions, but, subsequently, the existence of a global event that affected the Paleoproterozoic carbon cycle has been confirmed by numerous studies from different continents. Isotope ages from the Fennoscandian sedimentary sequences indicate that the minimum duration of the excursion was 100 Ma, from 2.21 Ga to 2.11 Ga. The start of the excursion is relatively poorly constrained between 2.32 Ga and 2.21 Ga, based on Fennoscandian and South African data, but the end of the event is well-defined between 2.11 Ga and 2.06 Ga. New data from the Väystäjä Formation, northern Finland, confirm the end of the excursion at >2.05 Ga. The Paleoproterozoic carbon isotope excursion indicates a major perturbation in the carbon cycle, and it has been related to the increase of the atmospheric O2 content. Recently published carbon isotope records from the Duitschland Formation, South Africa, give indications of a separate positive carbon isotope shift of undetermined duration and of unknown global extent, at ~2.32 Ga. In contrast to the lithologic record of the major excursion, the Duitschland carbonate sediments are closely associated with glacial sediments, analogously with the relationship observed between Neoproterozoic 13C-enriched carbonate sediments and glaciations.
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1. Introduction The Paleoproterozoic oceans were affected by a dramatic carbon isotope excursion between 2.2 Ga and 2.1 Ga (Karhu and Holland, 1996). Evidence for the excursion is recorded in carbonate sediments deposited during that time. The observed δ13C values are systematically higher than 4‰, and values exceeding 10‰ are not uncommon [δ13C gives the 13C/12C ratio R of the sample as a per mil difference relative to the V-PDB standard: δ13C = (Rsample/ Rstandard – 1) x 103]. This is in strong contrast to the record from preceding or following time periods, where the δ13C values remained close to a value of 0‰ (Strauss et al., 1992). Because the oceans are a major reservoir of carbon on the surface of the Earth, the excursion implies a significant perturbation in the global carbon cycle. In addition to the positive excursion in marine carbon isotope ratios, the Paleoproterozoic Earth was affected by other major environmental changes. These include a major magmatic event at 2.45 Ga, associated with a peak in the deposition of banded iron-formations (Heaman, 1997; Barley et al., 1997) and possibly three separate glacial intervals between 2.45 Ga and 2.22 Ga (Eyles and Young, 1994). The most important environmental change, however, was the rise of atmospheric oxygen at 2.3–2.0 Ga (Holland, 1994). For understanding the relationships between the carbon isotope excursion and the environmental changes, it is of great importance to know the form and the timing of shifts of the secular carbon isotope curve. The compilation of carbon isotope analyses from Paleoproterozoic sedimentary carbonate rocks by Karhu and Holland (1996) showed a single positive excursion in δ13C at ~2.2–2.1 Ga, but recently more complex secular patterns have also been suggested. Melezhik et al. (1999) presented a curve that included four separate positive excursions separated by distinctive minima in the in672
terval ~2.3–1.9 Ga. This multimodal pattern resembles the secular carbon isotope curve for the Neoproterozoic seawater, characterized by several long-lasting maxima punctuated by sharp minima during global glacial periods (Jacobsen and Kaufman, 1999). Most of the critical data points defining the beginning and end of the carbon isotope excursion represent sedimentary carbonate units of the Fennoscandian Shield. These units were deposited in basins associated with the rifting of the Archean craton and the formation of passive margins at ~2.2–2.1 Ga. In recent years, new isotope and age data have been published in Finland and elsewhere, offering a way to test and refine the details of the Paleoproterozoic secular carbon isotope curve.
2. Early records First indications of a major shift in the marine carbon isotope ratios during the Paleoproterozoic Era were obtained by Galimov et al. (1968) from Fennoscandian supracrustal successions. They reported two carbon isotope analyses of dolomite from the Keivy belt of the Kola Peninsula yielding δ13C values of 8.0‰ and 8.1‰ and two analyses from the Karelian sedimentary units of the Raiguba locality with δ13C of 6.3‰ and 9.1‰. Galimov et al. (1975) complemented these data by publishing a mean of 7.4‰ for five sedimentary carbonate samples from the Keivy belt in the Kola Peninsula and a mean of 7.5‰ for 20 samples from the Central Karelian region. A comprehensive study of Precambrian sedimentary carbonates by Schidlowski et al. (1975) included eight sedimentary dolomite samples from the Peräpohja belt in northern Finland, which is geographically distinct from the sedimentary units sampled by Galimov et al. (1968, 1975). The dolomite samples yielded δ13C values between 3.1‰ and 8.6‰, with a mean at 5.2‰. In the same paper, Schidlowski et al. (1975) also reported the first δ13C data from the
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Lomagundi Group in Zimbabwe. Excluding four samples with an uncertain stratigraphic setting, the eleven Lomagundi samples showed highly positive δ13C values between 7.3‰ and 13.4‰, with a mean at 9.7‰. The authors noticed the high δ13C values, but contended that “there is no reason to invoke special environmental conditions for the deposition of the Lomagundi suite”. The unusual isotopic characteristics of the Lomagundi carbonates led Schidlowski et al. (1976) to resample the Lomagundi Group for a new detailed study. They analyzed 67 dolomite samples, that yielded a mean δ13C of 8.2‰. In addition, they demonstrated that the Lomagundi carbonates were well preserved and had most probably retained their primary carbon isotope compositions. Motivated by problems in applying a global carbon cycle model to explain the observed record, the authors changed their earlier view and suggested a local unusual sedimentary environment as the cause of the enrichment of heavy carbon isotopes in the Lomagundi carbonates. The notion of local conditions as the cause of the enrichment in 13C was challenged when new Paleoproterozoic carbonate successions with 13C-enriched carbon isotope signatures were found locally in Norway (Baker and Fallick, 1989a), Scotland (Baker and Fallick, 1989b), the Ukraine (Zagnitko and Lugovaya, 1989), and regionally over an extensive area in the Fennoscandian Shield (Karhu, 1989, 1993; Yudovich et al., 1991). Since then, the existence of a global event affecting the Paleoproterozoic carbon cycle has been confirmed by numerous studies from different continents and shields, including North America (Melezhik et al., 1997; Bekker et al., 2003), South America (Bekker et al., 2003), Fennoscandia (Karhu and Melezhik, 1992; Melezhik et al., 1999); Africa (Master et al., 1993; Buick et al., 1998; Bekker et al., 2001), Australia (Lindsay and Brasier, 2002), and India (Maheshwari et al., 1999; Sreenivas et al., 2001).
3. Fennoscandian δ13C data The δ13C analyses of the Paleoproterozoic sedimentary carbonates from the Fennoscandian Shield show a bimodal distribution with pronounced maxima at about 1‰ and 10‰ (Karhu, 1993). The sedimentary carbonate units with δ 13 C at about 1‰ are distributed over the whole shield area. In contrast, sedimentary carbonates with δ13C > 4‰ are, without exceptions, restricted to the Paleoproterozoic successions deposited on the Archean craton. Many of these carbonate units have been associated with the informal Jatulian Group, deposited during the Paleoproterozoic rifting of the Archean craton. Clearly, the formation of these carbonate sediments cannot be attributed to the operation of fractionating processes in an unusual sedimentary environment, as the carbonate sediments with δ13C > 4‰ were deposited regionally in an area of ~1200 km by 600 km (Figure 16.1). Within individual supracrustal belts of the Fennoscandian Shield, the carbon isotope values show a continuous pattern of 13C enrichment and depletion with a single positive excursion. Examples are the Peräpohja and Kuusamo belts in northern Finland (Karhu, 1993), the Pechenga and Imandra–Varzuga belts in the Kola Peninsula (Karhu and Melezhik, 1992), and the Tulomozerskaya Formation of the Onega basin in the eastern part of the shield (Melezhik et al., 1999). In the Tulomozerskaya Formation, the complete thickness of the highly 13C-enriched dolomites reaches ~800 m (Melezhik et al., 1999). Although individual stratigraphic sequences are incomplete by nature, the combined evidence strongly suggests the existence of a single long-lasting positive carbon isotope excursion. The Fennoscandian sedimentary successions provide chronostratigraphic constraints for the excursion (Karhu, 1993). The isotope data from the Kuusamo and Peräpohja belts suggest that the excursion started before 2.21
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36°
24°
NORWAY 68° 68°
RUSSIA
SWEDEN
FINLAND
60°
60°
200 km 36°
24°
δ13C ≥ 4 per mil
Predominantly Phanerozoic
Archean
δ13C < 4 per mil
Paleoproterozoic
Margin of the Archean craton
Fig. 16.1. Areal distribution of 13C-enriched, Paleoproterozoic sedimentary carbonates in the Fennoscandian Shield. Circles designate analytical data from Karhu (1993), triangles refer to data of Karhu and Melezhik (1992) from the Pechenga and Imandra–Varzuga belts in Russia and of Baker and Fallick (1989b) from the Lofoten–Vesterålen area in Norway.
Ga. The carbonate units with highly positive δ13C values overlie an unconformity and an interval of intensive weathering at ~2.3 Ga, but a definitive lower age limit for these suc674
cessions is given by a U-Pb zircon age of 2405 ± 6 Ma from an igneous pebble in the basal conglomerate underlying the carbonate units at Kuusamo (Silvennoinen, 1991). This date,
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Table 16.1 Carbon and oxygen isotope values and major and trace element concentrations of carbonates from the Väystäjä Formation, northern Finland. Sample code
δ13C δ18O ‰, VPDB‰,
67B-BES-76 –1.59 11D-HAS-74 –0.65 53-BES-76 –0.15 77A-BES-77 0.30
–16.11 –17.63 –11.81 -15.63
Caa VPDB
Mga %
Fea %
Mna %
Sra %
Mn/Sr ppm
Dolb wt.%
23.92 23.27 23.47 25.71
11.45 9.91 11.33 8.03
1.45 4.20 1.22 4.02
0.43 0.29 0.19 0.29
125 273 69 157
34 11 28 18
98 95 98 84
Grid coordinates x y 7352.45 7355.30 7352.65 7353.18
2496.17 2517.78 2492.75 2505.95
a) 0.5 M acetic acid soluble fraction measured by ICP-AES at the Geological Survey of Finland (GTK) b) Wt.% dolomite in total carbonate fraction by XRD
however, cannot be used as a verification of the higher age limit of the excursion, because even the lowermost sedimentary carbonates have elevated δ13C values of > 4‰. The evidence for the end of the excursion is based on isotopic records from two sedimentary successions, one at Kiihtelysvaara and the other at Siilinjärvi. The lavas of the Koljola Formation at Kiihtelysvaara have been dated by U-Pb zircon to 2113 ± 4 Ma and they underlie a dolomite sequence showing a decreasing trend in δ13C. In contrast, the felsic lavas of the Koivusaari Formation at Siilinjärvi, dated by U-Pb zircon to 2062 ± 2 Ma, overlie a sedimentary carbonate unit showing a drop in δ13C. Combined, these results lead to a conclusion that the carbon isotope excursion ended between 2113 ± 4 and 2062 ± 2 Ma (see Karhu, 1993, for age references). New evidence from northern Finland confirms the age brackets for the end of the excursion. Perttunen and Vaasjoki (2001) reported a U-Pb zircon age of 2050 ± 8 Ma (A643, Keinokangas) for a felsic effusive porphyry in the Väystäjä Formation of the Peräpohja belt. The formation comprises black shales and dolomites interlayered with mafic and minor felsic volcanic rocks. The contact with the underlying Kivalo Group is not exposed. Four dolomite samples were collected from outcrops and analyzed for the isotope composition of carbon and oxygen at the Geological
Survey of Finland, using methods described by Karhu (1993). Isotope and chemical data for the sedimentary carbonates of the Väystäjä Formation are given in Table 16.1. The carbon isotope values show very little variation from –1.6‰ to +0.3‰, with a mean at –0.52 ± 0.81‰. No correlation is evident between δ13C and δ18O or δ13C and the Mn/Sr ratio, suggesting that the carbon isotope ratios of dolomites have largely retained their primary isotopic signatures. The Väystäjä Formation overlies the carbonate units of the Kivalo Group with distinctly higher δ13C values. Accordingly, the new data from the Väystäjä Formation provide critical information corroborating the earlier conclusion about the end of the carbon isotope excursion. By 2050 ± 8 Ma, the δ13C value of marine bicarbonate had returned back to a value of ~0‰.
4. Global δ13C data Karhu and Holland (1996) extended the Fennoscandian carbon isotope curve presented by Karhu (1993) to include carbonates from other continents. Their secular isotope curve is shown in Figure 16.2, supplemented by a few new data points from recent literature. Only formations with depositional ages constrained better than ±125 Ma have been included in the
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compilation. Most ages are based on U-Pb zircon dating of magmatic units, but in the absence of these, less reliable carbonate Pb-Pb ages have been accepted. Schidlowski and Todt (1998) reported a Pb-Pb carbonate age of 2150 ± 50 Ma for the Lomagundi carbonates of the Magondi belt in Zimbabwe, which replaces the old less precise estimate. Melezhik et al. (1997) reported highly positive δ13C values with a mean of 9.7 ± 2.8‰ for the Pistolet and Seward subgroups of the Labrador Trough, Canada. On the basis of U-Pb zircon data (Rohon et al., 1993), these units were deposited between 2169 ± 2 Ma and 2142 ± 4 Ma; analytical data for the latter age have not been published yet. Sedimentary carbonates in the Juderina Formation of the Yerrida Basin in Western Australia were investigated by Lindsay and Brasier (2002), who analyzed eleven carbonate samples yielding a mean δ13C value of 7.2 ± 0.7‰. Woodhead and Hergt (1997) reported a Pb-Pb carbonate age of 2170 ± 60 Ma for the Juderina Formation. Two data points from the compilation of Karhu and Holland (1996) were excluded on the basis of new geological evidence (Figure 16.2). Dolomites addressed to the >2.2 Ga Sompujärvi Formation of the Peräpohja belt in northern Finland were removed. New geological observations suggest that these dolomites may actually represent the overlying Palokivalo Formation, with an undefined depositional age with respect to 2.2 Ga (Vesa Perttu nen, pers. comm., 2002). Also the data point representing the Pretoria Group of South Africa was discarded. The Pretoria Group is a thick supracrustal succession covering several 100 Ma of geologic time. The age for the Hekpoort Formation of the Pretoria Group was derived on the basis of stratigraphical correlation with the Ongeluk andesite of the Griqualand West Basin (Karhu and Holland, 1996). The latter volcanic unit has been dated to 2222 ± 13 Ma on the basis of Pb-Pb whole-rock analyses (Cornell et al., 1996). A new carbonate Pb-Pb age from the 676
Mooidraai Formation, overlying the Ongeluk andesite suggests that the depositional age of the volcanic units may be significantly older than 2.2 Ga (Bau et al., 1999). Recently Hannah et al. (2003) reported a Re-Os age of 2322 ± 15 Ma for the Rooihoogte Formation of the Pretoria Group. This date provides the definitive upper age limit for the Paleoproterozoic carbon isotope excursion.
5. Discussion Based on data from the Fennoscandian and South African sedimentary successions, the Paleoproterozoic carbon isotope excursion started between 2.32 Ga and 2.21 Ga and ended between 2.11 Ga and 2.06 Ga. The minimum duration of the excursion is 100 Ma, and during this time the δ13C signatures in marine carbonates reached values of 10‰ or even higher (Figure 16.2). The carbon isotope data and isotope age constraints from other shield areas are roughly compatible with these age limits and the general form of the secular δ13C curve, determined largely by the Fennoscandian sedimentary successions. It is notable that no Paleoproterozoic sedimentary carbonate units with δ13C exceeding +4‰ are known to have been deposited after 2.06 Ga. The Lucknow Formation of the Griqualand West Basin in South Africa is characterized by highly 13C-enriched carbonates, and it has been suggested to represent a separate positive excursion at 1.9 Ga (Buick et al., 1998; Melezhik et al., 1999). New field observations, however, support an interpretation that the Lucknow Formation is older than the 2.06 Ga Bushveld complex and correlative to the carbonate successions deposited during the major excursion at 2.2–2.1 Ga in the Transvaal basin (Bekker et al., 2001). The general form of the Paleoproterozoic secular carbon isotope curve appears to be fundamentally different from that of the Neoproterozoic curve. The latter is character-
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15 Labrador Trough Lomagundi Group
δ13C (‰, PDB)
10 Juderina Formation 5
Väystäjä Formation
0
-5 2.6
2.4
2.0
1.8
1.6
1.4
Age (Ga) Fig. 16.2. Variation in the isotope composition of carbon in sedimentary carbonates during the Paleoproterozoic time. Open circles represent data from Fennoscandian successions from Karhu (1993), closed circles indicate carbonate units outside Fennoscandia from Karhu and Holland (1996). Stars indicate new data from the Labrador Trough (Melezhik et al., 1997), the Juderina Formation, Australia (Lindsay and Brasier, 2002), and the Väystäjä Formation in northern Finland (this study). In addition, the old isotope age estimate for the Lomagundi Group has been replaced by the Pb-Pb carbonate age from Schidlowski and Todt (1998). The envelope including all data points matches exactly the one drawn by Karhu and Holland (1996).
ized by long periods of positive δ13C values punctuated by sharp minima associated with glaciations (Jacobsen and Kaufman, 1999). In contrast, the Paleoproterozoic carbonate units deposited during the ~2.2–2.1 Ga excursion followed the three Paleoproterozoic glacial events. In the Fennoscandian Shield, glaciallyinfluenced sedimentary successions occur in the Sariolian Group, but they are separated by a ~2.3 Ga unconformity (Marmo and Ojakangas, 1984) from the overlying dolomite units with high δ13C values. The carbon isotope data presented by Buick et al. (1998) and Bekker et al. (2001) have supplied evidence for a separate positive δ13C peak that predates the major excursion at 2.2–2.1 Ga and is significantly younger than the unconformably underlying 2.48 Ga Penge BIF. The data come from the Duitschland Formation of the Transvaal Basin in South
Africa, where carbonate units with negative δ13C values overlie glacial diamictites and are separated by a sequence boundary from the overlying carbonates with highly positive δ13C values. The 2322 ± 15 Ma Rooihoogte Formation of the Transvaal Basin (Re-Os; Hannah et al., 2003) is considered to be correlative with the Duitschland Formation on the basis of sequence stratigraphic arguments. The meaning of these observations is open for two reasons. First, while the Duitschland Formation clearly underlies carbonate units formed during the 2.2–2.1 Ga excursion, independent evidence for the duration of the Duitschland carbon isotope excursion is lacking. Second, the global significance of this carbon isotope shift is unknown. Bekker et al. (2001) suggested deposition in open marine conditions, but comparable carbon isotope data from other depositional basins would be needed to
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confirm the existence of a global change in δ13C. The marine carbon isotope record from the period of Paleoproterozoic glaciations is still poorly defined. The long-term record of the carbon isotope composition of sedimentary carbonates is related to the operation of the geochemical carbon cycle. Carbon cycle mass balance considerations indicate that the excursion was associated with an increase in the fractional burial rate of organic carbon relatively to carbonate carbon (e.g., Karhu, 1993). This in turn would imply a large quantity of oxygen released as a by-product of organic carbon burial. Many lines of geochemical and geological evidence have suggested a significant rise in the atmospheric oxygen levels at 2.3– 2.0 Ga (Holland, 1994). Recently, the finding of mass independent fractionation (MIF) of sulfur isotopes of sulfides and sulfates from sedimentary units older than 2.47 Ga and the lack of MIF in sulfur isotopes of sulfates and sulfides from sedimentary units younger than ~2.4 Ga has provided independent evidence for a major change in the oxidation state of the atmosphere (Farquhar et al., 2000). The cause and effect relationship between tectonic and environmental events leading to this fundamental change are still poorly known. More studies are needed to clarify these relationships and to illuminate the significance of the 2.2–2.1 Ga δ13C excursion for the irreversible oxidation of the atmosphere, oceans, and the surface environments in general.
6. Conclusions Carbon isotope records of Paleoproterozoic carbonate sediments provide strong evidence for a major positive excursion in the δ13C values of marine dissolved carbon. Isotopic ages from the Fennoscandian successions indicate that the minimum duration of the excursion was 100 Ma, from 2.21 Ga to 2.11 Ga. The start of the excursion is poorly constrained 678
between 2.32 Ga and 2.21 Ga, but the end of the event is well defined between 2.11 Ga and 2.06 Ga. New data from the Väystäjä Formation, northern Finland, confirm the end of the carbon isotope excursion to >2.05 Ga. Recently published carbon isotope record of the Duitschland Formation, South Africa, gives indication of a separate positive carbon isotope shift of undetermined duration, preceding the major excursion at 2.21–2.11 Ga, but the global extent of this event is not known.
Acknowledgment I thank A. Bekker for his critical review of the manuscript.
References Baker, A.J., Fallick, A.E., 1989a. Evidence from Lewisian limestones for isotopically heavy carbon in two-thousand-million-year-old sea water. Nature 337, 352–354. Baker, A.J., Fallick, A.E., 1989b. Heavy carbon in two-billion-year-old marbles from Lofoten-Vesterålen, Norway: Implications for the Precambrian carbon cycle. Geochim. Cosmochim. Acta 53, 1111–1115. Barley, M.E., Pickard, A.L., Sylvester, P.J., 1997. Emplacement of a large igneous province as a possible cause of banded iron formation 2.45 billion years ago. Nature 385, 55–58. Bau, M., Romer, R.L., Lüders, V., Beukes, N.J., 1999. Pb, O, and C isotopes in silicified Mooid raai dolomite (Transvaal Supergroup, South Africa): implications for the composition of Paleoproterozoic seawater and ‘dating’ the increase of oxygen in the Precambrian atmosphere. Earth Planet. Sci. Lett. 174, 43–57. Bekker, A., Kaufman, A.J., Karhu, J.A., Beukes, N.J., Swart, Q.D., Coetzee, L.L., Eriksson, K.A., 2001. Chemostratigraphy of the Paleoproterozoic Duitschland Formation, South Africa: Implications for coupled climate change and carbon cycling. Amer. J. Sci.
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301, 261–285. Bekker, A., Karhu, J.A., Eriksson, K.A., Kaufman, A.J., 2003. Chemostratigraphy of Paleoproterozoic carbonate successions of the Wyoming Craton: tectonic forcing of biogeochemical change? Precambrian Res. 120, 279–325. Buick, I.S., Uken, R., Gibson, R.L., Wallmach, T., 1998. High-δ13C Paleoproterozoic carbonates from the Transvaal Supergroup, South Africa. Geology 26, 875–878. Cornell, D.H., Schütte, S.S., Eglington, B.L., 1996. The Ongeluk basaltic andesite formation in Griqualand West, South Africa: submarine alteration in a 2222 Ma Proterozoic sea. Precambrian Res. 79, 101–123. Derry, L.A., Kaufman, A.J., Jacobsen, S.B., 1992. Sedimentary cycling and environmental change in the Late Proterozoic: Evidence from stable and radiogenic isotopes. Geochim. Cosmochim. Acta 56, 1317–1329. Eyles, N., Young, G.M., 1994. Geodynamic controls on glaciation in Earth history. In: M. Deynoux, J.M.G. Miller, E.W. Domack, N. Eyles, I.J. Fairchild, G.M. Young (Eds.), Earth’s Glacial Record. Cambridge University Press, 1–28. Farquhar, J., Bao, H., Thiemens, M., 2000. Atmospheric influence of Earth’s earliest sulfur cycle. Science 289, 756–758. Galimov, E.M., Kuznetsova, N.G., Prokhorov, V.S., 1968. On the problem of the old Earth atmosphere composition in connection with results of isotopic analysis of carbon from Precambrian carbonates. Geokhimia 11, 1376–1381. (In Russian with an English abstract) Galimov, E.M., Migdisov, A.A., Ronov, A.B., 1975. Variation in the isotopic composition of carbonate and organic carbon in sedimentary rocks during Earth’s history. Geochem. Intl. 12(2), 1–19. Hannah, J.L., Stein, H.J., Markey, R.J., Bekker, A., Holland, H.D., Re-Os dating and chondritic Os isotopic composition of Paleoproterozoic shale, Transvaal Supergroup, South Africa. Submitted. Heaman, L.M., 1997. Global mafic magmatism at 2.45 Ga: Remnants of an ancient large igneous province? Geology 25, 299–302. Holland, H. D., 1994. Early Proterozoic atmospheric change. In: S. Bengtson (Ed.), Early life
on Earth. New York, Columbia University Press, 237–244. Jacobsen, S.B., Kaufman, A.J., 1999. The Sr, C and O isotopic evolution of Neoproterozoic seawater. Chemical Geology 161, 37–57. Karhu, J.A., 1989. Extensive Early Proterozoic carbon isotope anomaly recorded in sedimentary carbonates and kerogens from the Fennoscandian Shield. Epstein 70th Birthday Symposium, Abstracts with Program. Pasadena: California Institute of Technology, 80–82. 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. Karhu, J.A., Melezhik, V.A., 1992. Carbon isotope systematics of early Proterozoic sedimentary carbonates in the Kola Peninsula, Russia: correlations with Jatulian formations in Karelia. In: V.V. Balagansky, F.P. Mitrofanov (Eds.), Correlation of Precambrian formations of the Kola-Karelia Region and Finland. Kola Scientific Centre of the Russian Academy of Sciences, Apatity, 48–53. Karhu, J.A., Holland, H.D., 1996. Carbon isotopes and the rise of atmospheric oxygen. Geology 24, 867–870. Lindsay, J.F., Brasier, M.D., 2002. Did global tectonics drive early biosphere evolution? Carbon isotope record from 2.6 to 1.9 Ga carbonates of Western Australian basins. Precambrian Res. 114, 1–34. Maheshwari, A., Sial, A.N., Chittora, V.K., 1999. High-δ13C Paleoproterozoic carbonates from the Aravalli Supergroup, Western India. Intl. Geol. Rev. 41, 949–954. Marmo, J.S., Ojakangas, R.W., 1984. Lower Proterozoic glaciogenic deposits, eastern Finland. Geol. Soc. Am., Bull. 95, 10551062. Master, S., Verhagen, B.T., Bassot, J.P., Beukes, N.J., Lemoine, S., 1993. Stable isotopic signatures of Paleoproterozoic carbonate rocks from Guinea, Senegal, South Africa and Zimbabwe: constrains on the timing of the ca. 2 Ga “Lomagundi” δ13C excursion. In: Symposium on Early Proterozoic Geochemical and Structural Constraints–Metallogeny. Publication Occasionnelle 1993/23, Dacar, Sénégal, 38–41.
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Melezhik, V.A., Fallick, A.E., Clark, T., 1997. Two billion year old isotopically heavy carbon: evidence from the Labrador Trough, Canada. Can. J. Earth Sci. 34, 271–285. Melezhik, V.A., Fallick, A.E., Medvedev, P.V., Makarikhin, V.V., 1999. Extreme 13Ccarb enrichment in ca. 2.0 Ga magnesite-stromatolite-dolomite-’red beds’ association in a global context: a case for the world-wide signal enhanced by a local environment. Earth-Sci. Rev. 48, 71–120. Perttunen, V., Vaasjoki, M., 2001. U-Pb geochronology of the Peräpohja Schist Belt, northwestern Finland. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Pap. 33, 45–84. Rohon, M.-L., Vialette, Y., Clark, T., Roger, G., Ohnenstetter, D., Vidal, Ph., 1993. Aphebian mafic-ultramafic magmatism in the Labrador Trough (New Quebec): its age and the nature of its mantle source. Can. J. Earth Sci. 30, 1582–1593. Schidlowski, M., Eichmann, R., Junge, C.E., 1975. Precambrian sedimentary carbonates: carbon and oxygen isotope geochemistry and implications for the terrestrial oxygen budget. Precambrian Res. 2, 1–69. Schidlowski, M., Eichmann, R., Junge, C.E., 1976. Carbon isotope geochemistry of the Precambrian Lomagundi carbonate province, Rhodesia. Geochim. Cosmochim. Acta 40, 449–455. Schidlowski, M., Todt, W., 1998. The Proterozoic
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Lomagundi carbonate province a paragon of a 13C-enriched carbonate facies: geology, radiometric age and geochemical significance. Abstracts of ICOG-9, 1998, Beijing. Chinese Sci. Bull. 43 (supplement), 114. Silvennoinen, A., 1991. Kuusamon ja Rukatunturin kartta-alueiden kallioperä. Summary: Pre-Quaternary rocks of the Kuusamo and Rukatunturi map-sheet areas. Explanation to the maps of Pre-Quaternary rocks, sheets 4524+4542 and 4613. Geological Map of Finland 1:100 000, 1–63. Sreenivas, B., Das Sharma, S., Kumar, B., Patil, D.J., Roy, A.B., Srinivasan, R., 2001. Positive δ13C excursion in carbonate and organic fractions from the Paleoproterozoic Aravalli Supergroup, Northwestern India. Precambrian Res. 106, 277–290. Strauss, H., Des Marais, D.J., Hayes, J.M., Summons, R.E., 1992. The carbon-isotopic record. In: J.W. Schopf, C. Klein (Eds.), The Proterozoic Biosphere: a Multidisciplinary Study. Cambridge University Press, New York, 117–127. Woodhead, J.D., Hergt, J.M., 1997. Application of the ‘double spike’ technique to Pbisotope geochronology. Chem. Geol. 138, 311–321. Yudovich, Ya.E., Makarikhin, V.V., Medvedev, P.V., Sukhanov, N.V. et al., 1991. Carbon isotope anomalies in carbonates of the Karelian Complex. Geochem. Intl. 28(2), 56–62. Zagnitko, V.N., Lugovaya, I.P., 1989. Isotope geochemistry of carbonate and banded iron formation rocks from the Ukraine Shield, Nauk. Dumka. Kiev, 1–316. (In Russian)
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Cover page: Geologists and undergraduates on an excursion at Kaivopuisto Park, Helsinki, about in the year 1890. Persons from left to right: Grigori Lisitzin, Benjamin Frosterus, Gustaf Komppa, K.A. Moberg (sitting), Th. Stolpe, J.J. Sederholm, and Hugo Berghell. In the foreground: A.J.Varén and T. Laitinen. Note the hats and hammers. Photo: See Aarne Laitakari, 1936. Suomen geologisen seuran historiikki 1886–1936. Summary: The history of the Geological Society of Finland 1886–1936. Bull. Comm. géol. Finlande 115, 5–64. •
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Haapala, I., 2005. History of Finnish Bedrock Research. 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. 681–702. © 2005 Elsevier B.V. All rights reserved.
Development of science in Finland, including the geosciences, reflects the history of the country. In the 18th century, during the Age of Utility, geology was taught at the Academy of Turku, and the government initiated projects in mineral exploration. After Finland had been separated from Sweden and attached to Russia as a Grand Duchy in 1809, a mining office was established in 1821 and exploration was activated. During the Russian regime, high-quality mineralogical research was conducted by Nils and A.E. Nordenskiöld, Axel Gadolin, and F.J. Wiik. At the University of Helsinki, the Chair of Geology and Mineralogy was established in 1852. The Geological Survey of Finland was founded in 1885 and the Geological Society of Finland in 1886. These arrangements and the advent of polarizing microscopy provided a new base for geological research. In 1880–1940, J.J. Sederholm, Wilhelm Ramsay, and Pentti Eskola promoted the Finnish geological research, petrology in particular, to the forefront of the discipline. In the 1940´s and 1950´s, Th.G. Sahama and Kalervo Rankama became internationally recognized authorities in geochemistry. Sahama also founded a renown mineralogical laboratory at the University of Helsinki. The Geological Survey of Finland (at present, 700 employees) has borne the responsibility for geological mapping of the country, and, after World War II, also for the airborne geophysical and geochemical mapping. Exploration and study of mineral resources have played an important and fruitful role at the Geological Survey, and comprehensive research has also been carried out in isotope geology, petrology, geophysics, and geochemistry. In the 20th century, geological units were founded also in the universities of Åbo Akademi, Turku, and Oulu, as well as in the technical universities of Helsinki and Tampere. During the last two decades, research carried out by the Survey and the universities has led to an thoroughly refined picture of the structure, composition, and evolution of the Precambrian crust of Finland.
