Iron oxide-copper-gold deposits in the Kolari region, Finland (PDF)

Ore Geology Reviews 30 (2007) 75–105 

www.elsevier.com/locate/oregeorev 

Geology, geochemistry, fluid inclusion characteristics, and U–Pb age 

studies on iron oxide–Cu–Au deposits in the Kolari region, 

northern Finland 

Tero Niiranen a,b, *, Matti Poutiainen a , Irmeli Mänttäri c 

a Department of Geology, P.O. Box 64, FIN-00014 University of Helsinki, Finland 

b Northland Resources Inc., Teknotie 14-16, 11, FIN-96930 Rovaniemi, Finland 

c Geological Survey of Finland, P.O. Box 96, FIN-02151 Espoo, Finland 

Received 21 February 2005; accepted 23 November 2005 

Available online 19 January 2006 

Abstract 

Several iron oxide–copper–gold deposits are known in the Kolari region, in the western part of the Central Lapland 

greenstone belt, northern Finland. They are hosted by clinopyroxene-dominated skarns that were formed near to contact zone 

between ca. 1860 Ma Haparanda Suite intrusions and N2050 Ma Savukoski Group supracrustal rocks. All deposits are located 

within or next to shear and fault zones forming parts of the major, NNE-trending, Kolari shear zone. Three of the Fe–Cu–Au 

deposits, Kuervitikko, Cu-Rautuvaara, and Laurinoja were studied; all contain significant amounts of Cu (0.1% to 4.5%) and Au 

(0.1 to 6.6 ppm). At Laurinoja and Kuervitikko, Cu and Au are hosted by ironstone and skarn. At Cu-Rautuvaara, the host rock 

is a magnetite-disseminated albitite. The deposits have a distinct metal association of Fe–Cu–AuFAg, Bi, Ba, Co, Mo, Sb, Se, 

Te, Th, U, LREE. The wall and host rocks are intensely altered and display a deposit-scale zonal pattern. The distal alteration 

zone is characterised by albiteFbiotite, K-feldspar, and scapolite and the proximal zone by clinopyroxene–magnetiteFamphibole, 

scapolite, calcite, and sulphides. Mass balance calculations indicate that Al 2 O 3 ,TiO 2 , and Zr were immobile 

during alteration. The calculations also indicate that significant quantities of Fe 2 O 3 , CaO, CO 2 , S, Cu, Au, Bi, and Te were 

added to the proximal altered rocks. The main gains in the distal altered rocks are in Na 2 O, K 2 O, and Ba. Fluid inclusion data 

suggest that the fluids which circulated in the rocks during the main mineralisation event and subsequent brittle fracturing were 

highly saline (V56 wt.% NaCl) H 2 OFCO 2 fluids. The temperature during the main mineralisation event was between 450 and 

550 8C and the pressure was 1.5 to 3.5 kbar. 

Based on U–Pb age data of magmatic zircon from altered hanging wall diorite and granite that brecciates the ore, the age of the 

deposits is between 1864F5 and 1766F5 Ma. The 1797F5 Ma age of zircon in skarn, combined with the 1810 to 1780 Ma ages 

of the metamorphic titanite in altered wall rocks and skarn, suggest that the deposits were most likely formed ca. 1800 Ma. This age 

has been interpreted to be broadly contemporaneous with the D 3 thrusting event in the Kolari region during which the Kolari shear 

zone was activated. 

The data presented in this work are inconsistent with previously proposed models for the Kolari deposits that they are 

metamorphosed syngenetic iron formations, or skarns related to ca. 1860 Ma monzonite intrusions. Instead, we propose that the 

Kolari Fe–Cu–Au deposits are of metasomatic replacement-type, and are controlled by the Kolari shear zone structures related to 

* Corresponding author. Northland Resources Inc., Teknotie 14-16, 11, FIN-96930 Rovaniemi, Finland. Fax: +358 16 362 620. 

