EUROGRANITES 2005— Proterozoic and
Archean Granites and Related Rocks of the
Finnish Precambrian
O.T. Rämö, J. Halla, M. Nironen,
L.S. Lauri, M.I. Kurhila, A. Käpyaho,
P. Sorjonen-Ward, O. Äikäs
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ISSN 1795-8946
ISBN 952-10-2603-0 paperback
ISBN 952-10-2604-9 pdf
Addresses:
O.T. Rämö, L.S. Lauri, M.I. Kurhila:
Department of Geology, P.O. Box 64, FI-00014 University of Helsinki, Finland
J. Halla:
Geological Museum, Arppeanum, P.O. Box 11, FI-00014 University of Helsinki, Finland
M. Nironen, A. Käpyaho:
Geological Survey of Finland, P.O. Box 96, FI-02151 Espoo, Finland
P. Sorjonen-Ward, O. Äikäs:
Geological Survey of Finland, P.O. Box 1237, FI-70211 Kuopio, Finland
This publication should be referred to as:
Rämö, O.T., Halla, J., Nironen, M., Lauri, L.S., Kurhila, M.I., Käpyaho, A., Sorjonen-Ward, P., Äikäs,
O., 2005. EUROGRANITES 2005— Proterozoic and Archean Granites and Related Rocks of the Finnish
Precambrian. Eurogranites 2005 Field Conference, September 11–17, 2005. Publications of the Department
of Geology A1, 130 p.
Layout and technical editing: Laura S. Lauri
Printed by: Gummerus Kirjapaino Oy,
Saarijärvi 2005
Cover photo: the lateorogenic (~1.80 Ga) Puruvesi granite in Puruvesi, southeastern Finland
Photo: O. Tapani Rämö
(Geological Survey of Finland)
(Geological Survey of Finland)
Back cover photo: Jukka Lehtinen
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Foreword
Didier (1971–1975), L. Lameyre (1976–1980), and B. Bonin (since 1981). Most of the past Eurogranites
Precambrian granitoids and related rocks have been at the focus only sporadically, the last time in
southern Norway in 1996. We are now more than pleased to have the opportunity to take the Eurogranites
to the heart of the Fennoscandian shield of northern Europe to examine a wide variety of Proterozoic and
Archean granites and related rocks.
an overview of the lithology and evolution of the Fennoscandian shield, with special emphasis on the
synorogenic, postorogenic, and anorogenic granitoid and related rocks of southern and central Finland.
of Helsinki and the Geological Survey of Finland. It will run for six days (September 12–17, 2005),
examining a total of 22 lithological targets in (1) the classic mid-Proterozoic rapakivi granite region
of southeastern Finland; (2) the synorogenic, lateorogenic, and postorogenic granitoid realms of the
Paleoproterozoic Svecofennian orogen; and (3) the Karelian craton hosting Meso- and Neoarchean TTG’s,
sanukitoids, and related rocks as well as early Paleoproterozoic A-type granites in east-central Finland.
No formal technical sessions are arranged, but a pre-conference seminar will be given on September 11,
delegates.
With great pleasure, we welcome Eurogranites to Finland. We are pleased to have the opportunity to
arrange this meeting and sincerely hope that your visit will be rewarding.
O. Tapani Rämö
Jaana Halla
Mikko Nironen
Laura S. Lauri
Matti I. Kurhila
Asko Käpyaho
Peter Sorjonen-Ward
Olli Äikäs
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Table of contents
Foreword ..................................................................................................................................................... 7
Table of contents ......................................................................................................................................... 9
PART I: GENERAL .................................................................................................................................. 13
....................................................................................................... 15
2. The Finnish Precambrian in a nutshell.................................................................................................. 16
2.1. The Archean domain ....................................................................................................................... 17
2.2. The Svecofennian domain .............................................................................................................. 17
2.3. Younger units .................................................................................................................................. 17
3. Finnish granitoids through time ............................................................................................................ 19
3.1. Archean TTGs ................................................................................................................................. 19
3.1.1. Introduction ......................................................................................................................... 19
3.1.2. Occurrence and geochemistry of selected TTG suites in Finland ....................................... 21
3.1.3. Concluding remarks ............................................................................................................. 24
3.2. Neoarchean sanukitoids (2.74–2.70 Ga) ......................................................................................... 24
3.2.1. Introduction ......................................................................................................................... 24
3.2.2. Distribution and age............................................................................................................. 24
3.2.3. Geochemistry ....................................................................................................................... 27
3.2.4. Petrogenetic remarks ........................................................................................................... 28
3.3. Neoarchean leucogranites ............................................................................................................... 29
3.3.1. Introduction ......................................................................................................................... 29
3.3.2. Key localities of Neoarchean leucogranite in the Finnish Archean..................................... 29
3.3.3. U–Pb ages ............................................................................................................................ 29
3.3.4. Geochemical constraints on origin ...................................................................................... 30
3.3.5. Concluding remarks ............................................................................................................. 32
3.4. Early Paleoproterozoic anorogenic granites ................................................................................... 33
3.4.1. Introduction ......................................................................................................................... 33
3.4.2. Description of the ~2.44-Ga silicic intrusions ..................................................................... 33
3.4.3. Elemental geochemistry ...................................................................................................... 35
3.4.4. Nd isotopes .......................................................................................................................... 37
3.4.5. Concluding remarks ............................................................................................................. 37
3.5. Intermediate magmatism at the craton margin, central Finland ..................................................... 38
3.5.1. Introduction ......................................................................................................................... 38
3.5.2. Field geology and petrography ............................................................................................ 38
3.5.3. Geochemistry ....................................................................................................................... 40
3.5.4. Concluding remarks ............................................................................................................. 40
3.6. Proterozoic synorogenic granitoids ................................................................................................ 41
3.6.1. Distribution and age............................................................................................................. 41
3.6.2. General mode of occurrence ................................................................................................ 41
3.6.3. Geochemical constraints on origin ...................................................................................... 43
3.7. Proterozoic lateorogenic granites ................................................................................................... 45
3.7.1. Introduction ......................................................................................................................... 45
3.7.2. Geochemical features .......................................................................................................... 45
3.7.3. Nd isotopes .......................................................................................................................... 46
3.7.4. U–Pb ages ............................................................................................................................ 48
3.7.5. Sources and origin of the lateorogenic granites................................................................... 48
3.8. Postorogenic intrusions................................................................................................................... 49
3.8.1. Mode of occurrence ............................................................................................................. 49
3.8.2. Geochemical constraints on origin ...................................................................................... 50
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3.9. Rapakivi granites ............................................................................................................................ 51
.................. 51
3.9.2. Distribution and age............................................................................................................. 51
3.9.3. Petrology .............................................................................................................................. 52
4. Tectonic evolution: the granitoid perspective ....................................................................................... 56
4.1. Introduction .................................................................................................................................... 56
4.2. Archean ........................................................................................................................................... 58
4.3. Proterozoic ...................................................................................................................................... 59
4.3.1. Rifting of the Archean craton .............................................................................................. 59
4.3.2. Proterozoic orogenic granitoids ........................................................................................... 60
PART II: FIELD TRIP STOP DESCRIPTIONS....................................................................................... 65
DAY 1 (Monday September 12, 2005) ..................................................................................................... 67
Stop 1/1: Topaz-bearing alkali feldspar granite and associated marginal pegmatite ............................. 67
Stop 1/2: Wiborgite at Summa, Vehkalahti, south-central part of the Wiborg batholith ....................... 69
Stop 1/3: Pyterlite at Virolahti, south-central part of the Wiborg batholith ........................................... 69
Stop 1/4: Anorthositic raft in wiborgite at Ylijärvi, Ylämaa, east-central part of the Wiborg
batholith ............................................................................................................................................. 70
DAY 2 (Tuesday September 13, 2005) ..................................................................................................... 73
Stop 2/1: Keittomäki lateorogenic granite, Juva ................................................................................... 73
Stop 2/2: The marginal type of the Puruvesi granite, Herttuansaari, Kerimäki..................................... 73
Stop 2/3: The central type of the Puruvesi granite, Rastiniemi, Kesälahti ............................................ 74
Stop 2/4: The Valkamo layered granite, Imatra ..................................................................................... 76
Stop 2/5: The Eräjärvi postorogenic granite dike, Ruokolahti .............................................................. 77
DAY 3 (Wednesday September 14, 2005) ................................................................................................ 79
Stop 3/1: Type 2 Puula pluton at Sokkasenmäki ................................................................................... 79
Stop 3/2: Synkinematic granodiorite at Kollinkangas ........................................................................... 79
Stop 3/3: Type 3 Jämsä pluton, margin.................................................................................................. 80
Stop 3/4: Type 3 Jämsä pluton, center ................................................................................................... 81
Stop 3/5: Type 3 Jämsä pluton, evolved center ..................................................................................... 81
DAY 4 (Thursday September 15, 2005) ................................................................................................... 83
Stop 4/1: Microtonalite dikes at Kivennapa .......................................................................................... 83
Stop 4/2: The Pisa augen gneisses— Paleoproterozoic deformation of K-feldspar megacrystic
Neoarchean sanukitoids .................................................................................................................... 84
Stop 4/3: Neoarchean sanukitoids of the Lieksa complex in the Ilomantsi terrain ............................... 88
DAY 5 (Friday September 16, 2005) ........................................................................................................ 91
Stop 5/1: Arola leucogranite .................................................................................................................. 91
Stop 5/2: Arola granodiorite .................................................................................................................. 91
Stop 5/3: Kuusamonkylä tonalitic gneiss .............................................................................................. 92
Stop 5/4: ~2.43-Ga A-type granite at Tuliniemet .................................................................................. 93
Stop 5/5: Kaihlankylä migmatite ........................................................................................................... 93
DAY 6 (Saturday September 17, 2005) .................................................................................................... 95
PART III: ACKNOWLEDGMENTS AND REFERENCES .................................................................... 97
ACKNOWLEDGMENTS ........................................................................................................................ 99
REFERENCES ....................................................................................................................................... 101
PART IV: APPENDIX ............................................................................................................................ 113
Field trip stop coordinates ....................................................................................................................... 115
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PART V: NOTES .................................................................................................................................... 117
Day 1 ....................................................................................................................................................... 119
Day 2 ....................................................................................................................................................... 121
Day 3 ....................................................................................................................................................... 123
Day 4 ....................................................................................................................................................... 125
Day 5 ....................................................................................................................................................... 127
Day 6 ....................................................................................................................................................... 129
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PART I: GENERAL
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(O.T. Rämö)
The Fennoscandian (or Baltic) shield of Finland,
Sweden, Norway, and northwestern Russia (Fig. 1)
has been at the focus of active petrological research
since the late 19th century and is one of the bestknown Precambrian shield areas in the world (e.g.,
Sederholm, 1932; Rankama, 1963; Gorbatschev,
1993; Vaasjoki et al., 2005). The glacially polished
bedrock of the shield has provided ample sample
material for petrological and geochemical studies.
Since the pioneering works of J.J. Sederholm (Director of the Geological Survey of Finland 1892–
1933) and Pentti Eskola (Professor of Geology and
Mineralogy of the University of Helsinki 1924–
1953), the knowledge of the Precambrian bedrock
of Finland has evolved into a remarkably detailed
picture of Paleoproterozoic, and also Archean, crust
(e.g., Korsman et al., 1997, 1999; Lehtinen et al.,
2005).
Granitoid rocks are important constituents of the
Fennoscandian shield, and form more than half of
the Finnish bedrock. The Proterozoic granites of
Ga), lateorogenic (~1.83 Ga), and postorogenic
Regarding Archean granitoids of east-central and
northeastern Finland, much still needs to be done
in order to reach the level of knowledge currently
available for the Proterozoic granitoid rocks. New
data are, however, continuously emerging and these
will form the essence of the Archean targets that
nalite–trondhjemite–granodiorite (TTG) associations and sanukitoid suites as well as Neoarchean
leucogranites.
The classic mid-Proterozoic rapakivi granites of
southern Finland have been the subject of active
research since the late 19th century (Sederholm,
1891), and during the last thirty years in particular
(see, e.g., Rämö and Haapala, 1995, 1996; Haapala
et al., 2005). Extensive mineralogical, geochemical, geochronological, and geophysical data are
rocks and they have allowed a comprehensive assessment of the petrogenesis of these rocks. Many
details regarding the interplay of crustal accretion
and granitoid magmatism in this part of the shield
are, however, still open and will be addressed dur-
as synkinematic, latekinematic, and postkinematic
according to their relation to tectonic movements
during the Svecofennian (or Svecokarelian) orogeny. Sederholm (1932) suggested a four-fold division, with the two earlier granitoid suites (Group
I and Group II) corresponding to Eskola’s synkinematic and latekinematic granites; Group III
consisting of small discordant plutons, and the
rapakivi granites comprising Group IV. Simonen
The principal aim of Eurogranites 2005 in Finland is to bring together an enthusiastic group of
colleagues to examine and discuss the origin and
rocks along a circa 1500-km-long traverse (see inside covers) from the south coast (Helsinki area) to
east-central Finland (Kuhmo-Suomussalmi area).
postorogenic and considered the rapakivi granites
as anorogenic.
A remarkably comprehensive U–Pb zircon data
base (e.g., Vaasjoki , 1996) has constrained the ages
of the Proterozoic synorogenic plutonic rocks at
1.93–1.87 Ga, lateorogenic at 1.84–1.82 Ga, postorogenic at 1.81–1.79 Ga, and rapakivi granites
at 1.65–1.54 Ga (Vaasjoki, 1996). Syn-, late-, and
postorogenic may still be used as age groups in the
Finnish Svecofennian although these terms cannot
be applied to the whole Svecofennian orogen because of overlapping ages (e.g., Nironen, 1997). The
current division (Nironen, 2005) divides the Finnish Proterozoic orogenic granitoid rocks into four
groups: preorogenic (~1.93 Ga), synorogenic (~1.88
der, (1) the classic ~1.64-Ga rapakivi granites and
syn-, late-, and postorogenic granites of the ~1.9Ga Svecofennian orogen; and (3) the K-feldspar
megacrystic sanukitoids (high-Mg granitoids),
TTG suites, and A-type granites of the late Archean Karelian craton. All of the latest data from
geochronologic, isotope geological, geochemical,
petrological, and regional tectonic studies will be
available for discussion, with a particular focus on
comparisons and contrasts with Phanerozoic granitoid-forming processes.
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2. The Finnish Precambrian in a nutshell
(L.S. Lauri, M. Nironen, O.T. Rämö)
Norway. Most of the Proterozoic rocks in Finland
and Sweden were formed between 1.89–1.82 Ga.
The crustal processes that caused this magmatism
have traditionally been ascribed to the Svecofen-
The Fennoscandian shield (Fig. 1) is one of the
largest Precambrian shield areas in the world. It
forms the northernmost part of the Precambrian
East European craton that is mostly covered by
Paleozoic sedimentary rocks. The shield consists of
both Archean and Proterozoic rocks. Archean rocks
are exposed in northwestern Russia, eastern and
northern Finland, and northern Sweden. They form
the Archean domain that consists of three Meso- to
Neoarchean terrains, the Kola province, the Belomorian province, and the Karelian province. Juvenile Proterozoic rocks form the bedrock of central and southern Finland, Sweden, and southern
the Svecofennian domain (Gaál and Gorbatschev,
1987). Younger Proterozoic rocks in the shield comprise the 1.85–1.67-Ga Transscandinavian igneous
belt and the 1.75–1.50-Ga Southwest Scandinavian
domain in Sweden and Norway. The northwestern
margin of the shield is buried beneath and tectonically interleaved with the Early to Middle Paleozoic
Caledonides.
Fig. 1. Geology and major structural units of the Fennoscandian shield. PAC— Primitive arc complex; WAC— Arc complex
of western Finland; SAC— Arc complex of southern Finland; CFGC— Central Finland granitoid complex; CLGC— Central
Lapland granite complex; LGB— Lapland granulite belt. Right-diagonal ruling marks the northern edge of platform sediments.
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2.1. The Archean domain
Archean rocks underlie the Proterozoic supracrustal rocks, the Central Lapland granite complex
(CLGC), and the Lapland granulite belt (LGB) that
form the bedrock in northern and eastern Finland (Fig.
1). The oldest Proterozoic rocks within the Archean
The Archean domain comprises extensive tonalite–
trondhjemite–granodiorite (TTG) areas, remnants of
granite–greenstone complexes, and Proterozoic rocks
that have deposited on or intruded into the Archean
crust. Archean rocks crop out in the northeastern half
of Finland, northeast of the line connecting the tip
of the Gulf of Bothnia and the Lake Ladoga in Russia (Fig. 1). The largest continuous Archean areas are
found in the central and eastern parts of the country
but smaller, scattered occurrences are present also
in the northern parts. The ages of the Archean rocks
in Finland range from rare occurrences of >3.0 Ga
old rocks (e.g., the 3.5-Ga Siurua gneiss— the oldest
rock so far found in Europe; Mutanen and Huhma,
2003) to the major population of ~2.9 Ga to 2.7 Ga
old rocks (e.g., Vaasjoki et al., 1999, and references
therein).
in north-central Finland (e.g., Alapieti et al., 1990).
The supracrustal belts that have been deposited on
the Archean craton range in age from ~2.4 Ga to 1.9
Ga and contain both volcanic and sedimentary sequences (e.g., Lehtonen et al., 1998; Vaasjoki, 2001;
Kohonen and Marmo, 1992; Kohonen, 1995). The
central Lapland granites and the granulite belt are the
result of continent–continent collision between the
Karelian province and the Kola province at ~1.9–1.8
Ga (Daly et al., 2001).
2.2. The Svecofennian domain
The main feature in the Svecofennian domain is
the large (40,000 km2) Central Finland granitoid
complex (CFGC) that is surrounded by volcanic–
sedimentary belts (Fig. 1). The 1.92-Ga primitive
island arc rocks in central Finland, between the
Archean craton and the Central Finland granitoid
complex, are the oldest documented rocks in the
Svecofennian domain of Finland. However, isotope geological data suggest that older (~2.1–2.0
Ga) rocks may have formed the nucleus of the
complex (Lahtinen and Huhma, 1997; Rämö et al.,
2001a). Metagraywacke-dominated sedimentary
rocks and island arc-type volcanic rocks (1.90–
1.87 Ga), intruded by calc-alkaline granitoids,
comprise the schist belts in central and western
Finland whereas metapelite-dominated sedimentary rocks, quartzites, and carbonate rocks characterize the supracrustal rocks of southern Finland.
The mainly felsic to intermediate plutonic rocks
of the Central Finland granitoid complex and adjacent supracrustal belts have yielded ages in the
range of 1.89–1.87 Ga (Vaasjoki, 1996). Migmatites with tonalite leucosome were formed from
immature psammites 1.89–1.88 Ga ago, whereas
younger migmatization, associated with S-type
granites, took place at ~1.84–1.82 Ga in southern
Finland and formed the Late Svecofennian granite–migmatite zone (Korsman et al., 1999).
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). The present concept (Korsman et al., 1997)
is that the Finnish Svecofennian consists of three
arc complexes (Fig. 1): (1) the Primitive arc complex (PAC); (2) the Arc complex of western Finland (WAC); and (3) the Arc complex of southern
Finland (SAC).
2.3. Younger units
The youngest major rock units in the Finnish bedrock are the mid-Proterozoic rapakivi granites that
comprise four large batholiths and several smaller
batholiths or stocks, as shown in Fig. 2. The U–Pb
zircon ages of the rapakivi granites of Finland range
from 1.65 Ga to 1.54 Ga (Vaasjoki, 1977; Vaasjoki et
al., 1991; Suominen, 1991). The intrusions cut sharply across the Svecofennian medium- to high-grade
15
Fig. 2. Geological sketch map of southern Finland and vicinity showing the distribution of various Precambrian lithologic units
for dolerite dikes that cut sandstones in a fault-bounded basin in the Satakunta area (Fig. 3A). However, in
the clastic sequence in the Lake Ladoga basin of Russian Karelia (Fig. 3B) a monzodioritic sill has been
dated at 1457 ± 3 Ma (Rämö et al., 2001b). These two
chronologic piercing points yield maximum (~1.46
Ga) and minimum (~1.26 Ga) ages for the duration
of post-rapakivi basin sedimentation.
Small-scale igneous activity occurred in northern
Finland at ~1.1–1.0 Ga when several diabase dikes
were emplaced in the Salla and Laanila areas (Lauerma, 1987; Pihlaja, 1987). The ~400-Ma Caledonian domain also extends to the northwesternmost tip
of Finland (Fig. 1). The Caledonian area in Finland,
although very small, contains the highest point of
the country (Halti, 1328 m a.s.l.). The present erosional level of the Fennoscandian shield was attained
already in the Neoproterozoic times. The shield was
covered by sedimentary rocks in the Paleozoic but
only a few small, scattered occurrences of Cambrian
sandstones and Ordovician limestones have survived
the recent glaciations that again exposed and polished
the ancient bedrock of the shield.
metamorphic and plutonic rocks and produce contact
metamorphic aureoles (Vorma, 1972). Lithologic,
structural, and textural features (e.g., sharp intrusive
contacts, local presence of volcanic or subvolcanic
members, roof pendants, roof breccia outcrops) suggest that the present erosion level generally represents the upper parts of epizonal intrusive complexes
(Wahl, 1925, 1947; Vorma, 1975; Haapala, 1977a;
Bergman, 1986).
The rapakivi granite intrusions are spatially and
bro–anorthosite complexes, basaltic dikes, and rare
monzonitic/ferrodioritic rocks). They are also spatially intimately associated with Mesoproterozoic
(Jotnian/Riphean) clastic basins of redbed-type sandmatism (gabbroic sills and dikes, basaltic lavas; Fig.
2). The time of the onset of the basin formation is not
known in detail and thus the temporal relation of the
rapakivi granite intrusions and basin sedimentation
and basaltic magmatism is therefore unconstrained.
U–Pb mineral data indicate ages in the 1268–1256
Ma range (Suominen, 1991; Söderlund et al., 2004)
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al. (1996).
3. Finnish granitoids through time
3.1. Archean TTGs
(A. Käpyaho, L.S. Lauri)
3.1.1. Introduction
models were put forward by Martin (1987a), who
concluded that the TTGs of the Kuhmo district,
eastern Finland, were formed by partial melting of
Archean tholeiites and that garnet and hornblende
were necessary residual minerals. The genesis of
TTG magmas has also been attributed to melting of
The oldest known felsic crustal rocks in Finland
are mainly members of the tonalite–trondhjemite–
granodiorite (TTG) association. In general, these
rocks provide one of the most informative data
sources of the early Earth. The oldest TTGs so
far reported are the ~4.0-Ga Acasta gneisses from
the Slave craton, Canada (Bowring and Williams,
1999). Early Archean (>3.5 Ga) TTGs are also
known from, e.g., Greenland (Nutman et al., 1996),
South Africa (Kröner et al., 1996), the United States
(Mueller at al., 1996), and Finland (Mutanen and
Huhma, 2003).