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1. Introduction Geology reached the status of an independent science about 200 years ago, when other natural sciences, chemistry, physics, and biology had developed to such a level that they could provide a methodological base for a science studying the structure, composition, and evolution of the Earth. Characteristic of the critical years of geology, at the turn of the 18th and 19th century, was the competition of different doctrines (neptunism, plutonism, catastrophism, uniformitarianism), until the comprehensive works of Charles Lyell in the 1830’s established the general principles of modern geology. Before these times, geological and mineralogical research and teaching had been carried out in the framework of other sciences. Geological and mineralogical research in Finland has been governed by the international development of the science as well as the history of the country, with its economic, cultural, and political trends and opportunities. Economic considerations, exploration for mineral resources and their exploitation, were the important factors until the 19th century, and they still influence geological research. During the Swedish rule (1150–1809), first government-induced yet haphazard mineral exploration was carried out in Finland in the 16th and 17th centuries. In 1542, mining was started at the Ojamo iron ore deposit in Lohja. To develop mineral exploration and mining in Finland, the office of Commissioner of Mines was founded in 1638 and may be regarded as a predecessor of the present mining office at the Ministry of Trade and Industry. The possibilities for higher geological and mineralogical research in Finland were significantly improved when the first university, Academia Aboensis, was founded in Turku in 1640. Partly for political reasons, the university was moved in 1828 with its teachers and students to Helsinki and became the Imperial Alexander University, to be renamed Univer684
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sity of Helsinki after Finland had declared its independence in 1917. After the Great Nordic war (1700–1721), during the Age of Utility, both academic research and teaching in geology as well as mineral exploration by the government were intensified. As a province of Sweden, Finland had been impoverished by lost wars, and sought compensation from rising economic and educational standards. During this time the significant zinc-lead-copper deposit at Orijärvi and the iron ore at Sillböle in Helsinki were found. Under the direction of professor of economics, Pehr Kalm, and chemistry professor, Pehr Adrian Gadd, the students of the Turku Academy prepared in 1754–1795 a number of Master’s theses with geologic-mineralogical topics, such as occurrence of minerals and various rock and soil types, processing of metals, as well as mineral histories of different provinces and parishes. Chemistry professors Johan Gadolin, Pehr Adolf von Bonsdorff, and Adolf Edvard Arppe continued the tradition of research in mineralogy, especially in mineral chemistry, until the Chair of Geology and Mineralogy was established at the University of Helsinki in 1852.
2. Finnish geology in the 19th century After the Finnish War of 1808–1809, Finland was separated from Sweden and attached to Russia as a Grand Duchy. The Governor General was in 1810–1823 count Fabian Steinheil, an enthusiastic amateur mineralogist, who stressed the importance of the development of mineral exploration and mining industry. Especially, iron ore was needed for the iron works of Finland. On Steinheil’s initiative, the Mining Office (the name was changed in 1858 to Bureau of Mines) was founded in 1821, and his protégée, Nils Nordenskiöld, an internationally known mineralogist, acted as its superintendent from 1823 to 1855. Active FINNISH
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mineral exploration led to the discovery of several, unfortunately mostly uneconomic, iron ore fields in southern Finland. Nordenskiöld’s conclusion from the meager successes was that, to support mineral exploration, the entire country should be mapped geologically. On the basis of his proposals (1856, 1860, 1864), geological mapping in Finland was started in 1865, but this activity was interrupted in southern Finland by the gold discovery at Ivalojoki River in Lapland in 1868. The first published geological map was the 1:800,000 bedrock map of Inari and Utsjoki in Lapland, which was attached to A.M. Jernström’s doctoral thesis in 1874. A geological research office with its own directive and a 10-year budgetary endowment was established as part of the Bureau of Mines in 1877, and mapping interrupted by the gold rush to Lapland was resumed. By 1885, nine 1:200,000-scale combined bedrock and soil maps had been published from southern Finland, and Czar Aleksander III approved a motion by the Finnish Senate to establish an independent and permanent Geological Commission to continue geological research under the direction of the newly founded Board of Industry. The Commission commenced its work in 1886, and, as the Geological Society of Finland was also founded in the same year, 1886 became a true landmark of Finnish geology. Mining engineer and Master of Arts, K.A. Moberg, was appointed as the first Director of the Geological Commission (nowadays the Geological Survey of Finland). In the early and middle 19th century, several mineralogists of international standing were active in Finland: superintendent Nils Nordenskiöld (1792–1866), regarded as the Father of Mineralogy in Finland; his son A.E. Nordenskiöld (1832–1901), later an arctic explorer of world fame in Sweden; artillery general Axel Gadolin (1828–1892); and professor F.J. Wiik (1839–1909). But there was a dearth of well-known geologists until the end of the 19th century. This has a CHAPTER
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Nils Gustaf Nordenskiöld (1892–1866) is known as the “Father of Mineralogy in Finland”. He first studied law in Academy of Turku and thereafter mining, mineralogy, and chemistry in Stockholm and Uppsala, obtaining the degree of Mining Engineer in 1817. Nordenskiöld made a three year research journey to Denmark, Germany, France, England, and Scotland, and was in charge of the mineral exploration and geological-mineralogical studies in Finland as the head (superintendent) of the Bureau of Mines. He described and named more than 20 new minerals, of which a couple (phenakite, neotocite) are still accepted mineral species. On Nordenskiöld’s initiative, systematic geological mapping was started in Finland in 1865. Picture from the Archives for Prints and Photographs, the National Museum of Finland.
natural explanation. The bedrock of southern Finland contains a wealth of granitic pegmatites and skarns with large mineral crystals, some of them mineralogical rarities, and their investigation provided a challenge to mineralogists well versed in the crystallographic and mineral chemical research methods of the day. For actual geology, the situation was different. Finland lacks features demonstrating geological processes such as active volcanoes,
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Nils Adolf Erik (A.E.) Nordenskiöld (1832– 1901), son of N.G. Nordenskiöld, was trained in mineralogy and chemistry at home. He completed his Master’s studies at the University of Helsinki in 1853, and defended his mineralogical doctoral thesis in 1855. Plans existed to recruit him as the first ordinary Professor of Geology and Mineralogy of the University of Helsinki, but for political reasons he had to go into exile to Sweden in 1857. In Sweden he became a professor of the Natural History Museum and a famous polar explorer – he found the Northeast Passage in 1878–1879. Of the new minerals described by A.E. Nordenskiöld twelve are still accepted mineral species. Photo from the Archives for Prints and Photographs, the National Museum of Finland.
high mountain chains, deep canyons, and fossiliferous sedimentary strata, the investigation and interpretation of which would have been realistically possible. The flat, low-lying bedrock consisting mainly of granite and gneiss was far too tough a problem for the geologists of that time. But a change was on the way. The use of the polarizing microscope in the study of rocks, innovated by Henry Sorby in 1858, 686
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became rapidly popular especially in Germany during the second half of the 19th century, and Harry Rosenbusch of Heidelberg became the beacon of petrology, a new branch of geology. Microscopic petrography provided a method for investigation of mineral composition and texture of rock types and for their systematic classification, essentially improving genetic interpretations. This opened new possibilities also for Finnish geologists. Professor F.J. Wiik was, in the 1870’s, the first one to utilize the polarizing microscope. Self-taught, he did not get far with his studies but anyway gave basic petrological education to his students, who were then able to utilize it in their publications. Several of Wiik’s talented pupils (J.J. Sederholm, Wilhelm Ramsay, Victor Hackman, Walter Wahl) continued their studies with famous Rosenbusch in Heidelberg, and applied the acquired knowledge to their research and teaching. J.J. Sederholm, Wilhelm Ramsay, and Ramsay’s student Pentti Eskola became, at the end of the 19th century and the beginning of the 20th century, internationally acclaimed geologists promoting Finland to the league of foremost nations in Precambrian geology, petrology in particular.
3. Research organizations 3.1. From the Geological Commission to the Geological Survey The Geological Commission (in Finnish, Geologinen komissioni), founded in 1885, has undergone several changes in name (1925 Geologinen toimikunta or Geological Commission, 1945 Geologinen tutkimuslaitos, 1984 Geologian tutkimuskeskus; the last two translated as the Geological Survey), and has evolved into the present geoscientific organization with 700 employees. Geological Survey plays a leading role in the geological research and exploration for mineral deposits in Finland. The institution has been led by FINNISH
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K.A. Moberg in 1886–1893, J.J. Sederholm 1893–1933, Aarne Laitakari 1935–1960, Vladi Mar mo 1960–1969, Herman Stigzelius 1970–1980, Kalevi Kauranne 1980–1991, Veikko Lappalainen 1992–1997, Raimo Matikainen 1997–2003, and Elias Ekdahl from 2004 on. The permanent staff of the institution remained for long time rather small, 4–6 persons, but after Finland became independent, a steady, at times pulsating, growth and development began. The institution was at its largest in the 1980’s, ~1000 persons, but has been reduced to ~700 employers in the course of the current rationalization of the national administration. The headquarters of the Geological Survey is located in Espoo, with regional offices in Kuopio (central Finland) and Rovaniemi (northern Finland). Main activities of the Survey are geoscientific mapping, exploration for mineral deposits and other geological resources, as well as general research and development in geosciences. The 1:200,000 geological mapping program was abandoned already towards the end of the 19th century and replaced by mapping at a scale of 1:400,000 which has been completed and published for both bedrock and soils, and the 1:100,000 program is well advanced. In 1951, a country-wide airborne geophysical survey from an altitude of 150 m and a 400 m line spacing was commenced by the activities of Aarno Kahma and Mauno Puranen; this was completed in 1972 as the first country in the world. The airborne program has been continued with a line spacing of 200 m from 30–40 m altitude, and will be completed within the next few years. Systematic geochemical mapping of soils was commenced in 1973; this entailed much development and led to a high international level of expertise. Geophysical and geochemical studies have given information on the structure of bedrock, extent of various rock types, and indications on potential ore deposits. Geochemical mapping also provides a foundation for monitoring the chemical state CHAPTER
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of the environment. The laboratory for isotope geology has provided a detailed geochronological framework for petrogenesis of the Finnish Precambrian. Petrological studies have focused on the metamorphic and magmatic evolution of the Proterozoic bedrock in Finland. In economic geology, emphasis has been on metallogenetic and ore mineralogical studies and in industrial mineral studies. Mineral exploration and related research have always formed an important part of the activities of the Geological Survey. The long-term results have been excellent: the Geological Survey and the mining companies have during the 20th century discovered more than 30 metallic ore deposits that have been exploited. Also exploration for and research on industrial minerals have been successful and the significance of this branch has markedly increased in recent years. Four atomic power plants have been built in Finland, and since the 1970’s the Geological Survey has been involved in nuclear waste disposal studies. The research has been carried out in collaboration with several other organizations, and has produced a large amount of versatile information on fracturing and waterrock interaction in the subsurface bedrock.
3.2. Universities The Chair of Geology and Mineralogy was established at the University of Helsinki notably early, in 1852, and F.J. Wiik was appointed as its first holder in 1877. Wiik carried out meritorious mineralogical research on the interdependency of physical and chemical properties of amphiboles and pyroxenes and also tried to clarify the origin and stratigraphy of the Finnish bedrock. Wiik’s successors, Wilhelm Ramsay (1865–1928) and Pentti Eskola (1883–1964), promoted the geological research and teaching at the university to a high international level. Eskola’s students Th.G. Sahama (1910–1983) and Kalervo
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Wilhelm Ramsay (1865–1928) was the Professor of Geology and Mineralogy at the University of Helsinki in 1899–1928. He was a talented all-round geologist, who mastered mineralogy, petrology, physical geology, and Quaternary geology of his time. He found and described the alkaline rocks of the Kola Peninsula, where he and his students made seven research trips between 1887 and 1914. He formulated in 1898 the concept of Fennoscandia, and his Swedish textbook Geologins Grunder was used for a long time in the Nordic countries. Wilhelm Ramsay (standing) on his first expedition to the Kola Peninsula in 1887. Sitting on the left Dr. A.O. Kihlman (Kairamo), on the right prof. J.A. Palmén, the director of the expedition. Photo from prof. Franciska Sundholm’s private collection.
Rankama (1913–1995), both personal extraordinary professors of the university, became internationally leading geochemists, Sahama also a superb mineralogist. During professor Martti Saksela’s (1898–1977) tenure both teaching and research in ore geology and ore mineralogy were emphasized, and professor Heikki Tuominen’s (1914–1985) time was one for development in structural geology. Since 1982 (Ilmari Haapala, 1939–) the “hard rock” research projects have focused on the study of the petrogenesis of granites, rapakivi granites in particular, in Finland and elsewhere (China, 688
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United States, Namibia, Brazil), the geology of the Antarctica, ore geology, and bedrock ground water. The Institute of Seismology forms a separate research institute at the University of Helsinki and carries out high standard lithosphere research and other seismological studies. The main emphasis of studies at the Division of Geophysics has been on the hydrosphere, but with the recently established chair in Solid Earth’s Physics, teaching and research within this sector has started to develop as well. The Swedish speaking university at Turku, Åbo Akademi, was founded at the onset of Finland’s independence in 1918. Geology has been taught from the beginning, the Chair of Geology and Mineralogy has been held by Helge Backlund (1878–1958), Hans Hausen (1884–1979), Gunnar Pehrman, (1895–1980), Nils Edelman (1918–2005), and Carl Ehlers (1944–). In 1952–74, the department had as professor of general and applied geology Adolf Metzger (1896–1965) and Rudyard Frietsch (1927–). The research has focused especially on the structure and genesis of bedrock in the Åland Islands and southwestern Finland, with a recent emphasis on rapakivi granites and volcanogenic rocks. The Chair of Geology and Mineralogy was founded in the Finnish speaking University of Turku in 1958. Professor K.J. Neuvonen’s (1918–) research has been directed towards paleomagnetism of the Precambrian. His successor, professor Heikki Papunen (1936–), has led ore geological and ore mineralogical studies, especially on mineral deposits associated with mafic magmatism in various parts of Finland. The University of Oulu received the Chair of Geology and Mineralogy in 1961. It has been occupied by Juhani Seitsaari (1913–1976) and Kauko Laajoki (1940–). Associate professor Tauno Piirainen and his successor professor Tuomo Alapieti have led work on mafic magmatism and associated ore deposits in northern and eastern Finland, while FINNISH
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Laajoki and his students have concentrated on the sedimentological, structural, and metamorphic evolution of northern and eastern Finland. At the Department of Geophysics of the University of Oulu the emphasis has been on lithosphere geophysics, of which the studies of deep electromagnetics deserve special mention. In 1994, the Departments of Geology and Geophysics of the University of Oulu were amalgamated to form the Department of Geosciences, and ten years later the Division of Geophysics was joined to the Department of Physical Sciences. The Helsinki University of Technology has had a Chair of Mineralogy and Geology since 1937, Economic Geology since 1969, and Engineering Geology since 2003. The Chair has been occupied by Heikki Väyrynen (1888–1956), Aimo Mikkola (1917–), and Heikki Niini (1937–). In addition, the university has an (associate) professorship for applied geophysics. The research has focused on economic geology and lately on potential nuclear waste disposal sites and geophysical modeling. At the Tampere University of Technology there has been an (associate) professorship in engineering geology since 1967. The associate professorships of the Finnish universities were changed to full professorships in 1998. A new generation of professors has been appointed since 2001: Juha Karhu (1951−) and O.Tapani Rämö (1959−) in Helsinki, Krister Sundblad (1952−) and Olov Eklund (1960−) in Turku, Eero Hanski (1954−) in Oulu, and Kirsti LoukolaRuskeeniemi (1957−) in Helsinki University of Technology. The geology departments in the Finnish universities are relatively small, with two or three professors in geology and mineralogy. The geology departments of the Universities of Turku and Oulu have two chairs of in Quaternary geology and Helsinki three chairs in the field of geology and paleontology. In spite of the small size, their impact on Finnish geological research has been significant, and in the recent times domestic CHAPTER
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and international research cooperation has increased markedly.
3. 3. Mining enterprises Mining companies have made geological and mineralogical research in connection with exploration and exploitation of mineral resources, and the results have often been presented in the form of scientific articles and academic theses. Outokumpu Oy was founded on a discovery of copper ore in 1910 from a hill called Outokumpu (a Finnish word meaning “Odd Hill”) in eastern Finland. Under Eero Mäkinen’s (1886–1953) direction it developed into a large mining corporation with several mines and factories. A long-time chief geologist of the company was Paavo Haapala (1906–2002). During the last decades the company has evolved into an international conglomerate of metal industry, which has since 1995 gradually centered on stainless steel production and finished ore expioration in 2003. Since 1937 the company has supported academic research through Outokumpu Oyj Foundation. Otanmäki Oy (in 1968 attached to Rautaruukki Oy) was founded in 1960 on the iron-titanium-vanadium deposit of Otanmäki in north-central Finland. The company has exploited several iron and vanadium deposits and focused on steel industry. Active exploration under Heikki Paarma’s (1920–2001) leadership led to the discovery of several iron deposits, and the large Paleozoic apatite and pyrochlore-bearing Sokli carbonatite stock in northeastern Finland. The petrology and mineralogy of the carbonatite was studied in detail by Heikki Vartiainen. As the iron ore deposits become exhausted and mining uneconomic in the 1980’s, the company ceased active prospecting. Malmikaivos Oy, a subsidiary of the paper company Myllykoski Oyj, has carried out exploration and mining mainly in eastern Finland. In the 1990’s, the company centered
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on prospecting for and study of diamond-bearing kimberlites – a new branch of economic geology that has also stimulated lithospheric research in Finland. The structure of mining industry has since 1990’s changed markedly (see p. 698), and several operating mines and smelters are now owned by international companies. The industrial mineral companies Partek Oy Ab and Oy Lohja Ab exploited long calcitic and dolomitic marbles of Finland, and their work is now continued by Nordcalc Oyj. The fertilizer company Kemira Oy has exploited the large Paleoproterozoic apatite-bearing carbonatite of Siilinjärvi. Mondo Minerals Oy is currently exploiting talc deposits in eastern Finland. Several active enterprises quarry dimension stones, mainly granites, utilizing also geological expertise in this work.
3.4. Other research organizations Eero Mäkinen (1886–1953) is best known as the developer of the Finnish mining industry, but he also achieved significant results during his 6-year active spell as a geologist. His doctoral dissertation (1912) of the granitic pegmatites of the Tammela area in southern Finland is a classic of pegmatite research. His studies 1917 on alkali feldspars and their monoclinic-triclinic transformation was for decades a basic work in this field. Mäkinen accomplished the degree of Mining Engineer in Stockholm in 1918, and gave up personal scientific research thereafter. As the Managing Director of Outokumpu Oy during 1921–1953 he promoted a small, barely managing copper mine into a large mining corporation with several mines and factories. Building on the foundation laid by Mäkinen, the company has evolved into the present international conglomerate of metal industry. The input of Outokumpu Oy to geological research in Finland has been significant not only through the company’s own research but also through the Outokumpu Oyj Foundation, founded in 1937. In this photo from 1912, Eero Mäkinen is a young geologist. Photo from the archives of Outokumpu Oyj.
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The need for geological information of bedrock and its physical properties has increased along with the growth of volume and versatility of rock engineering during the second half of the 20th century. Different state, communal, and private organizations have been involved in geological studies related to planning and construction of railways, highways, underground shelters, subways and multipurpose tunnels, and water-conveyance tunnels. Since the 1970’s, the site-selection and planning of the disposal of nuclear waste produced by Finland’s four atomic power plants have been a challenging task for the Geological Survey, the Technical Research Centre of Finland, several consulting companies, and power plant enterprises. Results have been published in numerous reports, articles, and academic theses.
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Jakob Johannes (J.J.) Sederholm (1863–1934) was the Director of the Geological Survey (Geological Commission) of Finland in 1892–1933. He rose to fame by applying the uniformitarian doctrine to the Precambrian of Finland in the 1890’s. He compiled the first geological map of Finland in 1897, and published syntheses of the geology of Finland and Fennoscandia. His studies on the origin of migmatites and granites belong to classics of geology, and many widely used petrologic terms are coined by him. In the 1920’s his diplomatic and organizational skills were utilized by the League of Nations in solving international problems in Europe. Sederholm was Honorary Doctor of the universities of Oslo, Toronto, Kingston, and Uppsala. In this photograph from 1933, Sederholm (with the peace pipe and a bow) is decorated the Honorary Chief of the Winnebago Indians of Canada. Photo from Dr. Barbro Scheinin’s private collection.
4. Main fields of research 4.1. Petrology and physical geology Geologically, Finland consists of Precambrian crust covered by thin Quaternary soil. The Precambrian crust can be divided into two main domains: (1) the Archean ~3.1–2.6 Ga granite gneisses and greenstone belts in eastern Finland, and (2) the Paleoproterozoic ~1.93–1.80 Ga Svecofennian orogenic crust CHAPTER
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southwest of the Archean terrain. Intracontinental magmatic episodes are represented by the 2.45 Ga layered mafic intrusions of northern Finland and the voluminous 1.65–1.54 Ga rapakivi granites and associated rocks in southern Finland. The first overall account of the Precambrian of Finland was presented by J.J. Sederholm in 1897 in a 1:2,500,000-scale bedrock map. He later published several further overviews, the last one in 1932. In the absence of suit-
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able fossiliferous sedimentary sequences and direct age determinations, Sederholm (1932) divided the Precambrian supracrustal rocks, nowadays known to be Paleoproterozoic, into three cycles of sedimentation delineated by three epochs of diastrophism and granite intrusion: (1) The oldest complex of Katarchean Svionian sedimentary and volcanic rocks of southern Finland, the upper limit given by intrusion of post-Svionian gneissose (1st group) granites; (2) The younger Archean Bothnian and Lapponian sedimentary and volcanic rocks penetrated by post-Bothnian (2nd group) granites; and (3) The Proterozoic Karelian cycle comprising Jatulian and Kalevian sedimentary rocks that are penetrated by post-Kalevian (3rd group) granites. These were followed by the undeformed rapakivi granites, Jotnian sandstones, and associated diabase dike swarms. Sederholm emphasized the tectonicstratigraphic significance of conglomerates containing granitic pebbles – such conglomerates indicated deep erosion and major unconformity separating different cycles of sedimentation. He also used the intensity of metamorphism as an age criteria – sometimes with wrong conclusions. Since Sederholm’s time, more detailed geological mapping, application of modern methods of structural geology and sedimentary petrology, and isotope age determinations have markedly changed the picture. Sederholm assumed that every sedimentary or orogenic cycle was associated with the intrusion of one granite group; nowadays a minimum of two or three granite groups are distinguished in an orogenic cycle. The independent second cycle of his schema was soon eliminated; its components were included into the two other cycles to form two orogenic belts, the older 692
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Svecofennian belt in southern Finland and the younger Karelian belt in eastern and northern Finland. This model prevailed until the end of the 1950’s. The U-Pb zircon age determinations of Olavi Kouvo in 1958 brought about fundamental changes in the concepts of the Precambrian geology in Finland. It turned out that there was no age difference between the plutonic rocks of Svecofennian and Karelian belts, they both were 1.9 to 1.8 Ga. This led to the conclusion that the plutonism of both belts belongs to the same orogenic cycle; this was coined Svecokarelian orogeny by Simonen in 1971. Subsequent sedimentary petrological studies by Gabor Gaál and others showed that the Karelian (Jatulian) quartzites and conglomerates in eastern Finland do not represent orogenic geotectonic environments but rather had been deposited on the Archean craton. Consequently, the term “Svecokarelian orogeny” was abandoned and replaced by the Svecofennian orogeny. The study of granites and migmatites, the most characteristics rocks of southern Finland, was one of the cornerstones in Sederholm’s wide and versatile production. Many of his studies are classics of Precambrian research. He introduced new geological concepts and a number of widely used petrological terms (e.g., migmatite, anatexis, palingenesis, syntexis, arterite, ptygmatic folding, nebulitic, migmatite, myrmekite). His studies were based on accurate field observations and microscopic studies and they were written logically. Sederholm’s selected works of the years 1907–1934, Granites and Migmatites, were published as ~600-page new edition in the United Kingdom in 1967. Since Sederholm’s time one of the profound questions in the Finnish geology has been the origin of granites and their tectonic classification. Regarding origin, it was argued whether the granites are magmatic or formed from other rocks by metasomatic granitization. Sederholm himself was, as a student of Harry FINNISH
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Pentti Eskola (1883–1964) was the Professor of Geology and Mineralogy of the University of Helsinki from 1924 to 1953. He gained fame already with his doctoral thesis “On the petrology of the Orijärvi region in southwestern Finland” in 1914. In the next year, he published his famous metamorphic facies doctrine, a basic theme of modern metamorphic petrology, and supplemented it in later publications. Wellknown are also Eskola’s studies on the Norwegian eclogites and his eclogite theory, his studies on the origin and tectonic classification of granites, and his many papers dealing with origin and evolution of the bedrock of Finland. Although many of Eskola’s papers were based on local field studies, their discussions generally displayed an approach of a physical chemist rising above mere locale. For students and the public he wrote several popular geological textbooks. Eskola received many international scientific honors, the last of them was the Vetlesen Prize, which was awarded in 1963 in New York. He was an Honorary Doctor of Oslo, Padua, Bonn, and Prague universities. In this picture, Eskola is working at the Geophysical Laboratory in Washington, D.C., in 1921.
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Rosenbusch from Heidelberg, a magmatist, but when studying the migmatites of southern Finland he adopted also metasomatic views. Several other Finnish granite petrologists, Pentti Eskola, Maunu Härme, and Ahti Simonen accepted both magmatic and metasomatic models, but they stressed the importance of these concepts differently. The strongest magmatic standpoint was taken by Martti Saksela. Perhaps even more heated than the disputation on the origin of granites was the discussion of their tectonic classification and magmatic differentiation series. Sederholm divided granites into four groups that represented three orogenic (Svecofennian) and one younger anorogenic (rapakivi granites) cycles. In 1931, Eskola divided the orogenic granites into three groups: syn-, late-, and postkinematic. This widely used classification was modified by Walter Wahl, Martti Saksela, and Ahti Simonen. During the last quarter of the 20th century the arguing on the origin of the granites, differentiation series, and tectonic classification has changed its form. The origin of granites and related rocks has been linked to plate tectonic models, and by utilizing geochemical and isotope studies it has been possible to make far-reaching conclusions about the origin and evolution of the Finnish granites. The origin of the rapakivi granite–diabase association has been explained by the “mafic underplate” model: mantle-derived hot mafic magmas have caused partial melting of the lower crust producing the granite magmas. Pentti Eskola’s monumental 1914 dissertation On the petrology of the Orijärvi region in southwestern Finland led him to pioneering studies in metamorphic petrology. He recognized regular changes in mineral assemblages of metamorphic rocks and interpreted these to be related to changes in temperature and pressure. In 1915, he formulated the principles of the metamorphic facies doctrine, and in 1920 expanded it to a more general mineral facies doctrine.
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Geological mapping of the wide wilds of northern and eastern Finland was hard work requiring also skills of a woodcraftsman and ability to collaborate with the local people. In this picture, Dr. Erkki Mikkola (1904–1940), the famous mapper of Lapland, is sitting with old Mr. Hirvonen in front of the Ränttilä farm house in Viitasaari. The photo was taken in the summer of 1927 by W.W. Wilkman, Mikkola’s older colleague, who did a considerable job in mapping the bedrock of eastern Finland. Essentially in 1930–35, Mikkola mapped an area of ~150 km by 280 km in Lapland, and his maps and their explanatory texts are known of their high quality. Erkki Mikkola was killed in action in the Winter War at the age of 35. Photo from the Archives for old photographs, Geological Survey of Finland.
The effect of metamorphic conditions to mineral assemblages of rocks had been discussed earlier by George Barrow in 1893, Ulrich Grubenmann in 1904, and V.M. Goldschmidt in 1911, but Eskola’s facies doctrine gave a clear rule and model, based on the Gibb’s phase rule, for the mutual relations between the stable mineral associations, chemical composition of rocks, and the temperaturepressure conditions during metamorphism. The metamorphic facies doctrine is one of the basic principles of metamorphic petrology, and has soon prevailed for 90 years as a practical classification scheme depicting the P-T conditions of metamorphic rocks. Eskola’s authority in metamorphic petrology was so profound that only a few of his 694
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students had the courage to carry out further studies in that field. Two of these were Maunu Härme, who studied cordierite-anthophyllite gneisses, migmatites, and granitization, and Anna Hietanen, who later emigrated to the United States and studied the metamorphic rocks of California and Idaho. Heikki Tuominen and Toivo Mikkola presented reinterpretations of the Orijärvi area emphasizing the role of tectonic mobilization, but Eskola’s model has stood the criticism. Since the 1970’s, Kalevi Korsman and his team at the Geological Survey have carried out systematic studies on metamorphic geology of Finland correlating tectonic evolution and metamorphism. The Alpine-style folding and overthrust tectonics was first applied to the Finnish Precambrian by Swiss Eugène Wegmann at the end of the 1920’s, and his studies inspired several Finnish geologists to undertake tectonic studies. After World War II, Heikki Tuominen and his coworkers emphasized the role of faulting, especially as a control of ore deposition. These untraditional interpretations have generally been regarded as exaggerations, as was clearly the case with Tuominen’s 1957 map of the Orijärvi area, but so was the role of faulting generally underestimated in earlier studies. The principles of polyphasic deformation were applied to the Precambrian of Finland in the 1970’s and the 1980’s by Donald R. Bowes (the United Kingdom), Gabor Gaál, and Tapio Koistinen. It is nowadays agreed that our Paleoproterozoic and Archean rocks were deformed and metamorphosed in several stages, and the main stages are commonly observed during the routine mapping. The first serious plate tectonic model for the Svecofennian domain was presented by Anna Hietanen in 1975. She compared the evolution of igneous activity of Sierra Nevada and the Svecofennian of Finland, and suggested that the Svecofennian belt represents a Cordilleran-type orogeny, with subduction from the southwest (Sweden) to the northeast under the Archean continent. The model has FINNISH
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been developed further by Gabor Gaál, Raimo Lahtinen, and Mikko Nironen, who have assumed at least two successive Paleoproterozoic subductions and accretions of volcanic arcs to the Archean continent. An important additional evidence for the operation of plate tectonics during the Paleoproterozoic time was obtained in 1979 when Asko Kontinen described a well-preserved 1950-Ma ophiolite association from the Archean–Proterozoic boundary zone in eastern Finland. Interesting information of the drift history and the tectonic units of the Fennoscandian Shield has been provided by the paleomagnetic studies started in the 1960’s by K.J. Neuvonen at the University of Turku and continued by Lauri J. Pesonen and his team at the Geological Survey and the University of Helsinki. By measuring the natural remanent magnetization of precisely dated rocks it has been possible to construct the latitudinal drift and rotation of the shield and its components from the Archean to the present. Recently, the method has also been used in reconstruction of ancient supercontinents with Fennoscandia as a part of them. After J.J. Sederholm, syntheses of the Finnish Precambrian have been presented by Heikki Väyrynen in 1954, Ahti Simonen (e.g., 1960), and Pentti Eskola (1963), as well by Martti Lehtinen, Pekka Nurmi, and Tapani Rämö (1998, the Finnish predecessor of this book).