E-mail address: tero.n@nrmine.com (T. Niiranen). 

0169-1368/$ - see front matter D 2005 Elsevier B.V. All rights reserved. 

doi:10.1016/j.oregeorev.2005.11.002

76 

T. Niiranen et al. / Ore Geology Reviews 30 (2007) 75–105 

post-peak metamorphic D 3 thrusting event in northern Finland. Our data suggest that the Kolari deposits best fit into the category of 

epigenetic iron oxide–copper–gold mineralisation. 

D 2005 Elsevier B.V. All rights reserved. 

Keywords: IOCG deposits; Hydrothermal alteration; Proterozoic; Fluid inclusions; Kolari; Finland 

1. Introduction 

Several Fe FCu, Au deposits are known in the 

Kolari region in the western part of the Proterozoic 

Central Lapland greenstone belt (CLGB), Finland 

(Fig. 1). Deposits at Hannukainen and Rautuvaara 

were mined for iron between 1974 and 1990 with a 

total of 16 Mt iron ore produced (Puustinen, 2003). The 

total inferred amount of iron ore in the known deposits 

is several tens of Mt (Hiltunen, 1982). Besides iron, 

four deposits in the region contain significant amounts 

of copper and gold. However, up until now, Au and Cu 

have only been extracted from the Laurinoja ore body 

of the Hannukainen mine (Fig. 1; Hiltunen, 1982; 

Puustinen, 2003). 

Almost all of the known Fe–Cu–Au deposits are 

located near to contact zones between ca. 1860 Ma 

monzonite–diorite intrusions and N2050 Ma Savukoski 

Fig. 1. General geological features of northern Finland (a), location and general geological features of the Kolari region (b), and location of 

FeFCu–Au deposits (c). Location of the Kolari shear zone (KSZ) (c), modified after Korsman et al. (1997), lineation and fold axis data after 

Hiltunen (1982) and Väisänen (2002); stratigraphy groups after Lehtonen et al. (1998). The coordinate system is the Finnish national system (YKJ). 

CLGB=Central Lapland greenstone belt.

T. Niiranen et al. / Ore Geology Reviews 30 (2007) 75–105 77 

Group supracrustal rocks (Fig. 1; Hiltunen, 1982; Lehtonen 

et al., 1998). Nearly all of the deposits are located 

in, or next, to thrust or other fault zones that form the 

complex N- to NNE-trending Kolari Shear System (e.g., 

Väisänen, 2002). Hiltunen (1982) suggests that the 

Kolari deposits are stratabound skarns that were formed 

by contact metasomatic events related to intrusion of 

monzonites. Mäkelä and Tammenmaa (1978) and 

Frietsch et al. (1995) proposed, based on the sulphur 

isotope composition of sulphides, that the Kolari deposits 

are metamorphosed syngenetic iron formations. However, 

the deposits display characteristics that are 

inconsistent with both these models. These include a 

ca. 65 Ma time difference between the U–Pb zircon 

ages of the monzonite and skarn that hosts the ore 

(Hiltunen, 1982) and the lack of any textural evidence 

that would imply a sedimentary origin for the iron. 

Instead, many features of the Kolari deposits are typical 

of iron oxide–copper–gold (IOCG) deposits that are now 

widely recognised as a global class of ore deposits (e.g., 

Hitzman et al., 1992; Barton and Johnson, 1996; Hitzman, 

2000). 

The purpose of this work is to present U–Pb, fluid 

inclusion, and new geochemical data on the Kolari 

deposits in order to establish whether the Kolari deposits 

are skarns sensu stricto, metamorphosed syngenetic 

iron formations, or if the IOCG model better explains 

their characteristics. 