The petrogenesis of the TTGs is controversial
and has received considerable attention in recent
years. Experimental petrology has shown that the
Archean TTGs may be produced by partial melting
of partially hydrated metabasalts, leaving a garnetbearing residue (Rapp and Watson, 1995). Similar
(Kröner, 1985; Condie, 2005). Foley et al. (2002)
confronted this view on the basis of amphibole trace
element systematics and pointed out that the Nb/Ta
and Zr/Sm ratios in the Archean TTGs are inconsistent with an eclogite residue. However, experimental evidence shown by Rapp et al. (2003) indicate that eglocite residue may be stable with melts
resembling the TTGs.
Martin (1999) noted that modern adakites and
Archean TTGs are remarkably similar in geochemistry and thus could have been produced in a similar
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Some locations mentioned in the text are denoted. The Kuhmo greenstone belt is shown in black. Sanukitoid zones (see Section
3.2) are marked with dotted lines: K— Kuusamo (Halla, in prep.); N— Nilsiä (Halla, 2005); ESZ and WSZ— eastern and west-
manner. Modern adakites are found in a subduction
setting where a young and hot oceanic slab melts,
and the melt is further contaminated with mantle
wedge material (Kay, 1978). However, high Mg#
and elevated Ni and Cr contents that would be
expected in these melts owing to interaction with
mantle wedge are seldom observed in the Archean
TTGs (cf. Smithies, 2000).
The ongoing debate on the origin of the TTGs (cf.
Rollinson and Martin, 2005) certainly renders these
rocks a healthy ground for further elaborations.
This review of Archean TTGs is far from comprehensive— our aim is to summarize some geochemical data from selected key localities in the Finnish
Archean, concentrating on the Karelian province
of the Fennoscandian shield. Some granodiorites,
18
tonalites, and enderbites (e.g., from Ilomantsi and
Halla in Section 3.2.
3.1.2. Occurrence and geochemistry of selected
TTG suites in Finland
Siurua
The trondhjemite gneiss from Siurua, north-central Finland, is the oldest rock unit so far recognized
in the Fennoscandian shield (Mutanen and Huhma,
2003; cf. Fig. 4). It is a gneissose tonalite within
a granulite-facies block in the Archean Pudasjärvi
complex. The chemical composition of the Siurua
trondhjemite gneiss (Table 1; Sample A1602; Mutanen and Huhma, 2003) has SiO2 and Mg# close to
the average Early Archean TTG of Condie (2005)
(Fig. 5). The rock is slightly peraluminous (Fig. 6)
and shows strong enrichment in the LREE compared to the HREE (Fig. 7). The single grain U–Pb
zircon age of the sample (3.5 Ga) is close to the
(DePaolo, 1981) TDM model age (3.48 Ga) and initial εNd (at 3.5 Ga) value is +1.4 (Fig. 8) (Mutanen
and Huhma, 2003).
Fig. 5. SiO2 (wt.%) vs. Mg# [Mg2+/(Mg2++Fetot)*100, Fetot as
Fe2+] diagram showing the composition of the Siurua tonalite
gneiss (Mutanen and Huhma, 2003) and pyroxene tonalite
(sample A1661) from Koillismaa (Lauri et al., 2005), and
gneisses (Jahn et al., 1984). Mean values of Early and Late
after Smithies (2000).
Tojottamanselkä
Tojottamanselkä in northern Finland is another
occurrence of >3.0-Ga rocks in the Fennoscandian
shield (Fig. 4). It is a small (1 km by 2 km) Archean
gneiss dome within Paleoproterozoic supracrustal
rocks (Puustinen, 1977) and consists of tonalitic
and trondhjemitic gneisses. The Tojottamanselkä
gneisses have relatively low Mg number and plot
Fig. 6. SiO2 (wt.%) vs. A/CNK [molar Al2O3/
(CaO+Na2O+K2O)] diagram showing the composition of
TTGs of the Karelian province. See Fig. 5 for data sources.
metaluminous to peraluminous with increasing
SiO2 values (Fig. 6) and show strong enrichment of
the LREE compared to the HREE and commonly
have a small positive Eu anomaly (Fig. 7). Singlegrain U–Pb zircon datings on the Tojottamanselkä
gneisses yielded an intrusive age of 3115 ± 29 Ma
(Kröner and Compston, 1990). Initial εNd values of
the Tojottamanselkä gneisses are mostly negative
(~ –3.5) but more radiogenic (depleted mantle-like)
values are also observed (Fig. 8; Jahn et al., 1984;
Hanski et al., 2001). Jahn et al. (op. cit.) suggested
that the compositional variation in Tojottamanselkä
implies several source components; some tonalites
may represent partial melts from basaltic sources
that have nearly chondritic REE compositions,
whereas some (those with a distinctly negative ini-
Fig. 7. Chondrite-normalized REE patterns of selected TTGs
from the Karelian province compared to the mean values
reported for the Early and Late Archean TTGs and adakites
by Condie (2005). See Table 1 for references. Normalizing
values (C1 chondrite) from Sun and McDonough (1989).
19
older crustal components similar to the Siurua or
Koillismaa samples (Fig. 8). There are indications
of TDM model ages in excess of 3.3 Ga from the
Naavala tonalite in the Kuhmo district (Luukkonen,
2001).
tial Nd) represent partial melts of older continental
crustal material.
Kuhmo
In the Kuhmo district, eastern Finland (Fig. 4),
TTGs are found as distinct, relatively homogeneous
plutons and metatexitic and diatexitic migmatitic
gneisses. Martin (1987a) summarized geochemical
and age data on the gray TTG gneisses of Kuhmo.
The ages he reported are, however, based on the
Rb–Sr whole-rock method and, as indicated by Vaasjoki (1988), do not represent crystallisation ages of
the plutons. Vaasjoki et al. (1999) and Luukkonen
(2001) reported U–Pb zircon ages between 2.83 Ga
and 2.75 Ga for tonalites in the Kuhmo district.
The TTGs in Kuhmo have SiO2 between 64 wt.%
and 74 wt.% and Mg number between 59 and 31
(Table 1; Fig. 5). They are mostly peraluminous,
samples with the lowest SiO2 tend to be metaluminous (Fig. 6). The REE patterns of the rocks from
Kuhmo resemble average adakite, and show moderate enrichment of the LREE relative to the HREE
(Fig. 7), thus implying garnet-bearing residue. The
2.83 Ga and 2.75 Ga tonalitic gneisses analysed
from the Kuhmo district indicate no contribution of
Koillismaa
TTGs form a large part of the Archean bedrock in
Koillismaa, eastern Finland (Fig. 4). They are variably deformed granitoids and gneisses in which the
metamorphic grade extends up to granulite facies
(Lauri et al., 2005). A granulite-facies pyroxene
tonalite from Koillismaa (Table 1; sample A1661;
Lauri et al., 2005) plots on the border of adakite
0.93, corresponding to the samples from Kuhmo
that have similar SiO2 values (Fig. 6). TTGs in
Koillismaa are enriched in the LREE with (La/Yb)N
of 54 (Fig. 7). Conventional U–Pb zircon analyses yielded an age of 2808 ± 20 Ma for the sample
A1661 (Lauri et al., 2005). The ~2.8-Ga TTGs from
Koillismaa have initial Nd values from –1.9 to –0.8
bution from an older crustal component (Fig. 8). A
sample of migmatite paleosome from the Kuhmo
Fig. 8. εNd vs. age diagram for the TTGs from the Karelian province. Key to references: 1— Mutanen and Huhma (2003);
2— Käpyaho et al. (in review); 3— Luukkonen (2001); 4— Lauri et al. (2005); 5— Hanski et al. (2001). DM (depleted mantle)
evolution line is from DePaolo (1981), CHUR denotes the Chondritic Uniform Reservoir (DePaolo and Wasserburg, 1976).
20
Table 1
Selected analyses of TTGs
Siurua (1) Kuusamon- Naavala (2) Kivijärvi (2) Koillismaa
kylä (2)
(3)
Tojottamanselkä (4)
Early Archean Late Archean
TTG (5)
TTG (5)
A1602
H52
H413
H42
A1661
average
(n = 7)
average
(n = 212)
average
(n = 831)
SiO2
70.3
70.5
72.9
72.1
64.6
68.8
70.4
68.3
TiO2
0.45
0.37
0.20
0.15
0.71
0.48
0.31
0.42
Sample
Al2O3
15.6
14.6
15.0
14.4
16.3
15.4
15.2
15.5
Fe2O3tot
2.33
2.46
1.52
1.94
4.93
3.38
2.79
3.42
MnO
0.02
0.04
0.01
0.03
0.07
0.04
0.06
0.07
MgO
0.92
1.47
0.47
0.70
1.87
0.85
0.96
1.39
CaO
3.17
2.23
2.26
2.09
4.64
3.14
2.74
3.26
Na2O
4.48
4.85
4.30
4.03
4.44
4.80
4.71
4.51
K2 O
1.83
1.81
2.72
2.82
1.66
2.10
2.22
2.20
P2O5
0.03
0.10
0.06
0.07
0.19
0.12
0.10
0.14
S
<
-
-
-
67
-
-
-
F
600
-
-
-
-
-
-
-
Cl
170
-
-
-
200
-
-
-
Cu
<
-
-
-
30
-
-
-
Ga
22
-
-
-
24
-
-
-
Zn
52
-
-
-
66
-
-
-
Cr
<
-
-
-
31
-
45
35
Ni
8
-
6
3
24
-
17
22
Co
5.6
-
4
5
-
-
-
-
V
32
-
-
-
73
-
-
-
Sc
3.6
-
-
-
9.77
-
-
-
Sr
311
255
320
238
516
-
362
515
Ba
503
408
740
-
830
-
500
769
Rb
57
92
89
90
41
-
76
67
Zr
294
136
108
122
270
-
152
154
Hf
6.9
-
-
-
-
-
3.8
4.7
Ta
<
-
-
-
-
-
0.41
0.84
Nb
5.2
-
-
-
3
-
6.1
6.2
Pb
32
-
-
-
21
-
-
-
U
2.0
-
-
-
0.22
-
1.20
1.5
Th
46
-
-
-
5.11
-
4.10
8.1
Y
7.5
7
3
9
11.8
-
8.5
9.1
La
123.0
21.5
7.5
56.2
61.1
15.66
22
36
65
Ce
213.0
39.4
14.3
92
105.0
30.32
40
Pr
22.2
-
-
-
10.8
-
-
-
Nd
75.8
13.69
-
-
36.0
10.52
16
25
Sm
10.0
1.8
0.95
4.13
5.00
1.86
2.9
4.2
Eu
1.35
0.48
0.56
0.94
1.27
0.63
0.82
1.07
Gd
6.77
1.24
-
-
3.88
1.46
2.2
2.9
Tb
0.59
-
0.09
0.35
0.51
-
0.31
0.38
Dy
1.90
0.73
-
-
2.43
1.20
-
-
Ho
0.27
-
-
-
0.41
-
-
Er
0.60
0.37
-
-
1.28
0.62
-
-
Tm
<
-
-
-
0.15
-
-
-
Yb
0.55
0.38
0.32
0.69
0.81
0.62
0.82
0.71
Lu
<
0.063
0.08
0.076
0.14
0.10
0.14
0.11
Mg#
43.9
54.2
38.0
41.7
42.9
33.20
40.8
46.2
A/CNK
1.03
1.04
1.06
1.07
0.93
0.97
1.00
0.99
References: (1) Mutanen and Huhma (2003); (2) Martin (1987a); (3) Lauri et al. (2005); (4) Jahn et al. (1984); (5) Condie (2005).
Oxides in wt.%, trace elements in ppm. < below detection limit, - not analyzed.
21
jor and trace element and Nd isotope composition it
seems likely that these TTGs were produced from
varying sources in various geotectonic environments. Figure 5 shows that most of the TTGs in the
district has a rather similar Nd (at 2.8 Ga) value and
could thus represent one possible source for the
Koillismaa tonalites.
3.1.3. Concluding remarks
supporting the argument of Smithies (2000) that
most Late Archean TTGs do not show evidence of
adakite-like mantle wedge contamination (see also
Section 3.2).
The Karelian province of the Fennoscandian
shield implies TTG emplacement between 3.5 Ga
and 2.75 Ga. Because of large variation in the ma-
3.2. Neoarchean sanukitoids (2.74–2.70 Ga)
(J. Halla)
3.2.1. Introduction
luminous TTGs in Archean granite–greenstone terrains. Sanukitoids have been recognized from the
Superior province, Canada (Shirey and Hanson,
1984, 1986; Stern et al., 1989; Stern and Hanson,
1991; Beakhouse et al., 1999; Stevenson et al.,
1999), the Fennoscandian shield (Lobach-Zhuchenko et al., 2005; Kovalenko et al., 2005; Halla, 2002,
2005), the Dharwar craton, South India (Sarvothaman, 2001; Moyen et al., 2003), the Ukrainian
shield (Artemenko et al., 2003), the Pilbara craton,
Western Australia (Smithies and Champion, 2000),
the Zimbabwe craton (Kampunzu et al., 2003), and
Greenland (Steenfelt et al., 2005).
The Archean Karelian craton of the Fennoscandian shield is divided into three domains with different crust-formation ages (Lobach-Zhuchenko et al.,
2000a): West Karelian (~3.0 Ga), Central Karelian
(~2.85 Ga), and southeastern Vodlozero (~3.0 Ga).
Lobach-Zhuchenko et al. (2005) divided sanukitoids
of the Karelian granite–greenstone terrain into two
distinct zones of different age and composition,
roughly coinciding with the domain boundaries: the
western sanukitoid zone (WSZ) in the West Karelian domain and the eastern sanukitoid zone (ESZ)
in the Central Karelian domain (Fig. 4). Bibikova et
al. (2005) showed that the sanukitoids in the Karelian craton were formed in two narrow time intervals, at ~2705 Ma (western part) and ~2745 Ma
(eastern part). Thus there is a 30-Ma age difference
between the western and eastern sanukitoid zones
of the Karelian craton. The time gap between TTG
and sanukitoid formation is shorter in the west (60
Ma) than in the east (100 Ma).
In the westernmost part of the Karelian craton in
The rate of magma production and continental
growth was especially rapid during the Neoarchean
time. At around 2.7 Ga, large amounts of new continental crust were formed and accreted into a single
supercontinent, Kenorland, or several supercratons
such as Superia and Sclavia (e.g., Bleeker, 2003).
One of the key factors for understanding crustal
evolution 2.7 b.y. ago is the petrogenesis of sanukitoids, a series of high-Mg granitoids found in Archean terrains. Sanukitoids are narrowly restricted
in time (generally 2.74–2.72 Ga) and were formed
during the later stages of Neoarchean cratonization.
Knowing the origin of sanukitoids is fundamental
to the understanding of the genesis of late Archean
cratons as they provide a link between crust formation and mantle processes.
The term sanukitoid is used as a synonym for
Neoarchean high-Mg granitoids referring to a series
of granitoid rocks having relatively high Mg numbers and high Ni, Cr, LILE (Sr, Ba, P), and LREE
abundances at any given silica content (sanukitoid
nadian shield by Shirey and Hanson (1984), who introduced the term sanukitoid as the rocks resemble
high-Mg andesites termed sanukites in the Setouchi
area of Japan.
3.2.2. Distribution and age
The oldest sanukitoids (2.95 Ga) have been found
in the granite–greenstone terrain of the Pilbara craton, Western Australia (Smithies and Champion,
2000) and most of the known sanukitoids were
formed at ~2.7 Ga. Sanukitoids are late- to postkinematic (e.g., Shirey and Hanson, 1984; Stern et
al., 1989; Beakhouse et al., 1999) and intrude vo-
series granodiorites and monzodiorites are found
in the Nilsiä area of the Iisalmi terrain and in the
22
Fig. 9. Plots of selected element pairs for sanukitoids from eastern Finland (Nilsiä, Lieksa, Kuittila, Arola), the western (WSZ)
and eastern (ESZ) sanukitoid zones of the Karelian craton, and the Western Superior province, Canada. Data for TTGs in Finland are plotted for comparison. References as in Table 2.
23
Table 2
Comparison of selected major and trace element analyses of Karelian sanukitoids (SiO2 62–64 wt.%) with sanukitoids from the
Superior province and TTGs in Finland
Nilsiä1
Lieksa2
Kuittila3
Arola4
Kurgelampi
(WSZ)5
Elmus
(ESZ)6
Superior7
TTG Fi8
69.60
SiO2
62.90
63.50
63.00
61.91
62.25
62.51
66.40
TiO2
0.63
0.54
0.45
-
0.77
0.32
0.47
0.38
Al2O3
16.20
16.70
16.70
15.61
15.73
15.86
15.50
15.20
Fe2O3tot
4.76
4.63
3.68
5.02
5.68
3.43
3.25
2.79
MnO
0.07
0.07
0.07
0.08
0.09
0.06
0.06
0.04
MgO
2.27
2.32
2.24
2.97
4.07
2.65
1.23
1.17
CaO
3.66
3.83
4.42
3.84
4.72
3.74
2.75
2.63
Na2O
5.06
4.52
4.54
5.24
3.92
4.52
4.75
4.45
K2O
3.41
3.23
2.37
2.72
2.36
4.52
3.83
2.23
P2O5
0.41
0.27
0.14
0.26
0.40
0.37
0.23
0.10
70
70
86
92
99
45
43.4
-
V
Cr
50
70
64
131
66
224
27
-
Ni
30
20
25
27
24
65
23.1
11
Rb
97.6
75.7
110
70
73
122
113
83
Sr
610
840
990
893
668
1200
1020
351
Y
26.8
12.2
20
10
12
16
13.1
9
152
Zr
289
170
100
156
132
242
198
Nb
9.46
3.93
20
-
5
8
8
-
Ba
1200
1730
820
1307
1074
2259
1430
700
Hf
6.7
4.24
-
-
3.5
6.2
4.79
-
Ta
0.45
0.20
0.30
-
0.25
0.50
0.76
-
Pb
12.7
23.2
8
-
20
51
30.3
-
Th
9.12
6.12
4
-
10
36
17.3
-
U
0.90
0.31
-
-
-
-
3.25
-
La
66.1
47.7
21.7
55.4
36.3
111.2
78.2
34.7
Ce
135.0
90.4
46.0
107.0
73.7
223.8
153
61.4
Pr
15.5
10.5
8.18
24.7
-
-
Nd
54.2
38.6
18.0
44.0
35.64
93.6
52.1
17.4
Sm
8.46
6.41
3.4
7.6
6.03
12.5
7.86
2.5
Eu
1.64
1.53
1.00
1.59
1.49
3.00
1.68
0.8
Gd
6.07
4.78
-
3.57
4.08
8.90
5.13
1.74
Tb
0.74
0.6
0.30
0.40
0.53
0.80
-
-
Dy
2.64
2.33
-
-
2.30
3.10
2.36
0.91
Ho
0.47
0.4
-
-
0.44
0.50
0.43
0.39
Er
1.14
1.08
-
-
1.13
1.10
1.17
Tm
0.15
0.13
-
-
0.17
0.10
0.14
-
Yb
1.01
0.92
0.80
0.95
0.99
0.90
0.90
0.35
Lu
0.14
0.13
0.20
0.12
0.14
0.10
-
0.06
Mg#
0.49
0.50
0.52
0.54
0.56
0.58
0.43
0.45
Fe#
0.65
0.64
0.60
0.60
0.56
0.54
0.70
0.68
A/CNK
0.87
0.93
0.92
0.84
0.89
0.83
0.91
1.05
1
Sample PK-121 (Halla, 2005)
Sample PK-50 (Halla, 2005)
3
Sample P498/6.2 (O’Brien et al., 1993)
4
Sample G 214 (Querré, 1985)
5
Sample 39/1-98 (sample 36/3-98 for REE), Kurgelampi, WSZ (Lobach-Zhuchenko et al., 2005)
6
Sample 189, Elmus, ESZ (Lobach-Zhuchenko et al., 2005)
7
Average of eight samples from the Western Superior Province, Canada (Stevenson et al., 1999)
8
Average of 48 samples from eastern Finland (Martin, 1987a)
- not analyzed
2
24
Ilomantsi terrain. Sanukitoids are also found in the
Kuhmo–Suomussalmi granite–greenstone belt in the
Kianta terrain and in the Kuusamo area in the Koillismaa terrain. Sanukitoids of the Ilomantsi terrain
comprise both a lower crustal level, high metamorphic grade sanukitoid complex in the Lieksa area
and smaller, upper crustal level, lower metamorphic
grade sanukitoid plutons, e.g., the Kuittila pluton
(O’Brien et al., 1993) in the easternmost part of the
Ilomantsi terrain. The Lieksa sanukitoids consist of
porphyritic K-feldspar megacrystic granodiorites
and gneisses, which produce positive anomalies on
the aeromagnetic maps. The rare occurrence of pyroxene and granulite-facies supracrustal enclaves
together with the low U content and low radiogenic
Pb isotopic composition of the granitoids imply
an extensive high-grade granulite–gneiss terrain.
The conventional U–Pb zircon age for the Nilsiä
sanukitoids in the Iisalmi terrain is 2727 ± 34 Ga
(Paavola 1984), and the age obtained for the Lieksa
sanukitoids is 2733 ± 29 Ga (Halla, 2002). The Kuittila tonalites in the eastern Ilomantsi terrain have a
U–Pb zircon age of 2745 ± 10 Ma (O’Brien et al.,
1993; Vaasjoki et al., 1993).
lected major and trace element analyses of Karelian sanukitoids (SiO2 62–64 wt.%) from eastern
Finland (Nilsiä, Lieksa, Kuittila), the Western sanukitoid zone (Kurgelampi), and Eastern sanukitoid
zone (Elmus) with sanukitoids from the Western
Superior Province, Canada, and TTGs in Finland.
between sanukitoids and TTGs. The sanukitoids in
the Karelian craton are similar to the sanukitoids of
the Superior Province, Canada. Their A/CNK index
is lower than 1 indicating a metaluminous composition. The western sanukitoid group in Karelia comprises more homogeneous single-phase intrusions
higher in SiO2 and lower in K, Na, Sr, Ba, and LREE
relative to the eastern sanukitoids, whereas the latter are strongly differentiated poly-phase intrusions
showing lower SiO2 contents and a more profound
enrichment in K and Na compared with the western sanukitoids. These geochemical differences are
attributed to different degrees of enrichment of the
source regions in the mantle wedge. Figure 10 compares the REE patterns of the two granitoid groups
and shows that the sanukitoids from eastern Finland
(except for the Kuittila pluton) have REE patterns
very similar to those of Kurgelampi (WSZ) and the
Superior province, differing clearly from the Elmus
(ESZ) sanukitoids with higher LREE contents and
TTGs with lower HREE contents.