4.2. Geochemistry and isotope geology The evolution of instrumental analytical chemistry laid the foundation for the rise of geochemistry as an important new branch of geological sciences in the 1930’s, with V.M. Goldschmidt as the leading figure. Young Thure Georg (Th.G.) Sahama, who had in 1936 defended his doctoral thesis on the microtectonics of the granulites in Lapland, joined wholeheartedly this rapidly expanding field of science, and so did soon his friend Kalervo Rankama. In 1937–38 Sahama built CHAPTER
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in the Department of Geology, University of Helsinki, a geochemical laboratory with an X-ray spectrograph and optical emission spectrometer as the main instruments. In the late 1930’s and 1940’s, Sahama and his collaborators made a great amount of top quality analyses from rocks and minerals. Sahama studied the general geochemistry and rare earth element distributions in various rocks, whereas Rankama focused on the geochemistry of Nb and Ta. In 1945, Sahama published in Finnish a textbook Geokemia, and five years later, together with Rankama, an updated version Geochemistry in English. This book combined the principles of geochemistry by V.M. Goldschmidt and the results of new research. The book became a great success, it was translated in several languages, and new English editions were taken (6th edition in 1968). It is obvious that this book contributed markedly to the rapidly growing international interest in geochemistry during the 1950’s to 1970’s. The progress in geochemistry has continued, the analytical methods have improved, and new applications have been developed. Geochemistry and isotope geology (or isotope geochemistry) are in key position in studies related to petrological, ore geological, and environmental questions. Especially useful such studies have been in deciphering the geotectonic setting, origin, and evolution of igneous rocks. For example, the plate tectonic models of the Svecofennian orogenic belt are largely based on geochemical and isotope studies. First attempts to apply geochemistry of till to prospecting in the glaciated terrain of Finland in the 1940’s were promising, and in 1952 a permanent geochemical prospecting group was established in the Geological Survey. Soon also other prospecting organizations adopted till geochemistry as a routine exploration technique; this method has generally been used together with geophysical and geological methods, and in several cases with good success. In 1970 a geochemical department,
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Thure Georg (Th.G.) Sahama (1910–1983) was a Personal Extraordinary Professor of Geochemistry at the University of Helsinki in 1945–1977 and Academician (1972–1983). He got interested in geology and mineralogy as a school boy, and his first mineralogical papers were published one year after matriculation at the University of Helsinki. His doctoral thesis (1936) dealt with microfabrics of the granulites of Lapland, but at the end of the 1930’s he went on to geochemistry. Together with Kalervo Rankama he wrote the textbook Geochemistry, which became a success and activated geochemical research world-wide. In the 1950’s, Sahama’s research focused first on volcanology and then on mineralogy. Together with his assistants and coworkers, he described 15 new mineral species. In this picture from 1956, Sahama is sitting on an instrument box in the upper platform (3267 m above sea level) of Mt. Nyiragongo caldera in the Virunga volcanic area. Photo: M.-E. Denaeyer (from the archives of Geological Museum, University of Helsinki).
led by Kalevi Kauranne, was established at the Geological Survey, and it has carried out systematic regional mapping, using till, hard rock, stream sediment, humus, and water samples. The main results were published in 1990–96 in the impressive Geochemical Atlas, parts 1–3. The contribution of Finnish geologists to isotope geological research has been significant. Actually, the term isotope geology was introduced by Kalervo Rankama in 1954 in his classic text book Isotope Geology. By rapid evolution in this field of science, he wrote in 1960 its sequel, Progress in Isotope Geology. Chemist and geologist Walter Wahl carried out in the early 1940’s at the Department on Chemistry, University of Helsinki, isotopic analyses of U, Th, and Pb of some minerals from Finland and Sweden. In 1948, Kalervo Rankama used carbon isotopes analyses, made at the University of Minnesota, to verify the origin of carbon in the “carbon sacks” (Corycium enigmaticum; according to the current view, algae-covered mud balls in turbidites) described by J.J. Sederholm from the phyllites of the Tampere area. Olavi Kouvo (1920–) has made an important career in dating the crustal units of the Finnish Precambrian. He first determined the isotope ages (U-Pb, Pb-Pb, Rb-Sr, K-Ar methods) in research centers of the United States, and since the 1960’s at the isotope laboratory of the Geological Survey. Nowadays, the laboratory utilizes the U-Pb, Sm-Nd, Pb-Pb, and Rb-Sr methods for datings and petrogenetic studies, and stable isotope analyses (C, O, and H) to decipher past environmental changes. The work done in this laboratory has greatly contributed to our understanding of the evolution of the Earth’s lithospere and Precambrian atmosphere.
4.3. Mineralogy During the first half of the 20th century, 696
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Kalervo Rankama (1913–1995) was a Personal Extraordinary Professor of Mineral Chemistry at the University of Helsinki in 1950–1980. Rankama’s first important studies dealt with the geochemistry of Ta (Ph. D. thesis in 1944) and Nb. His book Isotope Geology (1954) was a pioneering work in the field, and later he edited book series The Precambrian I–IV and The Quaternary I–II. Photo from the archives of Geological Museum, University of Helsinki.
mineralogical research comprised traditional mineralogical–petrographic description of rocks and mineral deposits. Well-known Finnish mineralogists from this period are Leon. H. Borgström, Walter Wahl, Pentti Eskola, Eero Mäkinen, Aarne Laitakari, and Gunnar Pehrman. A new era in Finnish mineralogical research started when Th.G. Sahama switched from geochemistry to mineralogy. The geochemical laboratory of the University of Helsinki changed in 1952 to a mineralogical laboratory where X-ray diffraction (powder and single crystal methods) and soon also infrared spectroscopy and scanning electron microscopy were used to study, from the early morning to the night, minerals collected on the expeditions to Africa. First expeditions were directed to the Virunga volcanic field in CHAPTER
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the East African rift zone, in the border area between Uganda, Congo, and Ruanda, where the alkaline lavas of the Nyiragongo volcano were of especial interest. Later, Sahama’s main interest was in the granitic pegmatites of Africa. Together with his research assistants and collaborators, Sahama made systematic basic studies of several minerals and mineral groups (humite group, nepheline and kalsilite, beryl, kornerupine, sapphirine) and found 15 new mineral species. Sahama’s closest collaborators were his research assistents and Oleg von Knorring (1915–94), a Finnish mineralogist who had moved to Leeds, United Kingdom. In the 1950’s to 1970’s, several young geologists followed Sahama’s example and turned to mineralogical research, and this time may be regarded as the golden age of Finnish mineralogy. During the years 1954–83, Finnish mineralogists described 40 new mineral species, which is a remarkable achievement for a small professional group of a small country. The first crystal structure analysis was made by Atso Vorma in 1963; he showed that the mineral stokesite has a new type of chain silicate structure. The ore microscope became an important mineralogical research instrument with the growing exploration and ore exploitation activities after World War II. Instructions in ore mineralogy and ore microscopy became an important subject in university training, and ore microscope was adopted as a standard tool by exploration and mining geologists. It has been useful especially in identification of ore minerals and in mineral processing, where the mode of occurrence and intergrowths of ore minerals are of great practical importance. Later, the ore microscope has been supplemented and gradually partly replaced by electron microprobe. The first electron microprobe, operated by Jaakko Siivola, was implemented at the Geological Survey in 1964. The 670 mineral species found in Finland so far are described with their characteristics
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and properties, many also with color pictures, in a 400-page book Suomen Mineraalit (Minerals of Finland) compiled by Kai Hytönen (1999). Several Finnish mineralogists and chemists have also studied mineralogical and chemical composition of meteorites. Research in this field was started in 1821 by Nils Nordenskiöld. During the 20th century Leon. H. Borgström, Walter Wahl, and Birger Wiik continued studies on meteorites, and Wiik was one of the scientists invited to analyze the first lunar samples brought by Apollo 11. In 1969, Martti Lehtinen showed by mineralogical and petrological studies that the “dacite” of Lappajärvi is actually an impact lava rock and the circular Lake Lappajärvi is a 75 Ma-old meteorite impact crater, not a volcanic crater or caldera as previously assumed. After that, ten other impact craters and sites have been identified in Finland.
4.4. Economic geology At the turn of the 19th and 20th century, there were three periodically operating mines in Finland: the old Orijärvi Cu-Zn mine (discovered in 1757) in southwestern Finland and the Pitkäranta Cu-Zn-Sn mining field and the Välimäki iron deposit in Ladogan (Russian) Karelia. The gold rush to Lapland, promoted by the discovery of placer gold in the gravels of Ivalojoki River, had already slowed down, and the ore potential of Finland was generally regarded small. The discovery of the Outokumpu copper deposit in 1910 by Otto Trüstedt was an example of the successful application of geological knowledge to exploration; the prospecting was started in 1908 on the basis of a glacial boulder and the deposit was found two years later about 50 km northwest of the discovery site of the boulder. The exploitation of this large high-grade Cu-Zn deposit was for several years operating mainly with loss, until Eero Mäkinen (1886–1953) led the 698
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mining company in the 1920’s into profitable production. On the peace of Tartu, in 1920, Finland received Petsamo (Pechenga) region from Russia, and in the 1920’s–1930’s nickel deposits were found in this area. The exploitation was conducted as Finnish-Canadian cooperation, until the area was reclaimed back to Soviet Union in 1944. At the end of the 1930’s, the large Fe-TiV deposit of Otanmäki, the Cu-W deposit of Ylöjärvi, and the Ni deposit of Makola were discovered. After World War II, discoveries continued with the important Vihanti Cu-ZnPb, Pyhäsalmi Cu-Zn-pyrite, and Kemi Cr deposits, the Ni-Cu deposit of Kotalahti, Hitura and Vammala, Virtasalmi Cu, Luikonlahti Cu-Zn-Co-pyrite, and Hammaslahti Cu-Zn, as well as a number of smaller ones. Altogether, 33 exploitable ore deposits were found in the 20th century and a strong metal industry was built on the discoveries. When several base metal mines became exhausted in the 1970’s and later, Outokumpu Oy brought raw material also from abroad. Rautaruukki Oy finished exploitation of Finnish mines in 1990, and is nowadays manufacturing imported iron ore. Since Rautaruukki and Outokumpu withdrew from ore exploration, international companies and some Finnish junior companies have increased their exploration activities, which has led to discoveries of new precious metal, diamond and industrial mineral occurrences. Since Rautaruukki and Outokumpu withdrew from ore exploration, international companies and some Finnish junior companies have increased their exploration activities, which have led to discoveries of new precious metal, diamond and industrial mineral occurrences. The production of industrial minerals increased strongly during the 20th century, especially since 1979. The main deposits include calcitic and dolomitic marbles in southern Finland (Parainen, Tytyri, Ihalainen, Mustio), the Siilinjärvi apatite-flogopite carbonatite FINNISH
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(since 1979), several talc (soapstone) deposits, and deposits containing wollastonite, quartz, and feldspar, as well as silicate raw materials for concrete and rock wool. Since the 1980’s, the exploration carried out by the Geological Survey has centered on kaolin, marble, and ilmenite, and important new research results have been published creditably. The discovery of diamond-bearing kimberlites by Malmikaivos Oy in the 1990’s in eastern Finland opened new views for economic geology, and further exploration and research is continued by other organizations. When Finland joined the European Economic Area (EEA) and European Union (EU) in 1994-95, the mining legislation was formulated more open for foreign residents, companies and foundations according to EEA laws and regulations. This has led to arrival of many international companies in Finland and they now own several mines. Increased exploration activities of the international companies and some Finnish junior companies as well as the work of the Geological Survey of Finland has led to discoveries of new precious metal, industrial mineral and diamond occurrences Important geological and mineralogical data is commonly obtained during exploration and exploitation of mineral deposits. For example, five doctoral dissertations and a number of smaller articles have been published on the Outokumpu ore field. The ore geological research projects of the universities have also produced important basic data, and one gold deposit at Orivesi was found in a universitycompany joint project. The Geological Survey of Finland has for several decades systematically collected ore geological data and published metallogenic maps and their explanatory texts, including the synthesis of Kahma (1973) and the detailed up-to-date work of Saltikoff et al. (2005).
5. Synopsis Geological research and teaching took their CHAPTER
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first serious steps in Finland around the middle of the 18th century, during the Age of Utility. The government activated prospecting for ore deposits, in part successfully, and geological subjects were taught at the Academy of Turku. During the 19th century, when Finland was a Grand Duchy of Russia, the country had several internationally well-known mineralogists (Nils and A.E. Nordenskiöld, Axel Gadolin, F.J. Wiik) who studied mineral occurrences, mainly pegmatites and skarns containing euhedral crystals, suitable for crystallographic measurements and chemical analyses. The deeply eroded flat bedrock, consisting mainly of gneisses and granites, did not provide realistic possibility for ordinary geological research at that time. During 1880–1940, strong progress took place in the geological research. This evolution was in part related to development in the educational and scientific organization in the country. Teaching of geological subjects was established when the Chair of Geology and Mineralogy (founded in 1852) was finally filled in 1877 in the University of Helsinki. The Geological Commission (Geological Survey) was established in 1885 and the Geological Society in 1886. Another important prerequisite for the rise of geology in Finland was the international methodological evolution. The new discovery, to study thin sections of rocks by polarizing microscope, provided an excellent possibility to classify and infer the origin of the crystalline schists and plutonic rocks of the Finnish Precambrian. Combination of skilled field work and microscopic petrography opened new avenues for geological research and created possibilities to apply the uniformitarian doctrine to the “primitive” (Precambrian) crust. Evidently also the overall atmosphere in a country that had became conscious of its national identity was favorable for science and arts. J.J. Sederholm, Wilhelm Ramsay, and Pentti Eskola raised the geological, especially petrological, research and education in Finland to a high
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international level. As stated by Akiho Miyashiro, Bulletin de la Commission géologique de Finlande was in the first half of the twentieth century among the most important journals in metamorphic geology of the world. The decades after World War II were a time of steady evolution and laying of new foundation for both sophisticated geological research and mining industry. Geological Survey of Finland made pioneering work in carrying out airborne geophysical mapping and geochemical mapping over the whole country, simultaneously with geological mapping and active exploration. This was one reason for the numerous discoveries of new ore deposits in the 1950’s–1970’s. The role of Aarno Kahma, the head of the exploration department, must be acknowledged for planning and leading these activities. The higher scientific research was for a long time in the hands of relatively few internationally wellknown scientists. Th.G. Sahama and Kalervo Rankama rose in the 1940’s and 1950’s to fame in the field of geochemistry, Rankama also in isotope geology. Sahama moved in the 1950’s from geochemistry to mineralogy, becoming an international figure also in that field. When also many of Sahama’s students and assistants specialized in mineralogy, was mineralogical research in Finland on remarkably high level during the 1950’s–1970’s. Another example of the effect of one scientist devoted to research is found in isotope geology. The isotope laboratory, built at the Geological Survey under Olavi Kouvo’s leadership, has grown to a top-class research unit where 3–5 researchers study, by utilizing a variety of isotopic methods, the origin and evolution of the Earth’s crust and environmental issues. In recent years, the research made by the Geological Survey, universities, mining companies, and other research organizations have lifted the knowledge of the structure, composition, and evolution of the lithosphere 700
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and changing environment to a new level. Within the universities, important research has been made, not only in the geology departments, but also in the geophysical units: Institute of Seismology in the University of Helsinki and geophysical groups of Helsinki and Oulu Universities and in the Helsinki University of Technology. The collaboration of geologists and geophysicists has proven to be very fruitful. The research activity of young geologists has increased markedly, especially in the 1990’s. For example, in the 1990’s, 78 geologists defended their doctoral (Ph.D.) dissertations in Finnish universities, while the corresponding number in the 1960’s was 29, in the 1970’s 39, and in the 1980’s 37. Many of the dissertations have received international recognition. Although the crust of Finland is among the best known Precambrian shield areas, much is still to be done. New important findings are made continuously. These include the discoveries of diamond-bearing kimberlites, gold and platinum deposits, and meteoric impact craters. Improved models have been developed for crustal and environmental evolution, and these may again be utilized in practical issues, such as planning of exploration and environment protection. There is no end to the work of geologists in Finland.
Acknowledgments The manuscript was revised by Professors Heikki Niini, Tapani Rämö, Martti Lehtinen, and Pekka Nurmi. Their valuable suggestions and fruitful discussions are highly appreciated.
Selected literature This article is based mainly on the author’s earlier short histories of geological research and FINNISH
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sciences in Finland. For those who want to read more about the geological sciences in Finland, a short list of references is given here. Anonym, 2005. Finland. Mining Journal special publication, London, February 2005, 24 p. Eskola, P., 1963. The Precambrian of Finland. In: K. Rankama, (Ed.), The Geologic Systems, The Precambrian 1. Interscience Publishers, New York. 145–263. Haapala, I., 1986. Geologian yliopisto-opetuksen historia Suomessa.(History of Geology teaching in Finnish universities). Opusculum 6 (1), 3–63 Haapala, I., 2000. Geologia [Geology] In: P. Tommila, A. Tiitta, (Eds.), Suomen tieteen historia 3 (The history of Science in Finland 3), 268–305. Hanski, E., 2001. History of stratigraphical research in northern Finland. In: M. Vaasjoki (Ed.), Radiometric age determinations from Finnish Lapland and their bearing on the timing of Precambrian volcano-sedimentary sequences. Geol. Surv. Finland, Spec. Pap. 33, 15–43. Hausen, H., 1968. The history of geology and mineralogy in Finland 1828-1918. The history of Learning and Science in Finland 1828–1918. 7a. Helsinki 1968. Hytönen, K., 1999. Suomen Mineraalit (Minerals of Finland). Geologian tutkimuskeskus. Erillisjulkaisu (Geol. Surv. Finland, Spec. Publ.), 1–399. Niini, H., 1997. Maankamaran jalostus – ihmiskunnan perustarve, 60 vuotta geologian ja geofysiikan opetusta insinööreille (Refining of the bedrock – a fundamental necessity to mankind). Teknillinen korkeakoulu, insinöörigeologian ja geofysiikan laboratorio, Tiedonanto TKK–IGE–45/1997. Lehtinen, M., 1969. Meteoriitti-impaktiteoria ja Lappajärvi-muodostuma. Phil.Lic. Thesis, Univ. Helsinki. 1–140. (in Finnish) Niini, H., Uusinoka, R., 2000. Kalliorakennusgeologian historiaa Suomessa. Abstract:
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History of bedrock engineering geology in Finland. Vuoriteollisuus-Bergshanteringen 58 (1), 35–43. Saltikoff, B., Puustinen, K., Tontti, M., 2005. Explanation to the metallogenic map of Finland. Geological Survey of Finland, Special Paper 35 (in print). Sederholm, J.J., 1967. J.J. Sederholm, Selected Works, Granites and Migmatites. Oliver Boyd, Edinburgh and London. 608 p. Tanskanen, H. (Ed.), 1986. The development of geological sciences in Finland. Geological Survey of Finland, Bulletin 336, 1–344. The volume contains following articles: Kauranne, L.K.: Foreword, 5–7. Stigzelius, H.: Mineral exploration and geological surveys in Finland before 1886, 9–19. Simonen, A.: Stratigraphic studies on the Precambrian in Finland, 21–37. Härme, M.: The history of the petrologic study in Finland, 41–78. Hytönen, K.: The history of mineralogy in Finland 1918–1984, 79–100. Virkkala, K.: History of studies on Quaternary geology in Finland, 101– 163. Papunen, H.: One hundred years of ore exploration in Finland, 165–203. Ketola, M.: The development of exploration geophysics in Finland, 205–231. Kauranne, L.K.: Geochemical research in Finland, 233–258. Lappalainen, V.: The history of engineering geology in Finland, 259–271. Boström, R.: The history of the stone and mineral industry in Finland, 273–298. Haapala, I.: The history of geology teaching at Finnish universities, 299– 344. Virkkala, K., 1986. Geologian tutkimuskeskuksen 100-vuotishistoriikki. Summary: History of the Geologiacal Survey of Finland 18861986. Geologian tutkimuskeskus (The Geological Survey of Finland), Espoo, 1–93.
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CONTRIBUTORS Editors Dr. Martti Lehtinen is Professor and Head of the Geological Museum of the Finnish Museum of Natural History, University of Helsinki. He also lectures as a docent at the University of Helsinki. Email:
[email protected] Address: Department of Geology, P.O. Box 64, FI-00014 University of Helsinki, Finland Dr. O. Tapani Rämö is Professor of Petrology and Physical Geology at the Department of Geology, University of Helsinki.
Email:
[email protected] Address: Department of Geology, P.O. Box 64, FI-00014 University of Helsinki, Finland
Dr. Pekka A. Nurmi is Professor and Research Director (Bedrock Geology and Mineral Resources) at the Geological Survey of Finland. He also lectures as a docent at the University of Helsinki. Email:
[email protected] Address: Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland Mr. Sakari Haapaniemi, MA is visual planner at the the Inclus Communications in Helsinki.
Email:
[email protected] Address: Inclus Communications, FI-00520 Helsinki, Finland
Authors Dr. Ilmari Haapala is Professor Emeritus at the Department of Geology, University of Helsinki.
Dr. Eero Hanski is Professor of Geochemistry at the Department of Geosciences, University of Oulu.
Email:
[email protected] Address: Huvilakuja 2, FI-02730 Espoo, Finland
Email:
[email protected] Address: Department of Geosciences, P.O. Box 3000, FI-90014 University of Oulu, Finland
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Dr. Hannu Huhma is leading the Laboratory of Isotope Geology at the Geological Survey of Finland. He also lectures as a docent at the University of Turku.
Email:
[email protected] Address: Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland Dr. Yrjö Kähkönen is University Lecturer and docent at the Department of Geology, University of Helsinki.
Email:
[email protected] Address: Department of Geology, P.O. Box 64, FI-00014 University of Helsinki, Finland Dr. Jarmo Kohonen is Director of Information Management at the Geological Survey of Finland. He also lectures as a docent at the University of Helsinki.
Email:
[email protected] Address: Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland Dr. Annakaisa Korja is Seismologist at the Institute of Seismology, University of Helsinki. She also lectures as a docent at the University of Helsinki.
Email:
[email protected] Address: Institute of Seismology, P.O. Box 68, FI-00014 University of Helsinki, Finland
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Dr. Markku Iljina is Senior Geologist at the Geological Survey of Finland. He also lectures as a docent at the University of Oulu.
Email:
[email protected] Address: Geological Survey of Finland, P.O. Box 77, FI-96101 Rovaniemi, Finland Dr. Juha Karhu is Professor of Geochemistry and Hydrogeology at the Department of Geology, University of Helsinki.
Email:
[email protected] Address: Department of Geology, P.O. Box 64, FI-00014 University of Helsinki, Finland Dr. Tapio Koistinen is retired State Geologist of the Geological Survey of Finland.
Email:
[email protected]
Dr. Kalevi Korsman is retired Research Professor of the Geological Survey of Finland. He lectures as a docent at the University of Helsinki.
Email:
[email protected] Address: Pietiläntie 6 A, FI-03100 Nummela, Finland
Dr. Kauko Laajoki is Professor of Geology and Mineralogy at the Department of Geosciences, University of Oulu.
Email:
[email protected] Address: Department of Geosciences, P.O. Box 3000, FI-90014 University of Oulu, Finland Dr. Erkki J. Luukkonen is Manager of Bedrock Geology and Mineral Resources at the Eastern Finland Unit of the Geological Survey of Finland.
Email:
[email protected] Address: Geological Survey of Finland, P.O. Box 1237, FI-70211 Kuopio, Finland Dr. Mikko Nironen is Senior Research Scientist at the Geological Survey of Finland. He also lectures as a docent at the University of Helsinki.
Email:
[email protected] Address: Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland Dr. Petri Peltonen is Senior Research Scientist at the Geological Survey of Finland. He also lectures as a docent at the University of Helsinki.
Email:
[email protected] Address: Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland
Dr. Raimo Lahtinen is Manager of Bedrock and Mineral Resources at the Southern Finland Unit of the Geological Survey of Finland. He also lectures as a docent at the University of Helsinki. Email:
[email protected] Address: Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland Dr. Satu Mertanen is Senior Geophysicist at the Geological Survey of Finland.
Email:
[email protected] Address: Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland Dr. Hugh E. O’Brien is Senior Research Scientist at the Geological Survey of Finland.
Email:
[email protected] Address: Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland Dr. Lauri Pesonen is Professor of Solid Earth Geophysics at the Department of Physical Sciences, University of Helsinki. He also lectures as a docent at the Technical University of Helsinki. Email:
[email protected] Address: Department of Geophysics, P.O. Box 64, FI-00014 University of Helsinki, Finland
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Dr. Peter Sorjonen-Ward is Senior Research Scientist at the Geological Survey of Finland. He also lectures as a docent at the University of Helsinki.
Dr. Matti Vaasjoki was Senior Research Scientist at the Geological Survey of Finland. He also lectured as a docent at the University of Helsinki. He died on November 23rd, 2004, at the age of 58.
Email:
[email protected] Address: Geological Survey of Finland, P.O. Box 1237, FI-70211 Kuopio, Finland Dr. Heikki Vartiainen is retired Chief Inspector of Mines of the Ministry of Trade and Industry.
Address: Vanha Turuntie 4 F, FI-02940 Espoo, Finland
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Dr. Jouni Vuollo is Senior Geologist at the Geological Survey of Finland. He also lectures as a docent at the University of Oulu.