2. General geology 

The bedrock of northern Finland (Fig. 1) consists of 

an Archean basement, Palaeoproterozoic greenstone 

and schist belts, ca. 1900 Ma granulite belt, and Svecofennian 

1930 to 1770 Ma granitoids (e.g., Hanski et 

al., 1997). The CLGB was formed during prolonged 

stages of rifting of the Archean craton, with sedimentation, 

and magmatism in intracratonic and cratonic 

margin rift settings between 2500 and 1900 Ma (e.g., 

Hanski et al., 1997; Lehtonen et al., 1998), and was 

subjected to multiphase deformation and metamorphism 

during orogenic events between 1920 and 1770 

Ma (Ward et al., 1989; Lehtonen et al., 1998; Väisänen, 

2002; Lahtinen et al., 2003; Sorjonen-Ward et al., 2003; 

Hölttä et al., in press). 

Supracrustal rocks of the CLGB consist of terrestrial 

to deep marine sediments with thick intercalations of 

ultramafic to mafic volcanic rocks deposited on the 

metamorphosed Archean basement. The supracrustal 

sequence has been divided into seven lithostratigraphic 

units which are, from oldest to youngest: Salla, 

Onkamo, Sodankylä, Savukoski, Kittilä, Kumpu and 

Lainio Groups (Fig. 2; Lehtonen et al., 1998). Intrusives 

in the CLGB consist of ca. 2440 Ma layered 

intrusions, ca. 2220 Ma mafic sills and dykes, 2050 

Ma layered intrusions and diabases, 1930 to 1860 Ma 

dominantly intermediate to felsic intrusions, and ca. 

1820 to 1770 Ma felsic intrusions (Lehtonen et al., 

1998; Mutanen and Huhma, 2001; Hanski et al., 

2001; Rastas et al., 2001). 

In the Kolari region, the supracrustal rocks consist of 

Onkamo Group tholeiites and komatiites, Sodankylä 

Group quartzites, mica schists and gneisses, and conglomerates, 

Savukoski Group Fe–tholeiites, tuffites, 

dolomitic marbles and black schists, and Lainio 

Group quartzites, siltstones, conglomerates and minor 

volcanic rocks (Fig. 1; Hiltunen, 1982; Lehtonen et al., 

1998). The ca. 1930 to 1860 Ma diorites, tonalites, and 

monzonites constitute the main intrusives in the region 

(Hiltunen, 1982; Väänänen and Lehtonen, 2001). In 

addition to these, variably albitised 2200 to 2000 Ma 

diabases as well as ca. 1820 to 1770 Ma granitic rocks 

occur in the region (Hiltunen, 1982; Väänänen and 

Lehtonen, 2001; this work). 

The structural evolution of the CLGB during the 

Svecofennian orogeny is currently under revision, leading 

to somewhat confusing nomenclature between the 

older and the more recent reports. Lehtonen et al. 

(1998) divide the main deformation stages in the 

CLGB into five different events, whereas Väisänen 

(2002) and Hölttä et al. (in press) distinguish three 

ductile deformation stages and subsequent brittle stages 

which all took place within the period 1920 to 1770 

Ma. The deformation stages discussed in this work are 

after Väisänen (2002) and Hölttä et al. (in press), except 

that the completely brittle stages following the D 3 are 

referred here as D 4 . 

The D 1 structures comprise bedding parallel S 1 foliation 

that is only microscopically observable (Väisänen, 

2002; Hölttä et al., in press). Lehtonen et al. (1998) 

suggest that gentle folding was associated with D 1 . 

They also proposed that a static stage occurred between 

D 1 and D 2 during which the metamorphic conditions 

already reached greenschist to amphibolite facies conditions. 

The most prominent structural features in the 

CLGB were developed during the D 2 stage between 

1890 and 1860 Ma (Hiltunen, 1982; Sorjonen-Ward et 

al., 1992; Lehtonen et al., 1998; Väisänen, 2002; Hölttä 

et al., in press). 