According to Lobach-Zhuchenko et al. (2000b)
and Kovalenko et al. (2005), the sanukitoid intrusions from the younger Central Karelian domain of
the Karelian craton have positive initial εNd values
of +0.7 to +2.1 and TDM model ages of 2.70–2.85
Ga. The intrusions from the older West Karelian
3.2.3. Geochemistry
The sanukitoid series includes diorites, monzodiorites, granodiorites, monzonites, syenites, and
alkaline granites. The sanukitoids from the Iisalmi
and Ilomantsi terrains in eastern Finland are mainly
granodiorites or monzodiorites and show similar
sanukitoid-type geochemical characteristics: low
SiO2 contents (62.7–67.0 wt.%), high K2O contents (2.40–4.73 wt.%), high Mg numbers (45–52),
fractionated REE patterns [(La/Yb)N = 19–65)],
high HREE contents relative to TTGs, negative Eu
anomaly (Eu/Eu* = 0.53–0.81), and high Cr (40–80
ppm), Ba (1200–2300 ppm), Sr (610–850 ppm),
and P2O5 (0.25–0.42 wt.%) contents. They show
strong long-term depletion in U, high Th/U, low
U/Pb, high µ values for the source (~9), positive εNd
(at 2.7 Ga) values of +0.3 to +1.4, and TDM model
ages (DePaolo, 1981) of 2.75–2.86 Ga. The high µ
and apparently contradict the positive εNd (at 2.7
Ga) values that support a mantle origin.
The enrichment of sanukitoids in Mg, Cr, and Ni
suggests a peridotitic mantle-wedge source. Compared with TTGs, sanukitoids are also strongly
enriched in the LREE, Ba, Sr, and P, which points
to an enriched mantle source. Table 2 compares se-
Fig. 10. Chondrite-normalized REE patterns for sanukitoids
from eastern Finland (Nilsiä, Lieksa, Kuittila, Arola),
Kurgelampi (WSZ), Elmus (ESZ), the Western Superior
province, Canada, and for TTGs in Finland. See Table 2 for
references.
25
domain have initial εNd values of –1.7 to +0.7 and
TDM model ages of 2.80–2.92 Ga. Results from
eastern Finland overlap with those of the Central
and West Karelian domains and militate against a
prolonged crustal prehistory (Fig. 11). On the contrary, the Pb isotope composition of sanukitoids of
the Iisalmi and Ilomantsi terrains (model µ values
for the source ~9) indicates a substantial crustal Pb
component in the source (Halla, 2005). Similar results have been obtained from the Western Superior
province (Stevenson et al., 1999). Based on Pb isotopes, Halla (2005) suggested that sediment-derived
position of sanukitoids is also paradoxical: Nd, Sr,
and Hf isotopes point to a mantle origin (Corfu and
Stott, 1993, 1996), but Pb isotope compositions of
K-feldspars point to a crustal source (Stevenson et
al., 1999; Halla, 2005). The geochemical features of
wedge-derived sanukitoids are explained by partial
melting of a mantle-wedge source metasomatized
amounts of slab-derived (TTG) melts (e.g., Smithies and Champion, 2000), or by elements mobilized
processes (Kamber et al., 2002; Halla, 2005). The
mechanisms that caused the enrichment and partial
melting in the wedge are still unclear, as is the sig-
to the mantle wedge in slab dehydration processes
before melting. This conclusion is supported by a
number of studies that report high δ18O in zircons in
sanukitoids (e.g., King et al., 1998), also interpreted
as a crustal signature.
of sanukitoid series rocks.
Kovalenko et al. (2005) suggested a two-stage
model for the genesis of the Karelian sanukitoids.
3.2.4. Petrogenetic remarks
somatized during subduction or tectonic underplating (up to 200 Ma before melting), either by sig-
The somewhat paradoxical geochemical features
of the sanukitoids— low SiO2, high Mg number and
Mg, Ni, and Cr values— point to a mantle-wedge
peridotite source, whereas enrichment in the LILE
(Ba, Sr, and P) and the LREE indicate an enriched
(metasomatized) mantle source. The isotope com-
tion processes. In the second stage, at 2.74–2.70 Ga,
a thermal event, probably related to the collision of
the Belomorian mobile belt with the Karelian craton, caused melting in the previously metasomatized mantle. Kovalenko et al. (2005) presented
Fig. 11. εNd vs. age diagram for the Nilsiä and Lieksa sanukitoids. Data for Central and West Karelian domains (black and white
bars) from Lobach-Zhuchenko et al. (2000b) and Kovalenko et al. (2005). CHUR as in Fig. 8.
26
two alternative models for melting in the mantle
wedge: (1) delamination of the lower crust that
caused mantle upwelling and heating of the subcon-
and slab dehydration. At ~2.73 Ga, melting in the
mantle-wedge source produced high-Mg, high-K
granitoid (sanukitoid) magmas with high LILE, U,
Th, and Pb and radiogenic initial Pb isotope compositions. A strong U depletion event at 2.7–2.6
Ga during granulite-facies metamorphism further
enriched these rocks in Th relative to U and in U
relative to Pb.
Halla (2005) suggested that the Nilsiä and Lieksa
sanukitoids in eastern Finland originated from a
mantle-wedge source enriched in the LILE, U, Th,
and crust-derived Pb through sediment subduction
3.3. Neoarchean leucogranites
(A. Käpyaho, L.S. Lauri)
3.3.1. Introduction
One of the best-documented localities in the area is
the Arola (or Pohjajärvi) quarry (Hyppönen, 1983;
Martin and Querré, 1984; Querré, 1985; Käpyaho et
al., in review). This two-mica leucogranite is pink,
medium-grained, and slightly oriented. Main minerals are quartz, oligoclase, microcline, muscovite, and
biotite, with accessory magnetite, zircon, rutile, allanite, and rare garnet. Some of the related intrusions
contain magnetite crystals up to 0.5 cm in diameter,
and some of the granite bodies lack muscovite.
Neoarchaean leucogranites are found on all Archean cratons, yet the origin of this rock type is still
far from understood. These granites have been interpreted to represent crustal melts (e.g., Sylvester,
1994), but some studies (e.g., Moyen et al., 2001)
indicate that some high-K leucogranites could also
a link to the high-Mg/Fe sanukitoids (see Section
3.2). A common emplacement order in most Archean cratons is that the leucogranites often postdate
the much more voluminous TTG association rocks
and greenstone belts (for a review, see Sylvester,
1994).
In the Archean of the Fennoscandian shield, Neoarchean leucogranites are found, e.g., in northeastern Lapland (Juopperi and Vaasjoki, 2001) and in
the Pudasjärvi (Mutanen and Huhma, 2003), Ilomantsi (Vaasjoki et al., 1993), Koillismaa (Lauri et
al., 2005), and Kuhmo districts (Hyppönen, 1983;
Martin and Querré, 1984; Martin, 1985; Querré,
1985; Käpyaho et al., in review). In this summary,
we focus on the whole-rock geochemical composition, U–Pb ages, and Nd isotope characteristics of
the most extensively studied leucogranite occurrences in the Kuhmo and Koillismaa districts in the
Karelian province, and review the petrogenetic implications of these data.
Koillismaa district
The Archean part of Koillismaa forms the northernmost continuation of the Kuhmo district. The
area is still relatively poorly mapped, but some
leucogranite occurrences are known; these are assigned to the Harjavaara-type in the geologic map
(see Räsänen et al., 2004; Lauri et al., 2005). The
leucogranites are found as small bodies and dikes
within the migmatitic gneisses. They are mediumto coarse-grained and commonly show evidence of
deformation. Major minerals are quartz, plagioclase,
microcline, biotite, and in some cases muscovite.
3.3.3. U–Pb ages
The leucogranites of Kuhmo and Koillismaa have
U–Pb ages on the order of 2.71 Ga to 2.68 Ga. Ion
microprobe U–Pb zircon ages from the Kuhmo district vary between 2.70 and 2.68 Ga and the granites commonly contain inherited zircons (Käpyaho
et al., in review). The conventional analyses from
Koillismaa are rather imprecise and the most robust
age determination (on the Aholamminvaara granite,
sample A1656) yielded an age of 2711 ± 9 Ma (Lau-
3.3.2. Key localities of Neoarchean leucogranite
in the Finnish Archean
Kuhmo district
Small leucogranite bodies (<1 km2) and crosscutting granite dikes are common in the Archean bedrock west of the Kuhmo greenstone belt (Fig. 4).
27
ri et al., 2005). Concordant monazite ages of ~2.69
Ga in the Koillismaa region are consistent with the
ion microprobe results from Kuhmo. The zircon
population is heterogeneous and the monazite may
register the emplacement of the granites (Lauri et
al., 2005). Similar U–Pb monazite ages have also
been obtained from leucogranites in the Ilomantsi
terrain (Vaasjoki et al., 1993).
3.3.4. Geochemical constraints on origin
Whole-rock geochemistry
Fig. 12. SiO2 (wt.%) vs. A/CNK [molar Al2O3/
(CaO+Na2O+K2O)] diagram showing the composition of the
Archean leucogranites from Kuhmo and Koillismaa. Mean
values for Archean calc-alkaline granites (CA1 and CA2),
strongly peraluminous granites (SP3 and SP4) and alkaline
granites (ALK4) are after Sylvester (1994).
The Neoarchean leucogranites of Kuhmo and
Koillismaa are generally rich in silica, most of them
having >73 wt.% SiO2. They are moderately peraluminous, with A/CNK from 1.01 to 1.15 (Table 3).
of Chappell and White (1974), whereas the granites
from Kuhmo range from I-type to S-type (Fig. 12).
Most leucogranites in Koillismaa are magnesian acwhereas the leucogranites from Kuhmo span from
magnesian to ferroan with increasing SiO2 values
(Fig. 13). The LREE are strongly fractionated compared to the HREE— the latter are under detection
limit in most samples from Koillismaa (Fig. 14).
The leucogranites of Kuhmo generally show a
small negative Eu anomaly. In the granites of Koillismaa the Eu anomaly varies from slightly negative to slightly positive; positive values have been
registered in one sample (A1657; Lauri et al., 2005)
from a coarse-porphyritic granite that may register
some feldspar accumulation.
Fig. 13. FeOtot / (FeOtot + MgO) vs. SiO2 (wt.%) diagram for
the Archean leucogranites of Kuhmo and Koillismaa. Line
al. (2001). Symbols as in Fig. 12.
Nd isotopes
Leucogranites of the Kuhmo and Koillismaa districts are somewhat different in terms of Nd isotope
composition (Fig. 15). In both districts, the highest
values are ~+1. The lowest values in Kuhmo are
Ndi
around –2. However, in Koillismaa most values fall
between –1.5 and –4, and are thus slightly more negative (at 2.70 Ga) than the 2.84 Ga migmatite paleo-
Sylvester (1994) divided Archean granites into
three main categories, namely calc-alkaline granites,
strongly peraluminous granites, and alkaline granites,
each with two subgroups. According to Sylvester
(op. cit.), the calc-alkaline granites (CA) are partial
melts of a mainly tonalitic source, whereas the
strongly peraluminous granites (SP) are derived from
a dominantly sedimentary source. The source of the
alkaline granites (ALK) is not well constrained but,
according to Sylvester (1994), the Archean alkaline
granites differ from their Phanerozoic counterparts
more than the other two types (CA and SP) do.
component similar to the 3.5-Ga Siurua gneiss (Mutanen and Huhma, 2003) is present in the studied
leucogranites (Fig. 15). Therefore it is likely that the
leucogranites in eastern Finland were largely formed
by melting of the local Neoarchean crust (Käpyaho
et al., in review).
28
Table 3. Selected analyses of Neoarchean leucogranites
Koillismaa (1) Koillismaa (1)
Sample
A1656
A1657
Kuhmo (2)
Kuhmo (2)
CA1 (3)
CA2 (3)
SP3 (3)
SP4 (3)
ALK4 (3)
G16
G15
(n = 16)
(n = 12)
(n = 9)
(n = 14)
(n = 11)
SiO2
74.8
74.1
75.1
76.1
70.0
71.9
73.7
74.4
74.2
TiO2
0.13
0.08
0.19
0.13
0.40
0.23
0.21
0.12
0.14
Al2O3
14.8
14.4
13.11
12.55
14.6
14.7
14.5
14.2
13.5
Fe2O3tot
1.10
1.03
1.28
0.91
3.02
1.83
1.70
1.22
1.47
MnO
0.02
0.01
0.01
0.01
0.05
0.03
0.03
0.03
0.04
MgO
0.30
0.11
0.32
0.15
0.84
0.48
0.41
0.26
0.23
CaO
2.57
1.18
0.63
0.52
2.28
1.69
1.12
0.57
0.94
Na2O
5.15
3.78
3.92
3.54
3.89
4.45
2.94
3.91
3.74
K2O
0.89
4.53
4.32
4.67
3.58
3.69
4.69
4.69
4.68
P2O5
0.01
0.01
0.06
0.05
0.17
0.08
0.09
0.10
0.60
Cl
0.01
0.01
-
-
-
-
-
-
-
V
14
4
-
-
28
15
9
3.1
9
Cr
0
18
26
8
26
73
9.9
6.3
37
Ni
0
7
-
-
12
12
5
6.3
10
Cu
1
0
-
-
19
13
60
12
11
Zn
18
13
-
-
59
41
29
27
44
Ga
23
20
-
-
18
17
14
23
20
Rb
28
49
141
191
117
125
140
292
339
Sr
257
550
185
120
479
455
189
73
102
Zr
51
55
142
859
218
142
116
90
134
Nb
2
1
7
12
9
3.3
16
19
Ba
172
4732
1629
936
1300
1210
740
372
505
Pb
48
49
-
-
23
28
32
36
41
2.6
Sc
1.12
0.53
-
-
4.7
1.8
2.8
1.8
Th
29.5
0.92
-
-
21
22
22
21
41
U
0.3
<
-
-
2.4
3.3
11
5.4
6.8
Y
1.34
0.80
-
8
21
7
7.3
21
31
La
58.8
13.7
44.4
15.3
71
42
38.1
20.3
44
86
Ce
101
19.1
84.5
35
133
71
76.4
41.9
Pr
9.68
1.70
-
-
-
-
-
-
-
Nd
29.4
4.82
-
-
45
29
28
19
28
Sm
3.03
0.49
3.9
1.71
7.8
4.4
4.62
3.8
5.3
Eu
0.63
0.31
0.71
0.44
1.56
0.79
0.813
0.307
0.45
Gd
1.52
0.40
-
-
5.1
3.4
3.78
2.76
3.00
Tb
0.2
<
0.18
0.19
0.73
0.41
0.43
0.57
0.85
Dy
0.41
0.13
-
-
3.2
2.2
2.2
4.4
3.1
Ho
<
<
-
-
-
-
-
-
-
Er
<
<
-
-
-
-
-
-
-
Tm
<
<
-
-
-
-
-
-
-
Yb
<
<
0.63
0.58
1.55
0.51
0.82
2.05
1.9
Lu
<
<
0.12
0.09
0.24
0.09
0.14
0.32
0.37
A/CNK
1.05
1.08
1.07
1.06
1.01
1.02
1.21
1.13
1.04
Fe#
0.77
0.89
0.78
0.85
0.76
0.77
0.79
0.81
0.85
Oxides in wt.%, trace elements in ppm. < under detection limit, - not analyzed, 0 not detected.
References: (1) Lauri et al. (2005); (2) Querré (1985); (3) Sylvester (1994).
29
3.3.5. Concluding remarks
The ages thus far reported for the leucogranite
magmatism in eastern Finland indicate that the leucogranites were emplaced at ~2.70 Ga. Thus these
granites are coeval with high-grade metamorphism
in central Finland (Mänttäri and Hölttä, 2002); further studies are, however, needed to establish whether there is a connection between this metamorphic
event and leucogranite magmatism. Nd isotope
results from the leucogranites of eastern Finland
(Käpyaho et al., in review; Lauri et al., 2005) imply that these rocks most probably represent partial
melts of the local crust. Nevertheless, the positive
initial Nd values suggest that some of these granites could also be mixtures of juvenile magmas and
melts from pre-existing crust.
Based on whole rock chemical composition and
Sr isotope data Querré (1985) noted that there are
similarities between the leucogranites of Kuhmo
district and the younger S-type plutons, especially
the continental collision-related Himalayan leucogranites. The peraluminous nature and geochemical
characteristics are, in general, similar to average
Archean calc-alkaline and strongly peraluminous
granites of Sylvester (1994).
Fig. 14. Chondrite-normalized REE patterns for the Archean
leucogranites of Kuhmo (Querré, 1985; Asko Käpyaho, unpublished data) and Koillismaa (Lauri et al., 2005). Normalizing values (C1 chondrite) are from Sun and McDonough
(1989).
The leucogranites of Koillismaa fall between
the calc-alkaline (CA3, CA4) and alkaline (ALK4)
categories whereas the leucogranites of Kuhmo resemble the strongly peraluminous granites (SP4) of
source compositions and varying amounts of partial
melting as the Archean leucogranites in different
parts of the Karelian province were formed.
Fig. 15. εNd vs. age diagram for the Archean leucogranites of Kuhmo and Koillismaa. DM (depleted mantle) evolution line from
DePaolo (1981), CHUR denotes the Chondritic Uniform Reservoir of DePaolo and Wasserburg (1976).
30
3.4. Early Paleoproterozoic anorogenic granites
(L.S. Lauri)
3.4.1. Introduction
km (A in Fig. 16). It consists of coarse-grained
to pegmatitic, pink granite that had intruded into
Archean paragneisses. In the border zone, the granite shows a lineation but the center of the pluton is
homogeneous and unoriented. Major minerals are
quartz, K-feldspar, oligoclase, and biotite. Accessory minerals include muscovite, zircon, epidote,
At the beginning of the Proterozoic Eon, a period of more than 200 m.y. of tectonic quiescence in
the Fennoscandian Archean craton was broken by
a stage of intermittent rifting that lasted for several
hundred m.y. The onset of rifting has been attributed to mantle plume activity (e.g., Huhma et al.,
1990; Amelin et al., 1995). The initial extension at
~2.44 Ga was manifested by bimodal magmatism
Conventional U–Pb zircon age of the Rasimäki
pluton is 2352 ± 25 Ma (Horneman, 1990), which
implies that it is somewhat younger than the other
intrusions described here. However, the Rasimäki
granite geochemically resembles the ~2.44-Ga silicic intrusions and was probably formed by similar
processes. The age of the pluton is also uncertain, as
an alternative age of 2404 ± 155 Ma can be obtained
from the same data (Horneman, 1990). Horneman
(op. cit.) also associated the Rasimäki granite to
some nearby granite dikes that contain 2410 Ma old
titanite.
and several small, silicic plutons (e.g., Luukkonen,
1988; Alapieti et al., 1990; Huhma et al., 1990;
Amelin et al., 1995; Buyko et al., 1995; Lauri and
Mänttäri, 2002; Lauri et al., 2003, 2005). Subsequent extensional phases are registered by diabase
dike swarms and gabbroic intrusions at 2.3 Ga,
2.1 Ga, and 1.98 Ga (e.g., Vuollo et al., 2000) and
they were followed by the Svecofennian orogeny at
~1.93–1.87 Ga.
In Finland and adjacent Russia, a number of distinctive ~2.44-Ga silicic intrusions are found (Fig.
16). The Rasimäki granite pluton was emplaced
close to the southwestern margin of the Archean
craton (Horneman, 1990), while three small granite intrusions and associated granite porphyry dikes
are known in the Kuhmo–Suomussalmi area (Luukkonen, 1988). A quartz alkali feldspar syenite intrusion is found farther north in Koillismaa, close
Tuliniemet
Tuliniemet-type granites are found in the Kuhmo
district, on the eastern side of the Archean Kuhmo–
Suomussalmi greenstone belt (Luukkonen, 1988;
B and inset in Fig. 16). Three small intrusions and
associated granite porphyry dikes intrude Neoarchean TTG gneisses. The intrusions are almost
undeformed and clearly post-tectonic relative to the
Archean deformation seen in the host rocks. The
faint orientation seen in the brecciated border zone
(Lauri and Mänttäri, 2002). East of Koillismaa, in
northwestern Russia, is yet another granite pluton,
complex (e.g., Buyko et al., 1995). All these silicic
intrusions are within the Karelian province that also
emplacement. Major minerals in the Tuliniemettype granites are K-feldspar, plagioclase, quartz,
and biotite. Accessory minerals include sericite,
intrusions. Some granites of similar age have also
been reported from the Belomorian province (e.g.,
Lobach-Zhuchenko et al., 1998), but they have not
been considered in this review.
als. Central parts of the intrusions contain unmantled rapakivi-type K-feldspar ovoids, and mantled
ovoids were also found in one of the plutons. The
granite porphyry dikes contain microcline, quartz,
3.4.2. Description of the ~2.44-Ga silicic intrusions
crystallized groundmass consisting of quartz, plagiRasimäki
minerals.
The U–Pb zircon age of one of the granite porphyry dikes is 2435 ± 12 Ma (Luukkonen, 1988).
The Rasimäki granite pluton, described by
Horneman (1990), is a small stock of 8 km by 3
31
vicinity of the 2.44-Ga Koillismaa layered igneous
complex (Lauri and Mänttäri, 2002; C in Fig. 16).
The pluton has intruded into Archean TTG gneiss-
Recent U–Pb zircon results (Irmeli Mänttäri, unpublished data) imply an age of ~2425 Ma for one
of the granites. Luukkonen (1988) correlated the
Tuliniemet-type granites with the coeval Koillismaa
layered igneous complex and with an EW-trending
diabase dike swarm in the Kuhmo district.
and extrusive rocks were also found near the contact.
The Kynsijärvi pluton consists of homogeneous, undeformed, medium-grained, pink quartz alkali feldspar syenite. The main minerals present are mesoperthitic alkali feldspar, quartz and hornblende, with
feldspar clearly dominating. Minor and accessory
Kynsijärvi
The Kynsijärvi quartz alkali feldspar syenite is a
small (1 km by 5 km) EW-elongated intrusion in the
Fig. 16. Geological sketch map of eastern Finland showing the distribution of the early Paleoproterozoic (~2.44 Ga) anorogenic
32
16) dates back to 1929, when the intrusion was still
within the Finnish territory. Hackman and Wilkman (1929) described the intrusion and associated
granite porphyry dikes in their explanation to the
1:400,000 geological map of the Kuolajärvi area.
The Nuorunen granite is pink, medium- to coarsegrained subsolvus rock that mainly consists of microperthitic alkali feldspar, quartz, plagioclase, and
amphibole. Accessory minerals include zircon,
crographic intergrowths of quartz and alkali feldspar
are common. The rock is somewhat deformed and
slightly altered, with amphiboles partly replaced by
chlorite. Other secondary minerals include sericite,
stilpnomelane, and carbonate.
Buyko et al. (1995) reported a U–Pb zircon age
of 2450 ± 72 Ma for the Nuorunen granite. The
Fig. 17. FeOtot / (FeOtot + MgO) vs. SiO2 (wt.%) diagram
showing the composition of the early Paleoproterozoic anorogenic silicic intrusions of Finland and vicinity. Line sepa(2001). References: 1— Lauri and Mänttäri (2002); 2— Rämö
and Luukkonen (in prep.); 3— Horneman (1990); 4— Lauri et
al. (2005); 5— Luukkonen (1988).
show similar ages (2445 Ma to 2435 Ma; Balashov
et al., 1993; Amelin et al., 1995). At face value, the
age of the Nuorunen granite implies that the granintrusions.