Email:
[email protected] Address: Geological Survey of Finland, P.O. Box 77, FI-96101 Rovaniemi, Finland
INDEX OF PERSONS AND INSTITUTIONS Åbo Akademi University, 688 Academy Aboensis (Academy of Turku), 684, 685, 698 Alapieti, Tuomo, 688 Arppe, Adolf Edvard, 684 Backlund, Helge, 688 Barrow, George, 693 Bonsdorff, Pehr Adolf von, 684 Borgström, Johan Henrik Leonard (Leon. H.), 696, 697 Bowes, Donald R., 694 Bureau of Mines, 684, 685 Chair in Economic Geology, 689 Chair of Engineering Geology, 689 Chair of General and Applied Geology, 689 Chair of Geology and Mineralogy, 684, 687, 688, 689, 699 Chair in Quaternary Geology (Geology and Paleontology), 689 Chair in Solid Earth Physics, 688 Commissioner of Mines, 684 Department on Chemistry, 695 Department of Geology and Mineralogy 689 Department of Geophysics, 689 Department of Geology, 689 Department of Geosciences, 689 Edelman, Nils, 688 Ehlers, Carl, 688 Ekdahl, Elias, 687 Eklund, Olav, 689 Eskola, Pentti, 536, 686, 687, 692, 693–696, 699 European Diamonds PLC, 618 Finnminerals Oy, 689 Frietsch, Rudyart, 688 Gaál, Gabor, 692, 694 Gadd, Pehr Adrian, 684 Gadolin, Axel, 685, 698 Gadolin, Johan, 684 Geological Society of Finland, 682, 685, 699 Geological Survey of Finland, 4, 6, 30, 40, 41, 60, 104, 142, 198, 201, 413, 418, 427, 432, 448, 462, 492, 534, 577, 608, 614, 620, 629, 675, 685−687, 690, 691, 694−699 Geological Survey of Sweden, 492 Goldschmidt, V.M. (Victor Moritz), 695 Grubenmann, Ulrich, 693 Haapala, Ilmari, 688 Haapala, Paavo, 689 Hackman, Victor, 536, 686 Hanski, Eero, 689
Härme, Maunu, 692, 693 Hausen, Hans, 688 Helsinki University of Technology, 689, 699 Hietanen, Anna, 693, 694 Hjärne, Urban, 536 Hytönen, Kai, 697 Imperial Alexander University, 684 Institute of Seismology, 482, 552, 688, 699 Jernström, Anders Mauritz, 685 Kahma, Aarno, 687, 699 Kalm, Pehr, 684 Karhu, Juha, 689 Kauranne, Kalevi, 687, 695 Kemira Oy, 54, 689 Knorring, Oleg von, 697 Koistinen, Tapio, 694 Kontinen, Asko, 249, 694 Korsman, Kalevi, 694 Kouvo, Olavi, 4, 692, 696, 699 Laajoki, Kauko, 688, 689 Lahtinen, Raimo, 694 Laitakari, Aarne, 687, 696 Lappalainen, Veikko, 687 Lehtinen, Martti, 695, 697 Loukola-Ruskeeniemi, Kirsti, 689 Lyell, Charles, 684 Mäkinen, Eero, 689, 690, 696, 698 Malmikaivos Oy, 259, 618, 622, 623, 689, 698 Marmo, Vladi, 687 Matikainen, Raimo, 687 Metzger, Adolf, 688 Miyashiro, Akiho, 699 Mikkola, Aimo, 689 Mikkola, Erkki, 694 Mikkola, Toivo, 694 Mining Office, 684 Ministry of Trade and Industry, 684 Moberg, K.A. (Karl Adolf), 682, 685, 687 Moliis, Joseph, 536 Myllykoski Oy,j 689 Natural History Museum (Stockholm), 686 Neuvonen, K.J. (Kaarlo Juhana), 688, 694 Niini, Heikki, 689 Nironen, Mikko, 694 Nordenskiöld, Nils Adolf Erik (= A.E.), 685, 686, 698 Nordenskiöld, Nils Gustaf, 684−686, 697, 698 Nurmi, Pekka, 695 Otanmäki Oy, 689 Outokumpu Oy, 689, 690, 698 INDEX
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Outokumpu Oyj, 104, 690 Outokumpu Oyj Foundation, 689, 690 Outokumpu Mining Ltd, 261 Oy Lohja Ab, 689 Paarma, Heikki, 689 Palmén, J.A. (Johan Axel), 688 Papunen, Heikki, 688 Partek Oy, 689 Pehrman, Gunnar, 688, 696 Pesonen, Lauri J., 694 Piirainen, Tauno, 688 Posiva Co., 200, 223 Puranen, Mauno, 687 Rämö, Tapani, 695 Ramsay, Wilhelm, 686−688, 699 Rankama, Kalervo, 688, 695−697, 699 Rautaruukki Oy, 611, 617, 621, 689 Ridley, John, 67 Rosenbusch, Harry, 686, 692 Royal Ontario Museum (Canada), 198 Russian Academy of Sciences, 198 Sahama, Th.G. (Thure Georg), 536, 687, 688, 695-697, 699 Saksela, Martti, 688, 692 Savolahti, Antti, 536
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Sederholm, J.J. (Jakob Johannes), 6, 393, 446, 536, 682, 686, 687, 690–692, 695, 696, 699 Seitsaari, Juhani, 688 Siivola, Jaakko, 697 Simonen, Ahti, 692, 695 Sorby, Henry, 686 Steinheil, Fabian, 684 Stigzelius, Herman, 687 Sundblad, Krister, 689 Tampere University of Technology, 689 Technical Research Centre (Finland), 690 Tuominen, Heikki, 688, 694 University of Helsinki, 684, 686−688, 694−697, 699 University of Oulu, 104, 198, 688, 689, 699 University of Toronto, 198 University of Turku, 688, 694 Vaasjoki, Matti, 27, 536 Vartiainen, Heikki, 689 Väyrynen, Heikki, 689, 695 Vorma, Atso, 536, 697 Wahl, Walter, 536, 686, 692, 695−697 Wegmann, Eugène, 694 Wiik, Birger (H.B.), 697 Wiik, F.J. (Fredrik Johan), 685−687, 699
INDEX OF LOCALITIES Aakenustunturi, 147, 165−167 Ahmatunturi, 65, 67, 68 Ahmavaara, 107, 112, 114, 126, 127, 129−131 Ahola, 285, 289, 295, 300, 303, 326 Aholanvaara, 64 Ahvenisto, 537, 539, 542, 544− 546, 548 Ahvenlammi, 354, 355, 367− 370 Akanvaara, 116, 118, 120, 122, 146, 150, 153, 154, 167−171, 215, 226, 285, 291, 298, 302, 303, 305, 306, 323, 329, 330, 587 Ak’kanasvarri, 69 Alajärvi, 354 Alakylä, 312 Åland, 15, 460, 537, 539, 541, 545, 546, 548, 549, 556, 567, 573, 574, 575, 577, 581, 582, 584, 655, 688 Ala-Penikka, 107–109, 122, 123, 127, 128, 130, 131, 294 Alasenlahti, 368 Ala-Siikajärvi, 313 Alavieska, 363 Almesåkra, 567 Alnö, 615, 627 Antinneva, 363 Arkhangelsk, 619, 631, 634, 635 Åva, 450, 459−462, 470 Arvidsjaur, 182, 388 Aulanko, 348 Bamble, 656 Belomorian, 22−24, 26, 28, 64– 66, 74, 77, 78, 82, 175, 487, 489, 494, 496, 500, 501, 503, 507 Bergslagen, 5, 346, 351, 382, 388, 390−392, 396, 487,489, 491, 494, 499, 500, 503, 507−509, 511, 512, 519 Biennaroavvi, 69 Bjørnevatn, 71, 73 Bodom, 537, 539, 548 Burakovka, 201, 226, 652 Çaravarri, 182 Dala, 567, 572, 588 Dividal, 582–585, 591
Eckerö, 577 Edefors, 514 Ekojoki, 436 Enontekiö, 7 Eräjärvi, 450, 460 Erivaaransuo, 305, 306, 330 Eskosenvaara, 311 Eurajoki, 537, 539, 540, 543, 545, 549 Evijärvi, 346, 350−352, 357, 361−364, 389, 394, 395 Fedorova Tundra, 653 Finnmark, 656 Fjälskär, 537, 539 Föglö, 548 Forssa, 375−378, 426 Gardsjø, 73 Gävle, 567, 573 Gulf of Bothnia, 15, 330, 362, 482, 488, 494, 499, 507, 567, 579, 586 Haajainen, 317, 318 Haapajärvi, 355 Haapala, 305, 307 Haapalanmäki, 320, 321, 323 Haapamäki, 374 Haaparanta, 143, 165, 167, 181, 183, 290, 455 Haasianvaara, 51 Haaskalehto, 153, 171, 172 Hailuoto, 573, 580, 591 Haisuvuoma, 152 Hakasuo, 297 Hallakulma, 302, 312 Hallavaara, 302 Halmejärvi, 51 Halpanen, 617, 631 Halti, 585, 586 Häme, 12, 346, 348, 350, 351, 354, 375−378, 391, 392, 395, 396, 411, 417, 426, 437, 487, 490, 499, 507, 509−511, 539, 548 Hämeenkyrö, 453 Hämeenlinna, 346, 375, 377 Hamina, 539, 546 Hammaslahti, 313, 698 Hangasoja, 147, 152, 153, 167 Hanko, 470 Haparanda, 143, 455 Harhala, 370 Harjavalta, 572
Härnö, 514 Hassela, 487, 498 Hattu, 24, 28, 29, 31−33, 35− 38, 48, 52, 82 Haukilampi, 305, 310 Haukivesi, 615, 654, 658 Haukivuori, 346, 354, 355, 359, 361, 378, 379 Hautavaara, 306 Haveri, 352, 365−368, 370− 373, 394 Heinävesi, 309 Heinola, 377, 539 Helsinki, 2, 539, 548, 552, 682, 684, 686−689, 693,695−697, 699, 700 Hepoköngäs, 292, 293 Hetehongikko, 291, 292, 307 Hetta, 69, 449, 470, 472 Hietaharju, 44 Himanka, 353, 354 Himmerkinlahti, 289, 310, 311, 323, 324 Hirsilä, 374, 375 Hirsimaa, 290 Hirvaskoski, 59, 62, 282, 283, 285, 287, 326, 611, 614, 615 Hitura, 413, 419, 698 Hogland, 546 Hoivasvuori, 368, 373 Hokkalampi, 36, 291, 295, 301, 302, 305, 329 Honkajärvi, 297, 299 Honkala, 297 Honkavaara, 63 Horsmanaho, 243 Hosko, 33 Höytiäinen, 12, 282, 291, 307, 313, 315, 319 Humaljärvi, 321 Huosiuslampi, 305 Huuskonvaara, 51 Hyrynsalmi, 297, 311, 314, 315 Hyvinkää, 348, 353, 354, 375, 376, 380, 391, 413, 421, 426, 428, 429, 434, 435 Hyypiänmäki, 383 Iilijärvi, 348 Iisalmi, 11, 30, 51, 53, 54, 56− 58, 61, 65, 74−78, 81, 82, 201, 240, 273, 282, 285, 291, 301−303, 317, 325−327, INDEX
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329, 330, 455, 487, 489, 651, 653 Iitti, 542, 545, 548, 552 Iivaara, 16, 42, 593, 608, 628− 632, 634, 636 Ikaalinen, 346, 371 Ilkonahonkallio, 318 Ilomantsi, 11, 13, 20, 28−32, 35, 36, 38, 40, 45, 52−54, 56, 58, 60, 75, 78, 80−83, 228 Ilvesvaara, 292, 295 Imandra, 487, 488, 494, 501, 653, 673, 674 Inari, 7, 24, 70−72, 142, 175, 176, 449, 487, 494, 496, 501, 503, 654, 685 Ipatti, 28, 36 Iso-Kartano, 355 Iso-Naakkima, 583, 586, 589 Itämäki, 317 Ivalojoki, 685, 698 Jaala, 539, 542, 545, 548, 552 Jalka-aho, 298, 323 Jalokoski, 654, 658 Järvenpää, 348, 380, 381, 390− 392, 395, 396 Järvikäinen, 151 Jeesiörova, 153, 156, 157 Jer’gul, 69 Jero, 309, 310, 329 Jerta, 582, 583, 585 Jokijyrkkä, 320 Jönköping, 487, 491, 511−513 Jonsa, 57 Jormua, 6, 11, 13, 25, 74, 153, 162, 203, 238, 240−247, 249−255, 257, 263−273, 283, 285−287, 291, 314, 318, 319, 325, 487, 497, 517 Joroinen, 348, 352−354, 359, 360, 377, 379 Jouttiaapa, 153, 290, 309, 313 Juuanvaarat, 307, 313 Juurikka, 314 Juurikkaniemi, 44 Juva, 352, 379, 434 Kaarestunturi, 165 Kaavi, 50, 606, 616, 619−622, 631−639 Kainuu, 11, 53, 59, 201, 203, 215, 221, 226, 240, 282−284, 287, 288, 291−293, 295, 297, 299, 302, 303, 305−307, 309−311, 313−315, 317−320, 323−327, 329−331,586 Kaipola, 413, 423−425, 427, 430, 435 Kaivopuisto, 682
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INDEX
Kajaani, 50, 54, 238, 320, 325, 326 Kalak, 585 Kalanti, 468 Kalkku, 348 Kalkunmäki, 310 Kallinkangas, 306 Kalliomaa, 318 Kalliovaara, 323 Kangasala, 368 Kankaanpää, 346, 371 Kannus, 353, 354 Kapsajoki, 152, 163 Karakkalehto, 60 Karankaniemi, 196 Karasjok, 175, 656 Karelia, Russian, 26, 64, 78, 142, 154, 171, 177, 179, 196, 198, 201, 203, 204−210, 213, 214, 218, 226, 228, 290, 295, 310, 324, 537, 555, 556, 566, 574, 578, 618, 619, 698 Kärenvaara, 37 Karikkoselkä, 16, 584, 592 Karjakko-oja, 151 Kärki, 422, 425 Karkuvaara, 283, 285, 289, 295, 297, 300, 326 Karpinskiy, 631 Karstula, 584 Katajalampi, 307 Kauhajärvi, 428−431 Kaukua, 105, 114, 115 Kautokeino, 182, 578, 579, 656 Kautoselkä, 151, 153, 159, 161, 180, 181 Keinokangas, 290, 317, 318, 675 Keitele, 346, 351, 356, 361, 362, 365, 393, 394, 494, 498, 500, 503−505, 507−509, 511, 516, 517 Keivitsa, 153, 157, 158, 172− 174, 202, 203, 221 Keivitsansarvi, 174 Keivy, 672 Kelkkakangas, 45 Kellojärvi, 40, 46, 47 Kellostapuli, 147, 152, 153, 167 Kemi, 59, 104, 106, 107, 108, 116, 118, 120, 122, 132, 133, 291−294, 313, 331, 698 Kemi airport, 293 Kemihaara, 65, 67 Kemijärvi, 63, 282, 285, 287 Kemiö, 348, 380−382, 384− 387, 390−392, 395, 396 Kermavesi, 30 Keski-Penikat, 108, 294
Kettupeä,, 348 Khibina, 608 Kiannanniemi, 40 Kianta, 29, 30, 36, 38−45, 48, 51−53, 56, 58−60, 62, 75, 79−82, 282 Kiekki, 284, 285, 291, 299, 329 Kieksi, 573 Kiihtelysvaara, 284, 295, 297, 301, 305, 310, 315, 675 Kiikoinen, 346, 371 Kiimarova, 151, 163 Kiimaselkä, 152 Kiiminki, 11, 282, 284, 285, 287, 313, 314, 317, 318, 331 Kilvenjärvi, 110, 112, 129 Kirintökangas, 285, 305, 306, 313, 324, 326, 329, 330 Kirkkonummi, 2 Kirkkovuori, 539 Kisko, 348, 381, 383−387, 392 Kiskonkoski, 312 Kitka, 283, 289, 324 Kittilä, 140, 142−148, 151−167, 171, 175, 179−183, 240, 242, 262, 263, 271, 272, 487, 489, 494, 497, 500, 503, 504, 517, 518 Kiuasautonoja, 151 Kiukainen, 564 Kiuruvesi, 654, 658 Kivakka, 215, 226 Kivalo, 155, 177, 290, 301, 303, 306, 675 Kivesvaara, 320 Kiviaapa, 169 Kivijärvi, 49 Kivipurnuvaara, 151 Knaften, 487, 489, 494, 497, 498, 503, 504 Knapernummi, 572 Koillismaa, 43, 48, 59, 62, 63, 81, 104−108, 114−116, 118−124, 131, 132, 210, 291, 293, 611, 652 Koitelainen, 68, 114, 116, 118− 122, 146, 149, 150, 153, 167, 169−172, 215, 226 Koiteli, 318 Koivumäki, 44, 47 Koivusaarenneva, 413, 428, 430, 431, 433, 435−437 Koivusaari, 315, 317, 675 Kökarsfjärden, 537, 539, 548 Kokkola, 362 Kola Peninsula, 14, 64, 65, 175, 177, 198, 202, 203, 290, 292, 487, 567, 608, 653, 658, 672, 673, 688
Kolari, 144, 150 Koli, 36, 213, 220, 221, 223, 224, 284, 305, 307, 309, 310 Kolinummi, 546 Koljola, 310, 324, 675 Kolkonkangas, 309 Kolmiloukkonen, 284, 285, 289, 310, 323, 324 Kolunkylä, 344, 366–368, 370 Kolvitsa, 182 Kometto, 312 Konchozero, 654 Köngäs, 151, 160, 161 Konivaara, 52 Koitoiva, 151 Konttijärvi, 106, 107, 111−114, 118, 120, 122, 125, 126, 129−132, 205 Korkiavaara, 290, 323 Korpijärvi, 547 Korpivaara, 31, 39 Korppoo, 380, 382, 384−387, 395 Korsnäs, 576, 577 Kortejärvi, 611, 614, 615, 632−635 Kortevaara, 317 Korvuanjoki, 284, 286, 291, 303, 305, 307, 311, 329, 330 Koskenniemi, 622 Koskuenjärvi, 348, 367, 372 Kostamuksha, 82, 618, 619, 634, 635 Kotajärvi, 348, 352, 358 Kotalahti, 413-416, 419, 433− 436, 698 Kotila, 314 Koutoiva, 151 Kovasinvaara, 312 Kovdor, 625 Kovero, 24, 28, 36, 38, 40 Kuhmo, 11, 24, 40, 41, 43−45, 47−49, 51, 60, 75, 78, 81, 200, 201, 203−205, 207−209, 213−215, 221, 223, 224, 226, 228−230, 282, 284, 285, 291−293, 295, 297−300, 309, 324, 326, 329, 330, 575, 611, 614, 617−619, 631−635, 639 Kuittila, 32, 35, 38, 39 Kuljunki, 33, 35 Kuloi, 619 Kumisevanmäki, 54 Kummitsoiva, 145, 156 Kumpu, 140, 144, 146, 147, 152, 155, 164−167, 181−183, 324 Kumputunturi, 140, 160, 165 Kuntijärvi, 292
Kuopio, 285, 287, 292, 301, 313−315, 317, 319, 331, 606, 608, 619−622, 631−637, 639, 687 Kuorboaivi, 71, 175 Kuotko, 152 Kuovila, 348 Kurkikylä, 286, 291, 293, 297, 300, 303, 305, 329, 330 Kuru, 348, 374, 548 Kutsu, 82 Kuusaa, 356 Kuusamo, 11, 62, 64, 65, 105, 131, 142, 143, 154, 213, 217, 226, 228, 282, 283−285, 287, 289, 290, 292, 295, 297, 300, 303, 305, 306, 309, 311, 320, 323−326, 329, 330, 487, 501, 586, 608, 673, 674 Kuusijärvi, 105, 115−117, 123, 124, 284, 285, 290, 292, 300, 303, 306, 329 Kvartsimaa, 290, 309, 313, 330 Kylmäkoski, 368, 408, 421 Kylylahti, 258, 261, 262 Kymi, 537, 543 Kynsijärvi, 115 Kytö, 425 Kyykkä, 295 Laanhongikko, 293 Laanila, 15, 577−579, 589, 655, 656 Ladoga, 4, 15, 282, 283, 285, 319, 537, 566, 567, 569, 574, 577, 578, 588, 591, 654 Lähdemäki, 302, 307 Lahti, 375, 377 Lahtojoki, 638 Lainio, 145−148, 152, 153, 155, 164−167, 181−183 Laitila, 15, 537, 539, 541, 543, 545, 546, 548−550, 556 Laivajoki, 611, 614, 632−635 Lambina, 653 Lammaistenkoski , 572 Lapinlahti, 412−414, 416−418, 421, 435 Lappajärvi, 16, 574, 588, 589, 593, 698 Lappeenranta, 546 Latvajärvi, 152, 153, 163−167, 181, 182 Lauhanvuori, 15, 428, 431, 432, 580−582, 584, 586, 591 Laukunkangas, 413−415, 433− 436 Lautaporras, 376 Lavia, 346, 355 Lehtomäki, 305
Lemland, 450, 460, 461 Lempiänniemi, 348 Lentiira, 618 Lieksa, 29, 30, 35, 37, 38, 40, 82, 652 Lina, 76, 459, 514 Linkujoki, 152 Linkupalo, 147, 152, 153, 156, 166 Lipeävaara, 105, 114, 115, 117, 123 Lippumäki, 292, 301 Lofoten, 674 Lohja, 684 Loljunmaa, 119, 120, 122 Los, 392, 393 Losomäki, 257, 258, 267−270 Lovasjärvi, 548 Lovozero, 608 Luhanka, 371 Luikonlahti, 259, 265, 698 Lujaur Urt, 608 Lukkulaisvaara, 116, 118−121, 215, 226 Luleå, 389 Lumimäki, 54 Lumparn, 15, 16, 584 Luonteri, 450, 460, 470 Luossajavri, 72 Luvia, 572 Maarianvaara, 30, 259 Mäkipalo, 317 Makola, 436, 698 Mälaren, 567 Manamansalo, 50, 54, 282 Mantovaara, 147, 152, 153, 165−167 Mäntyharju, 539, 545, 547 Mäntykangas, 298 Mäntyvaara, 150, 151 Märket, 575–577, 655 Martimo, 290, 291, 318 Marttivaara, 703 Matinvaara, 299 Mauri, 354, 367 Melalahti, 311, 315 Metsäkulma, 572 Miihkali, 258, 265 Mikkeli, 617 Moisiovaara, 44, 49 Monchegorsk, 653 Moresveijohjkan, 71 Mosshaga, 450, 460 Mt. Generalskaya, 177, 653 Möykkelmä, 66, 68, 145, 150− 152, 154 Muhos, 15, 285, 567, 573, 574, 580, 589 Multivuori, 370 INDEX
•
711
Multsilta, 383 Muonio, 69, 70 Murtolampi, 105, 114, 115 Mustamaa, 318 Mustavaara, 123, 132, 323 Myllyniemi, 366−370, 372, 375, 394 Mynämäki, 537, 539 Naapurinvaara, 314 Naarva, 35, 38, 82 Näätämö, 73 Näätäniemi, 45 Nabar, 583, 585, 586 Nalganas, 583, 585 Napapiiri, 62, 63, 66, 81 Näränkävaara, 42, 101, 104− 107, 114, 116−120, 131, 132, 167, 201, 205, 226 Narkaus, 106, 107, 110, 111, 116, 118−120, 126, 128−130, 132, 292, 293, 301 Naruska, 65, 67, 68 Naruskajoki, 67 Naruskajärvi, 587 Näsijärvi, 344, 366, 370 Nattanen, 68, 76, 143, 461, 462, 471 Naulaperä, 312 Nauvo, 380, 382, 384−387, 395 Neiden, 73 Nenäkangas, 291, 307, 309, 311 Nilovaara, 285, 289, 303, 306, 326, 330 Nilsiä, 284, 301, 302, 307, 309, 313, 616 Nokia, 354, 367, 369 Nolppio, 152 Nordingrå, 567, 573 Norrbotten, 487, 489, 494, 497, 498, 500, 501, 503-505, 508, 509, 511, 517, 518 Norrgrynnan, 577 Norrlammala, 348 North Karelia, 201, 203, 213, 215, 217, 220, 221, 224, 228, 240, 255, 256, 282, 283, 287, 288, 291, 292, 295, 297, 301, 302, 305, 306, 309, 310, 313−315, 317, 323−326, 331 Nuasjärvi, 287, 291, 313, 314, 319, 320 Nunnanlahti, 24, 28, 36, 37, 56 Nuolusvaara, 68 Nurmela, 297 Nurmes, 29, 30, 51, 52, 75, 82 Nurmo, 362, 364 Nuttio, 151–153, 161, 162, 240, 241, 243, 244, 250, 262−269, 271, 272
712
•
INDEX
Nyssäkoski, 152, 153, 163, 164, 166, 167 Obbnäs, 537, 539, 548 Ohravaara, 297 Oijärvi, 24, 50, 59−61 Ojamo, 684 Olanga, 104, 116, 121, 652 Onas, 537, 539, 542, 548, 655 Onega, 177, 178, 226, 654, 673 Onkamo, 145−148, 150−156, 158, 171, 176, 178, 179, 263 Onkamonlehto, 150 Ontojärvi, 44 Opukasjärvi, 70–72 Oravisalo, 319, 321 Orijärvi, 348, 351−354, 380− 388, 390−392, 396, 684, 693, 694, 698 Oripää, 470 Orivesi, 367, 372, 373, 699 Osara, 366, 367, 370, 372, 394 Oskarshamn, 487, 491, 511, 512, 513 Ossaus, 309 Otanmäki, 202, 203, 221, 286, 317, 689, 698 Oulanka, 226 Oulu, 320, 459, 573 Oulujärvi, 53, 59, 282−285, 298, 299 Ounasvaara, 290 Outokumpu, 12, 161, 162, 203, 240, 241, 243, 244, 250, 252, 255−273, 282, 288, 291, 313, 318−321, 325, 327, 484, 487, 497, 689, 698, 699 Pääjärvi, 196, 201, 228, 652, 656 Paanajärvi, 285, 292, 297, 300, 306, 326 Paasivaara, 107−109, 122, 123, 130 Paasselkä, 16 Päästäispuro, 313, 323 Pahakangas, 45, 46 Paloinen, 444 Palokivalo, 290, 303, 306, 676 Paltamo, 287, 297, 301, 305, 311, 320 Pampalo, 31, 33, 35, 36, 39 Pampalonuurro, 31 Panelia, 564 Pärekangas, 285, 291, 298, 305, 307, 309, 323, 329, 330 Parikkala, 377, 378, 396 Parkano, 348, 374 Parkkila, 450, 460, 471 Pärnäjärvi, 539 Pasvik, 70, 73, 292, 295
Pechenga, 70, 73, 153, 162, 177, 178, 292, 487, 488, 494, 501, 673, 674, 698 Peipohja, 537, 539 Pekkarinen, 703 Pelkosenniemi, 165, 577 Pellinki, 348, 352, 380, 383, 385, 386, 388, 390, 391, 395, 396 Penikat, 59, 102, 104,106−111, 116, 118−120, 122, 123, 125, 127, 128, 130, 132, 153, 215, 291−294 Penikkajärvi, 151 Perämaa, 413, 428, 430, 432, 434, 435, 437 Peräpohja, 11, 59, 105, 106, 131, 142, 143, 153, 155, 171, 172, 177, 201, 211, 212, 217, 282−285, 287, 290, 292−295, 300, 303, 306, 309, 313, 314, 317, 318, 320, 323−326, 329, 330, 487, 501, 670, 672, 673, 675, 676 Perho, 374 Petäikkö, 310 Petonen, 315 Petsamo, 698 Peura-aho, 44, 45 Peuranpalo, 670 Peurasuvanto, 146 Pielavesi , 313, 317, 319, 348, 352−354, 356, 358, 359, 393, 395, 616 Pihlajavaara, 39 Pihtipudas, 348, 374 Piimäsjärvi, 421 Pirilä, 450, 460, 471 Pirivaara, 72, 73, 105, 114−116, 123 Pirkanmaa, 12, 346, 350, 351, 359, 365, 368, 371−374, 388−390, 394−396, 411, 412, 417, 419, 421−423, 426, 427, 433, 434, 437 Pirttimäki, 50, 54 Pirttiniemi, 348, 370 Piteå, 487 Pitkäranta, 698 Pitukansuo, 307, 311 Pogosta, 40 Pohjanmaa, 12, 346, 348, 350− 357, 361−363, 388, 389, 394−396, 654, 658 Pohtola, 370 Poikkimaa, 290 Pölkkylampi, 297 Pomokaira, 24, 61, 65, 68, 69, 71, 74, 76
Pomovaara, 462 Pori, 548, 569, 572, 576 Porkkalanniemi, 2 Porkonen, 158 Porrasniemi, 413, 422, 424, 425, 428, 433, 434, 435 Portimo, 59, 104, 106, 107, 111−114, 119, 120, 122, 123, 125, 126−130, 132 Porttivaara, 105, 107, 114, 116− 118, 120, 123 Porvoo, 383 Posio, 62, 282, 284, 285, 297, 300, 306, 310, 311, 323, 326, 329, 330 Pöyliöjärvi, 290, 323 Pudasjärvi, 11, 62, 81, 105, 131, 142, 201, 204, 205, 207, 226, 228, 230, 273, 282, 284, 285, 291, 292, 294, 317, 318, 325, 326, 329, 330, 487, 489, 611, 614 Puiroonmäki, 256 Pukala, 348, 366 Pulesjärvi, 344, 366−368, 370 Punkaharju, 377, 378, 396 Puolanka, 59, 81, 282, 284−287, 291, 298, 299, 302, 305−307, 309−312, 314, 315, 320, 321, 326, 330, 331 Puolankajärvi, 285, 291, 295, 297−300, 303, 305, 329, 330 Puruvesi, 30, 470 Puso, 310 Pyhäjoki, 363 Pyhäsalmi, 348, 352, 356, 358− 360, 395, 698 Pyhätunturi, 154, 165, 183 Pyhitys, 105, 114, 115, 117 Pylsynlahti, 370 Pyssykulju, 284, 286, 298, 321 Raahe, 4, 282, 283, 285, 319, 358, 363, 487, 654 Raatevaara, 288 Raiguba, 672 Raisædno, 69 Rajala, 151 Rantamaa, 290, 309, 313, 318, 330 Rantasalmi, 352–354, 377, 378 Ranua, 48, 51, 53, 59, 60, 63, 74, 75 Råstojaur, 70 Rautalampi, 358−360 Rautavaara, 50, 54, 56, 58, 74, 78, 81, 82 Reittiö, 301, 329 Renko, 450 Repolampi, 45, 49
Reposaari, 537, 539 Ridnitsohkka, 585, 586 Rieskavaara, 314 Riestovaara, 461, 462 Riitavuori, 348 Ristiina, 353, 354, 379 Ristijärvi, 284, 287, 311, 577− 579, 655, 656 Roninkangas, 317 Ronkonriutta, 309 Ropi, 24, 61, 69, 70 Ropitunturi, 69 Rotimojoki, 317 Rovaniemi, 306, 458, 462, 471, 687 Rukatunturi, 289, 309, 311, 330 Runkaus,201, 211, 290, 301, 303, 326 Runkausvaara, 153, 292 Ruoholampi, 546 Ruokonen, 314 Ruoppapalo, 152, 153, 164, 166 Ruossakero, 69 Ruukki, 358 Rybreka, 654 Ryönä, 606, 622 Rytikangas, 107, 112, 122, 123, 126−128, 130, 131 Sääksjärvi, 16, 588, 591 Sääperi, 291, 292 Saari, 284, 285, 291, 299, 329 Saarijärvi, 348, 374, 583 Saarikylä, 40, 42, 43, 44 Saarisenjärvi, 413, 414, 437 Saimaa, 12, 350, 352−355, 359, 375, 377−379, 395, 396 Sakulahti, 703 Salahmi, 284, 285, 287, 302, 303, 307, 313, 314, 317, 318 Salittu, 352, 380, 381, 383−387, 390−392, 396 Salla, 15, 65, 142, 143, 145− 156, 167, 169, 176, 178−180, 183, 204, 226, 263, 285, 287, 577, 579, 580, 589, 655 Salmi, 76, 537, 553, 556, 567, 569, 588 Salmijärvi, 287, 307, 315, 321, 330 Sammatinjärvi, 348, 352, 367 Säppi, 576 Särkilampi, 291, 292, 297 Sarvisoaivi, 69 Satakunta, 15, 567, 569, 571− 576, 579, 588, 589, 655 Sätkänävaara, 147, 166 Satovaara, 172 Sattasvaara, 145, 156 Säviä, 356, 358, 359
Savo, 12, 282, 287, 288, 318, 326, 346, 348, 350−362, 367, 380, 393−395, 487, 489, 494, 498, 500, 503−505, 508, 509, 517, 518 Savukoski, 64, 145−148, 152, 155−159, 166, 172, 177, 179, 181, 263, 586, 608 Seglinge, 450, 460, 462 Seinäjoki, 48, 361, 362, 364 Seitaperä, 618, 619 Selkäsenvuoma, 151 Sievi, 353, 354 Siika-Kämä, 107, 110, 122, 126, 128−131 Siikavaara, 309, 312 Siilinjärvi, 13, 50, 53, 54, 56, 65, 81, 608, 610−613, 617, 631, 632, 634, 635, 675, 690, 698 Siipyy, 361, 537, 539 Siivikkala, 354, 368, 369 Siivikko, 46 Siivikkovaara, 40, 45−47 Sileäkallio, 368, 370 Silisjoki, 72 Sillböle, 684 Silvevaara, 29, 30, 32, 37, 39 Simo, 50, 51, 59 Sirkka, 144, 156, 158, 183 Siurua, 13, 50, 60, 61 Siurunmaa, 586 Sivakkojoki, 33 Skellefte, 5, 388, 389, 394, 395, 487, 489, 497, 503, 504, 508, 514, 518 Sodankylä, 142−148, 150, 154− 156, 158, 165, 167, 171, 177, 179, 181, 263 Söderfjärden, 15, 16, 584 Sokli, 16, 67, 68, 143, 593, 608, 615, 621, 623, 625−628, 631−635, 689 Somerjärvi, 284, 286, 298, 307, 311, 320 Sompujärvi, 107−109, 122−124, 128, 130, 294, 301, 676 Sompuvaara, 290 Sonkajärvi, 54 Sørvaranger, 24, 70−73 Sotkamo, 311 Sotkaselkä, 156 Sotkuma , 295, 309 Soukkio,412, 413, 426 Sovasjoki, 158 Suhanko, 106, 107, 111−114, 118–122, 126, 127, 130−132, 205, 285, 292, 300, 326, 329 Sulva, 15 INDEX