The D 2 structures are characterised by tight to isoclinal 

F 2 folds and S 2 foliation. Generally the S 2 foliation 

is subparallel to bedding, gently dipping to 

horizontal and the F 2 folds are recumbed to reclined, 

although locally the foliation and axial planes are steep-

78 


Fig. 2. Stratigraphy of the Central Lapland greenstone belt together with age data for different igneous rocks in the area. After Lehtonen et al. (1998) 

and Hanski et al. (2001) and references therein. 

ly dipping or vertical. Kinematic indicators suggest that 

these structures were caused by horizontal movement 

related to thrust tectonics and that the movement direction 

was probably from SSW to NNE (Väisänen, 2002; 

Hölttä et al., in press). 

The thrusts, which were initiated in earlier stages, 

were re-activated during the D 3 stage when the D 1–2 

structures were overprinted by F 3 folds, S 3 foliation, 

and L 3 lineation with an extreme variation in direction, 

dip, plunge, and intensity (Lehtonen et al., 1998; 

Väisänen, 2002; Hölttä et al., in press). The variation 

in the direction of the D 3 structures has been 

explained by either simultaneous thrusting from opposing 

direction or with progressive thrusting that was 

accompanied with rotation (Sorjonen-Ward et al., 

1992; Lehtonen et al., 1998; Väisänen, 2002; Hölttä 

et al., in press). The absolute age of the D 3 probably 

varies among the different parts of the CLGB but it is 

between 1.86 and 1770 Ma (Väisänen, 2002; Hölttä et 

al., in press). 

The peak metamorphic conditions in the CLGB 

were probably reached during the D 2 stage (Lehtonen 

et al., 1998; Hölttä et al., in press). The metamorphic 

grade varies from greenschist to granulite 

facies conditions in the metamorphic zones that are 

separated by the D 3 thrust planes (Hölttä et al., in 

press). 

Structural features of the Kolari region are dominated 

by a number of thrust, shear, and fault zones that 

form the Kolari Shear System which is part of the N–S 

oriented, up to 50 to 100 km wide, Baltic-Bothnian 

Mega-shear (Fig. 1; Berthelsen and Marker, 1986; Väisänen, 

2002). The thrust and shear zones lie parallel or 

subparallel to the F 2 fold axis of the syncline–anticline 

structures developed during the NNE directed D 2 

thrusting (Hiltunen, 1982; Väisänen, 2002; Hölttä et 

al., in press). Thrusting during the D 3 stage led to 

reactivation of the D 2 thrust and fault zones, refolding 

of the D 2 structures, and development of strong L 3 

lineation in the vicinity of the thrust zones (Hiltunen, 

1982; Väisänen, 2002). The kinematic indicators suggest 

that, in the Kolari region, the D 3 structures were 

caused by W to E directed thrusting (Hiltunen, 1982; 

Väisänen, 2002; Hölttä et al., in press). As a result, the 

areas around the thrust zones are characterised by welldeveloped, 

gently to moderately W to SW plunging L 3 

lineation, and F 3 folds with axes plunging parallel to the 

lineation (Fig. 1).


3. Analytical procedures and mineral abbreviations 

Geochemical analyses were carried out at the laboratories 

of the Geological Survey of Finland (GSF). 

Whole-rock analyses were made using standard wavelength 

dispersive X-ray fluorescence spectrometry 

(XRF) from pressed powder pellets, and rare earth 

elements (REE) were analysed using inductively coupled 

plasma mass spectrometry (ICP-MS) after HF– 

HClO 4 digestion and lithium metaborate–sodium perborate 

fusion. Ag and Co were analysed with inductively 

coupled plasma atomic emission spectrometry 

(ICP-AES) after aqua regia digestion. Au, Bi, Sb, Se, 

and Te were analysed after aqua regia digestion using 

atomic absorption spectrometry and electrothermal atomisation 

in a graphite furnace (GFAAS). 

Fig. 3. Geology of the Hannukainen deposit. Note that the Kivivuopio ore body is not included into the surface geology map because its geometry 

and size is poorly known. After Hiltunen (1982).