3.4.3. Elemental geochemistry
The ~2.44-Ga silicic intrusions are all geochemically A-type, typically having, e.g., high alkali abundances, high Fe/Mg and Ga/Al ratios, high
contents of Zr, Nb, Ce, Y, and REE, and negative
Eu anomalies (Table 4). SiO2 varies from <72 wt.%
in the Kynsijärvi pluton to between 72 wt.% and
Fig. 18. SiO2 (wt.%) vs. A/CNK [molar Al2O3/
(CaO+Na2O+K2O)] diagram showing the composition of the
early Paleoproterozoic anorogenic silicic intrusions of Finland and vicinity. See Fig. 17 for numbered data references.
and stilpnomelane, which is probably of secondary
origin. The pluton has suffered from post-magmatic
alteration, manifested as intergranular albite rims
between the mesoperthite crystals. The 2442 ± 3 Ma
U–Pb zircon age of the Kynsijärvi pluton (Lauri and
Mänttäri, 2002) is close to the age (2436 ± 5 Ma;
are probably coeval.
Fig. 19. Zr vs. 10,000*Ga/Al diagram showing the composition of the early Paleoproterozoic anorogenic silicic intru-
Nuorunen
The best available description of the Nuorunen
granite pluton in northwestern Russia (D in Fig.
al., 1987). See Fig. 17 for numbered data references.
33
Fig. 20. Chondrite-normalized REE patters for the early Paleoproterozoic anorogenic silicic intrusions of Finland and vicinity.
Normalizing values (C1 chondrite) from Sun and McDonough (1989). See Fig. 17 for numbered data references.
Fig. 21. εNd vs. age diagram for the early Paleoproterozoic anorogenic intrusions of Finland and vicinity. DM (depleted mantle)
evolution line from DePaolo (1981), CHUR denotes the Chondritic Uniform Reservoir (DePaolo and Wasserburg, 1976). See
34
Table 4
Average compositions of the ~2.44 Ga A-type granites in the
Fennoscandian shield
Rasimäki1, 2
Tuliniemet2, 3
Kynsijärvi4
Nuorunen5
SiO2
75.7
75.6
70.8
75.4
TiO2
0.13
0.09
0.23
0.24
Al2O3
13.0
12.9
14.6
12.1
Fe2O3tot
1.43
1.41
3.06
1.71
MnO
0.02
0.04
0.05
0.03
MgO
0.12
0.28
0.11
0.18
CaO
0.80
0.59
0.58
0.69
Na2O
3.95
3.76
5.67
3.22
K2O
4.39
4.56
4.53
5.32
P2O5
0.02
0.03
0.03
0.03
Cl
-
<
100
100
V
-
-
<
14
20
Cr
-
30
17
Ni
-
5
<
6
Cu
-
7
<
7
Zn
32
44
75
23
Ga
22
24
31
22
Rb
181
398
102
168
Sr
56
51
35
77
Y
39
73
31.2
33.4
291
Zr
183
141
520
Nb
28
26
42
17
Ba
220
149
711
709
Sc
-
-
2.5
3.7
Th
27
48
11.1
9.9
U
18
8
1.06
1.4
Pb
-
40
36
27
La
51.9
33.9
61.9
69.2
Ce
106.1
70.0
108.3
144.7
Pr
19.7
9.19
11.4
16.9
Nd
39.8
32.6
39.6
58.4
Sm
6.03
7.63
6.85
9.39
Eu
0.48
0.29
0.6
0.72
Gd
8.32
7.87
6.01
7.42
Tb
1.10
1.37
0.91
1.07
Dy
6.30
8.40
4.89
5.56
Ho
1.35
1.95
1.02
1.06
Er
4.16
6.55
3.06
3.19
Tm
0.64
1.08
0.45
0.48
Yb
4.39
7.93
3.29
3.32
Lu
0.70
1.17
0.47
0.48
from magnesian to ferroan (Fig. 17). Kynsijärvi
and Nuorunen are metaluminous with A/CNK <1,
Rasimäki is mildly peraluminous with A/CNK between 0.97 and 1.09, and in Tuliniemet the highest
A/CNK values are >1.2 (Fig. 18). In the diagrams of
Whalen et al. (1987), all the ~2.44-Ga silicic intruterns of these rocks show enrichment in the LREE
compared to the HREE and a negative Eu anomaly
(Fig. 20).
3.4.4. Nd isotopes
Nd isotope compositions of the ~2.44-Ga granites are unique for each pluton or group and seem
(Fig. 21). Kynsijärvi shows the least radiogenic
values with initial εNd values between –4 and –5
(Lauri et al., 2005). The Nuorunen and Rasimäki
granites have a somewhat more radiogenic isotope
composition with εNd between –1.7 and –2.0 in the
former (Lauri et al., 2005) and between –1.2 and
–2.5 in the latter (Rämö and Luukkonen, in prep.).
Tuliniemet-type rocks show the largest spread of εNd
values, between –3.1 and +1.5 (Rämö and Luukkonen, in prep.). The Archean crust in Koillismaa has
εNd (at 2440 Ma) values between –5 and –8.5 (Lauri et al., 2005), whereas for the Archean rocks in
Kuhmo the corresponding values are between +1.0
and –4.5. (Käpyaho et al., in review; Section 3.3.).
The Archean basement surrounding the Nuorunen
pluton is somewhat less well documented but there
are indications that also there the εNd (at 2440 Ma)
values are close to zero or slightly negative (see discussion in Lauri et al., 2005).
3.4.5. Concluding remarks
The ~2.44-Ga granitoids in the Fennoscandian
shield are found as small plutons and dikes within
the Late Archean basement of the Karelian province. They are mostly undeformed and show geochemical features typical for anorogenic granites.
References: (1) Horneman (1990); (2) Rämö and Luukkonen (in
prep.); (3) Luukkonen (1988); (4) Lauri and Mänttäri (2002); (5)
Lauri et al. (2005). - not analyzed, < below detection limit.
a bimodal magmatic association that was emplaced
in the Archean craton in an extensional regime, the
77 wt.% in others. Rasimäki, Kynsijärvi, and Nuorunen have high Fe/Mg and are ferroan according
icic rocks probably representing lower crustal melts
Tuliniemet have somewhat lower Fe/Mg and span
35
3.5. Intermediate magmatism at the craton margin, central Finland
(O. Äikäs, L.S. Lauri)
3.5.1. Introduction
The western marginal zone of the Archean craton
and the overlying Proterozoic supracrustal rocks are
cut by a texturally, structurally, and compositionally
diate rocks, provisionally referred to as microtonalites (Huhma, 1981) form one distinct group among
these. The microtonalites occur mainly as dikes, but
also as larger intrusions in a relatively narrow zone
from the line Leppävirta–Tuusniemi–Kaavi up to
the area west of Oulujärvi (Fig. 22) (see also Rautiainen, 2000). One of the most numerous sets of SE–
NW-trending microtonalite dikes occurs within the
Archean paragneisses between the lakes Vuotjärvi,
near Juankoski and Syväri, near Nilsiä (Paavola,
1984). The location and direction of these microtonalite dikes coincide with a number of presumably
younger lamprophyre dikes (Huhma, 1981). Occasional observations and descriptions of these dikes
farther south and southeast are from Kangaslampi (map sheet 3234 in Fig. 22) and from Juojärvi
(Koistinen, 1993).
Larger occurrences include the horizontal and
gently dipping microtonalite sheets around Talvisalo, near Nilsiä, and in Kuopio (Äikäs, 2000). Some
other intrusive bodies, e.g., the diorites close to the
Kuopio (Rissala) airport (Äikäs, 2000; Lukkarinen,
2000) and the composite intrusion of Kaarakkala
farther north at Vieremä (Paavola, 2001, 2003), may
also be related to the microtonalite magmatism.
The age of the microtonalite magmatism is poorly
constrained because of zircon inheritance (Huhma,
1981; Paavola, 2003). Crosscutting relations indicate the presence of several pulses of microtonalite
dikes, although the emplacement ages of different
pulses may be close to each other. Huhma (1981)
reported a titanite age of 1850 Ma from a dike in
Kaavi, but it is not clear whether this age marks
the emplacement of the dike. A dike at Murtolahti,
Nilsiä, yielded discordant zircon ages between
1939 Ma and 1835 Ma and a titanite age of 1829
± 13 Ma (Irmeli Mänttäri, personal communicathe Kaarakkala intrusion was dated at 1864 ± 8 Ma
(Ruotoistenmäki et al., 2001; Paavola, 2003). The
1869 ± 5 Ma (Hannu Huhma, personal communication, 2005) Juurus tonalite is intruded by at least
36
one pulse of microtonalite dikes. According to the
primary concept of microtonalite dikes by Huhma
(1981), the microtonalites crosscut the 1857 ± 8 Ma
(Huhma, 1986) Maarianvaara granite, but at present this granite and its western counterparts are also
known to crosscut some microtonalite dikes. The
probable age span for microtonalite magmatism
extends at least from 1890 Ma to 1830 Ma and it
is possible that some conformable, deformed dikes
occurring within the Archean rocks are themselves
also Archean.
3.5.2. Field geology and petrography
can be divided into several groups: (1) dikes that
occur within the Archean rocks as conformable or
semi-conformable fragments; (2) dikes that crosscut
the Proterozoic mica gneisses but have been fragmented and deformed after emplacement; (3) less
deformed or undeformed dikes that crosscut both
Archean and Proterozoic rocks and are in turn cut by
~1860-Ma granites; and (4) dikes that crosscut the
~1860-Ma granites and other microtonalite dikes.
Dike widths vary from a few centimeters to several
meters and maximum length is probably only some
tens of meters, as the dikes usually cannot be traced
between outcrops. According to Rautiainen (2000),
to dark gray and homogeneous in appearance. They
have sharp, often jagged contacts to the wall rock
and commonly form intrusive breccias (e.g., Paavola, 1984). Apophyses and weak magmatic banding are also common. Some dikes show evidence of
Cloudy inclusions and net vein breccias between
ocellar texture is found in some places (e.g., Rautiainen, 2000).
Major minerals in the microtonalite dikes are subhedral plagioclase (>50 vol.%), biotite, and quartz
(~1–30 vol.%). Some dikes also contain hornblende.
Accessory minerals include titanite, apatite, epidote, zircon, rutile, chlorite, microcline, and oxide
minerals (Huhma, 1975; Rautiainen, 2000). Grain
size is small, between 0.05 and 0.5 cm. Dikes that
Archean bedrock shown in gray (Geological Survey of Finland GIS data, based on Korsman et al., 1997). Inset on the left: bedrock geology in the area north of Kuopio, extracted from the 1:1,000,000 bedrock map of Finland (Korsman et al., 1997). The
area was mapped by J. Paavola, H. Lukkarinen, A. Kontinen, and O. Äikäs. Map sheet numbers of the Finnish grid 1:100,000
are shown; the height of each sheet is 30 km. Raster images of the 1:100,000 geological maps can be downloaded at the GTK
37
lack hornblende resemble mica schists in the outcrop and— in some cases— can only be recognized
as microtonalites from thin sections.
3.5.3. Geochemistry
Rautiainen (2000) summarized the geochemical
data for the microtonalite dikes and concluded that
they were most probably generated by mixing and
to Rautiainen (op. cit.), the microtonalites can be
between 45
2
wt.% and 55 wt.%, a felsic group with SiO2 between
61 wt.% and 74 wt.%, and a hybrid group that falls
between the two end-members. All three groups are
metaluminous and I-type with A/CNK under 1.1
(Fig. 23). All microtonalite groups cross the magnesian-ferroan border in the FeOtot / (FeOtot + MgO)
vs. SiO2 diagram of Frost et al. (2001) (Fig. 24).
Microtonalites are enriched in the LREE compared
Fig. 23. SiO2 (wt.%) vs. A/CNK [molar Al2O3/
(CaO+Na2O+K2O)] diagram showing the composition of microtonalite samples depicted in Fig. 22. Data by the courtesy
of the Geological Survey of Finland (O. Äikäs).
ples commonly show a small negative Eu anomaly,
probably caused by hybridization. In felsic samples
the negative Eu anomaly is more pronounced. According to Rautiainen (2000), the microtonalites
Irvine and Baragar (1971). He also concluded that
gin arc basalts in trace element composition.
3.5.4. Concluding remarks
Fig. 24. FeOtot/(FeOtot + MgO) vs. SiO2 (wt.%) diagram of
microtonalite samples in Fig. 22. Data by the courtesy of the
Geological Survey of Finland (O. Äikäs).
Microtonalite dikes and related larger intrusive
bodies occur at the Archean craton margin in a geographically narrow zone. However, the time interval of the microtonalite magmatism is long, at least
60 m.y., and possibly more, if some dikes occurring
within the Archean rocks are themselves Archean.
Although the absolute ages of the microtonalite dikes
are still largely unknown, the magmatic episode
seems to consist of several pulses. It may be argued
that the microtonalite magmatism is a typical feature
of the Karelian craton margin and, for some reason,
the source was periodically tapped during Paleoproterozoic. The presence of two intermingled magma
observations, petrography, and geochemistry. Further isotope geochemical data would be needed in
order to constrain the age and origin of these dikes.
38
Fig. 25. Chondrite-normalized REE patterns of microtonalite samples: (A) Kivennapa area, Juankoski; (B) Honkamäki area,
Nilsiä. Normalizing values (C1 chondrite) from Sun and McDonough (1989). Data by the courtesy of the Geological Survey of
Finland (O. Äikäs).
3.6. Proterozoic synorogenic granitoids
(M. Nironen)
3.6.1. Distribution and age
Synorogenic granitoids, 1.89–1.87 Ga in age,
are ubiquitous in southern and central Finland
and are especially abundant in the ~40,000 km2
Central Finland granitoid complex (Fig. 26; see
also inside front cover). 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
has been dated at 1885–1880 Ma (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 subjected to penetrative deformation and
synkinematic magmatism, whereas in other areas
the postkinematic stage had already been reached.
granodiorites and granites with abundant subhedral
to anhedral potassium feldspar phenocrysts 1–4 cm
in diameter are located in the central and northeastern
parts of the complex. The granites contain biotite
are hornblende, titanite, apatite, magnetite, and
zircon. Both hornblende and biotite are found as
and the tonalites may also contain clinopyroxene
or orthopyroxene. Typical accessory minerals are
titanite, apatite, magnetite, and zircon. Abundant
suggest that mixing and mingling processes were
The felsic to intermediate synkinematic intrusive
rocks grade in places into subvolcanic quartz–
feldspar and feldspar porphyries.
Synkinematic granitoids 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. 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,
3.6.2. General mode of occurrence
Synkinematic rocks of the Gentral Finland
granitoid complex are typically medium-grained
granodiorites and granites (Fig. 27), but mediumgrained tonalites are also found, especially along
the western margin of the complex. The granitoids
exhibit a foliation that varies from slight orientation
to pervasive gneissic foliation. Coarse-porphyritic
39
Nironen, 2005). Abbreviations for crustal units (arc complexes) as in Fig. 1. Postorogenic intrusions are marked by stars:
1— Lemland, 2— Mosshaga, 3— Seglinge, 4— Åva, 5— Turku, 6— Renko, 7— Parkkila, 8— Luonteri, 9— Eräjärvi, 10— Pirilä.
the Archean craton. Microtonalite dikes (Section 3.5) are found west of the dashed line.
40
Fig. 27. Total alkali vs. SiO2 (TAS) diagram for Proterozoic synorogenic (synkinematic and postkinematic) granitoids in Finland (after Le Maitre et al., 1989). 1— trachybasalt, 2— basaltic trachyandesite, 3— trachyandesite (monzonite), 4— trachydacite
(quartz monzonite), 5— basalt, 6— basaltic andesite, 7— andesite (quartz diorite, tonalite), 8— dacite (granodiorite), 9— rhyolite
(granite).
interstitial and 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. The Type 3 plutons either have a pyroxenebearing margin or contain pyroxene throughout.
Large orthoclase megacrysts mantled by plagioclase
are typical of the Type 3 quartz monzonites.
Especially the Type 2 granites in the western part
of the Central Finland granitoid complex resemble
the rapakivi granites in their mineralogy and
geochemical characteristics (Nironen et al., 2000).
A couple of small monzodioritic to gabbroic bodies
are associated with the postkinematic plutons of
the Central Finland granitoid complex. Magma
1989). Biotite is generally more common than
hornblende, and accessory clinopyroxene occurs in
some phases. Other accessory minerals are titanite,
magnetite, apatite, and zircon. The granitoids
dioritic or dioritic phases. These features suggest
magma mixed producing a variety of hybrid rocks
(Nironen and Bateman, 1989; Lahtinen, 1996).
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 (Fig. 27; 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.
The postkinematic plutons are multiphase
intrusions that can be divided into three types (Elliott
et al., 1998; Elliott, 2003). The Type 1 plutons are
coarse-porphyritic biotite granodiorites and granites
with abundant orthoclase megacrysts and are found
along the southern margin of the Central Finland
coeval with the felsic rocks.
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. The postkinematic plutonic event
shifted from northeast toward west during 1885–
1870 Ma (Rämö et al., 2001a).
3.6.3. Geochemical constraints on origin
zircon, apatite, and ilmenite. The Type 2 plutons
are coarse-porphyritic or equigranular granites
that vary in grain size from medium to coarse. The
Geochemically, the granitoids show a shift
41
Fig. 28. Variation diagrams for Proterozoic synorogenic (synkinematic and postkinematic) granitoids in Finland. Fields in the
K2O vs. SiO2 diagrams after Rickwood (1989). Boundary of I-type and S-type granitoids in the molar Al2O3 / (CaO + Na2O +
K2O) [A/CNK] vs. SiO2
from Frost et al. (2001).
42
granitoid complex area (cf. Fig. 28); probably both
have a prominent arc-related source component.
Nironen et al. (2000) proposed that K-rich calcalkaline 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
synkinematic magmatism in the central Finland
granitoid complex. The synkinematic rocks
show a typical shift from the I-type to the S, but the rocks of
2
the supracrustal belts are generally more S-type
and more peraluminous than the rocks of the
Central Finland granitoid complex suggesting a
sedimentary component in the source area.
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. Overall, the postkinematic rocks are more
alkaline and evolved than the synkinematic rocks
(Figs. 27 and 28). The composition and similarity
to rapakivi granites of the postkinematic plutons
were derived from a granulitic residue left in the
lower crust after the extraction of the synkinematic
addition also contributed to the generation of
the postkinematic magmas. According to Elliott
(2003), the postkinematic magmas may consist of
crust with minor incorporation of restitic material
from lower crust.
The granitoids of the Central Finland granitoid
complex and granitoids of the supracrustal belts in
southernmost Finland have εNd (at 1880 Ma) values
from –1.5 to +1.1 and TDM model ages (DePaolo,
1981) from 2.05 Ga to 2.41 Ga, indicating that these
areas probably contain an older (~2.0 Ga) nucleus
(Lahtinen and Huhma, 1997; Rämö et al., 2001a).
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 and
TDM model ages of the rocks vary from 2.11 Ga to
2.27 Ga (Rämö et al., 2001a). The homogeneity of
the Nd isotope composition implies a homogeneous
source over the Central Finland granitoid complex
and that the crust-forming process was quite rapid
(Rämö et al., 2001a).
of relatively dry magmas derived from deep crust
(Elliott et al., 1998; Nironen et al., 2000). Nironen
et al. (2000) proposed that the postkinematic rocks
3.7. Proterozoic lateorogenic granites
(M.I. Kurhila)
3.7.1. Introduction
The lateorogenic granite zone of southern
Finland extends from the archipelago of
southwestern Finland east–northeast to Russia,
forming a semicontinuous belt about 500 km long
and 150 km wide (Fig. 29; see also inside front
cover). The zone is characterized by migmatizing,
high-K granites. The lateorogenic granites of
the Arc complex of southern Finland mostly do
not form batholiths sensu stricto, but usually
there is a gradual transition from microcline
granites through migmatites to subvolcanic
and metasedimentary rocks. The Puruvesi area
in eastern Finland (Fig. 29) is somewhat of
an exception in this respect, as it is a clearly
intrusive batholith with concentric lithological
variation. Although the lateorogenic granites
occupy large areas, their vertical thickness seems
to be rather limited. It is generally thought that
they form relatively thin undulating sheets within
the country rocks, at least in the western part of
the zone (e.g., Ehlers et al., 1993; Selonen et al.,
1996). Petrographically, the granites are quite
similar throughout the belt, the major minerals
being quartz, microcline, biotite, plagioclase,
cordierite, and garnet. Modal compositions plot
within alkali feldspar granite and syenogranite
show more variation, but they commonly include
apatite, muscovite, chlorite, monazite, zircon, and
rutile or anatase. At outcrop scale, these granites
are commonly heterogeneous: grain size and
distribution and the amount of garnet and biotite
decimeter-scale compositional layering.
3.7.2. Geochemical features
Geochemically, the lateorogenic granites
of southern Finland are rather heterogeneous.
Their SiO2 content varies from about 65 wt.%
43
3.7.3. Nd isotopes
to over 77 wt.%. They are peraluminous, but the
variation in A/CNK is relatively large, from 1.02
Along with the elemental geochemical composition, whole-rock Nd isotope composition varies
substantially along the zone. The least radiogenic
initial Nd isotope compositions ( Nd values around
–6) are found in the east, around Lake Puruvesi near
the Archean Karelian craton (Fig. 29). Such low Nd
variation; for example, the Eu anomaly ranges
from slightly positive to strongly negative.
The HREE are enriched in some samples but,
generally, the REE diagrams of these granites are
smoothly descending (Fig. 30). Variation in (La/
Yb)N
different amounts of residual garnet.
The commonly expressed notion that the late
Svecofennian granites are S-type (e.g., Ehlers
et al., 1993; Johannes et al., 2003) is not strictly
correct, as some of the rocks within the zone do
not meet with the criteria of the classic S-type
probe U–Pb data (Matti Kurhila, unpublished data)
that have revealed abundant inherited Archean and
older Paleoproterozoic zircons. In the northwestern
part of the lateorogenic granite zone, positive initial Nd values dominate, indicating a rather juvenile
source in that part of the zone. The majority of the
lateorogenic granites, however, have initial Nd values around –1. All granitoid rocks in the southern
(1974). Some of the granites of the belt have
virtually no garnet or monazite, their A/CNK
value may be below 1.1 and the Na content up to
5%. Also, muscovite is not always present.
less radiogenic initial Nd isotope composition than
den in the west and to Russia in the east. Generally, the lateorogenic granites are migmatizing and have gradational contacts
margins.
44
Fig. 30. Chondrite-normalized REE composition of a selection of lateorogenic granites across the Arc complex of southern Finland, displaying considerable heterogeneity. Normalization values from Boynton (1984). Unpublished data by Mikko Nironen
and Matti I. Kurhila.