•
713
Summa, 539 Suodenniemi, 346, 368, 369, 371, 373 Suomenniemi, 37, 539, 542, 545−548, 553 Suomu, 63, 64, 66 Suomujärvi, 64, 294 Suomussalmi, 13, 24, 40, 42− 45, 47 Suopelto, 539 Suoperä, 201 Suorre, 71, 72 Suur-Pellinki, 383 Suursaari, 546 Suvasvesi, 16 Svartälven, 567 Syväjoki, 297 Syöte, 105, 114−119, 123 Taalikkala, 546 Täilahti, 256 Taivaannaapuri, 539 Taivaljärvi, 40, 47, 48 Taivalkoski, 200, 201, 204, 207, 209, 226, 228, 230, 317, 318 Takamaa, 348, 366, 367, 370− 373 Talvivaara, 315 Tammela, 376, 690 Tampere, 12, 14, 346, 348, 350−355, 362, 365−375, 377, 388, 391, 393−396, 411, 412, 417, 423, 427, 484, 487, 489, 498, 505, 516, 696 Tanaelv, 175, 176, 182, 183 Tapanila, 703 Tarkki, 540, 545 Tarvasenvaara, 151 Tasanvaara, 31, 38, 39 Tepasto, 462 Terrinen, 425 Tersk, 175, 487, 494, 496, 500, 501, 567 Tervakivi, 348, 367, 370 Tervola, 670 Tesoma, 348, 365 Tetrinmäki, 358 Teuravuoma, 150 Tievjan, 71, 72 Tiiliharju, 581 Tiirismaa, 352−354, 375, 380, 391−393, 396 Tiittalanvaara, 31, 33, 35 Tikanmaa, 290, 309, 313, 330 Tilsa, 105, 114, 115 Tipasjärvi, 24, 30, 36, 40, 44, 47, 48, 51, 52 Tirmo, 385
714
•
INDEX
Tohmajärvi, 228, 313, 315 Toija, 381, 383, 384, 387 Toivakka, 444 Tojottamanselkä, 66, 68, 69, 146, 169 Torikylä, 314 Tornio, 101, 104−107, 113, 116, 118−120, 131, 132, 167, 201, 205, 226 Tsipringa, 210, 215, 226 Tšuomasvarri, 72, 654 Tulisaari, 653 Tulomozerskaya, 673 Tulppio, 65, 68, 625 Tuntsa, 24, 64−68, 71, 74, 81 Tuomivaara, 314, 317 Turjanniemi, 608 Turku, 450, 451, 458, 462, 548 Tuulijoki, 165, 182 Tuuliniemi, 368, 370 Tyypekinlampi, 413, 414, 437 Ukkolanvaara, 20 Uljaste, 581 Umba, 175, 182, 487, 494, 496, 500, 501, 503, 504 Umeå, 389, 394, 494, 499, 513, 518, 519 Umptek, 608 Unikumpu, 300 Urjala, 368 Urkkavaara, 295, 297 Utajärvi, 283, 285, 318 Utrio, 35 Utsjoki, 685 Uudiskorhola, 436 Uusimaa, 12, 346, 348, 350, 352−354, 375, 376, 380, 383−388, 390−392, 395, 396, 487, 489, 499 Väärälampi, 546 Vaasa, 15, 574, 576, 577, 655 Vaddas, 583, 585, 586 Vähä-Kassari, 368 Vähä-Lima, 355 Vainospää, 72, 462, 471 Vaivanen, 315, 317 Väkkärä, 539, 545 Valamo, 15, 578 Valijärvi, 348 Välimäki, 355, 698 Välivaara, 305 Vammala, 346, 368, 396, 413, 419, 421, 423, 424, 432−436, 698 Vanttauskoski, 306 Varissaari, 348 Värmälä, 453
Varpaisjärvi, 50, 54, 56−58, 217, 228, 230, 650−653 Varzuga, 487, 488, 494, 501, 673, 674 Vaskojoki, 176 Vätsäri, 73 Väyrylänkylä, 297, 298 Väystäjä, 290, 318, 675, 677, 678 Vazhinka, 654 Vehmaa, 15, 537, 539, 545, 556 Veikasenmaa, 151–153, 160, 163 Veitsivaara, 223 Veittijärvi, 370 Venejärvi, 63 Vesikkovaara, 147, 152, 153, 167 Vesivaara, 305 Vesmajärvi, 151, 153, 159−161, 163 Vesterålen, 674 Vestlax, 384−387 Veteli, 468 Vihajärvi, 284, 286, 291, 298, 307, 320, 323 Vihanti, 348, 698 Viianki, 119, 120, 122, 228 Viinaränninnotko, 370 Viistola, 310 Viljakkala, 366, 368, 370 Vimpeli, 354 Vintilänkaira, 65, 67 Virtasalmi, 348, 358−361, 375, 377−379, 394, 395, 586, 589, 698 Vitikkovaara, 303 Vittangi, 654, 658 Vittinki, 362, 364 Voche, 653 Vuokatti, 284, 287, 311, 314 Vuoriniemi, 47 Vuorivaara, 312 Vuosanka, 44 Vuotto, 318 Wiborg, 15, 534, 537, 539−546, 548−553, 556, 655 Ylikiiminki, 318 Yli-Penikat, 108 Ylivieska, 346, 350, 351, 356, 357, 361−364, 374, 388, 389, 394, 395 Ylämaa, 545 Ylläs, 165 Ylöjärvi, 354, 366−368, 370, 372, 373, 698 Yräjärvi, 152, 163
SUBJECT INDEX Abitibi belt, 81 Accretion, 58, 77, 79, 81, 82, 346, 365, 446−449, 473, 474, 488, 489, 501, 503, 505, 507, 512, 515−517, 520, 587, 637, 655, 660 Accretionary arc complex, central and western Finland, 5, 350 Accretionary arc complex, southern Finland, 5, 350, 359 Accretionary orogen, 346, 393, 488 Accretionary prism, 26, 52, 79, 351, 516, 518 Accretionary processes, 26−28 Accretionary stage, 507, 508, 520 Accretionary unit, 484, 499 Accretionary wedge, 82, 417, 433, 498, 504, 505, Actinolite, 31, 48, 262, 614 Aegirine(-augite), 547, 611, 625, 628−630 Aeromagnetic anomaly, 418, 427, 432, 489, 492 Aeromagnetic data, 39, 57, 59, 64, 223, 614 Aeromagnetic map, 8, 105, 147, 198, 200, 201, 203, 212, 228, 230, 323, 417, 418, 423, 426, 427, 432, 462, 489, 577, 614, 625, 629 AFM diagram, 149, 251, 252, 434, 435 AGDF (axial symmetry of the geocentric dipole field hypothesis), 650 Age group, Karelian metadiabases, 324 Age group, layered intrusions, 120 Age group, mafic dike swarms, 203 Age group, mafic−ultramafic rocks, 155 Age group, Mesoproterozoic, 574, 577, 586 Age group, Mesoproterozoic granitoids, 446, 448−452, 455, 456, 458, 459, 462, 468−470, 474 Ages, Archean, 14, 22, 27, 29, 38, 39, 43, 44, 49, 52, 53, 56−59, 63−65, 68, 69, 71, 81, 175, 243, 254 Ages, chronogram/histogram, 6, 27, 413, 448, 541, 567 Ages, depleted mantle model, 176, 467−471, 576, 577 Ages, K-Ar, 14, 53, 56, 74, 573, 586, 615 Ages, Lu-Hf, 76 Ages, Rb-Sr, 42, 43, 48, 68, 69, 619, 627 Ages, Sm-Nd, 5, 13, 14, 15, 39, 48, 56, 57, 59, 61, 68, 75, 76, 120, 142, 151, 153, 154, 157, 163, 174, 176, 181, 200, 206, 207, 211, 215, 221, 223, 226, 243, 251, 313, 315, 651, 652, 654, 655 Ages, U-Pb, 6, 14, 22, 26, 27, 29, 37−39, 44, 46, 48, 49, 53, 57, 58, 62, 63, 65, 68, 69, 71,73, 75, 120, 142, 149, 150, 153, 154, 157, 163−165, 167, 171, 172, 175, 176, 181,183, 203, 204, 206, 207, 211, 213, 217, 219, 221, 223, 226, 243, 257, 300, 307,310, 318, 351, 352, 365, 374, 380, 381, 386,
390, 412, 413, 426, 437, 484, 514,541, 545, 569, 574, 575, 586, 611, 615, 616, 619, 638, 651, 652, 674−676, 692 Ages, zircon age, 5, 6, 14, 29, 37−40, 43, 44, 46, 48, 49, 52, 53, 57−59, 62−65, 69, 71, 73, 81, 120, 149, 150, 152, 154, 157, 163−165, 167, 171, 172, 175, 176, 203, 206, 211, 219, 243, 257, 300, 307, 310, 318, 351, 353, 365, 374, 380, 381, 386, 390, 412, 413, 416, 426, 430, 437, 449, 452, 458, 459, 462, 498, 574, 575, 577, 611, 615, 616, 651, 652, 674−676, 692 Agmatite, 49 Ahvenisto (gabbro−anorthosite) complex, 544− 546 Ahvenisto pluton, 539, 545, 546, 548 Ahvenisto rapakivi area, 539, 548 Albite, 43, 45, 53, 116, 124, 171, 172, 473, 541, 628 Åland batholith, 15, 537, 541, 545, 546, 548, 549, 556 Alkali feldspar diabase, 539 Alkali feldspar syenite, 547 Alkaline complex(es), 143, 587, 625, 627, 628 Alkaline intrusions, 16, 22, 486, 566, 567, 569, 593 Alkaline rocks, 605, 608, 609, 628, 629, 631− 634, 688 Allanite, 454, 460, 541, 614 Allochthon, 75, 144, 240, 263, 271, 319, 585 Allochthonous, 11, 12, 25, 26, 36, 56, 58, 70, 75, 76, 180, 240, 246, 262, 263, 297, 318,319, 321, 323, 325, 326, 331, 350, 362, 374, 394, 489, 504, 566, 584 Alluvial braid, 311, 323 Alluvial deposits, 320, 588, 591 Alluvial fan, 165, 299, 303, 305, 321 Alluvial plain, 310, 312, 320, 573 Alluvial sediments, 303 Almandine, 60 Alpine style, 26, 694 Alteration, albite-sericite, 43 Alteration, calc-silicate, 258 Alteration, fenitic, 56, 617 Alteration, hydrothermal, 33, 38, 45, 47, 51, 58, 60, 63, 74, 75, 152, 588 Alteration, metasomatic, 257, 258 Alteration, sericitic, 33, 47 Alteration, talc-carbonate, 37, 253 Aluminous, 22, 68, 303, 419, 621 Amalgamation (amalgamated), 22, 70, 180, 181, 240, 346, 414, 437, 500, 507−509, 651, 653, 658, 659, 661, 689 Amazonia, 506, 507, 510, 513, 659−661 INDEX
•
715
Amazonian craton, 657−661 Amphiboles, 40, 57, 58, 64, 68, 128, 164, 172, 173, 205, 214, 223, 251, 252, 254, 259, 311, 330, 414, 417, 419, 424−426, 433, 454, 547, 548, 610, 611, 615, 625, 626, 628, 635, 687 Amphibolite, 7, 11, 45, 47−49, 54, 57, 64, 65, 69, 70, 175, 303, 314, 355, 359, 374, 585, 611, 625 Amphibolite facies, 28, 33, 38, 48, 51, 54, 56, 60, 68, 74, 80, 104, 144, 253, 282, 287, 349, 355, 361, 365, 375, 380, 394, 419, 519, 585 Anatexis, 45, 63, 79, 80, 84, 180, 458, 519, 692 Ancylite, 623 Andalusite, 35, 38, 58, 63, 301, 302, 381 Andean-type, 508, 511, 512, 520 Andesine, 536, 545 Andesite, 58, 149, 150, 152, 154, 155, 159, 175, 300, 348, 356, 360, 367, 371, 374, 376−378, 380, 381, 387−389, 392, 394, 395, 426, 449, 468, 656, 676 Andino-type, 182 Annelid, 580, 582 Anorogenic, 430, 446, 537, 552, 655, 693 Anorthosite, 71, 74, 106, 111, 114, 168, 170, 176, 417, 421, 430, 515, 520, 545, 546, 553, 659 Anthophyllite, 261, 694 Antigorite, 253, 257, 260−262 Apatite, 40, 54, 56, 61, 168, 212, 214, 215, 410, 417, 419, 425, 430−432, 451, 452, 454, 456, 460, 462, 541, 546, 549, 608, 610−618, 621, 623, 626−630, 689, 698 Apophyse, 38, 162, 259 APWP (apparent polar wander path), 650 Arc complex, accretionary, 5, 350, 359, 448, 449, 474 Arc complex, Central Finland, 5, 6, 350, 359, 450, 474 Arc complex, juveline, 175, 176 Arc complex, magmatic, 79 Arc complex, primitive, 5, 350, 359, 411, 412, 414, 434, 437, 447−449, 466, 467, 472, 474 498 Arc complex, southern Finland, 6, 350, 359, 412, 426, 434, 437, 447, 448, 450, 452, 459, 472−474 Arc complex, Svecofennian, 181, 256 Arc complex, western Finland, 5, 350, 411, 412, 437, 447−450, 472, 473 Archean ages, see Ages Archean bedrock, see Bedrock Arkose, 165, 177, 201, 329, 353, 390, 391, 567, 570 “Arkose conglomerate”, 306 Arkose quartzite, 165, 173 Arkosic conglomerate, 297, 299 Arkosic gneiss, 11, 69 Arkosic matrix, 301 Arkosic quartzite, 585 Arkosic rocks, 7, 12, 507 Arkosic sandstone, 295, 573, 588 Arsenopyrite, 543
716
•
INDEX
Asteroid, 16 Asthenosphere, 77, 633, 634, 638 Atmospheric changes, 13 Atmospheric oxygen, 672, 678 Augite, 111, 116, 123, 128, 205, 212, 215, 220, 223, 545, 547, 548, 575, 577 Augite, in cumulate, 106, 108−111, 116, 117, 123, 127, 421, 422, 425 Autochthonous, 11, 12, 75, 144, 158, 179, 240, 246, 262, 282, 314, 321, 325, 350, 489 BABEL (Working Group), 362, 388, 482, 484, 490, 491, 504, 518 BABEL lines (seismic reflection surveys), 78, 494, 495, 504, 511, 513, 518 Back-arc basin, 180, 246, 273, 325, 327, 392, 394, 484, 501, 503, 517, 520 Back-arc rifting, 501, 503 Baddeleyite, 5, 6, 204, 206, 223, 541, 546, 575, 586, 610, 623, 652 Baltic basin, 592 Baltic limestone, 582 Baltic-Bothnian megashear, 487, 497 Baltica, 557, 584, 585, 589, 657 Banded iron-formation, see also BIF, 20, 29, 31−33, 35, 71, 73, 158, 314, 315, 317, 349, 350, 380, 672 Barite, 610, 617 Barytocalcite, 623 Basalt, calc-alkaline, 356, 449 Basalt, (continental) flood basalt, 170, 176−178, 180, 209, 355, 557, 574 Basalt, EMORB, 159, 160, 183, 238, 240, 249, 250, 255, 262, 263, 267, 268, 270, 360, 361, 366, 380, 384, 394−396,497, 576 Basalt, high-Mg, 46, 82, 159, 224 Basalt, komatiitic, 44, 46, 47, 152, 171 Basalt, low-K, 355, 356, 358, 395, 449, 466 Basalt, low-Ti, 170, 171 Basalt, MORB, 51, 209, 212, 242, 246, 249, 262, 266, 267, 348, 350, 359, 361, 362, 364, 365, 371, 373, 374, 377−379, 381, 384, 387−389, 393, 497, 576 Basalt, NMORB, 159, 180, 249, 262, 267, 268, 270, 360, 362, 389, 394, 395 Basalt, oceanic island basalts, OIB, 160, 183, 240, 242, 243, 246, 247, 249−251, 255, 263, 266−268, 270, 273, 497, 575, 576 Basalt, olivine basalt, 212, 579 Basalt, siliceous high-magnesian basalts, SHMB, 118, 152 Basalt, tholeiitic, 33, 36, 40, 44, 46, 48, 49, 118, 132, 155, 159, 161, 204, 221, 301, 360, 422, 426 Basalt, volcanic arc basalts, VAB, 180, 240 Basalt, within-plate basalts, WPB, 160, 180, 209, 240, 395 Basalt, within-plate lavas, WPL, 348, 360−362, 365, 366, 374, 380, 384, 385, 394, 396 Basaltic dike, 250, 253, 258, 264, 577
Basin development, Karelia, 326 Basin inversion, 503, 505, 507, 508, 512, 513, 519 Basin, peritidal, 179 Bastnäsite, 541 Bedrock, 4, 6, 7, 13−16, 26, 28, 56, 142, 283, 321, 346, 348, 372, 415, 447, 449, 88−493, 499, 536, 539, 540, 542, 543, 546−549, 566, 569, 577, 586, 587, 608, 626, 627, 648, 685−688, 690, 693, 694, 699 Bedrock, Archean, 13, 26, 56, 321, 415, 449, 489 Belomorian mobile belt, 487, 489, 496, 500, 501 Belomorian terrain (terrane), 22, 24, 28, 64−66, 74, 78, 82, 175 Bergslagen area (district, field), 5, 388, 390−392, 396, 487, 489, 491, 494, 507, 509, 511, 512 Bergslagen microcontinent, 346, 351, 390, 392, 396, 494, 499, 500, 503, 507, 508, 512, 519 BIF, see also Banded iron-formation, 46, 47, 263, 291, 315, 318, 331, 677 Biotite, 14, 32, 33, 35, 37, 38, 40, 49, 52, 53, 56, 60, 64, 69, 71, 74, 110, 116, 164, 172, 173, 175, 205, 207, 212, 215, 223, 307, 417, 426, 449, 451, 452, 454, 456, 459, 460, 462, 470, 540, 541, 543, 545, 549, 616, 627−629 Black schist, 11, 12, 156, 158, 159, 172, 181, 203, 240, 258, 271, 298, 310, 311, 315, 317, 318, 360, 434, 436 Black shale, 179, 246, 291, 348, 349, 355, 356, 361, 362, 365−367, 371−373, 375, 380, 381, 389, 393−395, 675 Blueschist facies, 78 Boninite, 104, 157, 161, 209, 210, 212, 226, 264, 268 Boninite−gabbronorite, 208−215, 226 Boninite−norite dikes, 201, 202, 204, 205, 207, 208, 217, 226 Bothnian basin, 5, 350, 389, 392, 487, 498, 508, 514 Bothnia(n) microcontinent, 361, 494, 498−500, 503−505, 507, 508, 518 Bouguer anomaly, 417, 418, 432, 492, 494, 497−499 Bouma sequence, 317 Braidplain, 305, 310, 311, 320, 323 Bronzite, 106, 108, 109, 111, 116, 123, 127, 168, 205, 223, 421, 422, 424 Bronzite, in cumulate, 106, 108, 109, 111, 116, 123, 127, 168, 421, 422, 424 Bronzitite, 106, 116, 122 Burakovka intrusion, 201, 226, 652 Calc-alkaline, 24, 43, 52, 70, 71, 76, 143, 149, 152, 161, 162, 176, 182, 204, 208, 209, 244, 252, 263, 264, 271, 300, 356, 364, 377, 378, 380, 388, 426, 435, 615, 631, 654 Calcite, 56, 608, 610, 613, 614, 616−618, 620, 621, 623, 626−628, 635 Calc-silicate rocks, 64, 175, 256, 258, 259, 356, 360, 361
Caledonian, 4, 7, 23, 69, 486, 513, 569, 570, 582, 584, 585, 592, 609 Caledonian foreland, 591, 593 Caledonian orogen, 593 Caledonian orogeny, 16, 567, 582, 587, 588, 593 Caledonides, 14, 143, 174, 489, 566, 582−585, 592, 593 Cambrian, 4, 15, 16, 570, 579, 580, 582, 584, 585, 591, 592 Cambrian, basin, 592, 594 Cambrian sedimentary rocks, 570, 579, 580−582, 584, 591, 592, 672 Camptonite, 615, 616 Canrinite, 628, 629 Carbon cycle, 672, 673, 678 Carbon isotope excursion, 669, 672, 673, 675−678 Carbonate platform, 329, 331, 591, 592 Carbonate rock, 11−13, 154, 158, 165, 175, 179, 181, 246, 247, 257, 259, 271, 282, 310, 318, 357, 362, 490, 584, 592, 672 Carbonatite, 11, 13, 53, 54, 57, 68, 81, 143, 608, 610−614, 617, 621, 623, 625−628, 631, 634−636, 689, 698 Carbonates, magmatic, 610−612, 614, 616, 623 Carbonates, sedimentary, 310, 349, 350, 356, 358, 360, 361, 366, 371, 375, 376, 380, 381, 385, 388, 390, 392−394, 396, 501, 672−678 Cenozoic, 566, 586−588, 593 Cenozoic uplift, 593 Central Finland granitoid complex, 12, 346, 348, 350, 351, 365, 369, 374, 382, 389, 394, 411, 412, 417, 423, 428, 430, 435, 437, 450−455, 468, 469, 472, 474 Central Karelian complex, 487, 489 Central Lapland granitoid complex, 11, 63, 142−144, 155, 282, 285, 287, 458, 459, 462, 487, 489, 497, 513, 514 Central Lapland greenstone belt, 139, 142−144, 146, 154−156, 158, 164, 167, 171, 177, 178, 183 Central Puolanka Group, 59, 81, 282, 287, 298, 299, 302, 305, 306, 309, 320, 326, 330 Central Svecofennia, 346, 350−354, 361, 362, 364−367, 372, 379, 380, 382, 392−395 Chalcopyrite, 124, 127, 129, 253, 261, 415, 422, 424, 610 Charnockite, 22, 175, 446, 501, 503, 652 Chilled margin, 56, 106, 114, 118, 120−122, 172, 205, 207, 214, 215, 223, 224, 238, 249, 259, 422, 432 Chromian diopside, 260, 606, 620, 627, 638, 639 Chromitite, 106−108, 111, 116, 118, 122−124, 130, 154, 161, 167−171, 246, 249, 254, 257, 270, 271 Chronogram/Histogram, 6, 27, 413, 448, 541, 567 Chronostratigraphic, 566, 567, 673 CIA (Chemical index of alteration), 301, 302, 379 Classification, basins, Karelian domain, 282, 287 INDEX
•
717
Classification diagram, rapakivi granites, 551 Classification, Finnish granites, 692, 693 Classification, lithostratigraphic, 310 Classification, mafic−ultramafic plutonic rocks, 410, 411 Classification, magmatectonic, Proterozoic gran itoids, 446 Classification, Paleoproterozoic dike swarms, 204 Classification, tectofacies, 295, 300 Classification, Sokli carbonatites and alkaline rocks, 628, 631 Classification, volcanic rocks, 354 Clastic, 81, 156, 164, 165, 181−183, 299, 314, 317, 361, 579−581, 584, 591, 592 Clastic dike, 591 Climate, arid, 181, 291, 292, 329 Climate, humid, 301 Climate, semiarid, 291, 292 Climate, tropical, 301, 329 Clinohumite, 623 Clinopyroxene, 37, 46, 47, 116, 151, 152, 157, 172, 205, 207, 211, 212, 214, 215, 217, 223, 251, 254, 258, 261, 370, 414, 417, 419, 424, 428, 434, 451, 452, 456, 617, 618, 621, 628, 639, 640 Clinopyroxenite, 114, 170, 204, 214, 223, 243, 254, 255, 269 Collapse, 14, 80, 349, 473, 474, 484, 500, 503, 504, 507−510, 513−515, 518−520, 556 Collision, 14, 15, 24−26, 53, 61, 62, 71, 75, 77−79, 82, 84, 175, 179−181, 331, 349, 350, 361, 365, 392−394, 446−448, 484, 497, 500, 501, 503, 504, 507, 508, 511−513, 516, 517, 519, 520, 536, 584, 592, 639, 654, 658, 661 Collision, continent−continent, 484, 504, 507, 508, 511−513, 517, 584 Collision, oblique, 504, 507, 512, 519 Collision, Svecofennian, 25, 53, 77, 78 Colloform texture, 262 Columbite, 5, 541, 543 Complex deformation, 33, 36, 59, 62 Conductivity, 351, 493, 499, 520 Conglomerate, 7, 11, 12, 29, 33, 35, 68, 73, 105, 140, 146−148, 152, 153, 164−167, 177, 181, 182, 201, 203, 285, 291−293, 295, 297, 299−301, 303, 305−307, 309−311, 313−315, 317, 318, 320, 321, 323, 324, 329, 349, 350, 355, 356, 363, 364, 366−368, 370, 371, 376, 379, 385, 395, 546, 567, 570, 573, 582−584, 674, 692 Conglomerate, basal, 285, 292, 293, 300, 301, 309, 317, 323, 349, 582, 674 Conglomerate, polymictic, 33, 35, 68, 105, 165, 177, 291, 300, 307, 309, 313, 314, 324, 329 Conglomerate, volcaniclastic, 146, 363 Contamination, crustal, 76, 149, 152, 158, 171, 174, 176, 226, 415, 434, 437, 576, 577, 579 Contamination, sialic, 118, 159 Contamination, wall-rock, 48, 106, 113
718
•
INDEX
Continental breakup, 179, 241, 243, 252, 255, 270, 273, 325, 501, 517 Continental crust, 6, 14, 15, 29, 37, 45−47, 49, 78, 82, 170, 180, 198, 228, 246, 271, 273, 315, 331, 346, 446, 457, 517, 536, 537, 539, 540, 546, 552−554 Continental margin, 70, 71, 74, 182, 183, 241, 268, 271, 273, 325, 327, 331, 378, 379, 388−390, 414, 499, 505, 512, 517, 585, 587, 591, 592, 660 Continental reconstruction, 201, 557, 648, 650, 656, 657, 660 Cordierite, 12, 47, 54, 57, 58, 61, 63, 175, 385, 391, 414, 456, 458, 693 Corycium enigmaticum, 696 Cover rocks, 13, 36, 575, 584 Cover rocks, Paleoproterozoic, 13, 489, 501 Cover rocks, Proterozoic, 228 Cover rocks, Sariola, 292 Cover sequence, 26, 74, 263, 473, 497, 566, 567, 569, 579, 582, 587, 592 Craton, 15, 25, 28, 73, 80, 81, 131, 167, 175−178, 198, 203, 223, 228, 240−242, 256, 263, 268, 271, 273, 327, 331, 346, 349, 356, 388, 389, 393, 394, 412, 414, 416, 430, 434, 436, 437, 448−450, 455, 467, 468, 472, 484, 485, 487− 489, 494, 496−498, 500, 501, 503−505, 508, 509, 513, 515, 517, 518, 577, 587, 589, 591, 592, 618, 633−639, 648, 650, 652, 653, 656−661, 672, 673, 692 Craton, Archean, 80, 81, 131, 167, 175, 178, 198, 203, 223, 228, 242, 268, 346, 349, 356, 388, 389, 393, 394, 412, 414, 437, 448−450, 455, 467, 468, 472, 484, 489, 494, 498, 500, 501, 577, 633, 636, 652, 653, 672, 673, 692 Craton, East European, 484, 485, 489, 515, 589, 591, 592 Craton, Karelian, 25, 28, 240, 242, 256, 273, 349, 487−389, 494, 496, 497, 500, 501, 503−505, 508, 509, 513, 517, 518, 577, 633, 635, 638, 639, 656 Craton, Kola, 487, 489, 494, 496, 501, 517 Craton, Norrbotten, 487, 489, 494, 497, 500, 504 Cratonic sedimentary cover, 575 Cretaceous, 586, 593 Cross-bedding, 33, 148, 154, 368, 391, 571−573, 580 Cross-stratification, 299, 305, 320 Crust, oceanic, 7, 164, 179−181, 183, 241, 243, 244, 246, 251−253, 271, 331, 346, 350, 392, 433, 489, 494, 505, 585, 637, 660 Crustal growth, 28, 79, 346, 488, 493, 516, 520 Crustal (intracrustal) melting, 79, 80, 83, 514, 556 Crustal signature, 149, 155, 171, 180 Crustal thickening, 58, 518, 593 Crustal thinning, 80, 270, 510, 588 Crystal fractionation (see also Fractional crystallization), 47, 118, 149, 251, 267, 422, 424, 434,
632 Crystalline basement, 566, 574, 582−584, 589, 591 Crystalline bedrock, see Bedrock Cubanite, 261, 422 Cummingtonite, 60, 261 Cumulate complex, 43, 46, 47 Cumulate, chromite, 108, 123, 223, 419, 424 Cumulate, dunite-troctolite-olivine gabbro, 586 Cumulate, gabbro, 168 Cumulate, gabbroic, 123, 172, 244 Cumulate, komatiitic, 36, 44, 46 Cumulate, olivine, 106, 111, 114, 116, 123, 128, 172, 251, 422, 424, 432 Cumulate, olivine-clinopyroxene, 172, 223, 419 Cumulate, olivine-orthopyroxene, 414 Cumulate, olivine-pyroxene, 434 Cumulate, plagioclase-pyroxene, 431 Cumulate, pyroxene, 36, 44, 106, 113, 172, 223, 414, 223, 414, 419, 422, 424, 431, 434 Cumulate, tholeiitic, 44 Cumulate, two-pyroxene, 419, 422 Cumulate, ultramafic, 170, 172, 241, 245, 415, 417, 419, 421, 432 Cumulus, 45, 106, 109, 111, 112, 116, 118, 127, 128, 168, 172, 173, 205, 214, 215, 217, 220, 253, 254, 262, 410, 417, 419, 421, 422, 424, 425, 428, 430, 432 Cumulus texture, 205, 214, 217, 253, 262, 422, 430 Current bedding, 582 Dacite, 149, 152, 154, 163, 164, 348, 352, 356, 367, 371, 372, 374, 378, 380, 381, 386, 388, 394, 395, 698 Dacite, high-K, 181, 381 Dacite, low-K, 163 Decompression melting, 75 Deformation style, brittle−ductile, 56 Deformation, Cambrian, 585 Deformation, Caledonian, 584 Deformation, Proterozoic, 28, 57, 62, 65 Deformation, Svecofennian, 36, 62, 63, 70, 181 Deformation, synsedimentary, 363 Delamination, lithospheric, 83, 84, 507 Depletion, 39, 40, 49, 52, 132, 133, 157, 158, 367, 371, 374, 377, 378, 381, 384, 385, 387, 673 Depositional age, 52, 59, 282, 287, 299, 310, 319, 352, 510, 574, 583, 588, 589, 675, 676 Depositional basement, 29, 60, 73, 158, 291, 307, 319, 393, 566 Depositional basin, 154, 156, 179, 594, 677 Depositional environment, 81, 313, 391, 572, 581, 583, 587 Depositional evolution, 142 Depositional model, 325, 331, 573 Depositional sequence, 289 Depositional younging, 29 Deposit (see also Mineralization and Ore)