80 


Fluid-inclusion microthermometry was carried out 

using a Linkam THMSG 600 heating-freezing stage at 

the University of Helsinki. Accuracy levels of F0.1 and 

F2.0 8C were achieved for subzero and higher temperatures, 

respectively, by calibration using synthetic fluidinclusion 

standards (SynFlinc). FLINCOR (Brown, 

1989) and Linkam PVTX software were used to analyse 

laboratory microthermometric data. 

For conventional U–Pb age determinations, the decomposition 

of zircons and titanites and extraction of U 

and Pb follows mainly the procedure described by 

Krogh (1973, 1982). 235 U– 208 Pb-spiked and unspiked 

isotopic ratios were measured using a VG Sector 54 

thermal ionization multicollector mass spectrometer. 

The measured Pb and U isotopic ratios were normalised 

to the accepted ratios of SRM 981 and U500 standards. 

The Pb/U ratios were calculated using the PbDat-program 

(Ludwig, 1991) and fitting of the discordia lines as 

well as calculation of the ages were completed using the 

Isoplot/Ex 3 program (Ludwig, 2003). 

Mineral abbreviations used in figures, diagrams and 

tables in this work are: ab =albite, act=actinolite, 

am=amphibole, atp=anthophyllite, Au=native gold, 

bt=biotite, cc=calcite, cpx=clinopyroxene, cpy=chalcopyrite, 

ep=epidote, gr=garnet, hbl=hornblende, 

kfs=K-feldspar, mo=molybdenite, mgt =magnetite, 

po=pyrrhotite, py=pyrite, sca=scapolite, te=tellurides, 

tit=titanite, ura=uraninite. 

Fig. 4. Vertical sections across the Kuervitikko (a) and Cu-Rautuvaara (b) deposits. After Hiltunen (1982).


4. Deposit description 

Three Cu- and Au-rich iron oxide deposits were 

selected for this study: Cu-Rautuvaara, Kuervitikko, 

and Laurinoja (Figs. 1, 3, and 4); Table 1). Of 

these, Laurinoja is one of the five distinct ore bodies 

of the Hannukainen mine. All of the selected deposits 

are located in or next to the same thrust zone (Fig. 1). 

The host rocks and alteration sequence of the Laurinoja 

and Kuervitikko are very similar with each 

other (Figs. 2 and 3). In the Cu-Rautuvaara deposit, 

located ca. 7 km SW of the Hannukainen mine, the 

lithology is relatively similar to the Laurinoja and 

Kuervitikko deposits, but instead of ironstone and 

skarn, the Cu–Au mineralisation is hosted by magnetite 

disseminated albitite (Fig. 4). Nevertheless, the 

structural setting and general features of the Cu-Rautuvaara 

are so similar to the Laurinoja and Kuervitikko 

that it can be considered to have been formed by 

the same kind of processes that led to the formation of 

the latter two deposits. 

4.1. Laurinoja and Kuervitikko 

The Laurinoja and Kuervitikko deposits consist of 

massive to silicate-banded magnetite-rich lenses (ironstones) 

(Fig. 5), that are hosted by skarns formed near to 

the contact zone between the Savukoski Group supracrustal 

sequence and ca. 1860 Ma monzonite–diorite 

intrusions (Hiltunen, 1982). The footwall of these 

deposits consists of variably altered mafic metavolcanic 

rocks (Fig. 6d–g), quartzite quartz–feldspar schist and 

mica gneiss. Monzonite (Fig. 6a) and variably altered 

diorite (Fig. 6b–c) comprise the hanging wall (Figs. 3 

and 4). The skarns consist mainly of clinopyroxene 

(diopside–hedenbergite) and amphibole (actinolite– 

hornblende) dominated rocks. Also, clinopyroxenebearing 

scapolite (marialite) skarns (Fig. 6i) are abundant, 

and thin (b0.5 m) garnet-rich (andradite) horizons 

occur locally in the sequence. In addition, rocks consisting 

mainly of albite with varying amounts of biotite, 

amphibole, K-feldspar and quartz (Fig. 6h), and referred 

to as albitites in this work, are common in the sequence. 