Fig. 31. Geological map of southwesternmost Finland showing the geographic position and initial
genic granites from Kurhila et al. (in press).
45
Nd
values of eight lateoro-
Fig. 32. Distribution of U–Pb ages obtained from zircons and monazites throughout the lateorogenic granite zone. Error bars
are at 2 level. Unpublished data by Matti I. Kurhila.
between otherwise indistinguishable granites less
than 10 km apart has been documented (Kurhila et
al., in press). Also, the amount and ages of inherited
zircons vary a great deal. In the eastern part of the
granite zone, Archean and older Paleoproterozoic
zircons are fairly common, whereas in the west and
south, virtually no inherited zircons with preserved
U–Pb isotope ratios have been recorded (Kurhila et
al., in press).
those in the northern part of the Arc complex of
southern Finland (cf. Fig. 31), where the granites
record initial Nd values close to zero. This might indicate an EW-trending terrane boundary within the
Arc complex of southern Finland.
3.7.4. U–Pb ages
Traditionally, the lateorogenic granites have been
considered to be 1.84–1.82 Ga in age (e.g., Korsman
et al., 1997, 1999). However, we have recently
made numerous, mostly still unpublished U–Pb age
determinations of these granites, using both zircon
and monazite. Commonly, both minerals display
similar ages for individual samples. The collective
results are shown in chronological order in Fig. 32.
The age range of the lateorogenic granites is at least
from ~1.80 Ga to 1.85 Ga, i.e., clearly larger than
previously thought. On a broad scale, the ages show a
notable younging trend towards the east, but adjacent
samples within the zone may show considerably
different ages. For example, a ~25–30-Ma gap
3.7.5. Sources and origin of the lateorogenic
granites
The available elemental geochemical data, Nd
isotopes, and detailed U–Pb geochronology show
that the lateorogenic granites of southern Finland
were derived from varying sources. The wide
age range of inherited zircons, abundant garnet
and cordierite, and negative initial Nd values of
some of the granites indicate a major sedimentary
source component, whereas the lack of preserved
inherited zircons, positive Nd values, and extremely
46
granites is associated with local generation of melt,
possibly connected with formation of pressure
minima after the cessation of the main phase of the
orogeny (Kurhila et al., in press).
The various Svecofennian crustal terranes (cf.
Lahtinen and Huhma, 1997; Rämö et al., 2001a)
were already amalgamated by ~1.87 Ga and,
subsequently, high-grade metamorphism took place
throughout southern Finland over a period of ~50
m.y.. The ages of the lateorogenic granites record
roughly the same time span as the granulite facies
metamorphism within the arc complex of southern
Finland (Suominen, 1991; Van Duin, 1992; Väisänen
et al., 2002; Kurhila et al., in press). However, this
possible link between metamorphism and granite
magmatism needs further research.
leucocratic mineralogy in others are in favor of an
igneous source.
Selonen et al. (1996) proposed that the lateorogenic
granites were emplaced at middle crustal levels but
the tectonic regime under which this took place
is still subject to controversy. The granites are
thought to record either a transpressional intraplate
environment (Ehlers et al., 1993) or an extensional
zone after orogenic collapse (Korja and Heikkinen,
1995; Lahtinen et al., 2003). The typical features
of the granites, i.e., generally low-angle layering,
preferred magmatic orientation, migmatization,
and melt segregation, could be attributed to both
environments (Solar et al., 1998). In any case,
considering the capricious distribution of their ages,
it seems that the emplacement of the lateorogenic
3.8. Postorogenic intrusions
(M. Nironen, O.T. Rämö)
phase and younger granodiorite and granite
The postorogenic rocks of southern Finland are
found as ten relatively small intrusions that roughly
follow the northern boundary of the of lateorogenic
granite zone to Russia (Fig. 26). These intrusions
cut sharply the surrounding Paleoproterozoic
metamorphic crust and their U–Pb zircon ages range
from 1815 Ma to 1760 Ma (Vaasjoki and Sakko,
1988; Suominen, 1991; Vaasjoki, 1996; Väisänen et
al., 2000; Eklund and Shebanov, 2005).
hornblende is abundant in some of the granodioritic
varieties. Typical accessory minerals are titanite,
apatite, magnetite, zircon, and allanite. Chlorite and
Lamprophyric dikes are associated with the Åva
and Seglinge ring complexes. Bimodal lamprophyre–
granite magmatism has resulted in magma mixing
and mingling structures (Hubbard and Branigan,
1987; Branigan, 1989; Eklund et al., 1998; Eklund
3.8.1. Mode of occurrence
The intrusions are generally rounded with a
diameter of 2−15 km. In the far southwest part of
the country (Åland Islands), the Åva, Seglinge, and
Mosshaga intrusions are ring complexes and the
Lemland intrusion also has concentric compositional
and structural features (Eklund and Shebanov, 2005).
In contrast to the other intrusions, the Parkkila and
Eräjärvi (Nykänen, 1988) plutons in southeastern
Finland are dike-like bodies with lengths of several
kilometers. All intrusions sharply crosscut their host
rocks.
Large compositional variations (monzodiorite to
granite) are present in the intrusions of the Åland
Islands and at Luonteri whereas the others are more
homogeneous— Renko is quartz monzodioritic,
Parkkila granodioritic, and Pirilä and Eräjärvi
granitic. The Luonteri intrusion is a funnel-shaped
multiphase pluton that consists of an early tonalitic
Fig. 33. Total alkali versus SiO2 (TAS) diagram for postorogenic intrusions in southern Finland (after Le Maitre et
al., 1989). Fields are as in Fig. 27. Data from Eräjärvi are
marked by heavy dots.
47
Fig. 34. Variation diagrams for the postorogenic intrusions in southern Finland. Fields in the K2O vs. SiO2 diagrams are after
Rickwood (1989). Boundary of I- and S-type granitoids in the A/CNK [molar Al2O3 / (CaO + Na2O + K2O)] vs. SiO2 (wt.%)
from Eräjärvi are marked by heavy dots.
3.8.2. Geochemical constraints on origin
and Shebanov, 2005). Lindberg and Eklund (1988)
compared the geochemical features and contact
Geochemically, the postorogenic rocks of
southern Finland cover a wide range in SiO2 (from
32 wt.% to 78 wt.%; Eklund et al., 1998) and they
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. Prior to emplacement, the
magmatic evolution of the complex may have
been controlled by fractionation processes in a
midcrustal chamber (Eklund and Shebanov, 2005).
According to Bergman (1986), the Åva monzonite
intruded a previously emplaced lateorogenic granite
as branching concentric dikes by stoping, and the
subsequent granite widened the funnel laterally.
The intrusion mechanism of the Åva ring complex
have high K, Ti, and Ba contents and are strongly
enriched in the incompatible elements; the more
and the REE (Fig. 34; Nurmi and Haapala, 1986;
the postorogenic rocks of southern Finland as
shoshonitic, largely on the basis of 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 (cf.
Figs. 28 and 34). Lahtinen (1994) proposed fractional
crystallization and assimilation of mantle-derived
alkali basaltic magmas as the primary processes that
controlled the evolution of these rocks, and favored
an enriched subcontinental lithospheric mantle
source. Eklund et al. (1998) suggested carbonate
metasomatism as the reason for enrichment of
1987) imply emplacement of these ∼1800-Ma
rocks into rigid country rocks at a shallow depth (a
few km). In contrast, a thermobarometric study of
the 1815-Ma postorogenic rocks in the Turku area
indicates an emplacement pressure of ~4 kbar,
corresponding to a minimum depth of 14−15 km
(Väisänen et al., 2000).
48
the lithospheric mantle source and concluded that
metasomatism was more extensive in the east.
Nd isotope data on the postorogenic granitoids
are relatively few. The εNd (at 1800 Ma) value of
the granite of the Åva ring complex is +0.2 and the
depleted mantle model age is 2.02 Ga (Patchett and
Kouvo, 1986). The εNd (at 1800 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. Our recent (still
unpublished) data on the Renko quartz monzodiorite
stock and the Eräjärvi granite indicate that the former
is more juvenile [with an εNd (at 1815 Ma) value of
+1.3] than Åva, Parkkila, and Pirilä, whereas the
latter is quite similar [with an εNd (at 1800 Ma) value
of +0.4; see description of Stop 2/5] with them.
3.9. Rapakivi granites
(O.T. Rämö)
The Finnish word “rapakivi”means disintegrated
or crumbly rock and illustrates the tendency
of the rapakivi granites to weather more easily
than the other granitic rocks of Finland. In 1891,
J.J. Sederholm introduced rapakivi granite in
international geological literature in his pioneering
paper on the rapakivi granites of southern Finland.
Since then, southern Finland has been regarded as
the type area of rapakivi granite. In the 20th century,
a steadily increasing number of rapakivi granites
were described from around the world, especially
the Ukraine, South Greenland, eastern Canada,
mid-continental and southwestern United States,
and Brazil. Rapakivi granites are now known from
all continents (see reviews in Rämö and Haapala,
1995, and Haapala and Rämö, 1999). As currently
perceived, the majority of rapakivi granites are
mid-Proterozoic (1.8–1.0 Ga), but also some welldocumented Phanerozoic (cf. Calzia and Rämö,
2005; Haapala et al., 2005) and Archean (Sibiya,
1988; Moore et al., 1993) examples are known.
alkali feldspar megacrysts mantled by sodic plagioclase, two generations of alkali feldspar and quartz;
Vorma, 1976) is ubiquitous especially in voluminous batholiths. In order to take into account these
rapakivi granites as “A-type granites characterized
by the presence, at least in the larger batholiths, of
granite varieties showing the rapakivi texture”. This
sitional peculiarities, and magmatic association of
the rapakivi granites, but does not restrict their age.
3.9.2. Distribution and age
The classic rapakivi granites of Finland form a
substantial part of the Fennoscandian rapakivi association. Overall, the Fennoscandian intrusions are
mid-Proterozoic (1.67–1.47 Ga) and fall into four
age zones with a gross east–west pattern (Fig. 36).
The Wiborg batholith and associated plutons and
dike swarms in southeastern Finland and Estonia are
1.67–1.62 Ga old. Those in southwestern Finland,
Latvia (the huge Riga batholith), and west-central
Sweden (the Nordingrå complex) are dated at 1.59–
1.54 Ga. The rapakivi granites in central Sweden
(west of Nordingrå) are 1.53–1.47 Ga, and those in
Russian Karelia (the Salmi complex) 1.56–1.53 Ga.
3.9.1. General mode of occurrence, geochemistry,
The rapakivi granites are typically found as discordant plutons that were intruded into a metamorphic crust that was differentiated from the mantle
a few hundred Ma earlier. Geochemically, the
rapakivi granites are subaluminous A-type (Fig. 35,
Table 5), with high contents of the HFSE and LREE
(except Eu), and a marginally metaluminous to peraluminous and reduced to oxidized character. Some
complexes include volcanic rock types and minor
peralkaline lithologic units. The magmatic association of rapakivi granites is bimodal, as featured by
AMCG (anorthosite–mangerite–charnockite–granite) complexes and spatially and temporally asso-
scarce at the present Precambrian erosional sur-
of the Wiborg batholith (the Ahvenisto complex),
in Russian Karelia, and in central Sweden (the Ragunda complex) (Fig. 36).
Geophysical studies (e.g., Elo and Korja, 1993;
Korja et al., 1993; Luosto, 1997; Korsman et al.,
1999) show that (1) the Fennoscandian rapakivi
granite batholiths are found as relative thin (~5–
10 km) sheet-like bodies in the upper part of the
49
Table 5
Representative chemical composition of the Finnish
rapakivi granites: Wiborgite and pyterlite from the Wiborg
batholith (southeastern Finland) and topaz-bearing granite
from the Eurajoki stock (southwestern Finland). Data from
Haapala et al. (2005)
sample
SiO2
Wiborgite
Pyterlite
Topaz granite
1A/IH/2001
2A/IH/2001
5/IH/2001
69.9
76.70
75.78
TiO 2
0.42
0.18
0.02
Al2O 3
13.41
11.70
13.62
FeO W
3.35
1.53
0.65
Fe2O 3
0.45
0.31
0.19
MnO
0.06
0.02
0.05
MgO
0.35
0.09
0.00
CaO
2.16
0.80
0.64
Na2O
3.09
2.49
3.70
K2 O
5.46
5.59
4.35
0.11
0.02
0.02
P 2 O5
H2 O
+
0.31
0.17
0.24
FW
0.24
0.43
1.19
Total
99.37
100.03
100.43
-O=F2
0.18
0.50
Total
99.27
99.85
Cl (ppm)
800
300
100
Rb
271
349
1050
crust, and (2) that the crust shows particularly steep
ovoid thinnings associated with the rapakivi intrusions. For example, beneath the Wiborg batholith,
the crust is ~40 km thick— 15–20 km thinner than
in the surrounding areas (Fig. 36). A further typical
feature is that the crust underlying the rapakivi inand thinned lower crust (Korja et al., 2001).
3.9.3. Petrology
All the large Finnish rapakivi batholiths and most
of the stocks are multiple intrusions consisting of
several granite types. Minor anorthositic and gabbroic bodies are found in several of the complexes.
The contact relations suggest that they are older than
the associated rapakivi granites, but there is no siget al., 1991; Suominen, 1991). The rapakivi granite types differ from each other in texture, mineral
composition, and chemical composition. Among
the oldest rapakivi varieties are fayalite-bearing
biotite–hornblende granites. The major intrusive
phases are commonly represented by wiborgite (a
biotite–hornblende granite that contains oligoclasemantled K-feldspar ovoids) and pyterlite (biotite
granite with unmantled ovoids) (Fig. 37). They are
not, however, invariably present in all the rapakivi
intrusions. The relative amounts of various evengrained and porphyritic biotite granites vary widely. The youngest intrusive phases are topaz-bearing microcline–albite granites that contain lithian
siderophyllite (or protolithionite) as the dark mica
(Haapala, 1997). Typical accessory minerals in the
granites of the early and main intrusion phases are
99.93
Sr
155
69
8
Ba
1144
541
28
Ga
27
24
60
Be
5
5
16
Zn
125
86
197
Sc
8
3
12
Zr
459
304
51
Hf
12.7
10
5.7
110
Sn
8
9
W
3
3
9
Nb
26.4
25.8
69.8
Ta
2.3
2.41
22.6
Th
24.7
38.6
27.6
U
8.08
14.7
6.34
Pb
38
55
88
Y
63
75
55
La
95.9
99
47.4
Ce
182
182
97.2
Pr
21.5
21.2
11.1
Nd
80.6
76.2
37.3
Sm
14.6
13.7
9.26
Eu
2.56
1.26
1.18
Gd
12.6
12.2
7.77
Tb
2.14
2.2
1.84
Dy
12.7
13.8
12.7
Ho
2.67
2.98
2.71
Er
7.26
8.51
8.6
Tm
1.07
1.3
1.77
Yb
7.31
8.56
15
Lu
1.05
1.23
2.3
and ilmenite (see Vorma, 1971, 1976). In the biotite granites, monazite is found instead of allanite.
In the late-stage granites, topaz, monazite, bastnaesite, ilmenite, cassiterite, and columbite are typical
heavy accessory minerals (Haapala, 1974). The rarity of pegmatites and miarolitic cavities, late crysreactions, and scarcity of hydrothermal veins and
mineral alterations indicate that the early intrusive
phases crystallized generally from water-undersaturated melts. The last intrusive phases crystallized
from water-saturated magmas (Haapala, 1997; Haapala and Lukkari, 2005).
The bimodal magmatic association of the rapakivi granites (granites, silicic dikes, anorthosites
50
ers to advocate magmatic underplating as the probable mechanism for the generation of these rocks
(Haapala and Rämö, 1999, and references therein).
This involves partial melting of the lower crust in
response to thermal perturbations associated with
the underplating. Petrologic studies of rapakivi
complexes in Finland and Sweden (Rämö, 1991;
Andersson, 1997; Kosunen, 1999, 2004) have favored dehydration melting of quartzofeldspathic
been discussed (Eklund et al., 1994; Salonsaari,
1995; Kosunen, 2004).
Overall, the rapakivi granites and the associated
of the composition of the unexposed lithosphere
in any one area. This is illustrated in Fig. 36 that
shows variation in the initial Nd isotope composidian rapakivi complexes. The Finnish, Estonian,
and Latvian complexes and the Nordingrå pluton in
Sweden (age groups 1.67–1.62 Ga and 1.59–1.54
Ga) have initial εNd values of –3 to 0. This is compatible with their derivation from the 1.9-Ga Svecofennian crust. The Salmi batholith in Russian
Karelia is quite different in having initial εNd values
of –9 to –5.5. This indicates a substantial Archean
component in the source and complies with the position of the complex between the Archean and Proterozoic crustal domains (Fig. 36). Similar, yet quite
surprising, is the Nd isotope composition of the
rapakivi complexes in central Sweden (Fig. 36; see
also Andersson et al., 2002). The initial εNd values
(–10 to –4.5) point to a considerable Archean domain at the lower crust–uppermost mantle level in
central Sweden. No Archean crust has been demonstrated from the exposed parts of the Precambrian
in this area.
Fig. 35. Composition of the Finnish rapakivi granites
(Rämö and Haapala, 1995) in (A) FeOtot/(FeOtot + MgO) vs.
SiO2
(B) (Na2O + K2O) / CaO vs. Zr + Nb + Ce + Y diagram of
Whalen et al. (1987); and (C) Chondrite-normalized REE
diagram. Composition range of ~500 A-type granites (as
referred to in Frost et al., 2001) is shown in (A). In (B) OGT
is unfractionated granites and FG is fractionated granites
(Whalen et al., 1987).
51
Fig. 36. (A) Map showing the distribution and ages of the rapakivi granite complexes and diabase dikes as well as contours of
crustal thickness in the south-central part of the Fennoscandian shield. The gray lines outline three rapakivi age zones (1.59–
1.54 Ga, 1.67–1.62 Ga, and 1.55–1.54 Ga). The inset shows the area in relation to the major crustal domains in the shield. TIB
is the Transscandinavian igneous belt. The map is from Rämö and Korja (2000) and Rämö et al. (2000), where the pertinent
references are given. (B) Diagram showing variation of initial Nd isotope composition (εNdi) in the rapakivi granites (open(DePaolo, 1981), the ~1.9-Ga Svecofennian rocks (Huhma, 1986; Patchett and Kouvo, 1986), and the Archean crust of the Fen-
52
Fig. 37. Photographs of two main textural types of rapakivi granite from the Wiborg batholith: (A) wiborgite, (B) pyterlite.
Adopted from Rämö and Haapala (1990).
53
4. Tectonic evolution: the granitoid perspective
(M. Nironen, P. Sorjonen-Ward, L.S. Lauri, O.T. Rämö, A. Käpyaho, J. Halla, M.I. Kurhila)
4.1. Introduction
of granitoids. Moreover, Frosterus and Wilkman
some plutonic rocks in southern Finland as
synkinematic, latekinematic, and postkinematic
(with respect to the Svecofennian orogeny), based
on degree of deformation, as well as compositional
differences. At the same time, Sederholm (1932)
divided the Precambrian rocks of Fennoscandia
into four sedimentation cycles, each related to a
distinct plutonic event. The plutonic rocks of the
the basement gneisses within the cover rocks, and
Wegmann (1927) presented a tectonic synthesis in
terms of Alpine collisional tectonics which remains
essentially valid. These studies formed the basis for
(1949), which considered the thermal and rheological
aspects of orogenic processes and granite genesis
through the concept of mantle gneiss domes.
Here we present a current overview of the tectonic
setting of the granitoid rocks in the Fennoscandian
shield and their relation to the evolution of the
craton (see Fig. 38). This will hopefully provide a
synkinematic rocks of Eskola (1932). Some granite
stocks in southern Finland and Lapland belong to
the third cycle, and the rapakivi granites in southern
Finland are part of the fourth cycle. Sederholm was a
strong advocate of actualist principles, yet it must be
understood that prior to the advent of isotopic dating
trip, particularly in comparing and contrasting
Archean and Proterozoic granitoids within their
respective orogenic frameworks, and whether or
not the sequential magmatic evolution in Finland is
cycles were regarded as Archean. Following the
pioneering isotopic studies of Kouvo and Gast
structure and boundary conditions peculiar to the
by Simonen (1960) distinguished a belt of latekinematic microcline granites in southern Finland;
he later divided the Svecofennian plutonic rocks into
synorogenic, lateorogenic, and postorogenic and
considered that the rapakivi granites are anorogenic
(Simonen, 1980). Later still, in a review of granitoid
petrogenesis and metallogeny, Nurmi and Haapala
generic processes applicable to other terrains.
Terminology for the structural and lithic units
issue complicated by the need to standardize names
derived from diverse linguistic backgrounds! It is
Gorbatschev (1987), to consider the Archean and
Paleoproterozoic history of the Finnish part of the
Fennoscandian shield in terms of three large crustal
domains— the Kola, Karelian and Svecofennian
domains (Fig. 1). 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 Archean (3.5–2.6 Ga)
cratonic nucleus exceeding 200,000 km2 in area in
eastern Finland and adjacent Russia (Figs. 1 and
latekinematic, postkinematic, and anorogenic.
Meanwhile, in eastern Finland, which will form
Frosterus and Wilkman (1920, 1924) had implicitly
and correctly recognized the presence of two
superimposed orogenic cycles, separated by a major
unconformity. The older granitoid group, which is
truly Archean granitoids and gneisses, together
with remnants of even older low-grade supracrustal
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.
basement to the overlying (Paleoproterozoic)
sedimentary sequences, which were in turn deformed
and metamorphosed and intruded by a younger suite
54
Fig. 38. Time scale for the main tectonic events and associated lithologic units in Finland.
55
4.2. Archean
The Karelian domain in Finland is characterized by
a number of narrow northerly trending low-pressure
greenstone and metasedimentary belts, 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 Kola domain is only represented in Finland by
a small, but complex area in the far northeast of the
country, including granitoid gneisses, migmatites,
charnockites, aluminous metasedimentary rocks,
and iron formations (Meriläinen, 1976; Gaál et al.,
1989; Kesola, 1995).
The Kola and Karelian domains are separated in
Finland by the Paleoproterozoic Lapland granulite
belt, which was tectonically emplaced over the
Karelian domain some time after 1.9 Ga (cf. Gaál
et al., 1989). In the adjacent Kola Peninsula, there
shows both greater complexity and diversity,
as indicated for example, by a distinctive suite
of Neoarchean alkaline intrusions and gabbro–
anorthosite intrusions (Zozulya et al., 2001).
However, relationships between major crustal units
are also somewhat better constrained by detailed
structural analysis and careful isotopic dating and
recent studies have demonstrated (Daly et al., 2001)
that the correlatives of the Lapland granulite belt
can also be traced across the Kola Peninsula towards
the White Sea, where they merge with a highly
complex gneiss terrain, along the northern margin
of this Belomorian province have long been
controversial, but recent studies have demonstrated
that while there is a strong thermal overprint of
Svecofennian age (Bibikova et al., 2001; Skiöld et
al., 2001), it originated as a Neoarchean collisional
the discovery of Archean eclogite facies gneisses
at Gridino on the White Sea coast (Volodichev
et al., 2004) provides evidence of the ability of
Archean lithosphere to sustain lithospheric loading.