Deposits, alluvial, 165, 303, 305, 320, 588 Deposits, Au/gold, 22, 25, 172, 698 Deposits, brained river, 165, 303, 305 Deposits, Cambrian, 591 Deposits, cover, 240, 349 Deposits, Cr, 120, 698 Deposits, Cr-oxide, 106 Deposits, chromitite, 122 Deposits, Cu-Co-Zn-Ni±Au (sulfide), 261, 262 Deposits, Cu-Ni-PGE sulfide, 172, 174 Deposits, Cu sulfide, 273 Deposits, Cu-Zn-Pb, 381, 698 Deposits, debris flow, 317, 321, 370 Deposits, deep-water, 368 Deposits, epiclastic, 47, 48, 144, 146, 154, 303, 331 Deposits, Fe-Ti-V oxide, 104, 122 Deposits, flash flood, 305 Deposits, (foreland) basin, 181−183, 324, 593 Deposits, fluvial, 292, 581 Deposits, glacigenic, 285, 291 Deposits, hydrothermal, 36 Deposits, iron, 689, 698 Deposits, kaolin, 586 Deposits, (magmatic/massive) sulfide, 126, 130, 131, 261, 355, 356, 358, 366, 380, 390, 394, 422, 425, 434, 437 Deposits, mass flow, 81, 305, 319 Deposits, molasse, 182, 353, 364, 508, 510, 513 Deposits, mid-fan, 367 Deposits, nickel, 698 Deposits, Ni-Cu-Fe sulfide, 424 Deposits, Ni-Cu-PGE sulfide, 104, 410 Deposits, Ni-Cu±PGE (sulfide), 432, 435 Deposits, Ni-Cu (sulfide), 408, 412−415, 417, 421, 424, 434, 436, 437, 698 Deposits, Ni-Cu-Zn, 315 Deposits, Ni sulfide, 435 Deposits, oxide, 120, 428 Deposits, Paleozoic, 579 Deposits, PGE, 110, 120, 122, 129, 130, 132 Deposits, platinum, 700 Deposits, podiform, 122 Deposits, pyrite, 36 Deposits, pyroclastic (flow), 29, 31, 33, 47, 62, 374, 386 Deposits, pyrrhotite, 126, 127 Deposits, Quaternary glacial, 586 Deposits, resedimented, 81 Deposits, shallow-water, 364, 367, 388 Deposits, siliciclastic, 81 Deposits, soapstone, 37, 699 Deposits, Ti, 436 Deposits, Zn-Cu, 449 Deposits, V, 116, 698 Deposits, VHMS (volcanic-hosted massive sulfide), 389 Deposits, volcanic, 146 Deposits, volcaniclastic, 35, 36, 59 INDEX
•
719
Deposits, volcanogenic sedimentary, 380, 393 Denudation, 16, 177, 593, 594 Detachment (fault), 24, 61, 71, 80, 269−271, 273 Diabase, 11, 13, 15, 156, 157, 162, 165, 166, 171, 203, 311, 428, 495, 539, 546−548, 553, 554, 556, 566, 569, 575−580, 650, 653, 655, 692, 693 Diabase dike, 11, 13, 15, 157, 162, 166, 203, 539, 547, 548, 553, 554, 556, 566, 569, 575, 578, 579, 650, 653, 655, 692 Diabase sill, 15, 566, 577 Diallage-augite, 106 Diamond, 608, 617, 621−623, 627, 628, 638−640 Diamondiferous, 608, 619, 634, 636−638, 640 Diapir, asthenospheric, 249, 254, 255, 270 Diapir, mantle, 243, 271–273, 556 Diatexite (diatexitic granite), 458 Diatomaceous earth, 587 Differentiation index, modified (MDI), 214, 221 Dike, basaltic, 250, 253, 258, 264, 577 Dike, basaltic, calc-alkaline, 264 Dike, boninite−norite, 204, 205, 207, 209, 217 Dike, boninitic, 180, 241, 263, 267, 271, 324 Dike, calc-alkaline, 152, 162, 263, 264 Dike, diabase (see Diabase dike) Dike, felsic, 152, 164, 166, 167, 176, 181, 324, 541 Dike, Fe-tholeiitic, 201, 203, 204, 207−209, 215, 217, 221, 226, 228 Dike, gabbronorite, 196, 204, 205, 209, 217 Dike, granitic, 37, 259 Dike-in-dike, 244, 246, 247, 259 Dike, komatiitic, 44, 45 Dike, mafic, 37, 54, 56−58, 74, 119, 143, 151, 152, 159, 178, 203, 204, 229, 246, 264, 541, 566, 567, 569, 570, 574, 579, 593, 650 Dike, porphyritic, 39, 60, 374, 393 Dike, porphyry, 40, 497, 537, 539, 546, 547, 549, 554, 655, 659 Dike, tholeiitic, 37, 201, 203−205, 207−210, 217, 224, 226, 264, 271, 324, 497, 577, 652 Dike swarm, 178, 198−201, 203−205, 207, 208, 211, 212, 218, 223, 228, 324, 455, 539, 547, 548, 655, 656 Dike swarm, classification, 204 Dike swarm, mafic, 62, 74, 178, 198, 201, 203, 207, 209, 223, 226, 570, 572, 589, 652, 653, 660, 661 Dike, ultramafic, 49, 112, 246, 264, 267 Diopside, 252, 258–261, 606, 614−618, 620, 622, 627, 638, 639 Diorite, 57, 348, 358, 423, 452, 654 Dividal Group, 582, 584, 585, 591 Docking, 503−505, 507, 513, 515, 518, 519 Dolerite Group, 655, 660 Dolomite, 12, 56, 154, 203, 285, 291, 298, 306, 309−311, 313, 315, 317, 318, 324, 330, 395, 585, 608, 610, 613, 614, 623, 670, 672, 673, 675−677, 672, 673, 675−677 Dome, 54, 69, 73, 142, 146, 150, 158, 169, 285,
720
•
INDEX
295, 309, 313, 319, 349, 389, 457, 553 Dome, basement, 69, 146, 150, 285, 309, 319 Dome, gneiss, 69, 142, 146, 158 Dome, thermal, 349 Dropstone, 295, 297 Dunite, 161, 170, 174, 241, 253, 257, 263, 264, 428, 586 Earthquake, 16 East European craton, see Craton Eclogite, 77, 640 Eclogite, bimineralic, 636, 637 Eclogite, diamondiferous, 636, 638, 640 Eclogite, mantle eclogite, 621, 639 Eclogite detritus, 617 Eclogite facies, 49, 175 Eclogite seams, 640 Eclogite xenoliths, 637, 639, 640 Enderbite, 57, 58, 83, 175, 494, 501, 651 Electrotelluric studies, 61 EMORB, see Basalt Eocambrian, 566 Epiclastic, 31, 47, 48, 144, 146, 154, 303, 331 Epicontinental, 179, 183, 309, 324, 349 Epidote-amphibolite facies, 297 Erosional level, 14, 16, 65, 228, 591 Eskolaite, 258, 260 Eu anomaly, 49, 250 Eurajoki stock, 539, 540, 543, 545, 549 EUROBRIDGE, 488, 490, 512 EUROPROBE, 78 Evijärvi field, 346, 350, 351, 361−364, 389, 394, 395 Evolution scheme, Paleoproterozoic, 202 Exhumation, 38, 53, 57, 74, 80, 82, 83, 176, 182, 500, 504, 515, 587−589, 593 Exsolution, 205, 217, 551, 614 Extension, 25, 79, 81, 154, 182, 198, 228, 257, 267, 290, 302, 307, 311, 326, 376, 384, 385, 388, 389, 446, 457, 474, 503−505, 507, 509− 511, 514, 515, 518−520, 555, 573, 587, 589, 592, 650 Fan, 291, 299, 303, 305, 317, 318, 321, 329, 331, 350, 367, 395, 589 Faulting, listric, 57, 62, 254, 553 Fayalite, 174, 540, 543, 545 Fe-tholeiite, 207−209, 228, 324 Feeder channel, 159, 198 Feeder dike, 117, 122, 133, 243, 244, 249, 251−253, 376, 574 Fenite, 608, 610, 611, 617, 625, 627−629, 636 Fennian orogen, 500, 507 Fennian orogeny, 500, 503, 505, 507−510, 518 Fennoscandia, 15, 26, 179, 201, 226, 446, 449, 488, 507, 511−513, 515, 519, 520, 536, 537, 540, 544, 556, 566, 587, 591, 648, 656, 658− 661, 673, 677, 688, 691, 695 Fennoscandian Shield, 5, 7, 14, 16, 22, 23, 25,
28, 53, 57, 61, 68, 76, 77, 81, 83, 104, 142, 158, 167, 171, 176−179, 201, 203, 204, 211, 218, 224, 226, 228, 282, 292, 313, 331, 471, 484, 485, 487−490, 492, 493, 496, 499, 500, 503, 507, 509, 511, 515−518, 520, 554, 557, 566, 567, 569, 570, 574, 584, 587, 588, 592−594, 637, 646, 648−661, 672−674, 677, 694 Ferropicrite, 153, 162, 177 Finnmarkian phase, 584, 585 FIRE (reflection seismic survey), 78 Fission track, 587, 588 Fissure filling, 584 Flood basalt, see Basalt Flood-plain, 573 Fluid−rock interaction, 75 Fluorite, 454, 460, 462, 472, 541, 617, 628 Fluvial, 165, 291, 292, 295, 299, 305−307, 310, 311, 320, 321, 326, 329, 330, 349, 364, 366− 368, 390, 391, 393−395, 564, 571, 572, 580, 581, 591 Fluvial plain, 580 Föglö diabase dike swarm, 548 Fold-and-thrust belt, 114, 508 Foredeep, 325, 327, 505, 592 Foreland, 16, 26, 28, 65, 68, 74, 81, 144, 319, 508, 588, 589, 592−594 Foreland basin, 181−183, 588, 591, 593 Fossil, 580, 582 Fossil trace, 580 Forssa Group, 376−378, 425 Fractional burial rate, 678 Fractional crystallization (see also Crystal frac tionation), 49, 106, 111, 122, 149, 164, 249, 251, 269, 415, 430, 433−436, 470, 576 Fractional crystallization, closed-system, 417, 430 Fractional crystallization, open-system, 436 Fractionation, 47, 49, 106, 118, 132, 133, 149, 154, 180, 214, 249, 251, 263, 267, 321, 414, 415, 417, 422, 424, 434, 462, 576 632, 678 Fractionation, chromatographic, 263 Fractionation, closed-system, 251, 254 Fractionation, magmatic, 106, 133, 214, 249, 251, 414, 415, 417, 424, 576 Fractionation, mass independent (MIF), 678 Fractionation, metamorphic, 154 Fractionation, open-system, 132, 180 Fractionation, REE, 49, 321 Fractionation, reverse, 414, 422 Francolite, 626 Funnel-shaped, 39, 116, 172, 417, 460, 625 Gabbro−anorthosite, 22, 123, 412, 415, 417, 418, 421, 539, 544−546, 554 Gabbro−wehrlite association, 171, 172, 177, 178, 203, 212 Gabbronorite (see also Leuco- and Microgabbronorite), 123, 196, 204, 205, 207, 209, 217,
421, 428, 430 Galena, 48, 120, 261, 422, 591 Garnetite, 254, 269 Garsjø and Bjørnevatn belts, 73 Geochemistry, basalt (lava), Jormua, 266−270 Geochemistry, basalt (lava), Outokumpu, 266− 270 Geochemistry, discrimination diagrams, 159, 160, 163, 209, 355, 447, 466, 469, 549 Geochemistry, diagrams, Al2O3 vs. TiO2, 156, 157 Geochemistry, diagrams, Al2O3/TiO2 vs. Ti/Zr, 212 Geochemistry, diagrams, AFM, 251, 252, 435 Geochemistry, diagrams, box-and-whisker, 265 Geochemistry, diagrams, CMA, 435 Geochemistry, diagrams, Cr vs. Zr, 269 Geochemistry, diagrams, εNd vs. age, 153, 162, 352, 354, 467, 555 Geochemistry, diagrams, FeO*/MgO vs. SiO2, 210 Geochemistry, diagrams, Jensen’s cation plot, 209 Geochemistry, diagrams, La/Yb vs. Th/Ta, 213 Geochemistry, diagrams, MgO vs. TiO2, 208 Geochemistry, diagrams, MgO vs. Cr, 208 Geochemistry, diagrams, multielement, 466, 576 Geochemistry, diagrams, 143Nd/144Nd vs. 147Sm/ 144 Nd, 161 Geochemistry, diagrams, Sm-Nd data, 215 Geochemistry, diagrams, Th vs. TiO2, 155 Geochemistry, diagrams, Ti vs. Zr, 268, 355, 359, 362, 364, 371, 374, 377, 378, 381, 384, 385 Geochemistry, diagrams, (Y+Nb) vs. Rb, 550 Geochemistry, diagrams, Zr vs. Nb, 270, 633 Geochemistry, diagrams, (Zr+Nb+Ce+Y) vs. (K 2O+Na2O)/CaO, 550 Geochemistry, dike swarms, 208−213 Geochemistry, felsic and intermediate plutonic rocks, 463−466 Geochemistry, HFSE (high-field strength elements), 132, 152, 157, 161, 170, 381 Geochemistry, history in Finland, 198, 695−697 Geochemistry, immobile (trace) elements, 355, 389 Geochemistry, incompatible (trace) elements, 51, 149, 157, 159−161, 170, 172, 251, 264, 267, 432, 433, 466, 470, 617, 631, 632 Geochemistry, Jormua ophiolite, 252 Geochemistry, layered intrusions (Tornio− Näränkävaara belt), 118−122 Geochemistry, low-Al magnesian basalt, 172 Geochemistry, low-Al tholeiite, 215, 224 Geochemistry, mafic dike swarms, 207−214 Geochemistry, Nuttio ophiolite, 252 Geochemistry, ophiolites, 264, 265 Geochemistry, Outokumpu gabbros, 252 Geochemistry, PGE, 126−131 Geochemistry, rapakivi granites, 550, 551 INDEX
•
721
Geochemistry, REE, 39, 40, 46, 49, 51, 52, 116, 118, 120−122, 155, 157, 159−163, 172, 174, 209, 214, 224, 249−252, 254, 255, 263−266, 282, 318, 321, 355, 360, 362, 364, 367, 372, 374, 377, 380, 385, 386, 396, 549, 551, 577, 608, 616, 631, 632, 695 Geochemistry, REE, dike swarms, 214, 224 Geochemistry, REE, felsic rocks, Kittilä area, 164 Geochemistry, REE, mantle dikes, 255 Geochemistry, REE, layered intrusions (Tornio−Näränkävaara belt), 121, 122 Geochemistry, REE, mafic metavolcanic rocks, Kittilä Group, 160 Geochemistry, REE, ophiolites (Jormua, Nuttio, Outokumpu), 162, 250, 252, 255, 266 Geochemistry, REE, ophiolites/serpentinites, 266 Geochemistry, REE pattern, chondrite-normalized, 118, 121, 122, 130, 131, 155, 157, 159, 160, 172, 209, 214, 249, 251, 264, 266, 360, 362, 364, 367, 372, 377, 378, 380, 386, 549, 631, 632 Geochemistry, REE pattern, primitive mantlenormalized, 149, 152, 159, 162, 164, 254, 255, 265, 266 Geochemistry, REE, rapakivi granites, 551 Geochemistry, REE, volcanic rocks, Evijärvi field, 364 Geochemistry, REE, volcanic rocks, Häme belt, 377, 378 Geochemistry, REE, volcanic rocks, Pirkanmaa belt, 372, 373 Geochemistry, REE, volcanic rocks, Savo belt, 360 Geochemistry, REE, volcanic rocks, Tampere belt, 372, 373 Geochemistry, REE, volcanic rocks, Uusimaa belt, 386, 387 Geochemistry, REE, ultramafic metavolcanic rocks, Savukoski Group, 157 Geochemistry, Ru anomaly, 129, 130, 133 Geochemistry, trace elements, 149, 152, 155, 157, 159−161, 163−165, 167, 170, 575 Geochemistry, trace elements, (arc, MORB, and WPL lavas), 359, 364, 371, 377, 384 Geochemistry, trace elements, mafic volcanic rocks, Evijärvi field, 364 Geochemistry, trace elements, volcanic rocks, Häme belt, 377, 378 Geochemistry, trace elements, volcanic rocks, Pirkanmaa belt, 371, 373 Geochemistry, trace elements, volcanic rocks, Savo belt, 359 Geochemistry, trace elements, volcanic rocks, Tampere belt, 371, 373 Geochemistry, trace elements, volcanic rocks, Uusimaa belt, 384, 387 Geochemistry, variation diagrams, 463−465, 551
722
•
INDEX
Geological evolution, 648, 659, 661 Geological evolution, Lapland granulite belt, 175 Geological evolution, Neoproterozoic, 589 Geological evolution, Northern Finland, 63, 587 Geological evolution, Paleoproterozoic, 201, 202 Gersdorffite, 253, 415, 422 Geotherm, 53, 57, 79, 81, 639 Geothermal gradient, 49, 79 GGT/SVEKA (transect), 488 GIS database, 198, 200, 201 Glacial erosion, Quaternary, 580, 586 Glaciation, Paleoproterozoic, 677, 678 Glaciation, Pleistocene, 16 Glaciogenic rock, 295, 297, 301, 349 Glimmerite, 54, 610, 612−615, 631 Gondwana(land), 179, 661 Gothian belt, 658−660 Gothian evolution, 488, 499, 500, 515, 519 Gothian orogen (orogeny), 496, 500, 567 Graben, 4, 291, 315, 327, 567, 573, 574, 588, 615 Graben, half-graben, 291 Graded bedding, 154, 305, 320, 368 Granite, A-type, 469, 472, 514, 537, 549, 551, 552 Granite, I-type, 463, 469, 514, 550 Granite, coarse-porphyritic, 451, 452, 454, 462 Granite, lateorogenic, northern Finland, 458, 462, 469−471, 473 Granite, lateorogenic, southern Finland, 14, 448, 456−460, 462, 469−472 Granite, Lina-type, 76, 459, 514 Granite, M-type, 550 Granite, Nattanen-type, 68, 76, 143, 461, 462, 471 Granite, postorogenic, northern Finland, 143, 183, 458, 461, 462, 470, 473, 474 Granite, postorogenic, southern Finland, 15, 428, 460, 472, 474 Granite, Rovaniemi, 458, 462, 471 Granite, S-type, 76, 463, 468, 469, 471, 474, 489, 491, 513, 550 Granite, Vainospää, 72, 462, 471 Granite-migmatite (zone), Svecofennian, 349, 375, 380, 391, 393 Granitoid, Archean, 11, 13, 14, 36, 54, 143, 153, 292, 293, 299−302, 489 Granitoid, A-type, 489, 514 Granitoid, calc-alkaline, 24, 70, 76, 468, 489, 491, 509, 511 Granitoid, I-type, 463, 466, 468−471, 491, 511, 514 Granitoid, petrogenesis, 48, 466−472 Granitoid, preorogenic, 446, 448, 449, 466, 471, 474 Granitoid, postkinematic, 423, 472, 474 Granitoid, Svecofennian, 29, 53, 285, 574 Granitoid, synorogenic, 182, 472, 474 Granitoid, synkinematic, 351, 424, 468, 469, 472, 474 Granitoid, syntectonic, 29, 491, 511
Granitoid rocks, 13, 26, 351, 446−448, 471, 473, 474, 494, 509 Granodiorite, 2, 7, 12, 13, 29, 30, 37, 39, 40, 49, 52, 70, 82, 143, 152, 164, 166, 271, 348, 377, 381, 384, 385, 390, 391, 428, 582, 584, 628 Granodiorite, coarse-porphyritic, 452, 454 Granulite (rock), 14, 37, 53, 56−58, 61, 65, 71, 74, 75, 77, 81−83, 175, 176, 182, 228, 230, 380, 469, 494, 517, 651, 695, 696 Granulite complex, Lapland, (see also Lapland granulite belt), 14 Granulite complex, Siurua, 60, 61 Granulite complex, Varpaisjärvi, 50, 53, 54, 56−58, 228, 230, 651 Granulite facies, 7, 12, 29, 37, 38, 43, 53, 56−58, 61, 71, 74, 75, 80, 83, 144, 175, 176, 183, 349, 355, 361, 380, 394, 419, 496, 501, 512, 650−652 Granulite terrane (terrain), Tersk, 175, 487, 494, 496, 500, 501 Granulite terrane, Umba, 175, 182, 487, 494, 496, 500, 501, 503, 504 Graywacke, 12, 30−32, 41, 165, 175, 246, 271, 314, 315, 317, 319, 344, 349, 350, 353−356, 361, 362, 365−367, 369, 370, 375, 376, 378− 381, 389−391, 395, 396, 417, 421, 449, 498, 504, 508 Gravimetric anomaly, 105 Gravimetric survey, 546, 569 Gravitational instability (disequilibrium), 47, 78 Gravity data, (see also Map, gravity anomaly), 38, 83, 117, 417 Gravity differentiated sill, 171 Gravity investigation, 73 Gravity flow, 329 Gravity low, 417 Gravity studies, 61 Gravity survey, 71, 417 Greenschist facies, 53, 56, 57, 61, 80, 144, 146, 285, 287, 301 Greisen, 541, 543, 544 Grenvillian orogeny, 488, 499, 515, 389, 661 Grunerite, 31, 35, 46, 73 Haaparanta (Haparanda) suite (plutonic rocks), 143, 166, 167, 181, 183, 290, 455 Hailuoto Formation, 573, 580, 591 Halti-Ridnitshohkka igneous complex, 585 Häme belt, 12, 346, 348, 350, 351, 375, 377, 378, 391, 392, 395, 396, 411, 417, 426, 487, 490, 499, 507, 509, 511 Häme diabase dike swarm, 539, 548 Hämeenkyrö pluton (batholith), 453 Hanko granite, 470 Harzburgite, 25, 74, 77, 116, 241, 246, 253, 257, 263, 264, 621, 636, 637, 639 Hassela shear zone, 487, 498 Hattu schist belt, 28, 29, 33, 36−38, 48, 52, 82 Hattu supracrustal belt, 31, 32, 35, 36, 38
Häme Group, 376−378, 426 Heterolith, 305, 306, 311, 330 Hetta (granite) complex, 69, 449, 470, 472 HFSE (high field strength elements), 132, 152, 157, 161, 170, 381 High-grade terrane, Archean, 203, 228 High-grade, terrain, 29, 78, 83 Highstand, 291, 305, 306, 311, 312, 330 Hinterland, 507, 509 Hirsilä belt, 374, 375 Hirvaskoski shear zone, 59, 62, 282, 287, 326, 611, 614, 615 Histogram/Chronogram, 6, 27, 413, 448, 541, 567 Honkajärvi Group, 297, 299 Hornblende, 39, 46, 48, 49, 51, 56, 57, 69, 71, 75, 172, 214, 215, 220, 254, 370, 415, 417, 426, 449, 451, 452, 456, 460, 540, 543, 545, 546, 549, 615, 617, 625, 628, 629 Hotspot, 226 Höytiäinen basin, 282, 307, 313, 315, 319 Höytiäinen belt, 12 Hudsonian orogeny, 658, 661 Hyaloclastic, 45, 249 Hyaloclastite, 148, 159, 240, 246 Hybridization, 549, 552 Hybrid rock, 46, 116, 180, 419, 428, 452, 539, 548 Hybrid zone, 417, 426, 428 Hydrothermal alteration, see Alteration Hydrothermal venting, 271 Hyrynsalmi Group, 297, 311 Hyvinkää layered intrusion, 413, 426, 428, 429, 434 Hyypiä Group, 288 Iapetus Ocean, 584, 587, 592 IAT, see Tholeiite Ignimbrite, 368, 370, 385, 549 Iisalmi block, 201, 225, 285, 303, 327, 330 Iisalmi complex, 11, 240, 273, 282, 285, 301, 302, 326, 329, 455, 587, 489, 651, 653 Iisalmi terrain, 51, 53, 54, 56−59, 61, 68, 74−78, 81, 82 Iivaara alkaline complex, 16, 593, 608, 628, 629, 631 Ijolite, 608, 628−632, 636 Ilmenite, 114, 214, 251, 415, 454, 462, 541, 545, 549, 606, 610, 614, 616−618, 620−622, 696 Ilomantsi area, 13, 228 Ilomantsi belt, 11 Ilomantsi terrain, 28−32, 35, 36, 38, 40, 42, 45, 52−54, 56, 58−60, 75, 78, 80−83 Imandra intrusion, 653 Imandra−Varzuga(−Pechenga) belt, 487, 488, 494, 673, 674, 501 Immiscibility, 424 Immiscible sulfide liquid, 421, 432, 436 Impact crater, 16, 574, 584, 697 Impact structure, 583, 584, 587−589, 591, 592 INDEX
•
723
Inari arc, 494, 501 Inari area, 7, 142, 175, 449, 487, 494, 496, 501, 505 Inari terrain, 70−72 Inari terrane, 176, 654 Industrial minerals, 687, 689, 698 Intercumulus, 109, 123, 127, 173, 214, 220, 251, 414, 417, 419, 422, 424, 425, 428, 430, 433 Intracratonic, 28, 131, 176, 179, 183, 315, 327, 331, 349, 473, 497, 514, 588, 589, 591 Intracratonic basin, 176, 179, 331, 588 Intracratonic rift, 183, 327, 497 Intracratonic rift basin, 315, 473, 588, 591 Intracrustal reworking process, 26 Intracrustal melting, 79, 83 Intrusion, emplacement mechanism, 457, 460 Intrusion, funnel-shaped, 39, 116, 172, 417, 460, 625 Intrusion, gabbro, 251, 257, 271, 428 Intrusion, granitoid, 29, 35, 60 Intrusion, layered, 7, 11, 25, 59, 62, 104, 105, 107, 108, 118, 120, 122, 125, 126, 131, 142, 143, 146, 154, 157, 167, 169−172, 176−178, 183, 201, 203−205, 210, 211, 214, 215, 219, 226, 285, 291, 292, 294, 300, 324, 421, 423−430, 433, 434, 436 Intrusion, mafic−ultramafic, 7, 105, 142, 143, 155, 168, 214, 410− 412, 415, 417, 422, 423, 432, 434, 435, 437 Intrusion, polyphase, 415 Intrusion, postorogenic, 450 Intrusion sequence, 455, 462 Intrusion, synvolcanic, 411, 412, 426, 434 Island arc, 6, 14, 160, 161, 180, 182, 209, 224, 241, 256, 263, 264, 267, 271, 325, 331, 346, 350, 356, 361, 364, 377, 379, 384, 388, 389, 434, 446, 448, 466, 472, 484, 489, 490, 494, 497−500, 503, 504, 507, 508, 512, 517, 576, 655 Island arc, evolved, 14, 490 Island arc, intra-oceanic, 180, 264, 271, 325 Island arc, oceanic, 180, 264, 271, 325, 388, 497, 499 Island arc, primitive, 14, 489 Island arc complex, 14, 256 Island arc system, 6, 331 Island arc tholeiites, 160, 180, 267, 361, 384, 576 Isostatic equilibration, 15 Isostatic uplift, 16, 447 Isotope composition, Pb-Pb galena, 48 Isotope composition, Pb-Pb whole rock, 48 Isotope geology (history in Finland), 4, 687, 695−697, 699 Isotopes, C, 177, 179, 282, 285, 310, 311, 326, 329−331, 350, 501, 517, 635, 672−676, 695 Isotopes, δ13C value, 154, 158, 310, 318, 350, 358, 501, 635, 636, 672−678 Isotopes, δ34S value, 158, 159, 174 Isotopes, εHf value, 120, 471
724
•
INDEX
Isotopes, εNd value, 53, 68, 75, 76, 153, 160−164, 167, 171, 174, 180, 207, 210, 226, 320, 351, 384, 389−391, 411, 437, 467, 468, 471−474, 553, 575−579, 635 Isotopes, εSr value, 635 Isotopes, γOs value, 169, 171, 174 Isotopes, Lu-Hf, 76 Isotopes, Nd, 149, 150, 165−167, 351, 354, 369, 389, 471, 489, 497, 577, 635 Isotopes, Os, 161, 169, 171, 174, 257 Isotopes, Pb (Pb-Pb), 37, 38, 44, 48, 52, 56, 58, 69, 76, 120, 355, 366, 394, 435, 498, 576, 695, 696 Isotopes, Rb-Sr, 43, 660, 696 Isotopes, Re-Os, 254, 270, 636, 638, 676, 677 Isotopes, S, 158, 159, 174, 696 Isotopes, Sm-Nd, 46, 57, 151, 152, 176, 209, 226, 243, 315 Isotopes, Sr, 49, 51, 77, 634−636 Isotopes, Sr-Nd, 634, 635 Isotopes, stable, 158, 633, 696 Isotopes, U-Pb, 22, 58, 179, 181, 575 Isotopic age, 22, 26, 43, 59, 61, 104, 204, 228, 244, 271, 362, 410, 548, 653, 654, 676−678, 692 Isotopic age, chronogram, U-P zircon and bad deleyite ages, 6 Isotopic age (dating), K-Ar, 53, 56, 57, 75, 573, 586, 615, 660, 696 Isotopic age, histogram, Paleoproterozoic granitoids, 448 Isotopic age, histogram, Archean ages in Finland, 27 Isotopic age, Pb-Pb, 38, 44, 64, 301, 318, 353, 355, 435, 676, 677, 696 Isotopic age, Pb-Pb, titanite, 301 Isotopic age, Pb-Pb, zircon, 76, 353 Isotopic age, Rb-Sr, 43, 44, 48, 68, 69, 619, 627 Isotopic age, Sm-Nd, 151, 154, 157, 158, 160, 162, 163, 174, 176, 223 Isotopic age, TDM model age, 39, 61, 151, 152, 154, 162, 320, 471, 494, 496 Isotopic age, U-Pb, 153, 176, 181, 182, 201, 203, 204, 207, 213, 226, 514, 539, 545, 565, 574, 575, 619, 638, 648, 650, 696 Isotopic age, U-Pb, baddeleyite, 6, 204, 206, 223, 541, 574, 575, 586, 652 Isotopic age, U-Pb, dike swarm, 207, 218, 219, 221, 223−225, 227, 229 Isotopic age, U-Pb, mafic−ultramafic intrusion, 412, 413 Isotopic age, U-Pb, monazite, 38, 57, 175, 458, 503, 617, 651 Isotopic age, U-Pb, ophiolites, 243 Isotopic age, U-Pb, perovskite, 575, 619 Isotopic age, U-Pb, titanite, 22, 38, 57, 65, 68, 69, 71, 75, 175, 183, 380, 458, 503, 577, 616 Isotopic age, U-Pb, zircon, 6, 27, 29, 37−40, 43, 44, 46, 48, 49, 53, 57, 58, 61−65, 68, 69, 71,
73, 81, 120, 142, 149, 150, 152, 154, 157, 163, 167, 171, 172, 175, 176, 181, 203, 206, 243, 257, 300, 307, 310, 318, 351, 352, 365, 367, 374, 375, 380, 381, 386, 390, 392, 412, 413, 416, 426, 437, 449, 458, 459, 462, 484, 498, 541, 574, 575, 611, 615, 616, 651, 652, 674−676, 692 Isotopic studies, 25, 27−29, 57, 59, 65, 77, 142, 176, 209, 313, 351, 352, 470, 693, 694 Isotopic studies, conglomerate clasts, 165 Isotopic studies, layered intrusion, 118, 120 Itämäki belt, 317 Jaala−Iitti complex, 542, 545, 548, 552 Jasper, 165 Jatuli, 203, 228, 240, 287, 291, 298, 300, 303, 305−307, 309−311, 313−315, 317, 320, 321, 324−327, 329−331 Jatuli, Lower, 287, 305, 327 Jatuli, Marine, 228, 310, 311, 313, 315, 317, 326, 327 Jatuli, pre-Jatuli, 297, 327, 330 Jatuli, sub-Jatuli, 291, 307, 309 Jatuli, Upper, 287, 300, 309, 327 Jatuli quartzite, 309, 314, 317, 320, 324 Jatuli tectofacies, 303, 307, 309, 310, 313, 320, 321, 329 Jatulian, 13, 144, 145, 154, 155, 171, 177, 179, 201, 203, 291, 305, 307, 309, 314, 315, 331, 349, 350, 358, 653, 673, 691, 692 Jatulian Group, 201, 203, 673 Jatulian quartzite, 13, 171, 305, 314, 315, 692 Jatulian, ’marine’, 203 Jatulian, Post- Jatulian, 13, 144, 146, 154, 156, 171, 177, 179, 201, 203, 306, 315, 317, 332, 350, 358, 653, 673, 691, 692 Jatulian, Pre-Jatulian, 150 Jatulian, Sub-Jatulian, 309, 310 Jensen cation plot, 149, 208, 215 Jerta Nappe, 582, 583, 585 Jormua ophiolite, 6, 11, 13, 25, 238, 240, 242−247, 249−255, 263−265, 269, 270, 272, 285, 314, 319, 325, 497, 517, 694 Jotnian, 566, 567, 572, 655, 692 Jotnian period, 15 Jotunite (ferrodiorite), 539 Juuanvaarat quartzite belt, 288, 313 Juurikkaniemi Group, 44 Kainuan, 302, 303, 305, 306, 323, 324, 349 Kainuu (schist) belt, 11, 54, 59, 215, 240, 282, 287, 288, 293, 297, 299, 302, 303, 305, 309, 311, 314, 317, 320, 323, 325, 326, 329−331 Kalahari craton, 656, 657, 661 Kalak Nappe, 585 Kaleva, 201, 203, 240, 246, 273, 285, 287, 291, 298, 313−315, 317−321, 323−327, 331, 350, 379 Kaleva, basal, 314 Kaleva, Lower, 240, 246, 285, 287, 291, 298, 313−315, 317−319, 321, 325, 327, 331, 350