A small lens of carbonate rock occurs within the clinopyroxene–amphibole 

skarn at Kuervitikko (Fig. 4), 

whereas no carbonate rocks have been found at Laurinoja. 

Pegmatitic granite dykes, b1 to several m in 

thickness, are abundant in both areas and, in places, 

ca. 1780 Ma medium-grained granite dykes crosscut 

and brecciate the ore and wall rocks (Fig. 5c–d). 

The main oxide mineral in both deposits is magnetite; 

pyrite, pyrrhotite, and chalcopyrite are the dominant 

sulphides (Fig. 7). Locally, minor amounts of 

molybdenite occur in association with chalcopyrite 

(Fig. 7g). Native gold has been detected in silicate 

gangue, chalcopyrite, and magnetite in both the Laurinoja 

and Kuervitikko occurrences (Fig. 7i–l). Clinopyroxene, 

amphibole, albite, scapolite, biotite, calcite, 

and quartz comprise the gangue in the ironstones. 

Sulphides mainly occur as disseminations, but locally 

also as veins in the ironstones, skarns and altered 

wall rocks. In places, pyrrhotite and chalcopyrite form 

narrow (b30 cm wide), massive lenses or veins in the 

ironstone, and locally chalcopyrite and pyrite, together 

with variable amounts of quartz, scapolite, K-feldspar 

and calcite occur as fracture infill in the ironstone (Fig. 

Table 1 

Grades, tonnages and general characteristics of the Laurinoja, Kuervitikko, and Cu-Rautuvaara deposits 

Deposit/prospect Laurinoja Kuervitikko Cu-Rautavaara 

Size and grade 

4.56 Mt* at 43% Fe, 

1 ppm Au, 0.88% Cu; 

33 Mt** at 36–53% Fe, 

b0.1–11.0% Cu, 

b0.1–6.6 ppm Au, 

b0.1–17.7 ppm Ag, 

20–1000 ppm Co 

Monzonite, diorite, 

mafic metavolcanic rock 

1.2 Mt** at 36–53% Fe, 

b0.1–8.3% Cu, b0.1–6.0 ppm Au, 

b0.1–2.1 ppm Ag, 50–830 ppm Co 

N4 Mt** weakly to moderately 

Cu–Au mineralised rock, 

b0.1–1.5% Cu, b0.1–2.6 ppm Au, 

b0.5–1.2 ppm Ag 

Wall rocks 

Monzonite, diorite, 

Monzonite, diorite, mafic 

mafic metavolcanic rock, albitite metavolcanic rock 

Main host rock(s) Ironstone, cpx–afb skarn Ironstone, cpx–afb skarn Albitite 

Opaques mgt, py, po, cpyFmo, Au, te mgt, py, cpy, poFAu, te mgt, po, cpyFpy, ura 

Gangue 

cpx, hbl–act, ab, bt, Fsca, cpx, hbl–act, ab, Fsca, bt, 

ab, atf, bt, Fcpx, qtz, tit 

qtz, cc, kfs, gr, ep 

cc, kfs, qtz, gr, ep 

Metal association Fe, Cu, Au, SFAg, Bi, Ba, Fe, Cu, Au, SFAg, Ba, Bi, 

Fe, Cu, Au, SFAg, Ba, Bi, 

Co, Mo, Sb, Te, LREE 

Co, Mo, Se, Te, LREE 

Mo, Se, Te, Th, U, LREE 

Data from Kinnunen (1980), Hiltunen (1982), Keinänen (1995), Puustinen (2003) and this work. See Analytical procedures and mineral 

abbreviations for mineral abbreviations. *Production 1978–1990. **Inferred.