Moreover, the presence of the intervening Inigora
ophiolitic assemblage (Shchipansky et al., 2004),
and the similarity in age (2.74–2.72 Ga) between
the Gridino eclogites and granitoid magmatism in
56
eastern Finland provides a new basis for assessing
the timing and nature of Archean crustal formation
processes.
The Archean history of the Fennoscandian shield
in Finland extends back into the Paleoarchean with
gneisses of trondhjemitic composition (Mutanen
and Huhma, 2003). Isolated occurrences of 3.0–3.2Ga tonalitic gneisses (Kröner et al., 1981; Paavola,
1986) or xenocrystic zircon grains in younger
plutons (Sorjonen-Ward and Claoué-Long, 1993)
are known from a number of locations throughout
the shield. However, much of crustal growth and
subsequent tectonic and thermal reworking of the
Archean of Finland can be constrained to within the
interval 2.88–2.63 Ga (e.g., Vaasjoki et al., 1993,
1999, and references therein; Hölttä et al., 2000a;
Evins et al., 2002; Lauri et al., 2005). So far it
between discrete terrains, whether on the basis of
age, metamorphic grade, or intensity and polarity
of deformation, although a general subdivision
is possible (cf. Sorjonen-Ward and Luukkonen,
be made between greenschist to lower amphibolite
facies supracrustal terrains, characterized by steep
structural enveloping surfaces of fabric lineations,
intruded by discrete plutons, which contrast with
generally gently to moderately dipping higher grade
gneiss terrains. This may represent a fundamental
rheological decoupling between upper and lower
crust, as observed in geological and seismic
interpretations of other late Archean terrains
(Drummond et al., 2000; Sorjonen-Ward et al.,
2002), but it is still unclear whether the high-grade
terrains record an extensional as well as contractional
kinematic history (cf. Sandiford, 1989).
Recent characterization of the thermal history of
medium-pressure (8–11 kbar) granulites in central
Finland (Hölttä, 1997; Hölttä and Paavola, 2000;
Hölttä et al., 2000a), combined with data from
lower crustal and mantle xenolith suites accessed
by Neoproterozoic kimberlites (Hölttä et al., 2000b;
Lehtonen et al., 2004), marks important progress
in constraining the lateorogenic thermal regime
and timing of crustally derived granite genesis and
mantle depletion events. Enderbitic rocks are dated
to 2.68 Ga, somewhat younger than the predominant
2.74–2.72 Ga ages for tonalitic to granodioritic
et al., 1999; Lauri et al., 2005; Käpyaho et al., in
review). Sanukitoid attributes are increasingly
recognized in plutonic rocks in the Archean of
eastern Finland, with varying modal compositions
and host-rock metamorphic grade (Halla, 1998;
Käpyaho et al., in review), as well as in adjacent
Russia (Lobach-Zhuchenko et al., 2000a, b,
2005). In some places, such as the Hattu schist
belt, in easternmost Finland, a close spatial and
temporal relationship between felsic volcanism,
sedimentation, deformation, and emplacement of
these granitoids has been demonstrated (O’Brien
et al., 1993; Sorjonen-Ward, 1993). Although there
rocks emplaced at higher crustal levels, whereas
some high-grade gneisses record ages as young as
2.63 Ga. The timing of postorogenic exhumation is
not constrained, except that the 2.61-Ga Siilinjärvi
carbonate complex, representing the youngest
Archean magmatic event and having been emplaced
in a brittle–ductile transitional environment, appears
to be discordant with respect to country rock isograd
trends, implying that metamorphic zonation at the
present erosion level is a latest Archean, rather than
younger phenomenon.
Chemical characteristics of volcanic sequences
in greenstone belts and granitoids intruding them
have been used for comparison with modern arc
magmatism (Martin et al., 1984; O’Brien et al.,
1993) and for proposing plate tectonic models
(Taipale, 1983, 1988; Martin et al., 1984; Piirainen,
1988). Martin (1986, 1987a,b) modeled source
compositions and melting processes for tonalites
and trondhjemites intruding and adjacent to the
Kuhmo greenstone belt, to support the hypothesis
that partial melting of amphibolite in subducting
slabs was feasible under an Archean geotherm. In
general, granitoid magmatism in the Kuhmo region
record a systematic evolution from tonalites to
high-Mg/Fe granitoid rocks and are followed by
leucogranites (Querré, 1985; Martin, 1985; Vaasjoki
distribution and polarity of supposed subduction
zones, the data are certainly consistent with rapid
crustal growth through construction of volcanic
compelling evidence for eruption of greenstones
in proximity to an older substrate is indicated by
the abundance of partially assimilated granitic
1999; Halkoaho et al., 2000), and the presence of
komatiitic dikes and layered sills in migmatites
adjoining the Kuhmo greenstone belt (Luukkonen,
1992).
4.3. Proterozoic
4.3.1. Rifting of the Archean craton
Rifting of the Archean craton continued intermittently to ~1.98 Ga as shown by a spread of high
precision ages for sills and dike swarms between
2.3–1.98 Ga ages (e.g., Vuollo, 1994; Vuollo et al.,
2000). At 2.1–2.0 Ga, rifting of the Archean craton
In the early Proterozoic (commencing from ~2.44
Ga), the Fennoscandian Archean craton experienced
a period of extension that was possibly initiated
by a mantle plume (e.g., Huhma et al., 1990). The
extension resulted in widespread incipient rifting
and faulting of the craton and created pathways for
sporadically felsic volcanism and deposition within
a marine setting. Final breakup occurred, probably
diachronously from north to south along the present western margin of the Karelian craton, between
2.0–1.95 Ga, culminating in the formation of the
Jormua ophiolite complex (Kontinen, 1987; Peltonen et al., 1998) and coeval picritic continental
~2.44 Ga age have been found scattered throughout
the northern Fennoscandian shield in Sweden, Finland, and northern Russia (e.g., Alapieti et al., 1990,
mas also resulted in generation of felsic magmas by
partial melting of the lower crust and, possibly, by
an AFC process (Lauri et al., 2005). The 2.44-Ga
silicic intrusions typically show little or no deformation during emplacement and are geochemically
of A-type (Luukkonen, 1988; Lauri and Mänttäri,
2002; Lauri et al., 2005; Section 3.4).
plume-related magmatism likely accompanied, if
not initiated continental breakup.
ration in that Archean zircons have recently been
tle section of the Jormua ophiolite complex (Pel57
tonen et al., 2003). By implication, the mantle rocks
represent late Archean subcontinental lithosphere,
exhumed from beneath the Karelian craton during
continental breakup. Accordingly, their highly depleted nature (Peltonen et al., 1998) is consistent
with their being complementary reservoirs to crust
formed during the late Archean, and also with the
depleted character of Paleoproterozoic tholeiitic
dike swarms.
ary, partially melted the lower crust, and generated
the synkinematic magmas by mixing and mingling of the crustal and mantle magmas. The thick,
derplate generated high-temperature low-pressure
metamorphism that culminated coevally with the
emplacement of the synkinematic magmas. The
hot underplate subsequently triggered partial melting of the lower crustal granulite that was left after
the extraction of the synkinematic magmas, thus
generating the postkinematic magmas of central
Finland.
4.3.2. Proterozoic orogenic granitoids
The Svecofennian orogeny
oid rocks of Finland (Nironen, 2005) is based on
new geochronological data and recent concepts of
the evolution of the Svecofennian but retains the
general terminology of Simonen (1980):
In Finland, the Svecofennian crust has three main
features that need attention when attempting to explain the tectonic evolution: (1) the magmatic rocks
have a narrow age range of 1.90–1.87 Ga, excepting
the 1.93–1.92-Ga tonalitic gneisses; (2) the metamorphism is of the high-temperature low-pressure
type which culminated at 1.88 Ga in central Finland
and at 1.82 Ga in southern Finland; and (3) the crust
is anomalously thick, ~60 km in central Finland, diminishing to 50 km toward southern and western
Finland.
The generation and emplacement of Paleoproterozoic Svecofennian granitoids in Finland has
been attributed to the progressive accretion of at
least three arc complexes against the Archean craton (Lahtinen, 1994; Nironen, 1997; Korsman et al.,
1999). According to these models, the oldest granitoids, the 1.93–1.92-Ga tonalitic gneisses, were
probably generated when a rather primitive island
arc accreted to a more evolved arc with a ~2.0-Ga
nucleus in an oceanic environment. These arcs accreted to the Archean margin ~1.91 Ga ago; now
they form the Primitive arc complex and Arc complex of western Finland (see Figs. 1 and 26).
According to the models of Lahtinen (1994) and
Nironen (1997), another arc complex approached
from the (present) south where an oceanic basin
was narrowing by subduction in opposite directions (Fig. 39). By 1.89 Ga, this basin had closed
by subduction, and the arc complex (Arc complex
of southern Finland) accreted to the previously
accreted one (Arc complex of western Finland).
Continued convergence caused thickening of the
lithosphere which became gravitationally unstable.
The base of the lithosphere was delaminated and
compensated by hot asthenosphere. Decompression
melting of the asthenospheric material produced
(1) Preorogenic rocks were generated in an island arc environment and were placed in their
present location during the Svecofennian orogeny. Preorogenic (1.93–1.92 Ga) granitoid rocks
are found within the Primitive arc complex.
(2) The synorogenic stage lasted from 1.89 Ga
to 1.87 Ga as deduced from zircons dated from
rocks of southern Finland (Vaasjoki, 1996). The
emplacement and deformation of the synorogenic rocks is assigned to the accretion of three
arc complexes to the Archean craton. 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.
(3) The lateorogenic granites of southern Finland are located within the southern Finland arc
complex and are associated with metamorphism
and low-angle crustal movements 1.84–1.82
Ga (or 1.85–1.80 Ga; see Section 3.6) ago. The
granites and associated migmatites form the
Late Svecofennian granite–migmatite zone.
(4) The postorogenic plutons of southern Finland
are located within the late Svecofennian granitemigmatite zone but, contrary to the lateorogenic
granites, these rocks were emplaced along faults
and shear zones within a largely stabilized crust
at 1.81–1.76 Ga.
58
Initial accretion of two amalgamated arc complexes 1.91 Ga ago to the Archean craton margin caused thickening of the crust
and the underlying mantle lithosphere. Collision of another arc complex against the accreted ones 1.89 Ga ago caused continued lithospheric thickening and eventually led to delamination of the base of the lithosphere and compensation by hot asthe-
prisms and caused subsequent partial melting of these rocks (lateorogenic magmatism). Thick, dense lower crust maintained
the thickened crust in isostatic balance. CFGC— Central Finland granitoid complex.
Zircon data from plutonic rocks indicate an
age gap at 1.86–1.85 Ga (Vaasjoki, 1996) and the
youngest detrital zircons in quartzites of southern
Finland are around 1.87 Ga (Lahtinen et al., 2002).
Ehlers et al. (1993) concluded that the lateorogenic
granites of southwestern Finland were emplaced at
1.84–1.83 Ga during transpressional deformation
with thrusting from the south–southeast, whereas
Korja and Heikkinen (1995) presented an extensional model for the lateorogenic granites.
craton (Karelian and Belomorian provinces) collided with the Archean Kola province, producing
the high-grade Lapland granulite belt in the suture zone and inducing extensive melting in the
central Lapland area. It has also been proposed
that the northwestern part of the Karelian province was previously (before 2.0 Ga) rifted from
the main continent and collided back to the present location slightly before the collision with the
Kola province occurred (see, e.g., Lehtonen et
al., 1998).
The granitoids of northern Finland can also be
divided into preorogenic, presumably preorogenic, synorogenic, lateorogenic, and postorogenic
groups, although the age groups differ from those
in the Svecofennian domain:
Collisional tectonics in northern Finland
While the Svecofennian arc accretion was taking place in the present central and southern Finland, the present northeastern part of the Archean
59
greenstone belt and the exposed Archean rocks
of the Ranua and Kianta terrains of the Karelian
craton. It is apparent from isotopic studies (Lauerma, 1982; Huhma, 1986) that there is an Archean source component to these rocks, whereas
emplacement ages may be as young as 1.82 Ga,
comparable with extensive granitoid complexes
in adjacent Sweden (Ahtonen and Mellqvist,
1997; Sorjonen-Ward and Luukkonen, 2005).
Potassic to monzogranitic dike networks within
an Archean substrate are typical, and recent studies have focused on attempting to characterize
chemical compositions with petrophysical attributes (Airo and Ahtonen, 1999). Similarly
Puranen (1989) speculated that the ubiquitous
ferromagnetic magnetite, and relatively low Fe
contents of these granitoids hinted at a distinctly
oxidized source terrain.
(with respect to Svecofennian orogeny) are found
in both the Lapland granulite belt, which is in
large part derived from Paleoproterozoic sources
(Meriläinen, 1976; Sorjonen-Ward et al., 1994),
as well as within the adjacent Neoarchean Inari
terrain. This plutonic activity, dated at around
1.95–1.91 Ga, has been interpreted as a magmatic
arc recording closure of an ocean basin, although
opinions concerning polarity, precise timing and
relationships between formation and exhumation of the Lapland granulite belt diverge widely
(Hörmann et al., 1980; Barbey et al., 1984; Gaál
et al., 1989; Daly et al., 2001).
(2) In northwestern Finland, in the Kittilä–Enontekiö area, the so-called Hetta complex includes
a range of rock types and represent the rifted part
of the northwestermost Karelia. The intrusive
(5) Plutonic complexes that are clearly discordant
and postorogenic with respect to deformation in
Lapland include the 1.78-Ga Nattanen suite and
Vainospää granite (Haapala et al., 1987). These
although the most common rock types are tonalplutons. Very little is known of their tectonic setting but the heterogeneity of zircon populations
and diverse nature of country rocks suggests that
widespread inheritance frustrates the determination of precise emplacement ages (cf. Huhma,
partial derivation from Archean lithosphere and
were emplaced at high crustal levels, effectively
delineating the distribution of Archean rocks at
depth and providing a constraint on rates of postorogenic denudation, including exhumation of
the Lapland granulite belt.
(3) Synorogenic rocks are characterized by the
1.89–1.86-Ga Haaparanta suite. They are found
in western Lapland on both sides of the Finnish–Swedish national border. The Haaparanta
suite consists of two groups, monzonitic (gabbro,
quartz monzodiorite, monzodiorite, monzonite,
and quartz monzonite) and trondhjemitic (quartz
diorite, tonalite, trondhjemite, and granodiorite;
Lehtonen, 1984). Although emplacement ages
are well constrained, both isotopically and from
Mid-Proterozoic rifting
Another rifting event of the Fennoscandian
shield at ~1.65–1.57 Ga, long after the tectonics of
the Svecofennian orogeny had ceased, resulted in
the emplacement of voluminous rapakivi granites
in southern Finland (e.g., Rämö and Haapala, 1995,
and references therein). This process was related
to a long-lasting magmatic underplating episode
in the subcontinental mantle involving generation
of both mantle and crustal melts. The ultimate tectonic cause of the magmatic underplating remains
controversial, however. Plausible mechanisms include active or passive rifting, extensional collapse
of orogen, and deep mantle plumes. Petrologically
unstable domains in the lithospheric mantle (related to earlier or contemporaneous distant subduction zones) could also have controlled the loci
of magmatism (Haapala and Rämö, 1992; Rämö
and Haapala, 1995). The Fennoscandian rapakivi
granites have recently been related to intermittent
strong isotopic evidence of derivation from an
Archean source (Hiltunen, 1982; Lehtonen,
1984; Huhma, 1986; Perttunen, 1991; Väänänen,
1998; Perttunen and Vaasjoki, 2001; Väänänen
and Lehtonen, 2001). These rocks therefore have
evidence for calc-alkaline magmatism within the
former Karelian cratonic margin in northern Finland, in contrast to the situation farther south.
(4) Lateorogenic granites are present throughout
much of southern Lapland, between the Lapland
60
Fennoscandian shield (Åhäll et al., 2000). Roughly
north–south trending magmatic arcs related to this
postulated mechanism are located on the eastern
Fennoscandian rapakivi granite complexes (Fig.
36). Thus a model of extensive long-time mantle
upwelling and resultant periodic, migrating melting
Fig. 36). This “inboard model”is, however, unable
to account for the non-linear age distribution of the
and deep crustal structures, remains the most plausible scenario (cf. Haapala et al., 2005).
61
62
PART II: FIELD TRIP STOP DESCRIPTIONS
63
64
DAY 1 (Monday September 12, 2005)
Guide: O.T. Rämö
Rämö, 1991; Vaasjoki et al., 1991; Jaala–Iitti—
Salonsaari and Haapala, 1994; Salonsaari, 1995;
Ahvenisto— Alviola et al., 1999).
The Wiborg batholith itself (Fig. 40) is
key rock types in the classic Wiborg rapakivi granite
batholith of southeastern Finland. The batholith
covers circa 19,000 km2 and forms a relatively thin
of the crust. Roughly two-thirds of the batholith is
situated in the Finnish territory, the remainder is
in northwesternmost Russia, including the town of
Vyborg (or Wiborg) according to which the batholith
was originally named.
Roughly one third of the batholith is located beneath
the Gulf of Finland (Fig. 2; Koistinen, 1994). The
Finnish mainland part of the batholith was mapped
on 1:100,000 scale in the 1960’s and 1970’
s by the
Geological Survey of Finland (e.g., Vorma, 1972;
Simonen, 1987) and is currently the petrologically
best-known part of it. Subsequently, detailed
petrographic and geochemical studies have been
carried out on three separate rapakivi complexes on
rocks (gabbroids, anorthosite, diabase dikes), of the
rapakivi association are subordinate. Roughly 80%
of the Finnish part consists of wiborgite (or dark
wiborgite). The remainder is pyterlite, porphyritic
granite (biotite granite with angular alkali feldspar
megacrysts), equigranular fayalite–biotite granite
(tirilite), equigranular hornblende, hornblende–
biotite, and biotite granite, porphyry aplite (leucocratic
granite with occasional alkali feldspar megacrysts),
and topaz-bearing alkali feldspar granite.
On this trip, we will examine four rock types in the
central part of the batholith: a topaz-bearing alkali
feldspar granite and associated pegmatite, wiborgite,
pyterlite, and anorthosite and leucogabbronorite
including iridescent plagioclase (spectrolite).
Stop 1/1: Topaz-bearing alkali feldspar granite and associated marginal pegmatite
least slightly) older than the marginal equigranular
granite and stockscheider pegmatite (Haapala and
Lukkari, 2005). In terms of Nd isotope composition
(Rämö, 1991), the granites are identical: εNd (at 1640
Ma) values are –0.2 ± 0.5 (porphyritic granite) and
–0.3 ± 0.6 (equigranular granite). These initial ratios
Kymi stock ~10 km north of the town of Kotka
(Fig. 40). The Kymi stock is an oval, relatively
small (6 km by 2.5 km) cupola-shaped intrusion of
topaz-bearing alkali feldspar granite that sharply
cuts the wiborgite, pyterlite, and porphyry aplite of
the Wiborg batholith. The stock (Fig. 41) consists
of two rock types (Haapala, 1977; Haapala and
Lukkari, 2005): a porphyritic central granite and an
equigranular marginal granite. Both are leucocratic
alkali feldspar granites with Li-enriched dark mica
and accessory topaz; the tin content of the micas is
250–300 ppm (Haapala, 1977). Accessory minerals
include monazite, columbite, bastnaesite, thorite,
molybdenite, and pyrochlore. Geochemically, the
two granites are peraluminous and anomalously
high in F, Li, Be, Ga, Rb, Sn, and Nb and low in
Ti, Fe, Mg, Ba, Sr, and Zr (Haapala, 1977; Rieder
et al., 1996; Haapala and Lukkari, 2005), with
the equigranular granite showing a clearly more
anomalous character than the porphyritic granite.
The two granites do show a sharp mutual contact
and the central porphyritic granite appears to be (at
been measured for wiborgite in the south-central
part of the batholith (e.g., Rämö, 1991).
At the contact between the Kymi stock and the
surrounding wiborgite and pyterlite, a <5m-thick
zone of topaz-rich pegmatite granite and pegmatite
(stockscheider) is present (Haapala, 1977; Kaartamo,
1996; Kaartamo et al., 1996). The main minerals in
the stockscheider are K-feldspar (often amazonitic),
quartz, albite (in two generations), biotite (mainly
as extensive dendritic clusters), topaz, tourmaline,
pegmatite belongs to the NYF-class of erný (1991)
(Kaartamo et al., 1996) and has been quarried for
gem-quality topaz.
Both the granites of the Kymi stock and the
65
The Suomenniemi (Rämö, 1991), Ahvenisto (Alviola et al., 1999), Jaala–Iitti (Salonsaari, 1995), and Onas complexes are also
66
with the most intensively mineralized veins farthest
away from the stock (e.g., Haapala, 1977).
On the trip we will examine both granites and the
stockscheider that are quarried in the southeastern
part of the stock just east of highway 357.
surrounding older rapakivi granites host greisen and
quartz veins as well as irregular greisen bodies. The
veins and bodies are associated with arsenopyrite,
wolframite, beryl, genthelvite, as well as Pb, Zn,
and Cu sulphides, and show sings of lateral zoning
Stop 1/2: Wiborgite at Summa, Vehkalahti, south-central part of the Wiborg batholith
A glacially-polished outcrop of wiborgite (the
rapakivi granite proper) by a freeway ramp 3 km east
of downtown Hamina. The rock is loaded with alkali
feldspar megacrysts 1 cm to 10 cm in diameter; most
of them are mantled by an oligoclase rim. There is,
however, a tendency for the largest ovoids to lack
mantles. Some megacrysts show multiple rims and,
occasionally, oligoclase mantles are also found
around rounded microgranite fragments. Quartz
occurs in two generations: (1) short-prismatic small
phenocrysts (pseudomorphs after high quartz) and
silicates are hornblende and biotite. The wiborgite
also contains narrow pegmatitic veins and vugs (~15
cm in diameter) with coarse quartz, alkali feldspar,
The rock is fresh and good for sampling, yet the
glacially-polished surface displays the rapakivi
texture in a superb way as the oligoclase mantles have
turned white owing to post-glacial weathering.
Stop 1/3: Pyterlite at Virolahti, south-central part of the Wiborg batholith
Extensive dimension stone quarry (Haikanvuori)
in pyterlite by the Bay of Virolahti ~5 km west of the
Russian border (Fig. 42). This pyterlite is texturally
similar to the wiborgite inasmuch as the size and
form of the alkali feldspar ovoids and the conspicuous drop quartz; only some of the alkali feldspar
67
Fig. 42. Geological sketch map of the Virolahti area in the southeastern part of the Wiborg batholith. Pyterlite quarries (including the target of STOP 1/3) are indicated.
megacrysts are mantled by oligoclase, however.