Kaleva, sub-Lower, 291, 313, 314, 317 Kaleva, sub-Upper, 318, 319 Kaleva, traditional, 313 Kaleva, Upper, 203, 240, 246, 273, 285, 287, 291, 314, 318−321, 323, 325−327, 331, 350, 379 Kaleva tectofacies, 240, 314, 319, 326 Kalevian, 144, 198, 203, 317, 321, 349, 350, 691 Kalevian Group, 198, 203 Kalix belt, 285 Kaolin, 329, 586, 698 Kaolinite, 301, 587 Kaolinitic, 581, 586, 589 Kaolinization, 301 Karelia−Kola−Kuloi cratonic block, 619 Karelia−Norrbotten boundary, 509, 511 Karelian craton, 25, 28, 240, 242, 256, 273, 349, 487−489, 494, 496, 497, 500, 501, 503−505, 508, 509, 513, 517, 518, 577, 633, 635, 638, 639, 656 Karelian domain, 22−26, 28, 43, 51, 53, 56, 57, 59, 61, 62, 65, 66, 69−72, 74−78, 80−83, 282, 283, 291, 499, 635 Karelian formations, 13, 144, 179, 215, 223, 282, 283, 287, 292, 295, 306, 320, 323, 325, 327, 331, 349 Karelian province, 26, 75, 201, 203, 488, 651−654 Karelian supergroup, 198, 201, 282, 287 Karjalite, 171, 203, 212, 214, 215, 310, 324 Kärki block, 422 Katarchean, 691 Kautokeino greenstone belt, 182 Keitele microcontinent, 346, 351, 356, 361, 362, 365, 393, 394, 498, 500, 503−505, 507, 508, 517 Keivy belt, 672 Kemiö−Järvenpää field, 348, 380, 381, 390, 392, 396 Kenorland (supercontinent), 178, 179, 656 Kersantite, 615, 616 Khondalite series, 175, 494 Kianta terrain, 29, 38−45, 48, 51−53, 56, 58, 59, 62, 75, 79, 81, 82 Kiiminki belt, 11, 285, 287, 313, 317 Kimberlite, 22, 25, 53, 57, 76, 566, 587, 592, 593, 606, 608, 609, 615−623, 627−640, 689, 700 Kimberlite, Group I, 617, 619, 621, 627, 630−637 Kimberlite, Group II, 617, 618, 631, 634 Kinoshitalite, 612, 616, 623 Kiruna−Arvidsjaur porphyry Group, 182 Kittilä allochthon, 240, 263, 271, 487, 489, 497, 500, 504, 517, 518 Kittilä arc, 494, 497, 504 Kittilä greenstone, 142, 144, 154, 156, 158, 160, 162, 164, 183, 262, 263, 271 Kittilä Group, 146−148, 151−153, 158−167, 179−182 Kivalo Group, 155, 177, 301, 675 Kola craton, 487, 489, 494, 496, 501, 505, 517 Kola domain, 22, 24, 62, 65, 70−72 INDEX
•
725
Kola province, 25, 201, 587, 652 Kola−Lapland orogeny, 28 Koli belt, 307 Koli quartzite, 309 Koli sill, 213, 220, 221, 223, 224 Komatiite, 22, 29, 33, 40, 45−47, 60, 76, 82, 145, 150−154, 156−158, 171, 177, 183, 224, 300, 436 Komatiite, basaltic, 60, 104, 224 Komatiite, cumulate, 36, 44, 46 Komatiite, dike, 44, 45 Komatiite−picrite association, 156, 177 Kornerupine, 58, 697 Kortejärvi carbonatite, 611, 614, 635 Korvuanjoki Group, 303, 305, 307, 311, 330 Kostamuksha greenstone belt, 81 Kotalahti Ni-belt, 415 Kotalahti-type (Ni-Cu sulfide deposit), 434 Kuhmo belt, 11, 40 Kuhmo block, 201, 203−205, 207−209, 213−215, 221, 223, 224, 226, 228−230, 295, 299, 330, 611, 614 Kuhmo greenstone belt, 11, 41, 43−45, 47−49, 51, 78, 81, 203, 215 Kuhmo terrain, 75 Kumpu Group, 140, 146, 147, 152, 164−167, 181−183, 324 Kuopio belt, 285 Kuorboaivi schist belt, 71 Kurkikylä Group, 297, 300, 303, 329 Kuusamo belt, 11, 62, 64, 65, 105, 131, 142, 143, 154, 283, 285, 287, 289, 290, 295, 297, 300, 303, 305, 306, 309, 311, 323, 324, 326, 487, 501, 673 Kuusamo block, 285 Kyanite, 33, 35, 38, 47, 58, 301, 302, 329 Kyanite quartzite, 302, 329 Kyykkä Group, 295 Labradorite, 223, 546 Lainio Group, 147, 148, 152, 153, 164−167, 181−183 Laitila batholith, 545, 546, 548, 549 Laivajoki carbonatite, 611, 614, 635 Lamproite, 575, 615−618, 631, 632, 634, 636 Lamprophyre, 459, 461, 615, 616, 627, 628, 630−633 Lamprophyre, ultramafic (UML), 246, 266, 267, 615, 616, 627, 628, 630, 632 Lapinlahti gabbro−anorthosite, 412, 415, 417, 418, 421 Lapland granulite belt, 6, 7, 14, 24, 28, 61, 65, 66, 68, 71, 72, 74, 76, 142−144, 158, 174−176, 182, 183, 287, 487−489, 494, 496, 500, 501, 503, 504, 505, 507, 513, 517 Lapland−Kola orogen, 14, 175, 176, 488, 500, 501, 503, 517 Lapland−Kola orogeny, 182, 504, 518 Lapland−Kola suture zone, 178 Lapponian, 144−146, 306, 691
726
•
INDEX
Lateorogenic granites, 14, 446, 448, 449, 456−460, 462, 469−474 Lauhanvuori granite, 428 Laurentia, 507, 513, 557, 584, 589, 656−661 Laurite, 131 Lava, basaltic, 70, 160, 299, 365, 366, 376, 381, 554 Lava, komatiitic, 46, 48, 68 Lava channel, 44 Layered intrusion, 7, 11, 25, 59, 62, 104, 105, 107, 108, 118, 120, 122, 125, 126, 131, 142, 143, 146, 154, 157, 167, 169−172, 176−178, 183, 201, 203−205, 210, 211, 214, 215, 219, 226, 285, 291, 292, 294, 300, 324, 421, 423−430, 433, 434, 436 Leucite, 616, 618, 633 Leucocratic, 51, 123, 263, 422, 425, 541, 543, 545 Leucogabbro, 106, 110, 116, 123, 417, 421, 430, 545 Leucogabbronorite, 539, 545 Leucogranite, 29, 35, 38, 82 Leuconorite, 545 Leucosome, 35, 52, 60, 67, 182, 349, 489, 491 Leucotonalite, 32. 41, 52 Leucotrondhjemite, 39 Leucoxene, 301 Lieksa complex, 29, 35, 37, 38, 40, 82 Lieksa complex, granulite, 37, 82 Ligurian-type ophiolite, 240, 243, 246 “Limestone, Baltic”, 582 Limestone, calcilutitic, 582 Limestone, glauconitic, 584 Limestone, Ordovician, 15, 584 Lithosphere, 22, 25, 47, 73−80, 83, 84, 120, 154, 178−180, 198, 240, 241, 257, 262, 263, 269−272, 497, 498, 510, 517, 630, 636−638 Lithosphere, Archean, 25, 73, 74, 76, 79, 80, 83 Lithosphere, continental, 76−78, 270, 498 Lithosphere, oceanic, 77, 79, 180, 198, 240, 241, 257, 262, 263, 269, 271, 497, 517 Lithospheric delamination, 83, 84, 507 Lithostratigraphic, 82, 144−146, 158, 165, 287, 310, 313, 320, 326, 331, 501, 567 Lizardite, 253, 257, 262, 421 Loljunmaa dike, 119, 120, 122 Lomagundi Group (Zimbabwe), 673, 677 Lonestone, 295, 297 Lopian cycle, 26 Lopolith, 426, 574 Lower Kaleva belt, 317 Lowstand, 291, 305, 312, 324, 329, 330 Luoma Group, 43, 44 Mackinawite, 422 Mafic dike, 37, 54, 56−58, 62, 74, 119, 143, 151, 159, 178, 198, 201, 203, 204, 207, 209, 223, 226, 229, 246, 264, 541, 566, 569, 570, 572, 574, 579, 589, 593, 650, 652, 653, 661 Magma, basaltic, 118, 132, 164, 198, 271, 470,
573, 576 Magma mixing and mingling, 429, 452, 461, 472, 548, 551, 552, 554 Magma type, 104, 107, 118, 120, 132, 157, 204, 210, 226, 618 Magmatism, alkaline, 81, 255, 508 Magmatism, arc, 79, 82, 83, 356, 362, 375, 392, 394, 446, 447, 472, 474, 496, 497, 501, 507, 517 Magmatism, bimodal, 47, 62, 82, 180, 356, 381, 412, 428, 429, 461, 537, 539, 546, 549, 552, 555, 556 Magmatism, calc-alkaline plutonic, 71, 473, 474, 654 Magmatism, carbonatite, 57, 81 Magmatism, granite, 356, 458, 461, 512, 518 Magmatism, mafic, 25, 77, 83, 104, 155, 226, 231, 414, 426, 520, 556, 688 Magmatism, obduction-related, 181 Magmatism, ocean-floor, 243 Magmatism, rapakivi granite, 515, 539, 552, 553, 555, 556 Magmatism, subduction-related, 179, 375, 484, 501, 504, 508, 509 Magmatism, TTG, 40 Magnetotelluric, 25, 73, 496 Mantle, asthenospheric, 271, 273, 576, 635, 636, 638 Mantle, depleted, 75, 76, 158, 160, 176, 246, 249, 263, 266, 267, 352, 362, 366, 378, 384, 394, 433, 435, 456, 467−471, 498, 555, 576−579 Mantle, lithospheric, 25, 73, 74, 76, 77, 176, 179, 209, 226, 243, 247, 251, 254, 255, 262, 269−271, 273, 435, 470, 473, 498, 509, 556, 577, 606, 608, 619, 620, 627, 631, 633−640 Mantle, lithospheric, oceanic, 243, 254, 262 Mantle, lithospheric, subcontinental, 25, 74, 77, 176, 209, 226, 243, 247, 251, 254, 255, 269, 270, 271, 273, 435, 470, 498, 509, 619, 633, 635, 636 Mantle, melt(ing), 69, 118, 158, 177, 265−267, 269, 273, 434, 515, 553, 619, 630, 634, 635 Mantle, primitive, 149, 152, 159, 162, 164, 226, 254, 255, 264−267, 269, 270 Mantle-derived, 37, 39, 180, 452, 468, 470, 509, 514, 553, 606, 620, 638, 693 Mantle diapir, 243, 271, 273, 556 Mantle metasomatism, 264, 265, 271 Mantle peridotite, 240, 246, 249, 251, 253, 255, 257, 265, 269, 271, 553, 617, 618, 636 Mantle plume, 47, 131, 154, 176−180, 183, 198, 209, 226, 556, 557, 588, 656 Mantle tectonites, 240, 241, 243, 244, 246, 247, 249, 251, 253, 257, 258, 262−264, 266, 271 Mantle xenocryst, 606, 616, 618, 620−622, 627, 638, 639 Mantle xenolith, 25, 75, 77, 243, 269, 616, 618, 619, 622, 636−639 Map, bedrock, 490−493, 540 Map, dike swarms, 199, 200, 206, 218, 219, 225,
227, 229, 230, 568, 570, 575 Map, geological, 5, 23, 24, 30−32, 41, 42, 50, 66, 67, 72, 105, 108, 110, 112, 115, 143, 145, 147, 168, 173, 242, 245, 256, 259, 263, 283, 286, 288−290, 294, 298, 328, 347, 357, 366, 411, 418, 423, 425, 427, 429, 431, 432, 447, 450, 453, 455, 457, 460, 461, 485, 486, 554, 568, 570, 571, 575, 583, 610, 612, 629, 674 Map, gravity anomaly, 9, 117, 418, 432, 492, 624 Map, lithological, 366, 382 Map, magnetic (anomaly), 8, 30, 41, 105, 147, 200, 230, 418, 427, 432, 487, 492, 614, 620, 624, 629 Map, petrological, 542−544, 547 Margin, passive, 77, 181, 241, 270, 272, 273, 319, 325, 327, 331, 508, 592, 672 Marginal chill zone, 173 Marginal series, 106, 107, 111−114, 116, 120, 122, 124−127, 129−133, 214, 414, 422, 426, 428 Megacyclic unit (mafic−ultramafic rocks), 102, 106−108, 110, 111, 118, 119, 123, 132, 133 Megacyclic unit (sedimentary rocks), 294 Melagabbro, 123 Melalahti Group, 311 Melanite (titanian andradite), 629, 630 Melanocratic, 422, 430 Melasyenite, 608 Melatroctolite, 548 Melt migration, 45, 51, 630 Melteigite, 628−630 Mesoproterozoic, 4, 7, 285, 488, 499, 500, 515, 537, 564, 566, 567, 569, 570, 573−577, 580, 581, 587−589, 591, 594, 648, 660 Mesozoic, 2, 241, 566, 586−588, 591, 593, 594 Metabasalt, 32, 72, 259, 496 Metadacite, 497 Metagabbro, 159, 165, 258, 261, 298, 306 Metagraywacke, 158, 181, 240, 313, 317, 318, 366, 497 Metaluminous, 76, 428, 466, 469−471, 474, 549 Metamorphism, amphibolite-facies, 28, 33, 38, 43, 48, 51, 54, 56, 60, 68, 74, 80, 83, 104, 144, 175, 253, 282, 287, 349, 355, 361, 365, 375, 380, 394, 419, 519, 585 Metamorphism, blueschist-facies, 78 Metamorphism, burial, 648 Metamorphism, contact, 58, 648 Metamorphism, eclogite-facies, 49, 175 Metamorphism, epidote-amphibolite facies, 78 Metamorphism, granulite belt, 175 Metamorphism, granulite-facies, 7, 12, 29, 37, 38, 43, 53, 56−58, 61, 71, 74, 75, 80, 83, 144, 175, 176, 183, 349, 355, 361, 380, 394, 419, 496, 501, 512, 650−652 Metamorphism, greenschist-facies, 53, 56, 57, 61, 80, 144, 146, 285, 287, 301 Metamorphism, Outokumpu nappe, 261 Metamorphism, Pirkanmaa belt, 365 Metamorphism, Tampere belt, 365 INDEX
•
727
Metamorphism, ultramafic rocks, 419 Metarhyolite, 508 Metasandstone, 285, 315, 319, 321 Metasedimentary rocks, 12, 14, 16, 22, 27, 142, 144, 146, 156, 287, 346, 354, 423, 496, 498 Metasomatic, 77, 116, 123, 161, 243, 257, 258, 473, 625, 628, 636, 692 Metasomatism, 51, 76, 264, 265, 271, 469, 470, 473, 610, 625, 626 Metasomatism, K, 626 Metasomatism, mantle, 243, 264, 265, 271 Metasomatism, Na, 469, 610, 625 Metavolcanic rock, felsic, 146, 149, 151−153, 165, 180, 226 Metavolcanic rock, intermediate, 155, 167 Metavolcanic rock, mafic, 105, 142, 147, 150, 151, 153−156, 158−160, 162, 163, 165, 179, 180, 484 Metavolcanic rock, ultramafic, 150, 153, 157 Mg-ilmenite, 605, 617, 620, 621 Microcline, 33, 451, 541, 610 Microcline granite, 14, 429, 446, 458 Microcline porphyroblast, 307 Microfossil, 566, 573, 574, 580, 583, 584, 587 Microgabbronorite, 116, 118−121, 125 Microtonalite dike, 450, 455, 456, 472 Mid-ocean ridge, 180, 240, 389, 517 Migmatite, 6, 11, 22, 26, 35, 38, 40, 43−45, 48, 49, 51, 52, 54, 58, 59, 67, 80, 81−83, 143, 175, 285, 287, 313, 319, 349, 351, 359, 361, 365, 375, 380, 381, 391, 393, 414, 416, 417, 419, 421, 422, 457, 491, 498, 504, 511−514, 518, 692, 693 Migmatite, agmatite, 49 Migmatite, granite, 349, 375, 380, 391, 393 Migmatite, nebulitic, 48, 51, 52, 692 Migmatite, neosome, 22, 49, 51, 58, 59, 61, 63, 419 Migmatite, paleosome, 49, 52, 53, 57, 63, 81 Migmatite, stromatic, 40, 48, 51, 52, 59, 67 Migmatite, TT (tonalite−trondhjemite), 11, 43, 48, 51, 57−59, 61, 73, 349, 416, 417 Migmatite, veined, 287 Mineralization, see also Deposit and Ore Mineralization, Ag-Pb-Zn (Taivaljärvi), 43 Mineralization, Au, 28, 29, 33, 60 Mineralization, chalcophile elements, 120 Mineralization, classification (Tornio−Näränkävaara belt), 120, 121 Mineralization, Cu-Ni-PGE, 124, 125 Mineralization, Cu-PGE, 126 Mineralization, disseminated oxides (ilmenite ore), 431 Mineralization, disseminated sulfides, 120, 122, 124, 126, 127, 129−132, 173, 174, 414, 415, 428 Mineralization, massive sulfide (deposits), 355, 356, 380, 390, 394, 414, 415, 421 Mineralization, Ni-Cu, 414 Mineralization, Ni, komatiite-hosted, 40 Mineralization, Ni-PGE, 173, 174 Mineralization, offset, 110, 128, 129
728
•
INDEX
Mineralization, PGE, 120, 125, 128, 130, 133 Mineralization, principal types (Tornio−Näränkävaara belt), 125, 126 Mineralization, reef-type, 124, 125, 132, 133 Mineralization, W-Mo, 39 Mineralogy, history (in Finland), 684−697 Minette, 615−617, 630, 632 Mingling, magma, 429, 452, 454, 455, 461, 472, 546−449, 551, 552, 554 Mingling, magma−sediment, 417 Mixing, magma, 116, 249, 267, 270, 447, 461, 621 Mixing, rocks, 362, 412, 428, 455, 613, 635, 636 Mixing, sulfides, 262 Moho, 53, 78, 203, 553 Moho-depth map 491, 494, 497, 553, 554, 556 Molasse (molasse-like, molasse-type), 164, 181, 182, 353, 364, 508, 510, 513 Molybdenite, 39, 415, 422, 543 Monazite, 14, 38, 39, 57, 176, 458, 462, 503, 541, 571, 610, 614, 617, 651 Monticellite, 615−617, 620 Monchegorsk intrusion, 653 Monzodiorite, 143, 444, 454, 456, 458, 514, 539, 546 Monzogranite, 30, 40–42, 51, 52, 63, 71, 76, 80, 82, 83, 452 , 548 Monzonite, 7, 11, 152, 165, 456, 460, 462, 469, 473, 474, 509 Monzonite, Åva, 460, 462 Mooidraai Formation, 676 MORB, (mid-ocean ridge basalt), see Basalt Moresveijohjkan complex, 71 Mudrock, 355, 367 Mudstone, 306, 309, 311, 569, 573, 579, 580, 582−584, 592 Muhos Formation, 285, 567, 573, 574, 580, 589 Nabar Nappe, 585, 586 Nalganas Nappe, 585 Napapiiri terrain, 63, 63, 64, 66, 81 Nappe, 15, 28, 69, 144, 182, 257, 261, 263, 273, 282, 291, 297, 309, 313, 319−321, 497, 582, 584−586 Nauvo−Korppoo field, 380, 385 Neogene, 567, 587, 593, 594 Neoproterozoic, 143, 566, 579, 580, 582, 586, 587, 589, 591, 592, 595, 619, 648, 660, 672, 676 Neosome, 22, 49, 51, 58, 59, 61, 63, 419 Nickeline, 422 Nilsiä belt, 307, 309 NMORB (normal mid-ocean ridge basalt), see Basalt Nordic orogen, 500, 513, 514 Nordic orogeny, 500, 505, 507, 513−515, 519, 520 NORDSIM (ion microprobe), 64, 181, 219, 243, 488 Norite, 175, 204, 414, 415, 422, 423, 652 North Karelia (schist) belt, 215, 224, 228, 240,
255, 256, 282, 297, 302, 317 Norrbotten craton, 487, 489, 494, 497, 498, 500, 501, 504 Norrbotten microcontinent, 500, 503−505, 508, 517 Nunnanlahti greenstone belt, 36, 56 Nunnanlahti shear zone, 36 Nuttio ophiolite, 240, 262, 263, 265, 266, 268, 271, 272 Nuttio serpentinite belt, 161, 162, 262, 263 Nuttio serpentinites, 152, 153, 161, 162, 240, 262, 265, 267 Obduction, 74, 181, 240, 257, 258, 262, 271, 273 Obduction, ophiolite, 273 Ocean floor arc, 446 Ocean floor basaltic (volcanic) rocks, 254, 271 Ocean floor environment, 358 Ocean floor magmatism, 243 Ocean floor ophiolites, 504 Ohravaara Group, 297 OIB (ocean island basalt), see Basalt Oijärvi greenstone belt, 50, 59−61 Oligoclase, 536 Olivine clinopyroxenite, 173 Olivine gabbro, 251, 415, 417, 426, 430, 586 Olivine gabbronorite, 115, 418, 430 Olivine leucogabbronorite, 539, 545 Olivine norite, 107, 415 Olivine pyroxenite, 172, 173 Olivine websterite, 173, 417, 422 Onkamo Group, 146, 148, 150−154, 156, 158, 171, 176 Ophiolite (see also Jormua, Outokumpu, and Nuttio ophiolites), 179, 180, 201, 202, 204, 223, 237, 240−244, 246, 251, 253, 257, 266, 268, 271, 273, 291, 325, 349, 484, 489, 504 Ophiolite complex, 11, 25, 161, 162, 179, 180, 285, 291, 318, 319 Ophiolite, Ligurian-type, 240, 243, 246 Ophiolite, tectonic setting, 272 Opukasjärvi Group, 70−72 Orangeite, 616 Ordovician, 15, 582, 584, 585, 591−593 Ore deposits (see also Mineralization), 261, 684, 687, 698, 699 Orijärvi area, 348, 380, 382, 384, 385, 388, 391, 392, 694 Orijärvi granodiorite, 348, 352, 381, 384, 385, 386, 390, 391 Orthoamphibole, 57 Orthoclase, 454 Orthocumulate, 46, 111, 112, 127, 128, 425 Orthoferrosilite, 540 Orthopyroxene, 37, 56, 57, 170, 172, 204, 205, 207, 209, 210, 214, 217, 223, 249, 253, 260, 414, 415, 417, 419, 424, 430, 434, 436, 452, 545, 548 Orthoquarzite, 154, 303, 305, 307, 309−311, 313,
321 Oskarshamn−Jönköping belt (OJB), 487, 491, 511, 512, 513 Otanmäki belt, 317 Otanmäki intrusion, 202, 203, 221 Oulanka complex, 266 Oulujärvi shear zone, 53, 59, 282−286, 298, 299 Outokumpu association (assemblage), 12, 259, 319, 327, 484 Outokumpu nappe (complex), 257, 261, 273, 282, 291, 313, 319−321 Outokumpu ophiolite, 161, 162, 203, 240, 241, 250, 257, 263−266, 268, 271−273, 291, 318, 319, 325 Outokumpu-type deposits/occurrences/ores, 261, 262, 319 Outokumpu-type (serpentinite/ultramafic) massifs, 240, 243, 250, 255−257, 260, 261, 264, 271, 273 Overprint, 25, 35, 38, 53, 54, 57, 58, 60, 63, 66, 74, 203, 636 Overprint, metamorphic, 53, 57, 60, 74 Overprint, metasomatic, 636 Overprint, Proterozoic, 25, 54, 58, 63, 66 Overprint, Svecofennian, 74, 203 Overprint, tectonic, 23, 54, 63 Overprint, thermal, 25, 54, 63, 66 Overthrusting, 74, 158, 181−183, 503 Oxygen fugacity, 47, 430 Pääjärvi bock, 228 Paanajärvi belt, 285, 297 Paleoclimate, (see also Climate), 586, 653 Paleocurrent, 183, 571, 572 Paleogene, 567, 587, 591, 593 Paleolatitude, 650−652, 654−656, 6559, 660 Paleolongitude, 650, 651 Paleomagnetism, 198, 228, 648, 688 Paleopole, 652, 655 Paleopole, key pole, 650, 655 Paleoproterozoic cover (rock), 170, 175, 181 Paleoregolith, 36 Paleosol, 295, 301−303, 349, 586−588, 591 Paleosome, 49, 52, 53, 57, 63, 81 Paleostress, 198, 226, 228, 229 Paleosuture, 450, 469 Paleozoic, 4, 15, 22, 25, 53, 77, 346, 566, 579, 582, 587, 588, 592−594, 648, 650, 689 Parautochthonous, 12, 144, 282, 314, 326 Parent(al) magma (melt), 118, 120, 129, 132, 149, 152, 156, 170, 172, 174, 180, 198, 210, 212, 215, 226, 251, 410, 415, 417, 422, 425, 429, 430, 432, 433, 436, 553, 640 Parent rock, 301, 586 Partial melting, 14, 22, 38, 46, 47, 49, 52, 53, 58, 61, 62, 69, 75, 76, 79, 84, 164, 175, 253, 265, 266−269, 468, 469, 553, 554, 556, 693 Pechenga belt, 70, 487, 488, 494, 501 Petchenga ferropicrite, 162, 177, 178 INDEX
•
729
Pegmatite (pegmatitic), granite, 2, 68, 75, 297, 424, 459, 514, 585 Pegmatite, gabbro, 116, 125, 128, 243, 244, 418 Pegmatite, ultramafic, 109, 114, 126, 254 Pellinki Group, 383, 385, 386, 388 Pellinki field, 348, 380, 385, 386 Pentlandite, 124, 127, 253, 261, 415, 422, 424 Peraluminous, 38, 49, 63, 71, 76, 428, 466, 468−471, 473, 474, 549 Peräpohja (schist) belt, 11, 59, 105, 131, 142, 143, 155, 171, 172, 177, 283, 285, 287, 290, 293−295, 300, 309, 317, 320, 323−326, 487, 501, 670, 672, 673, 675, 676 Peridotite, abyssal, 271 Peridotite, mantle, 249, 265, 553, 617, 618, 638 Peridotite massif, 257, 260, 271 Peridotite, orbicular, 408, 421 Peridotite zone, 414, 419, 422 Perovskite, 615−619, 621 Perthite, 509, 541, 547 Petrogenesis, alkaline rocks, 608 Petrogenesis, diabase dikes, 577 Petrogenesis, Finnish Precambrian, 687 Petrogenesis, granitoids, 48 Petrogenesis, layered mafic intrusions, 292 Petrogenesis, magmatic sulfide ores, 422, 435 Petrogenesis, rapakivi granites, 557, 688 PGE reef, 106−112, 122−124, 126−128, 130−133 Phanerozoic, 14, 29, 78, 79, 143, 198, 433, 446, 484, 488, 496, 520, 554, 569, 574, 587, 588, 591, 636 Phoscorite, 623, 628, 631 Phosphorite, 315, 317, 318, 362 Phosphorus ore, 608, 625, 626 Phlogopite, 172, 173, 214, 414, 419, 425, 608, 610, 612−618, 623, 625−628, 630 Picrite, 156−158, 162, 174, 177, 178, 183, 212, 374, 378, 380, 381, 383, 384, 386, 387, 390, 396 Picrite (“ferropicrite”), 162, 177, 178 Pielavesi−Pyhäsalmi region, 356 Pigeonite, 116, 174, 205, 217, 540 Pillow basalt, 35, 45, 385, 389 Pillow breccia, 150, 156, 159, 247, 363 Pillow lava, 148, 150, 151, 156, 159, 160, 244, 246, 249, 306, 313, 358, 360, 363, 368, 376, 378, 379, 381, 383, 391, 394−396 Pillow structure, 44, 156, 355, 360, 362, 374, 380 Pirkanmaa belt, 12, 346, 351, 359, 365, 371−374, 388−390, 393−396, 411, 412, 417, 419, 421−423, 426, 427, 433, 434, 437 Pilgujärvi Group, 177 Piteå−Raahe shear zone, 487 Plagiogranite, 180, 240, 241, 243, 250−252, 255 Plate tectonics, 13, 78, 198, 241, 242, 324, 346, 516, 694 Plate tectonics, tectonosedimentary model, 327 Plate tectonic model, 484, 488, 693−695 Plate tectonic (accretionary) processes, 6, 28,
730
•
INDEX
241, 516 Plate tectonics, 13, 78, 198, 241, 242, 324, 346, 516, 694 Platform, 81, 183, 313, 327, 329, 331, 566 Platform carbonate, 329, 331, 591, 592 Platform sedimentation, 591 Pleonaste, 621, 623 Plume, 26, 47, 77−79, 84, 104, 131, 154, 176−180, 183, 198, 209, 226, 556−558, 625, 636, 638, 652, 656 Podiform, 122, 246 Pohjanmaa belt, 12, 346, 348, 350−357, 361−363, 388, 389 Poikilitic, 106, 111, 112, 114, 116, 123, 127, 128, 172, 417, 425 Polmak−Pasvik−Pechenga belt, 70, 73 Pomokaira terrain 24, 61, 65, 68, 69, 71, 74, 76 Porphyry, 29, 40, 147, 148, 152, 153, 157, 162− 167, 179−182, 290, 300, 306, 317, 318, 348, 351, 367, 374, 376, 426, 452, 453, 462, 485, 497, 537, 539, 543, 545−549, 554, 615, 650, 655, 659, 675 Porphyry, feldspar, 348, 452, 453 Porphyry, felsic, 147, 148, 152, 153, 162−167, 180, 181, 318, 374 Porphyry, felsite, 546 Porphyry, granite, 157, 351, 462, 485 Porphyry, hornblende-plagioclase, 426 Porphyry, plagioclase, 348, 367, 376 Porphyry, quartz, 157, 300, 351, 462, 546, 549, 650, 655, 659 Porphyry, quartz-feldspar, 166, 348, 537, 539, 545−547, 548, 554 Porphyry, rhyolitic, 497 Portimo complex, 104, 106, 111, 122, 123, 125, 128−130, 132 Portimo dikes 111, 112, 114, 119, 125, 126, 129, 130 Post-collisional, 25, 53, 70, 78, 81, 519 Postcumulus, 251, 425 Postjotnian, 15, 203, 566, 569, 574−576, 578, 650, 655, 657, 660 Postorogenic granites, 15, 143, 183, 428, 446, 458, 460−462, 470−474, 514, 638 Pothole, 113, 421 Preorogenic rocks, 446, 448, 449, 466, 467, 471, 472, 474 Pretoria Group, 676 Primitive arc complex (of central Finland), 5, 350, 359, 412, 414, 434, 437, 447−449, 462, 467, 472, 474, 498 Primitive island arc, 14, 489 Primitive magma (or melt), 111, 156, 249 Primitive mantle, 264, 265, 267, 269, 270 Primitive mantle-normalized, 149, 152, 159, 162, 164, 254, 255, 266 Protolith, 54, 69, 176, 257, 287, 369, 585 Provenance, Archean, 56, 69, 155, 299, 314, 318, 320, 325