82 


Fig. 5. (a) Typical barren massive ironstone. Note the S 3 foliation in the rock (dotted white line). The S 3 is crosscut by chalcopyrite and scapolite– 

quartz–chalcopyrite–pyrite filled fractures along the direction of the white dashed line. (b) Close up photograph of the fractures from the area 

marked with white rectangle in (a). (c) Massive barren ironstone that is brecciated by ca. 1780 Ma granite. (d) Cu–Au-mineralised ironstone that is 

intruded by ca. 1780 Ma granite. The coin in the figures is 26 mm in diameter. All photos from Laurinoja. 

5a–b). Magnetite in the sulphide-bearing ironstones is 

typically coarser and more euhedral in form than in the 

barren ironstones (Fig. 7a–c). 

4.2. Cu-Rautuvaara 

In the Cu-Rautuvaara deposit, the bulk of the Cu–Au 

ore is hosted by magnetite-disseminated, biotite- and/or 

anthophyllite-bearing albitite which forms a ca. 45 m 

thick, elongated lens (Fig. 4). Ironstone lenses and clinopyroxene–amphibole 

skarns similar to those at Laurinoja 

and Kuervitikko are also present, but they are 

significantly smaller in size. As in Kuervitikko and 

Laurinoja, the hanging wall consists of diorite and monzonite, 

and altered mafic metavolcanic rock comprises 

the immediate footwall. Chalcopyrite and pyrrhotite are 

the dominant sulphides, although pyrite is locally abundant. 

Magnetite is the dominant oxide mineral, and small 

clusters of uraninite have been detected in the clinopyroxene–amphibole 

skarn (Fig. 7h). Albite, anthophyllite, 

biotite, and quartz are the main gangue minerals. 

4.3. Main structural features of the deposits 

Both Laurinoja and Kuervitikko are located in the 

easternmost thrust zone of the Kolari Shear System 

next to the Äkäsjoki fault (Fig. 1). The strong foliation 

and lineation developed into rocks in the vicinity 

of the thrust zone during the D 3 stage which 

overprints the structural features of the D 1 and D 2 

stages so that, for example at Hannukainen and 

Kuervitikko, the older structural features cannot be 

recognised. 

The S 3 foliation and the L 3 lineation are best developed 

in altered wall rocks and silicate-banded ironstones, 

and locally can be distinguished in barren 

massive ironstones (Figs. 5 and 6). However, the S 3 

foliation is less developed in sulphide-bearing ironstones 

which locally appear unfoliated. The S 3 foliation 

at Laurinoja and Kuervitikko dips at ca. 308 to the west 

and the L 3 lineation plunges gently to the southwest. In 

both areas, the elongated lens-shaped ironstone units 

plunge gently to the SW following the direction of the 

L 3 lineation (Figs. 3 and 4). The youngest structural 

features detected at Laurinoja and Kuervitikko are thin 

mylonite seams and brittle fractures that crosscut the S 3 

foliation at a steep angle (Fig. 5a, b). 

The Cu-Rautuvaara deposit is structurally located 

in the NW limb of a major SW opening syncline next 

to the same thrust zone in which Laurinoja and 

Kuervitikko are located (Figs. 1 and 4; Hiltunen, 

1982). Although the S 3 foliation and the L 3 lineation


Fig. 6. Photomicrographs of the wall rocks from Laurinoja, Kuervitikko and Cu-Rautuvaara. (a) Typical hanging wall monzonite from DDH79- 

76.70. (b) Typical distal altered diorite from DDH75-33.60. Hornblende with minor biotite comprises the dark minerals and the pale areas consist of 

albite with minor scapolite. (c) Albite–biotite–K-feldspar-altered diorite from DDH75-43.75. Biotite is the dominant mafic mineral, and the pale 

areas consist of albite and K-feldspar. In addition, there is a fine-grained magnetite dissemination in the rock (black spots). (d) Least-altered footwall 

metavolcanic rock from DDH79-217.69. Main minerals are hornblende (dark) and plagioclase (pale). (e) Intensely biotite-altered metavolcanic rock 

from DDH79-205.90. Dark minerals are biotite with small amounts of hornblende and actinolite. Pale areas consist of albite with traces of quartz. (f) 