The rock is also more silicic than the wiborgite
Kinnunen et al., 1987). The pyterlite is also cut by
pegmatitic dikes (a couple of meters wide) with
secondary calcite. Pyterlite at Virolahti has been
quarried for dimension stone purposes since the
16th century; in the 19th century much of the rock
was shipped to St. Petersburg. The current commercial name for the Virolahti pyterlite is Carmen
Red.
Characteristic features of the Virolahti pyterlite
include pegmatite bodies and miarolitic cavities
that contain, besides alkali feldspar and quartz,
and some topaz and beryl but no muscovite (e.g.,
Stop 1/4: Anorthositic raft in wiborgite at Ylijärvi, Ylämaa, east-central part of the
Wiborg batholith
In the east central part of the Wiborg batholith
at Ylämaa, a couple of large anorthosite rafts are
found in wiborgite (Fig. 43). The rafts consist of
anorthosite and leucogabbronoritic rock types and
characteristically contain spectrolite (iridescent lab-
radorite; e.g., Lahti, 1989), currently exploited as
gem stone and facing stone material. We will examine a spectrolite quarry in the westernmost raft
(please exercise extreme caution— the quarry walls
are very loose). In the quarry, three gabbroic/an68
Fig. 43. Geological sketch map of the Ylämaa area in the east-central part of the Wiborg batholith. Ylijärvi anorthositic raft (the
target of STOP 1/4) is indicated.
orthositic rock types are present: coarse-grained
anorthosite (plagioclase adcumulate) with minor orthopyroxene, clinopyroxene, oxide, and
apatite is cut by (1) a coarse-grained leucogabbronorite and (2) a plagioclase-porphyritic leu-
U–Pb zircon age of the anorthosite is 1633 ± 2
Ma (Suominen, 1991; Vaasjoki et al., 1991) and
the initial εNd value is –0.4 ± 0.6 (Rämö, 1991),
the associated granite has been dated at 1633 ±
5 Ma (Suominen, 1991) and it has an initial εNd
value of –2.0 ± 0.4.
After examination of the quarry, we will stop at
the Ylämaa gem museum by highway 387 to take
a closer look at spectrolite and different commercial brands of wiborgite.
groudmass including late K-feldspar replacing
plagioclase. The gabbroic and anorthositic rocks
are cut by granite pegmatite dikes presumably
related to the surrounding rapakivi granite. The
69
70
DAY 2 (Tuesday September 13, 2005)
Guides: M.I. Kurhila, M. Nironen, O.T. Rämö
tion in the physical appearance of the lateorogenic
the granites of the easternmost part of the lateorogenic granite zone (Fig. 44). In this part the youngest intrusions of the whole belt are present; the emplacement ages are about 1.80 Ga or even younger
(cf. Fig. 45). The granites illustrate well the varia-
the geochemical and isotopic characteristics (cf.
Figs. 46 through 48). The last stop of the day will
be a ~1.79-Ga postorogenic monzogranitic dike in
Ruokolahti.
Stop 2/1: Keittomäki lateorogenic granite, Juva
A roadcut about 20 km west of Sulkava town
along highway 436. This is a grayish garnet-bearing
microcline granite. Mineralogically, the Keittomäki
granite deviates from other lateorogenic microcline
granites in that it has no monazite and only a little
zircon. It is richer in plagioclase and its microcline
structure is less organized than in the other granites of the group. Anatase and spinel group oxides
are abundant accessory minerals. However, the
compositional, essentially low-angle layering is
very characteristic of the lateorogenic granites, as
are the abundant stratiform aggregates of relatively
large garnet grains. Orientation (due to biotite and
cordierite lineation) is generally weak and varying.
The granite contains gneissic xenoliths and garnet
clusters, and the grain size ranges from small to
very coarse. There are also cross-cutting pegmatite
dikes.
Source rock attributes cannot be unequivocally
determined on the basis of geochemical and mineral data alone. Although the rock has many features
typical of S-type granites (e.g., A/CNK value 1.13,
high SiO2 content, abundant garnet and muscovite),
many features militate against a metapelitic origin
(absence of monazite, low Th, elevated Ni and V
content).
An ion microprobe study of the zircons from this
granite (Matti Kurhila, unpublished data) shows bimodality in terms of age. The majority of the spots
form a somewhat scattered cluster between 1.87 Ga
and 1.86 Ga. The younger population yields an age
of about 1.79 Ga, which we interpret as the cooling
age of the batholith. The granite is associated with
widespread high-temperature, low-pressure metamorphism at 1.80–1.79 Ga and it is probable that the
older U–Pb zircon age represents the main phase of
the Svecofennian orogeny. The detrital zircons in
the Svecofennian metasedimentary rocks typically
show a wide age range extending into the Archean
(e.g., Claesson et al., 1993; Lahtinen et al., 2002),
whereas zircons from the Keittomäki granite do not
display such a broad range of ages.
Whole-rock Nd isotope compositions of the
granite (O. Tapani Rämö, unpublished data) accord
with those of the lateorogenic granites farther to
the west. The TDM model age (DePaolo, 1981) of
the rock is 2.19 Ga and the initial Nd (at 1790 Ma)
value –0.9.
Stop 2/2: The marginal phase of the Puruvesi granite, Herttuansaari, Kerimäki
Outcrop on the shore of Lake Puruvesi near a dirt
road. The Puruvesi granite batholith is located between the Archean Karelian craton and the Arc complex of southern Finland (labeled K and SAC, respectively, in the inset of Fig. 44). The batholith has
a concentric structure, the margin being quite leucocratic and inhomogeneous compared to the granite of
the central type (see description of Stop 2/3).
The granite is very felsic, consisting mainly of
quartz and microcline. The rock is generally me-
dium-grained, but in places gradually coarsens into
pegmatite grade; i.e., the pegmatite does not form
distinct cross-cutting dikes. Mica gneiss forms
patchy inclusions, with a preferred northwest orientation. Garnet is scarse, but locally forms rather
dense aggregates; biotite and muscovite are also
present. As much as 78% of the rock consists of
SiO2, and surprisingly, Na level is elevated, which
cordingly, there is a weak positive Eu anomaly (Fig.
71
47). The rock is only slightly peraluminous and its
Fe/Mg ratio is extremely high for an orogenic granite (Fig. 46).
Both the zircons and monazites record a ~1.80
Ga age for this granite. The zircons have inherited
cores with a variety of Neoarchean and Paleoproterozoic ages, which indicate that the granite (or its
precursor) was derived from a sedimentary source.
However, the prominent 2.0–1.9-Ga population of
many Svecofennian metasediments (Lahtinen et al.,
2002) is almost totally absent.
crustal origin for this granite. The initial
Nd
value
er crustal material. The TDM model age (DePaolo,
1981) is 2.53 Ga (Fig. 48).
of Day 2 are labeled with stars. The letters in the index map are: SAC— Arc complex of southern Finland, WAC— Arc complex
al. (1997).
Stop 2/3: The central phase of the Puruvesi granite, Rastiniemi, Kesälahti
A small outcrop on the east coast of Lake Puruvesi, next to a ferryboat wharf. The granite is
grayish, slightly porphyritic, homogeneous, fresh
and undeformed. The K-feldspar megacrysts display a weak magmatic orientation. Two micas are
present, biotite in much larger quantities. Garnet is
virtually absent, although some of it is found in the
each other and with the granite of the marginal
type (Stop 2/2), i.e., 1.80 Ga. Again, the inherited zircons display Archean and Paleoproterozoic
ages. The inheritance pattern is very similar to that
of the granite at Stop 2/2, which, together with the
same emplacement age, indicates that the granites
were derived essentially from the same source.
The Nd isotopes tell the same story. Initial εNd
values are on the order of –6.2, complying with
the marginal phase. The TDM model age (DePaolo,
1981) is 2.49 Ga, suggesting that Archean basement lies beneath the Proterozoic Puruvesi granite
area.
fairly rich in alkalis, A/CNK ratio being 1.08 (Fig.
46). There is a strong HREE depletion and a negative Eu anomaly, suggesting that both garnet and
plagioclase were retained in the source.
The zircon and monazite ages are identical with
72
Fig. 45. U–Pb ages of the granites of Day 2. Open symbols mark ion microprobe data and closed symbols multigrain isotope
dilution data. Error bars are at 2 level. Unpublished data by Matti I. Kurhila.
Fig. 46. FeOtot/(FeOtot+MgO) vs. SiO2 (wt.%) and A/CNK vs. SiO2 (wt.%) variation diagrams of the granites of Day 2. The
postorogenic Eräjärvi granite is petrologically clearly different from the lateorogenic granites. A/CNK is molar Al2O3 / (CaO +
Na2O + K2O). Unpublished data by Mikko Nironen and Matti I. Kurhila. Boundary of I- and S-type granitoids in the A/CNK
[molar Al2O3 / (CaO + Na2O + K2O)] vs. SiO2 (wt.%) diagrams after Chappell and White (1974). Line separating the ferroan
73
Fig. 47. Chondrite-normalized REE diagram for the granites of Day 2. Unpublished data by Mikko Nironen and Matti I.
Kurhila.
Stop 2/4: The Valkamo layered granite, Imatra
A roadcut near the Russian border showing a
good example of a layered lateorogenic granite
with alternating leucocratic and darker parts.
There are pegmatite dikes that deflect the weak
foliation in the granite, but the dikes are crosscutting rather than gradational in character. In
addition there are abundant garnet aggregates,
both stratiform and randomly dispersed. The
grain size varies along with the composition. The
darker layers show a preferred orientation caused
by pinitized cordierite grains, and the leucocratic
vein-like banding is often parallel to the orientation. The veins themselves have no orientation.
On the basis of field observations it seems that
the leucocratic parts represent low-temperature
in situ melting of the parent granite. Geochemical evidence supports this, as the veins are very
poor in REE (Fig. 47) and have a more evolved
major element composition (Fig. 46).
The monazites of the leucocratic parts record
an age of 1.80 Ga, similar to the granites of the
Lake Puruvesi (Stops 2/2 and 2/3). A bulk TIMS
U–Pb analysis of the zircons is to some extent
in accordance with this, although the results are
very discordant and some fractions indicate zircon inheritance with 207Pb/ 206Pb ages over 1.9
Ga.
The initial εNd value of the melanocratic granite is –1.4, and thus, the source rock was obviously different from that of the Puruvesi granites.
Although the zircon U–Pb heterogeneity clearly
indicates inheritance, the initial Nd isotope composition implies only a moderate input of older
crustal material. Due to a rather high Sm/Nd ratio, the TDM model age (DePaolo, 1981) of the
granite is also relatively high, 2.48 Ga.
74
Stop 2/5: The Eräjärvi postorogenic granite dike, Ruokolahti
46 and 47). It is slightly peraluminous and enriched
in Ti and P (Nykänen, 1988). The bimodal magmatism commonly associated with postorogenic intrusions (Eklund et al., 1998) is also prominent in the
area. Numerous small cogenetic lamprophyre dikes
cross-cut the gneissic country rocks, but we will not
Ourcrops and a roadcut by a sealed road number
438. In the early 1940’s, this rock was quarried for
anti-tank barriers. The intrusion is a small sheetlike body that cuts the synorogenic gneissic rocks
sharply; both the dike and the country rocks can
be examined at this stop. The dike is ~4.5 km long
and 300 m wide and trends northeast. The only sign
of deformation is minor faulting perpendicular to
strike. On aeromagnetic images, the dike appears
as a distinct positive anomaly. The contact with the
country rock is visible, and a ~0.5 m wide darker,
biotite- and plagioclase-rich contact variety can be
seen.
The rock is medium-grained and in places there
are scattered alkali feldspar and plagioclase phenocrysts. The color of the rock varies between red and
gray. Locally, small (2–5 cm in diameter) garnet–
cordierite gneiss xenoliths may be observed. Plagioclase (An20), alkali feldspar, quartz and biotite
are the main minerals, chlorite, muscovite, apatite,
A zircon age of 1792 ± 5 Ma (Nykänen, 1988;
recalculated by M. Vaasjoki) has been obtained for
this granite. The result partly overlaps with the ages
of the lateorogenic granites further northeast. A
whole-rock Nd isotope analysis gave an Nd (at 1792
Ma) value of +0.3 (Fig. 48). This demonstrates that
the postorogenic magmatism within the arc complex of southern Finland is slightly more juvenile
compared to the lateorogenic one in the same area.
The host rock is a migmatitic gneiss that has been
metamorphosed in high-T, low-P granulite facies.
The assemblage with abundant sillimanite and corgarnet = cordierite + K-feldspar + melt. The folding in the gneiss has deformed at least one foliation,
porphyroblasts, and leucosome veins.
Geochemically, the Eräjärvi granite displays
typical features of the postorogenic granites (Figs.
Fig. 48. Nd isotope evolution trends of the granites of Day 2. The depleted mantle line is from DePaolo (1981), CHUR denotes
Chondritic Uniform Reservoir (DePaolo and Wasserburg, 1976). Unpublished data by O. Tapani Rämö.
75
76
DAY 3 (Wednesday September 14, 2005)
Guides: M. Nironen, B.A. Elliott, O.T. Rämö
pyroxene-free hornblende quartz monzonite/granite forming the bulk of the interior part of the pluton (Fig. 49). In the east-central part of the pluton,
evolved, relatively felsic granite is found and cut
by aplitic microgranite dikes. The U–Pb age of
the Jämsä pluton is 1878 ± 15 Ma (Rämö et al.,
2001a), the εNd (at 1875 Ma) values range from
–0.1 to –0.4, and the TDM model ages (DePaolo,
1981) from 2.18 to 2.22 Ga. Rb–Sr data on four
samples from the pluton (two from the margin,
two from the center) indicate a remarkably low
initial 87Sr/86Sr of 0.7030 ± 0.0009 (Fig. 50). Furthermore, Rb–Sr data on the aplitic microgranite
cutting the monzogranite in the central part of the
pluton suggest that the Rb–Sr isotopic system was
stabilized ~1810 Ma ago.
ine various types of synorogenic rocks in the Central Finland granitoid complex: one synkinematic,
intensely foliated granodiorite, and two postkinematic intrusions. In particular, we will study the petrographic and geochemical changes in the postkinethe Central Finland granitoid complex (see Elliott
et al., 1998; Rämö et al., 2001a; Elliott, 2003; Fig.
49). The synkinematic country rock granitoid of the
monzonites and granites that constitute the pluton
itself thereafter from margin to center.
The Jämsä pluton covers an area of ~50 km2 and
consists of a 0.5–1 km wide margin of pyroxenebearing quartz monzonite that grades into coarse
Stop 3/1: Type 2 Puula pluton at Sokkasenmäki
The outcrop is a dimension stone quarry 200 m
east of a dirt road in the southeastern part of the
Island of Väisälänsaari. It represents the relatively
large (~450 km2) and homogeneous Puula pluton
dikes as well as relatively large plutonic bodies of
intermediate composition are found on the glacially polished outcrops. The main minerals are alkali
feldspar, plagioclase, quartz, and biotite, with accessory hornblende, apatite, and zircon. The SiO 2
of the Puula pluton is from 62 wt.% to 65 wt.%,
A/CNK ranges from 1.02 to 1.03 and Fe/(Fe+Mg)
from 0.82 to 0.85, the εNd (at 1875 Ma) value is
–0.5, and the TDM model age (DePaolo, 1981) is
2.17 Ga. The pluton has a heterogeneous zircon
polulation, and has yielded a 207Pb/206Pb age of
1891 ± 1 Ma (Rämö et al., 2001a).
granitoid complex. The gray, coarse-porphyritic
granite is non-foliated and contains abundant subhedral to anhedral alkali feldspar megacrysts 1–4
cm in diameter, and smaller (1–2 cm) plagioclase
megacrysts. Some of the alkali feldspar megadiate enclaves, schist xenoliths, and intermediate
Stop 3/2: Synkinematic granodiorite at Kollinkangas
The outcrop that has been quarried for road construction material is along a dirt road 2 km north
from Road 6031, 5 km west of the town of Jämsä.
The gray, medium-grained granodiorite is a typical
representative of a synkinematic rock of the CFGC
rock is strongly foliated and small shear zones deform the foliation. The main minerals are plagioclase, quartz, hornblende, and biotite, with accessory titanite, opaque minerals, apatite, and zircon.
The εNd (at 1875 Ma) value of the granodiorite is
–1.6 and the TDM model age (DePaolo, 1981) is
2.29 Ga.
these are clearly xenoliths of volcanic rock. The
77
Stop 3/3: Type 3 Jämsä pluton, margin
A roadcut by Road 604 just north of downtown
Jämsä showing the marginal assemblage of the Jämsä pluton, a dark brown, slightly porphyritic quartz
monzonite. The main minerals include perthitic or-
single grain plagioclase mantles have been observed
as in the classic rapakivi granites of southern Finland. The contact between the marginal assemblage
and the central Jämsä granite is gradational over a
small area (within meters), gradually losing olivine
and pyroxene to amphibole, and marked by a distinct
decrease in magnetic susceptibility from the margin
towards the center. The SiO2 content of the margin of
the Jämsä pluton is from 58 wt.% to 64 wt.%, the A/
CNK value is from 0.88 to 0.94, and the Fe/(Fe+Mg)
ratio is from 0.90 to 0.93. The εNd (at 1875 Ma) value
of this marginal quartz monzonite is –0.1 and the TDM
model age is 2.18 Ga.
silicates: ferro-edenitic to hastingsitic hornblende,
annite biotite, ferrosilite orthopyroxene, fayalitic olivine, and ferro-augite clinopyroxene. Olivine and
pyroxene are commonly seen as relicts surrounded
or being replaced by amphibole, but are also found as
independent grains. Biotite and amphibole are commonly interstitial, and sometimes found as small inclusions in larger alkali feldspar and plagioclase. Alkali feldspars are sometimes rimmed by a composite
78
Stop 3/4: Type 3 Jämsä pluton, center
in the marginal assemblage. Small aggregates of
A roadcut by Highway 9 ~5 km northeast of
downtown Jämsä. The central Jämsä granite is a
leucocratic (white to pink), porphyritic granite.
The main minerals are microcline, plagioclase,
quartz, ferro-edenitic to hastingsitic hornblende,
and annite biotite, with accessory titanite. Amphi-
out the central assemblage, becoming more common toward the evolved quartz-rich phase of the
pluton. The foliated central granite (presumably
magmatic foliation) is cut by aplitic microgranite dikes probably related to the pluton. The SiO 2
of the central granite ranges from 63 wt.% to 66
wt.%, the A/CNK value is from 0.90 to 1.01, and
the Fe/(Fe+Mg) ratio is from 0.88 to 0.89. The εNd
(at 1875 Ma) values of the central granite and microgranite are –0.4 and –0.6, respectively.
out the central granite, but biotite becomes more
common toward the center. Alkali feldspars surrounded by composite mantles of quartz and plagioclase are also found in the central granite, and
myrmekitic intergrowths are more common than
Fig. 50. Rb–Sr isochron
diagram for the Jämsä postkinematic pluton. See text for
details. Data from Rämö et al.
(2001a).
Stop 3/5: Type 3 Jämsä pluton, evolved center
An outcrop in the forest in the east-central part of
the pluton south of Juoksulahti, east of Highway 9.
The evolved central assemblage of the pluton is a leucocratic, coarse grained, quartz-rich granite. The main
minerals include microcline, plagioclase, quartz, biotite, and minor amounts of hornblende, with accessory
titanite, and allanite and epidote associated with late
stage fracturing. Feldspars are heavily sericitized, biotites are highly altered, and myrmekitic intergrowths
central assemblage. These bands are rich in amphibole
and biotite, and contain more zircon, apatite, and Fe–Ti
oxides than the other rock types of the Jämsä granite.
have separated from the quartz-rich evolved granite
through local differentiation.
79
80
DAY 4 (Thursday September 15, 2005)
Guides: J. Halla, P. Sorjonen-Ward, O. Äikäs
Fennoscandian shield in east-central Finland and
Archean sanukitoid-type granitoids on the western
amine Paleoproterozic tonalitic dikes at the border
zone of the Proterozoic and Archean domains of the
Stop 4/1: Microtonalite dikes at Kivennapa
The bedrock exposures at Kivennapa provide an
insight into the complex setting of the microtonalite dikes (Figs. 51 and 52). Typical composite dike
lithologic units ranging from migmatitic and folded/boudinaged Paleoproterozoic mica schist to late
granite pegmatite with successive intrusive phases
of Juurus tonalite, microtonalites, granodiorite, and
granite. A large erratic boulder of microtonalite
shows porphyritic texture typical of numerous dikes
in the Juankoski map sheet area.
parts, intruding and brecciating the Juurus tonalite.
A later dike of the leucotonalitic material crosscuts
the darker microtonalite. Complex outcrop with
Fig. 51. Detailed map of a microtonalite outcrop in Kivennapa (Juankoski). Map: Olli Äikäs (2003).
81
Fig. 52. Typical composite dike of microtonalite intruding the Juurus tonalite at Kivennapa, Juankoski. Photo: Olli Äikäs.
Stop 4/2: The Pisa augen gneisses— Paleoproterozoic deformation of K-feldspar megacrystic
Neoarchean sanukitoids
General description
gray or white K-feldspar megacrysts. In places the
megacrysts are abundant and form aggregates. Particularly in the northern part of the area, typically
reddish megacrysts often form aggregates that may
change gradually to more homogeneous, coarsegrained granite dikes or lenses. Occasionally the
augen are completely absent, especially near the
Säyneinen schist belt, in which case the rock is a
banded gneiss.
Recent studies (Halla, 2002, 2005) have shown
that these gneisses are deformed Neoarchean
sanukitoid series granodiorites and monzodiorites.
The petrographic features and Pb isotopic composition of the megacrysts indicate that they are
porphyroclasts representing deformed Neoarchean
phenocrysts of magmatic origin. The geochemical
and isotopic features and the genesis of sanukitoids
are discussed in Section 3.2 in this volume.
Archean K-feldspar megacrystic gneisses are
found in the southern part of the Karelian Domain
in the Rautavaara Archean area of the Iisalmi terrain (Fig. 53). These gneisses have been traditionally termed as the Pisa augen gneisses due to their
appearance and because they are found mainly in
the area (~20 km by 10 km) around the Paleoproterozoic Pisa schist belt. On the northern and western
borders of the area the augen gneisses change more
or less gradually to basement gneisses. Towards
the eastern border of the area they change in appearance to banded hornblende- and mica-bearing
Säyneinen schist belt. The gneisses vary considerably in appearance but are typically schistose, dark
gray rocks containing diverse amounts of red, light
82
K-feldspar porphyroclasts
sion indicate lower recrystallization rates with
respect to strain rate compared with the eastern
part. Based on this observation, the Nilsiä augen
gneisses are roughly divided into two groups:
western group with lower recrystallization rates
and eastern group with higher recrystallization
rates with respect to strain rate. The western
group includes also reddish augen gneisses from
the northern part of the area with indications of
local granitization and alteration processes.