Provenance, Proterozoic, 71, 314, 317, 320 Provenance, Proterozoic, volcanic, 13, 349, 364, 366−368, 376 Pudasjärvi block, 201, 204, 205, 207, 226, 228, 230, 318, 330, 611, 614 Pudasjärvi complex, 11, 105, 131, 142, 273, 291, 292, 294, 317, 318, 325, 326, 329, 487, 489 Pudasjärvi−Iisalmi block, 330 Pudasjärvi terrain, 62, 81 Puolanka Group, Central, 59, 81, 282, 287, 298, 299, 302, 305−307, 309, 320, 326, 330 Puolanka Group, East, 291, 198, 307, 309−312, 330, 331 Pyrite, 36, 48, 253, 261, 292, 415, 436, 610, 617, 698 Pyrochlore, 610, 623, 626, 689 Pyroclastic, 29, 33, 43, 47, 62, 150, 297, 300, 313, 362, 366, 367, 374, 376, 380, 385, 386, 426 Pyrope, 606, 617, 620−622, 627, 638, 639 Pyrrhotite, 32, 124, 126, 127, 253, 261, 415, 422, 424, 436, 610, 623 Pyterlite, 536, 539, 540, 543, 545, 549 Quartz arenite,7, 315, 317, 349, 352, 353, 375, 380, 390−393, 396, 570, 583 Quartz arenite, Los, 392, 393 Quartz arenite, Tiirismaa, 352, 375, 391−393, 396 Quartz diorite, 110, 448, 451, 456 Quartz gabbro, 173 Quartz-feldspar porphyry, see Porphyry Quartz monzodiorite, 448, 546 Quartz monzonite, 12, 444, 448, 454, 456, 548 Quartz syenite, 114, 547 Quartzite (see also Orthoquartzite, Sericite quartzite), 6, 7, 11−13, 58, 64, 68, 70, 105, 144, 146, 148, 150, 154, 155, 165, 166, 171, 203, 280, 282, 285, 288, 290, 291, 297, 301−303, 305−307, 309−311, 313−315, 317, 318, 320, 321, 323, 324, 329, 490, 497, 510, 582, 585, 586, 659, 692 Quaternary, 586, 587, 688−690 Quetico belt, 52 Raahe−Ladoga (zone) belt, 4, 282, 283, 285, 319, 654 Radiogenic heat, 79, 83 Rain drop imprint, 572 Ranua terrain, 48, 53, 59, 60, 64, 74, 75 Rapakivi granite, 4, 15, 16, 446, 454, 498, 520, 536, 537, 539−541, 543−546, 548−557, 566, 567, 569−571, 575, 584, 588, 637, 655, 660, 688, 692, 693 Rapakivi granite, age, 536, 537, 539, 541, 548, 554, 555, 556 Rapakivi granite, bimodality, 515, 539, 546, 549, 552, 555, 556 Rapakivi granite, chemical composition, 549−552 Rapakivi granite, definition, 636, 537
Rapakivi granite, distribution, 537, 539, 540, 542−544, 548, 554, 556 Rapakivi granite, (granite) magmatism, 5, 515, 553, 555, 588 Rapakivi granite, origin, 552−556 Rapakivi granite, paleomagnetic, 198, 656, 658− 660, 695 Rapakivi granite, stratigraphic, 247, 326 Rapakivi granite, texture, 48, 533, 536, 537, 550, 552 Rapakivi granite, texture, origin, 550−552 Rautalampi region, 358 Rautavaara complex, 54, 56, 58, 78, 81, 82 Rautavaara, terrain, 74 Reconstruction, Archean−Proterozoic boundary, 484 Reconstruction, continental, 201, 557, 648, 650, 656−661, 695 Reconstruction, global-scale, 656 Reconstruction, paleogeographic, 325, 326, 329 Reconstruction, stratigraphic, 247, 326 Recumbent folding, 29, 144, 349, 356, 361, 365, 414 Redbed, 181 Reef-type PGE-deposit (see also PGE reef), 120, 122, 132 Reflection, 25, 71, 73, 78, 80, 82, 203, 228, 482, 484, 488, 490, 491, 493, 494, 496, 499, 507, 511, 516, 518 Refraction, 25, 53, 71, 73, 77, 78, 83, 484, 488, 490, 491, 493, 494, 496, 497, 507, 517 Reittiö belt, 301 Remanent magnetization, 648−654, 694 Remanent magnetization, multicomponent analysis, 649, 650 Remanent magnetization, primary remanence, 648−655, Remanent magnetization, secondary remanence, 648−650 Reworking, complex, 62 Reworking, crustal, 26, 27, 82, 494 Reworking, tectonic, 36, 54, 59, 62, 70, 73, 78, 83, 587 Reworking, thermal, 22, 25, 28, 29, 43, 58, 69, 73, 74, 77 Rhyodacite, 164 Rhyolite, 149, 163, 175, 180, 183, 348, 349, 356, 360, 364, 367, 370, 371, 374, 376−378, 380, 381, 388−390, 394, 395, 449, 461, 462, 508 Rhyolite, high-K, 348, 367, 370, 371, 381, 394, 395 Rhyolite, spherulitic dike, 461, 462 Richterite, 610, 612−614, 616−618, 623, 627, 628 Rift basin, 154, 291, 297, 315, 326, 327, 329, 330, 473, 498, 508, 557, 588, 591, 653 Rift inversion, 291−293, 329 Rifting, 13, 25, 47, 57, 70, 76−78, 104, 167, 176, 178, 179, 198, 201, 226, 228, 249, 255, 271, INDEX
•
731
273, 291, 309, 315, 317, 327, 331, 349, 367, 377, 391, 395, 396, 426, 497, 501, 503, 556, 573, 587, 636, 637, 652, 653, 655, 656, 658, 672, 673 Rifting, continental, 104, 176, 201, 249, 255, 327, 652 Ring complex, Åva, 459, 460, 462, 470 Ring complex, emplacement (intrusion) mechanism, 460, 462 Ring complex, Seglinge, 460−462 Rio Negro−Juruena belt, 657, 659 Riphean, 566, 589 Ripple mark, 571, 572, 581, 582 Rodingite, 252 Rodinia, 592, 648, 661 Ropi terrain, 24, 61, 69, 70 Ruoppapalo granodiorite, 152, 164, 166 Russian platform, 589 Saamian cycle, 26 Saari−Kiekki belt, 285, 299 Saarikylä Group, 43, 44 Salahmi belt, 285, 302, 303, 307, 314, 317 Saimaa area, 12, 350, 353−355, 359, 375, 377−379, 395, 396 Salittu picrite, 384, 387, 390, 391, 396 Salla greenstone area (belt), 142, 143, 146, 150, 156, 167, 204 Salla Group, 146, 148−154, 167, 169, 180 Salmi rapakivi (granite), 76, 537, 553, 556, 569, 588 Sandur, 297 Sandstone, 15, 295, 297, 306, 310, 314, 317, 320, 323, 368, 390, 564, 567, 569, 571−574, 579−585, 588, 589, 591 Sandstone, arkosic, 573 Sandstone, clastic, 580, 584 Sandstone dike, 584 Sandstone, fluvial, 368, 564 Sandstone, graded, 295, 320 Sandstone, laminated, 317, 323 Sanidine, 616 Sapphirine, 58, 697 Saprolite, 586, 587, 589 Sariola, 285, 287, 289, 291−295, 297−303, 306, 307, 309, 324, 326, 329 Sariola−Jatuli, 326, 329 Sariolian Group, 677 Sarmatia, 484, 507, 511, 512, 515, 519, 520 Sarmatian, 484, 507, 515 Satakunta sandstone, 567, 569, 571−574, 588 Savo arc, 362, 393, 494 Savo belt, 12, 346, 348, 350, 351, 353, 355−361, 367, 380, 393−395, 487, 489, 498, 503, 504, 509 Savukoski Group, 148, 152, 156−159, 166, 172, 177, 179 Scandian phase, 584−586 Scapolite, 305, 473
732
•
INDEX
SCLM (subcontinental lithospheric mantle), 243, 254, 255, 269, 270, 273, 635−638 Sea-floor spreading, 179 Sedimentary cover, 26, 240, 263, 489, 566, 569, 573−575, 582−584, 586, 588, 589, 591, 593, 594 Sedimentary (rocks), 142, 156, 161, 171, 172, 175, 177, 178, 180, 181 Sedimentary sequence, 146, 154 Sedimentation, 12, 15, 25, 28, 81, 154, 167, 177, 181−183, 203, 287, 295, 303, 319, 320, 325, 327, 329, 331, 446, 501, 517, 582, 587−589, 591, 691, 692 Sedimentation age, 325 Sedimentation, deep-water, 501 Sedimentation, margin, 319, 327 Sedimentation, platform, 591 Sedimentation, rift, 325, 517 Segregation, 251, 252, 254, 432, 434, 619−621 Seismic data, 78, 494, 516, 573, 580, 582 Seismic profile, 75, 488, 511, 573 Seismic reflection, 25, 71, 80, 203, 228, 488, 493, 499, 507, 511 Seismic refraction, 53, 73, 77, 83, 484, 488 Seismic soundings, 625 Seismic structure, 499 Seismic studies, 76, 82, 271, 517, 546, 553 Sericite quartzite, 70, 154, 301−303, 305−307, 309, 310, 330, 585 Shale, 52, 179, 567, 573, 582−585 Shear zone, 12, 16, 33, 35, 36, 51, 53, 54, 59, 62, 75, 78, 80, 144, 175, 282, 285, 287, 298, 299, 326, 349, 355, 356, 381, 390, 423, 449, 452, 457, 458, 487, 496, 498, 499, 507, 508, 512, 515, 519, 539, 611, 615, 636, 640 Shear zone, brittle−ductile, 51 Sheeted dike, 159, 238, 240, 243, 244, 246, 247, 249, 251, 254, 257, 267, 271, 273 Shoshonitic, 355, 367, 394, 395, 463, 470, 474 SHRIMP (zircon studies), 37, 39, 69, 71 Silicocarbonatite, 610, 613, 631 Sill, 15, 33, 36, 43, 44, 47, 49, 64, 112, 114, 155, 159, 171, 172, 177, 198, 201, 203, 204, 211−215, 219−221, 223, 224, 291, 301, 303, 306, 307, 309, 310, 324, 362, 370, 385, 386, 419, 423, 425, 430, 449, 495, 566, 569, 574, 575, 577, 586, 615, 653−655, 660 Sill, (meta)diabase, 301, 303, 307, 495, 566, 577 Sill, differentiated, 43, 49, 171, 177, 212 Sill, karjalitic, 212−214 Sill, komatiitic, 47 Sill, layered, 44, 203, 204, 211−213, 215, 221, 223, 324 Sill, (ultra)mafic, 43, 44, 47, 49, 64, 155, 309, 423, 574 Sill, tholeiitic, 586 Sillimanite, 57, 58, 64, 69, 175, 300, 363, 375, 391, 586 Silvevaara granodiorite, 29, 30, 37, 39
“Sirkka line”, 144, 156, 158 Siurua granulite complex, 60, 61 Skarn, 64, 259, 260, 313, 362, 366, 685 Skellefte district (SD), 5, 487, 489, 497, 503, 508, 514, 518 Skellefte field, 388, 389, 394, 395 Slumping, 29, 320, 428 Snowball Earth, 658 Soapstone, 37, 161, 699 Sodankylä Group, 148, 150, 154−156, 158, 165, 171, 172, 177 Sokli carbonatite, 68, 143, 608, 621, 625, 626, 635, 689 Somerjärvi Group, 298, 307, 311, 320 Sompujärvi block, 108, 294 Sørvaranger terrain, 24, 70−74 Sphalerite, 48, 261, 415, 422 Spidergram, 159, 371 Sperrylite, 124, 125 Spinifex, 46, 48 Stannite, 261 Staurolite, 33, 38, 58, 63, 70, 258, 297, 317 Stratigraphy (see also Lithostratigraphy), 29, 36, 40, 80, 106, 109−111, 113, 114, 127−129, 131, 146, 156, 169, 214, 220, 243, 244, 246, 247, 249, 263, 267, 271, 282, 289, 306, 309, 310, 312, 320, 325, 391, 417, 421, 426, 433, 592, 639, 687 Stratigraphy, cumulus, 127, 128, 421 Stratigraphy, Karelian sequence, 300, 330 Stratigraphy, layered intrusion, 106, 109−111, 114, 131, 169, 421, 426, 433 Stratigraphy, magmatic, 129, 220, 433 Stratigraphy, ophiolite, 243, 247, 249, 267, 271 Stratigraphy, Tampere belt, 12, 362, 365, 367, 370, 373 Strike-slip, 33, 35, 56, 78, 508, 517 Strike-slip, dilatacy pumping, 457 Stromatolite, 154, 177, 179, 306, 309, 313, 388, 391, 670 Strontianite, 610 Subarkose (subarkosic), 309, 310, 570, 580, 581 Subdivision, Kaleva, 314 Subdivision, Mesoproterozoic, 566 Subdivision, Neoproterozoic, 566 Subduction, 14, 26, 47, 71, 76, 79, 176, 179, 198, 251, 273, 325, 346, 350, 351, 356, 362, 365, 367, 375, 377, 381, 388, 392, 394− 396, 410, 414, 482−484, 488, 494, 501, 503− 505, 507−509, 511−520, 556, 557, 588, 694 Subduction, locking, 503, 508 Subduction, opposite polarities, 508, 517 Subduction-related, 179, 346, 350, 381, 388, 392, 484, 501, 504, 505, 508, 509 Subduction, retreating, 507, 511, 512, 516 Subduction reversal, 484, 503, 505, 509 Subduction switch-over, 503, 505 Subduction zone, 47, 79, 161, 180, 273, 325, 346, 351, 362, 365, 388, 392, 394, 410, 482, 484,
488, 494, 504, 505, 507, 508, 511, 512, 516, 518, 519, 556 Subduction zone, migration, 507, 508 Subjotnian, 15, 203, 566, 567, 569, 589, 655 Subjotnian diabase dike, 15, 203, 566, 567, 569, 655 Subsolidus mineral, 220 Subsolidus (re)equilibration, 419 Suisaarian, 177 Sulfide assemblage, Vammala, 421 Sulfide saturation, 415, 436 Sumi, 226, 285, 287, 289, 290, 291, 293, 300, 306, 324, 326, 329 Sumi Group, 226 Sumi−Sariola Group, 201, 226 Sumi−Sariola rift, 285, 326, 329−331 Sumi−Sariolian, 154, 176 Suomenniemi diabase dike swarm, 539, 540, 542, 547, 547 Suomenniemi pluton, 537, 539, 545−548, 553 Suomu terrain, 64, 66 Suomussalmi greenstone belt, 40, 42−45, 47 Supercontinent, 178, 179, 226, 507, 517, 589, 592, 648, 656, 658, 661, 695 Supercontinent, breakup, 178, 179, 226, 517, 589, 592, 658 Supergroup, Karelian, 198, 201, 282, 287 Supergroup, Svecofennian, 201 Superior craton, 81, 656, 657 Supracrustal belt, 12, 28, 29, 32, 35, 36, 38, 40, 44, 59, 62, 65, 68, 142, 178, 282, 287, 346, 350, 374, 380, 390, 394, 395, 417, 468, 472, 673 Supracrustal belt, Svecofennian, 380, 390, 394, 395 Supracrustal gneiss, 24, 29, 56, 64, 459 Supracrustal rocks, 7, 12, 13, 26, 35, 36, 38, 51, 59, 63, 64, 68, 72, 74, 105, 142, 144, 146, 150, 175, 177, 179, 203, 282, 285, 287, 290, 294, 324, 329, 331, 346, 348−350, 358, 361−363, 368, 374, 375, 380, 381, 383, 385, 388, 393, 396, 417, 449, 452, 456, 458, 459, 489, 491, 498, 511, 554, 566, 588, 691 Supracrustal rocks, Karelian, 297, 290, 324 Supracrustal rocks, Proterozoic, 7, 146, 175, 331 Supracrustal sequence, 28, 29, 51, 62, 80, 81, 144, 164, 285, 356, 566 Svecobaltia, 494, 519 Svecobaltic orogen, 500, 507, 519, 520 Svecobaltic orogeny, 507, 511−513, 518, 519 Svecofennia, central, 346, 350−354, 361, 362, 364−367, 372, 379, 380, 382, 392−395 Svecofennia, southern, 346, 350−353, 359, 365, 374, 375, 382, 393, 395, 396 Svecofennian, 4, 5, 350 Svecofennian arc complex, 181 Svecofennian basement, 569, 570 Svecofennian bedrock, 6, 13, 14, 348, 447, 536, 539, 546−548 Svecofennian belt, 351, 355, 692, 694 INDEX
•
733
Svecofennian domain, 6, 14, 22, 25, 75, 83, 346, 348−350, 375, 390, 393, 410, 484, 488−490, 499, 694 Svecofennian island arc, 6, 14, 256, 325, 655 Svecofennian nappe, 28 Svecofennian orogen, 14, 181, 410, 412, 430, 488, 500 Svecofennian orogeny, 14, 15, 22, 25, 26, 53, 54, 56, 57, 62, 74, 165, 181, 201, 203, 268, 319, 329, 331, 412, 428, 448, 484, 500, 552, 567, 649−651, 653, 654, 659, 692 Svecofennian orogeny, lateorogenic stage, 446 Svecofennian orogeny, postorogenic stage, 446 Svecofennian orogeny, preorogenic stage, 446 Svecofennian orogeny, synorogenic stage, 412, 428, 437, 446 Svecofennian rocks, 4, 484, 539, 547, 552, 569 Svecofennian (sub)division, 4, 5, 350 Sveconorwegian, 488, 492, 515, 570, 587−589, 591, 650 Sveconorwegian orogeny, 492, 587, 589, 650 SVEKALAPKO (an Europrobe project), 488 SVEKA (seismic refraction profile), 77, 78, 488, 490 Svionian, 350, 393, 692 Syenite, 68, 114, 115, 498, 547, 608, 610−612, 628 Syenitoid, 514 Synorogenic rocks, 70, 182, 411, 412, 421, 430, 446, 448, 450, 455, 468, 469, 471−474, 654 Taivalkoski block, 207, 209, 228 Talus, 299 Tampere belt, 12, 348, 350−355, 362, 365−367, 369−375, 377, 388, 391, 393−396, 417, 423, 427, 487, 489, 498, 505, 516 Tampere belt, central, 365−367, 371 Tampere belt, eastern, 369 Tampere belt, lithological map, 366 Tampere belt, western, 369, 371 Tanaelv (Tana) belt, 175, 176, 182, 183 Tarkki granite, 540, 545 Tasanvaara tonalite, 36, 38, 39 Tavastia island arc, 499, 500, 503, 507, 508, 512 TDM model ages, see Isotopic age Tectofacies, 289, 291, 300 Tectofacies, Jatuli, 245, 303, 306, 307, 309, 310, 313, 320, 321, 329 Tectofacies, Kainuu, 299, 303, 305−307, 313, 326, 329, 330 Tectofacies, Kaleva, lower (eastern), 240, 245, 314, 326 Tectofacies, Kaleva, upper (western), 240, 245, 319, 320, 326 Tectofacies, Karelian, 289, 291, 320, 327 Tectofacies, Sariola, 289, 295, 299−301, 305, 307 Tectofacies, Sumi, 289, 290 Tectofacies, Sumi−Sariola, 329 Tectonic contact, 46, 51, 161, 282, 285, 291, 307, 311, 456
734
•
INDEX
Tectonic emplacement, 161, 164, 180, 181 Tectonic event, 40, 42, 52, 64, 78, 181, 428, 507, 587, 593, 648, 650 Tectonic (geotectonic) model, 78, 81, 82, 180, 324, 488, 500, 503, 507, 515, 520, 594 Temperature, blocking, 56, 380 Temperature, unblocking, 649, 652 Tempestite, 291, 299, 323 Tempestitic, 305, 306, 329, 330 Tersk (granulite) terrane, 175, 487, 494, 496, 500, 501 Tertiary, 567, 586 Tetraferriphlogopite, 610, 612, 614, 616, 618, 623, 626, 627 Thermochronometry, titanite-rutile, 513 Tholeiite, 15, 29, 40, 44−46, 49, 73, 149, 163, 201, 203−205, 207−210, 212, 215, 217, 224, 226, 228, 251, 252, 263, 378, 380, 430, 434, 435, 449, 472, 548, 575, 576, 586 Tholeiite, basaltic, 318 Tholeiite, continental, 204 Tholeiite, Fe-rich, 36, 48, 68, 159, 209, 224 Tholeiite, Fe-tholeiite, 60, 201, 203, 204, 207−210, 215, 217, 221, 224, 226, 228, 324 Tholeiite, IAT, 160, 204, 209, 224, 262, 576 Tholeiite, island arc tholeiite, 160, 161, 180, 224, 267, 356, 361, 384, 576 Tholeiite, low-Al, karjalite, 203, 204, 212, 215, 224, 310, 324 Tholeiite, low-Ti, 204, 205 Tholeiite, Mg-tholeiite, 60 Tholeiite, Mg-rich, 36 Tholeiite, MORB, 51 Tholeiite, olivine, 548, 569, 577, 579 Tholeiite, transitional, 575 Thrust, 36, 53, 61, 65, 68, 73, 175, 361, 362, 365, 417, 512, 582, 584−586 Thrust belt, 144, 182, 242, 508, 593 Thrust front, 68 Thrust plane, 457, 512, 514 Thrust sheet, 25, 175, 182, 319, 584 Thrust slice, 282, 514 Thrust wedge, 496, 501 Thrust zone, 158, 325 Thrusting, (see also Overthrusting, Underthrusing), 38, 40, 53, 54, 57, 82, 83, 144, 176, 181, 182, 501, 503, 504, 507−509, 512−514, 516−518, 585 Tidalite, 330 Tipasjärvi greenstone belt, 40, 47, 48, 51, 53 Tipasjärvi supracrustal belt, 44, 51 Ti-pyrope (titanian pyrope), 616, 638 Tirilite, 540, 543 Tirmo Group, 385 Tonalite, 32, 35, 38−40, 43, 51, 52, 57, 58, 70, 71, 79, 251, 416, 449, 451, 456, 489 Topaz-bearing granite, 541, 543, 545, 549 Topaz-bearing quartz porphyry, 549 Tornio−Näränkävaara belt, 104−107, 113, 116,
118−120, 131, 132, 167, 201, 205, 226, Tornquist line, 570, 592 Tornquist Ocean, 592 Trachyandesite, 165, 358, 378, Trachyandesite, high-K, 358 Trachyandesite, high-P, 358 Trachyte, 164, 165, 167, 367, 370, 378, 394 Transform fault, 503, 504, 507, 512, 518, 519 Transpression, 35, 38, 51, 53, 54, 82, 512, 519 Transpression, dextral, 38, 51, 54, 82 Transpressional, 15, 36, 349, 457, 473, 507 Transscandinavian igneous belt (TIB), 5, 487, 492, 499, 554 Tremolite, 36, 37, 46, 60, 253, 257−261, 614 Troctolite, 545, 548, 586 Trondhjemite, 13, 40, 43, 51, 52, 251, 349, 446, 456 Trondhjemitic, 39, 40, 48, 57, 58, 61, 62, 69, 73 Tsuomasvarri (gabbro−diorite) intrusion, 654 TTG (trondhjemite-tonalite-granodiorite), 39, 43, 47, 49, 51, 52, 57, 73, 79, 142, 251, 416, 417, 456 TTG, Archean, 39, 49, 51, 142 TTG, magmatism, 47, 49, 79 TTG, migmatites, 43, 51, 58, 73, 416, 417 Tuff breccia, 370, 376, 383 Tuffite, 156, 159, 165, 285, 298, 303, 309−311, 313, 318, 330, 348, 356 Tulppio supracrustal belt, 65, 68 Tundra intrusion, 649 Tuntsa supracrustal belt, 65, 68 Tuntsa terrain, 24, 64−68, 71, 74, 81 Turbidite, 33, 35, 36, 60, 203, 273, 285, 291, 297, 299, 313−315, 317−319, 325, 361, 364, 366−369, 371, 373, 375, 385, 388, 389, 393, 394, 494, 696 Turbiditic, 12, 32, 63, 203, 240, 297, 299, 300, 305, 309, 313−315, 317, 318, 321, 329, 331, 348−350, 361, 365, 366, 389, 393, 394, 417, 497, 504 Turbidity current, 295, 319, 367 Ukrainian Shield, 485, 515, 519, 520, 651, 659 Ultramafic dike, 49, 112, 161, 246, 264, 267, 269 Ultramafic lamprophyre, 246, 249, 266, 267, 410, 615, 627, 628, 630, 632 Ultramafic massif, 250, 255−259, 261−264, 267, 271, 273, 419 Ultramafic pipe, 415 Ultramafic rocks, 11, 12, 25, 36, 37, 45, 48, 60, 68, 70, 74, 82, 104, 106, 108, 116, 122, 132, 148, 150, 153, 156, 157, 161, 162, 173−175, 180, 183, 203, 244, 257, 260, 262, 298, 361, 365, 379, 385, 410, 415, 417, 484, 497, 566, 585, 615, 618, 625 Ulvöspinel, 616, 621 Umba granulite terrane (UGT), 175, 182, 487, 494, 496, 500, 501, 503, 504 Umeå allochthon, 499, 513, 518, 519
Umeå, (area) field, 389, 390, 394, (494) Unconformity, 36, 54, 59, 62, 64, 74, 144, 177, 203, 204, 213, 289, 291−294, 299, 301−303, 305, 307, 309, 310, 314, 317−319, 321, 393, 674, 677, 692 Unconformity, angular, 181, 320, 323, 393, 396, 566 Unconformity, erosional, 307, 314 Underplating, 53, 75, 77, 78, 83, 429, 484, 518, 554, 556, 558, 588, 639 Underplating, magmatic, 75, 78, 83, 518, 554, 556 Underthrusting, 71, 77 Uplift, 14, 16, 38, 74, 84, 105, 177−179, 181−183, 203, 228, 230, 447, 460, 505, 591, 593, 594, 636 Uplift, crustal, 179, 183 Uplift, isostatic, 16, 447 Urtite, 628 Uusimaa belt, 12, 346, 348, 350, 352−354, 375, 376, 380, 383−388, 390−392, 395, 396, 487, 489, 499 VAB (volcanic arc basalt), see Basalt Vaddas Nappe, 583, 585, 586 Väkkärä granite, 539, 545 Valleriite, 422 Vammala Ni-province, 419, 421, 423, 424, 432, 433, 436 Vammala-type (Ni-Cu sulfide deposit), 434 Variolitic, 46, 150, 156 Värmälä pluton, 453 Varpaisjärvi granulite complex, 53, 54, 57 Varpaisjärvi block, 228 Väyrylänkylä nappe, 297 Vehmaa batholith, 15, 537, 539, 556 Veittijärvi conglomerate, 370 Vendian, 566, 579, 591 Vendian period, 15 Vetreny belt, 154, 171 VHMS (volcanic-hosted massive sulfide) deposit, 389 Vihajärvi, Group, 298, 320, 321, 323 Viianki block, 228 Virtasalmi field, 359−361, 394 Virtasalmi region, 348, 358, 359, 375, 377 Vodlozero block, 652 Volcanic rocks (see also Metavolcanic rocks) Volcanic rocks, acid, 290, 300, 329 Volcanic rocks, bimodal, 12, 62, 356, 394, 395 Volcanic rocks, calc-alkaline, 152, 300, 356, 364, 377, 378, 380, 388, 426 Volcanic rocks, felsic, 7, 11, 12, 36, 37, 43, 44, 47, 48, 63, 152, 180, 183, 290, 350, 356, 360, 362, 374, 376, 384, 388−390, 393, 449, 507, 675 Volcanic rocks, intermediate, 11, 12, 290, 300, 385 Volcanic rocks, mafic, 7, 11, 12, 29, 44, 68, 73, INDEX
•
735
177, 318, 355, 356, 358−362, 364, 372, 374, 380, 384, 385, 387−391, 394, 449, 653 Volcanic rocks, rhyolitic, 62, 180, 388, 395 Volcanic rocks, ultramafic, 11, 48, 60, 635, 685, 396, 497 Volcaniclastic deposit, 35, 36, 59, 385 Volcaniclastic rocks, 33, 44, 146, 148, 156, 363, 385, 386, 389 Volcanism, 154, 158, 177, 179, 181, 182, 317 Volcanism, ocean floor, 179 Volcano-sedimentary sequence, 14, 144, 226, 240, 417 Volgo-Uralian, 484, 488, 519 Vuokatti Group, 311 Weathering, 13, 16, 36, 52, 150, 201, 249, 291− 293, 295, 297, 300−303, 305, 307, 327, 329, 330, 393, 581, 586−589, 591, 616, 626, 674 Weathering crust, 586 Weathering profile, 291, 305, 586, 591 Weathering, chemical, 52, 201, 291, 301–303, 307, 327, 329, 330, 586 Weathering, in situ, 292, 293, 307, 586 Weathering, kaolinitic, 581 Weathering, paleoweathering, 150, 305 Weathering, physical, 291, 295, 297, 329 Weathering, subareal, 591 Websterite, 106, 107, 114, 174, 417, 422 Wehrlite (see also Gabbro-wehrlite association), 77, 204, 214, 223, 422, 621, 635−637, 639 Wiborg batholith, 534, 539, 540, 542, 543, 545, 546, 548, 549, 551 Wiborgite, 534, 536, 539, 540, 542, 543, 545,
736
•
INDEX
549, 551, 552 WPB (within-plate basalt), see Basalt Xenocrysts, 29, 43, 243, 306, 606, 616−618, 620−622, 627, 634, 637−639 Xenocrystic (zircon), 68, 69, 244, 324, 498 Xenolith, 73, 77, 83, 114, 116, 118, 119, 121, 125, 150, 174, 414, 449, 459, 608, 616, 634, 637, 638, 639 Xenolith, crustal, 53, 57, 75, 77, 78, 434, 515, 620, 623 Xenolith, gneiss, 150, 414, 459 Xenolith, mantle, 25, 75, 77, 243, 269, 616, 618, 619, 636, 637, 639 Xenoliths, in kimberlite, 22, 25, 53, 57, 76, 77, 622, 638 Xenotime, 58, 75, 541 Yilgarn craton, 80, 81, 656−658 Ylivieska field, 346, 350, 351, 356, 361−364, 374, 388, 389, 394, 395 Zircon, 25, 27, 29, 39, 40, 43, 44, 46, 48, 52, 56, 59, 61, 63, 64, 68, 69, 76, 120, 163, 166, 243, 247, 254, 255, 269, 270, 324, 348, 376, 374, 379, 381, 392, 412, 425, 448, 451, 452, 454, 456, 460, 462, 467, 473, 498, 501, 515, 541, 546, 548, 575, 577, 610, 611, 613−615, 623, 638, 651 Zircon, detrital, 22, 62−64, 142, 155, 166, 167, 181, 273, 282, 291, 314, 319, 323, 325, 350, 352, 353, 362, 367, 375, 379, 381, 390, 392, 396, 493, 494, 498, 571