The variable appearance of the Nilsiä augen
gneisses is due to the different types of porphyroclasts developed in response to shear strain
(Fig. 54). In the northern and western parts (Stop
2), the larger augen commonly represent θ-type
porphyroclasts without real wings, or incipient overall δ-type porphyroclasts with relatively
short and narrow, curved wings suggesting slow
recrystallization rates with respect to strain rate
(Fig. 55A). Smaller porphyroclasts show more
developed σ-types of wings. Some of the large
megacrysts have retained their original rectangular shape.
In the eastern part of the Nilsiä area, the porphyroclasts represent overall δ-type and, more
commonly, overall σ- and φ-types with longer
and thicker wedge-shaped wings indicating faster recrystallization rates with respect to strain
rate compared with porphyroclasts in northern
Microstructures
Large porphyroclasts of the western group (Stop
4/2) of the Nilsiä augen gneisses are K-feldspar
megacrysts of Carlsbad-twinned perthitic microshowing sharp contacts between the core and the
mantle (Fig. 55C). Small myrmekitic intergrowths
of oligoclase and quartz along the margins of the
megacrysts are common. Relatively slow, smallscale grain boundary migration (GBM) recrystallization (for more detailed description of deformation mechanisms, see Passchier and Trouw, 1996)
in the temperature regime 1 of Hirth and Tullis
(1992) along the margins of the original megacrysts has produced an incipient core-and-mantle
gen-like appearance is outlined by granoblastic
quartz–feldspar aggregates of the recrystallized
mantle extending away from the core.
The types of K-feldspar porphyroclasts in the
northern and western part of the Nilsiä intru-
Fig. 53. Geological map showing the location of the Nilsiä and Lieksa sanukitoids. Inset: map of southern Finland showing the
Paleoproterozoic Svecofennian domain (gray), the Archean Karelian domain (white).
83
high 208Pb/204 Pb ratios with respect to 206Pb/204Pb ratios, developed in response to long term enrichment
of Th over U in the rocks, and (2) low 207 Pb/204Pb
and 206Pb/204Pb ratios due to the prolonged period of
time that has elapsed since the U-depletion. K-feldspars in the Nilsiä sanukitoids have high 208 Pb/204Pb
ratios with respect to the 206Pb/204 Pb ratios. The enrichment of 208Pb is thought not to be an initial feature of the K-feldspars; instead it seems to be a feature related to deformation. The 232Th–208Pb model
ages calculated for K-feldspar–whole rock pairs of
the Nilsiä eastern group range from 2.25 Ga to 1.84
Ga and include the oldest model ages in the Nilsiä
area. The youngest model age of 1.72 Ga is found
in the western part of the area where recrystallization rates are slower and Paleoproterozoic local
granitization and alteration is a common feature,
especially in the north. The 232 Th–208Pb model ages
for the Nilsiä western group range from 1.93 Ga to
1.72 Ga.
structure with a sharp boundary around the core
without transitional zones showing subgrain structures. Microcline is also abundant in the matrix,
where it occurs in recrystallized aggregates and as
recrystallized matrix consists of microcline, plagioclase, quartz, chloritized biotite, and minor hornblende and epidote. In the northern part, where local granitization is common, the matrix contains
also muscovite and the K-feldspar megacrysts are
reddish in colour due to alteration processes.
More deformed porphyroclasts, especially of the
eastern group, have thicker recrystallized mantles
indicating higher recrystallization rates with respect to strain rate. The cores of the porphyroclasts
are Carlsbad-twinned, perthitic microcline grains.
The margins of the original K-feldspar grains and
internal zones of deformation have commonly undergone extensive GBM recrystallization in the
temperature regime 1 of Hirth and Tullis (1992)
leading to the formation of thick, recrystallized
mantles around the original grains (Fig. 55D). The
recrystallized mantle extends away from the core
forming tails or wings in the matrix of microcline,
plagioclase, quartz, chloritized biotite, and minor
hornblende and epidote. In the most deformed
parts of the rock, altered ghosts of completely re-
Concluding remarks
The microstructural evidence for the Nilsiä
sanukitoids indicate that the K-feldspars have recrystallized by GBM (grain boundary migration) in
the low temperature regime of recrystallization-accommodated dislocation creep (Regime 1 of Hirth
and Tullis, 1992), indicating deformation temperatures of 400−500 °C. At 1.9 Ga, 208Pb-rich lead
evolved in high Th/U sites (probably grain boundaries and fractures), entered feldspar by GBM recrystallization of the original magmatic K-feldspar
grains during retrograde metamorphism related to
the Paleoproterozoic Svecofennian orogeny.
matrix.
Pb isotopes
Previous studies (Halla, 2005, 2002) have shown
that the Nilsiä and Lieksa granitoids have distinctive whole-rock Pb isotopic characteristics with (1)
θ) type porphyroclasts have round to
elliptic mantles but no real wings. Phi (φ) type porphyroclasts have straight centerlines and the wings are symmetrical with
respect to the porphyroclast. Sigma (σ) type porphyroclasts have wide mantles near the porphyroclast and parallel wings. The
wings have gently curved centerlines and they are asymmetric; the wing extends from the top of one side and from the bottom
of the opposite side (known as stair-stepping). Delta (δ) type porphyroclasts have narrow, strongly curved wings that are asymmetric with respect to the porphyroclasts. Complex porphyroclasts develop several sets of wings.
84
grained matrix. The wings of the porphyroclasts, which commonly represent overall θ- or incipient δ-type, are poorly developed indicating slow recrystallization rates with respect to strain rate. (B) The porphyroclasts in the eastern part of the area are
more developed winged porphyroclasts of overall δ-, σ-, or φ–types indicating faster recrystallization rates relative to strain
rate. (C) K-feldspar megacryst from the western part of the area with relatively thin recrystallized margin. Note the sharp
boundary around the core. (D) Porphyroclast showing a more developed mantle-and-core structure. Fine-grained recrystallized
wings extend away from the microcline core parallel to the matrix foliation. The matrix foliation wraps around the porphyrorecrystallization-accommodated dislocation creep (Regime 1 of Hirth and Tullis, 1992) indicating a deformation temperature of
400–500 °C (low- to medium-grade conditions).
85
Stop 4/3: Neoarchean sanukitoids of the Lieksa complex in the Ilomantsi terrain
(Fig. 56D) are commonly Carlsbad-twinned perthitic orthoclase, occasionally microcline, containing
euhedral inclusions of plagioclase and hornblende.
Smaller grains are antiperthitic subhedral plagiocla-
General description
The eastern part of the Lieksa complex of the
western Ilomantsi terrain (Fig. 53) includes undeformed porphyritic high-grade granodiorites (charno-enderbites) containing orthoclase megacrysts
and pyroxenes (Stop 4/3) and protomylonitic to mylonitic gneisses that have undergone deformation
and retrograde metamorphism. Porphyritic granodiorites are especially abundant around Lake Koitere
and are therefore also known as the Koitere granodiorites. Granulite-facies supracrustal enclaves and
very low U-content of both the granodiorites and
gneisses indicate the presence of an extensive highgrade terrain. The undeformed orthopyroxene-bearing granodiorites (charno-ederbites) and retrograde
gneisses exhibit similar geochemical characteristic,
which supports the assumption that undeformed
high-grade charno-enderbites represent the protolith of the retrograde gneisses. The geochemical
features of the Lieksa granitoids are similar to those
of the Nilsiä sanukitoid series granodiorites (Halla,
2005) indicating that both the Lieksa granodiorites
and gneisses belong to the sanukitoid-series granitoids. The geochemical and isotopic features and
the genesis of sanukitoids are discussed in Section
3.2.
green clinopyroxene and in places also orthopyroxene. Incipient recrystallization along grain boundaries of the megacrysts is common, but mantle-andcore structures are not developed yet.
Deformed granitoids of the protomylonitic stage
(Fig. 556E) contain Carlsbad-twinned perthitic microcline megacrysts and smaller, antiperthitic plagioclase megacrysts showing simple twinning and
containing abundant small epidote grains. In protomylonitic gneisses, the feldspar grains are separated
axes of recrystallized porphyroclasts. The matrix
consists of feldspars, quartz, dark green and often chloritized biotite, epidote, chlorite, and hornblende.
along original grain boundaries and along internal
deformation bands, progressively replacing original grains. Progressive strain resulted in grains
exhibiting a core-and-mantle structure. The mantle forms by small-scale grain boundary migration
(GBM) recrystallization of the K-feldspar rim in
the temperature regime 1 of Hirth and Tullis (1992),
possibly with incorporation of matrix material (by
grain boundary sliding and diffusion).
With increasing deformation and the resulting recrystallization-accommodated dislocation creep in
the temperature regime 1 of Hirth and Tullis (1992),
the original grains gradually decrease in size and
K-feldspar porphyroclasts
The dark-colored, undeformed charno-enderbites
of the Lieksa area (Fig. 56A) containing well-preserved, undeformed orthoclase megacrysts represent the protolith of the adjacent variably deformed,
lighter-colored gneisses. In the less deformed parts
of the shear zone between the high-grade blocks,
granitoids show abundant σ-type porphyroclasts
and S–C fabrics indicating a protomylonitic stage of
deformation (Fig. 56B). The more deformed parts
of the shear zone show strongly elongated recrystallized quartzo-feldspathic ribbons (Fig. 56C).
crystallized grains. In the more advanced stage of
deformation (Fig. 56F), the granitoids show strongly elongated recrystallized quartzo-feldspathic ribbons containing small cores of the original feldspar
grains. The size of the original grains has been further reduced by progressive recrystallization of the
porphyroclasts.
Recrystallization-accommodated
dislocation
creep produces strain softening of initially coarsegrained aggregates, and tends to partition strain
Microstructures
Three main stages of deformation can be distinguished in the progressive development of porphyritic granodiorites into strongly deformed gneisses:
(1) the initial stage, (2) the protomylonitic stage,
and (3) the advanced stage.
The megacrysts in the initial undeformed stage
86
Fig. 56. Hand specimens and photomicrographs of the Lieksa sanukitoids. (A) Undeformed high-grade charno-enderbite with
well-preserved orthoclase megacrysts. (B) Retrograde gneiss showing S–C fabrics and abundant overall σ-type porpyroclasts
tened quartz–feldspar ribbons indicating a more advanced stage of deformation. (D) Undeformed porphyritic granitoid. The
large grain is a Carlsbad-twinned perthitic orthoclase phenocryst containing euhedral inclusions of plagioclase and hornblende.
Smaller grains are antiperthitic plagioclase. (E) Deformed granitoid of the protomylonitic stage of deformation showing incipigrains are further reduced in size by progressive recrystallization of the porphyroclasts with increasing deformation and the
resulting recrystallization-accommodated dislocation creep (Regime 1 of Hirth and Tullis, 1992).
87
ratios due to the long period of U-depletion.
232
Th–208Pb model ages of 2.73 Ga and 2.71 Ga
calculated for the Lieksa K-feldspar megacrysts
of the initial undeformed stage (Stage 1) are consistent with the U–Pb zircon age of 2733 ± 29
Ma (Halla, 2002) obtained for the Lieksa granitoids. The protomylonitic gneisses (Stage 2) have
mixed model ages of 2.22–2.16 Ga, and the sample showing most advanced deformation (Stage 3)
has a model age of 1.85 Ga, which coincides with
the timing of the Paleoproterozoic Svecofennian
orogeny. The Th–Pb model ages for the Lieksa
granitoids seem to correlate with this stage of deformation.
grains enhances the access of water and progression of the replacement and softening reactions.
Original grains are gradually reduced in size by
progressive recrystallization. Strain softening may
continue until the aggregate is completely recrystallized and the rock appears as a banded gneiss.
All the variably deformed granitoids in the Nilsiä
and Lieksa areas show recrystallization in the temperature regime 1 of Hirth and Tullis (1992) indicating deformation temperatures of 400−500 °C.
The existence of undeformed granulitic protoliths between zones of intensively deformed
gneisses in Archean high-grade terrains may well
be explained by strain softening processes (e.g.,
grain-size reduction, rotation of grains, formation
Concluding remarks
ening). The balance between strain softening and
strain hardening, controlled by strain rate and temperature, determines whether a shear zone develops as a thin zone of very strained rocks or a wider
zone of less deformed rocks.
The microstructural evidence for the Lieksa
sanukitoids indicate that the K-feldspars have recrystallized by GBM (grain boundary migration)
in the low temperature regime of recrystallization-accommodated dislocation creep (Regime 1
of Hirth and Tullis, 1992) indicating deformation
temperatures of 400−500 °C (similar to the Nilsiä
augen gneisses). At 1.9 Ga, 208Pb-rich lead evolved
in high Th/U sites (probably grain boundaries and
fractures) in the rock entered the feldspar by GBM
recrystallization of the original magmatic K-feldspar grains during retrograde metamorphism related to the Paleoproterozoic Svecofennian orogeny.
Pb isotopes
The Lieksa granitoids have whole-rock Pb isotopic characteristics similar to those of the Nilsiä
augen gneisses with (1) high 208 Pb/204 Pb ratios
with respect to 206Pb/204Pb ratios, developed in response to long term enrichment of Th over U in
the rocks, and (2) low 207 Pb/204 Pb and 206 Pb/204Pb
88
DAY 5 (Friday September 16, 2005)
Guides: A. Käpyaho, L.S. Lauri, O.T. Rämö
Since the classic work of Kouvo (1958), several radiogenic isotopic methods have been applied
to solve the ages and origin of the crust in eastern
Finland. These include Rb–Sr (e.g., Martin and
Querré, 1984), Sm–Nd (e.g., Gruau et al., 1992),
Lu–Hf (Patchett et al., 1981), U–Pb (e.g., Hyppönen, 1983; Luukkonen, 1988, 2001), and K–Ar
(Kontinen et al., 1992) determinations. Some con-
ent plutonic phases determined on both methods
appears to be somewhat similar. Recently, the zircon ages and sources of some plutons have been
studied by using the secondary-ion mass spectrometry (SIMS) and Nd isotopes (Käpyaho et al.,
ages on plutonic rocks from Kuhmo district have
caused controversy and presently the U–Pb dat-
of the Neoarchean plutonism of the Kuhmo district from >2.8 Ga to 2.44 Ga will be outlined and
the similarities and contrasts between the modern
and Archean granitoid rocks will be discussed.
In addition, we will examine a 2.43-Ga A-type
granite that geochemically and petrographically
resembles the classic mid-Proterozoic rapakivi
granites of southern Finland.
sources have contributed to the plutonic rocks in
the Kuhmo district.
method for solving the actual crystallization and
source ages of the plutons, whereas the Rb–Sr
method in most cases records secondary effects
(Vaasjoki, 1988; Halliday et al., 1988). Nevertheless, the sequence of emplacement of the differ-
Stop 5/1: Arola leucogranite
Ion microprobe dating has revealed a heterogeneous zircon population and therefore only a rough
age estimate of 2.69 Ga is available (Käpyaho et
al., in review). The initial εNd value of this leucogranite is ~ –1. As demonstrated by Querré (1985),
this granite closely resembles the Himalayan leucogranites.
The Arola granite is pink, medium-grained and
equigranular leucogranite composed of quartz,
microcline, oligoclase, muscovite, and biotite. Accessories are zircon, rutile, magnetite, epidote, and
calcite. This granite is slightly peraluminous, relatively high in silica (up to 76 wt.%; Querré 1985),
and very depleted in the HREE (see Section 3.3).
Stop 5/2: Arola granodiorite
The Arola granodiorite is a microcline-porphyritic granodiorite that typically contains LILE-
elevated LILE contents as well. This intrusion has
been dated by using the conventional multigrain
U–Pb zircon method and the age of 2734 ± 3 Ma
is considered to represent the emplacement age of
this pluton (Hyppönen, 1983), whereas concordant
titanites register an age of ~2.69 Ga. The initial
εNd value of this pluton is around +1. On the basis
of geochemical data (Querré, 1985), Moyen et al.
(2001) suggested that this intrusion belongs to the
sanukitoid suite.
microcline grains (sometimes megacrysts) are often oriented and the groundmass that consists of
biotite and plagioclase shows a prominent schistosity. The granodiorite is metaluminous and the
Mg number [Mg/(Mg + Fe tot)×100] is generally
over 50. The silica range is typically between 66
wt.% and 69 wt.% and the rock shows generally
89
tive to the belt. In (B), three early Paleoproterozoic anorogenic A-type granite plutons (Tuliniemet, Kikonvaara, Iso-Kyllönen)
Stop 5/3: Kuusamonkylä tonalitic gneiss
Kuusamonkylä tonalitic gneiss is gray and equigranular. It comprises quartz, plagioclase, and biotite and accessory zircon, apatite, allanite, and epidote. On this outcrop, schistosity and felsic dikes
are folded and leucosome formation is in progress.
The Kuusamonkylä gneiss is slightly peraluminous
(A/CNK 1.0 to 1.1) and the silica ranges from 69
wt.% to 72 wt.% (Martin, 1987a). This tonalite is
depleted in HREE and, as concluded by Martin
(1985, 1987a), garnet in the residue is likely. The
pluton has been dated by ion microprobe (dating location 5 km north from the present outcrop) and has
a U–Pb age of 2.74 Ga (Käpyaho et al., in review),
age of 2.65 ± 0.3 Ga (Martin, 1985). The initial εNd
value of the gneiss is +1.
90
Stop 5/4: ~2.43-Ga A-type granite at Tuliniemet
The Tuliniemet granite belongs to a series of small
early Paleoproterozoic K-rich granite intrusions (Tuliniemet, Kyllönen, and Kikonvaara plutons) that cut
sharply the Archean metamorphic bedrock just east of
the Kuhmo–Suomussalmi greenstone belt (Fig. 57).
These plutons are associated with granite porphyry
dikes that cut both the granites and the surrounding
Archean rocks. U–Pb mineral data, however, indicate
similar crystallization ages for both, on the order of
2430–2420 Ma (Luukkonen, 1988; Irmeli Mänttäri,
unpublished data). This felsic magmatism is roughly
The Tuliniemet granite is a porphyritic, coarse-
Some of the alkali feldspar megacrysts (diameter
2–5 cm) are mantled by plagioclase (rapakivi texture). The marginal part of the Tuliniemet intrusion is characterized by an equigranular, mediumgrained biotite granite that is enriched in certain
trace elements (e.g., Rb, Th, U). The Nd isotope
systematics of these A-type plutons and the associated felsic dikes have been affected by a rather
strong Proterozoic (Svecofennian) overprint; initial (magmatic) εNd values of the coarse-grained
granites, however, probably range from –2.5 to
+0.7, averaging –1.3 ± 1.2 (1σ).
trusions farther to the north (Fig. 16) and may be related
to the rifting of the Archean craton (cf. Iljina and Hanski, 2005).
Stop 5/5: Kaihlankylä migmatite
and contains 72 wt.% silica; the Mg number is
49. Both leucosome and mesosome are depleted
in the HREE. Coarse-grained massive amphibole
batches contain 48 wt.% silica and their Mg number is 73. Unpublished ion-probe U–Pb data from
a nearby locality show an age of 2.94 Ga for the
mesosome.
Migmatite in Kaihlankylä quarry is polydeformed and metatexitic. Dark gray mesosome consists of biotite and amphibole and the thickness of
the layers varies from a few mm to several tens
of cm. The mesosome has 60 wt.% silica and an
A/CNK value of 0.85. The leucosome is leucogranodioritic or leucogranitic and peraluminous
91
92
DAY 6 (Saturday September 17, 2005)
Guides: P. Sorjonen-Ward, O.T. Rämö
back to Helsinki from the Kuhmo area, stopping on
the way at the Kuopio airport for participants with
described (Kontinen, 1987) and remains to serve as
a genuine example of the outcome of Precambrian
plate tectonic processes. We will also have ample
time for discussion on the bus as we drive south.
The estimated time of arrival in Helsinki is 6 PM on
Saturday evening.
stops will be in the program. However, if desired
and time permits, we may stop by the ~1.95-Ga Jormua mantle section peridotites at Kontiomäki. Jor-
93
94
PART III: ACKNOWLEDGMENTS AND REFERENCES
95
96
ACKNOWLEDGMENTS
We are thankful to the Coordinator of Eurogranites, Professor Bernard Bonin, for the opportunity of aremy of Finland, the Geological Survey of Finland, the University of Helsinki, Palin Granit Oy, and IGCP
Project 510 (A-type Granites and Related Rocks through Time)— this is gratefully acknowledged. We are
obliged to Drs. Hannu Huhma and Irmeli Mänttäri of the Geological Survey of Finland for sharing unpublished Nd and U–Pb isotope data with us, to Professor Martti Lehtinen of the Geological Museum, University of Helsinki for XRD determinations, and to Dr. Pentti Hölttä (Geological Survey of Finland), Dr.
Annakaisa Korja (Institute of Seismology, University of Helsinki), Ms. Elina Arponen,, and Mr. Hannu
Lauri for comments and discussions. We thank Dr. Tapio Ruotoistenmäki of the Geological Survey of
A. Elliott (University of North Alabama) for volunteering to act as a guide on Day 3. Mr. Jukka Lehtinen
assisted with photography, which is gratefully acknowledged. The late Matti Vaasjoki1 participated in the
team. This guide is a contribution to IGCP Project 510.
1
As a Geological Survey of Finland geochronologist, Matti Vaasjoki’
s contribution was instrumental in acquiring and interpret-
97
98
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109
110
PART IV: APPENDIX
111
112
Field trip stop coordinates
Stop #
Finnish
National Grid
Northing
Finnish
National Grid
Easting
Latitude
[oN]
Longitude
[oE]
1/1
6717.500
3493.600
60.56804o
26.88019o
1/2
6717.569
3507.226
60.56865o
27.12867o
1/3
6715.250
3538.000
60.54612o
27.68935o
1/4
6740.800
3546.250
60.77454o
27.84566o
2/1
6854.152
3561.457
61.78950o
28.16166o
2/2
6867.137
4472.389
61.90999o
29.47154o
2/3
6874.914
4485.405
61.98049o
29.71857o
2/4
6787.269
4442.539
61.19006o
28.92861o
2/5
6817.708
4434.642
61.46192o
28.77117o
3/1
6850.011
3492.114
61.75717o
26.84747o
3/2
6864.565
2558.019
61.88338o
25.09969o
3/3
6863.697
2564.262
61.87460o
25.21805o
3/4
6869.450
2571.508
61.92492o
25.35801o
3/5
6869.139
2570.780
61.92226o
25.34402o
4/1
6996.705
3550.566
63.06995o
27.99693o
4/2
7017.319
4413.580
63.24803o
28.27659o
4/3
7019.072
4519.735
63.27366o
30.38993o
5/1
7151.051
4450.780
64.45454o
28.97409o
5/2
7150.469
4452.564
64.44958o
29.01132o
5/3
7147.564
4464.578
64.42499o
29.26155o
5/4
7164.892
4468.259
64.58077o
29.33417o
5/5
7143.947
4475.220
64.39349o
29.48300o
113
114
PART V: NOTES
115
116
Day 1
117
118
Day 2
119
120
Day 3
121
122
Day 4
123
124
Day 5
125
126
Day 6
127
128