Biotite–K-feldspar-altered metavolcanic rock from DDH79-190.85. Dark minerals consist of biotite with traces of actinolite. Moderate magnetite 

dissemination (black spots). K-feldspar, albite and quartz comprise the pale minerals. (g) K-feldspar–biotite–clinopyroxene-altered footwall 

metavolcanic rock from DDH79-177.70. Pale minerals consist of K-feldspar porphyroblasts and minor amounts of quartz and albite. Biotite, 

actinolite and minor amounts of clinopyroxene and magnetite comprise the dark minerals. (h) Magnetite-disseminated albitite from DDH109- 

217.25. (i) Scapolite skarn with disseminated magnetite, chalcopyrite, and pyrrhotite (dark) from DDH109-236.00. Younger biotite–pyrite veins 

crosscut the skarn. 

are well developed in the wall and host rocks of the 

Cu-Rautuvaara deposit, they are less pronounced than 

those at Laurinoja and Kuervitikko. The dominant 

direction of the main structural trend is SE and the 

dip is steep (Fig. 4). As at Laurinoja and Kuervitikko, 

thin mylonite seams and brittle fractures crosscut the 

main structures and are the youngest deformational 

feature.

84 


Fig. 7. Reflected light photomicrographs of the ores and skarns from Kolari. (a) Typical fine-grained magnetite (pale grey) ore with clinopyroxene 

(dark grey) gangue, Laurinoja. (b, c) Typical chalcopyrite-bearing magnetite ore from Laurinoja. Note the significantly coarser grain size of the 

magnetite compared to the magnetite without sulphides in (a), and the well-developed crystal form of the magnetite grains in (c). (d) Cu–Aumineralised 

albitite from Cu-Rautuvaara. (e) Pyrite, magnetite, and chalcopyrite in clinopyroxene skarn, Kuervitikko. (f) Chalcopyrite and 

magnetite in clinopyroxene skarn from Laurinoja. Thin chalcopyrite veins or fracture infills crosscut the clinopyroxene and magnetite. (g) 

Molybdenite flakes in association with chalcopyrite and pyrite in clinopyroxene skarn at Laurinoja. (h) Uraninite and small pyrite grains in 

clinopyroxene, Cu-Rautuvaara. (i) A small gold grain in chalcopyrite next to magnetite, Laurinoja. (j) Native gold in amphibole, Kuervitikko. (k) 

Native gold in coarse-grained magnetite, Laurinoja; oil immersion. (l) Gold grains in a fracture in clinopyroxene, Laurinoja; oil immersion. 

4.4. Size, grade, and distribution of copper and gold 

The total inferred amount of iron ore in the Kuervaara, 

Laurinoja, Vuopio, and Lauku ore bodies of the 

Hannukainen deposit is ca. 68 Mt (Hiltunen, 1982). 

According to Hiltunen (1982), the Laurinoja ore body 

contained ca. 33 Mt of iron ore with Fe content varying 

between 36 and 53 wt.% and is the largest ore body at 

Hannukainen (Fig. 1; Table 1). The Cu and Au contents 

at Laurinoja are poorly constrained, but Hiltunen (1982) 

estimates that the average concentration of the whole 

ironstone lens is 0.36 wt.% and 0.15 ppm for Cu and Au, 

respectively. Copper and gold are, however, not evenly 

distributed throughout the ironstone. The richest parts of 

the ore are located in the northeastern part of the ironstone 

lens whereas the grades are lower in the margins 

and southwestern part (Hiltunen, 1982). For the 4.6 Mt 

of ore mined from Laurinoja, the grades were 43 wt.% 

Fe, 0.88 wt.% Cu, and 1 ppm Au (Puustinen, 2003). 

The distribution of Cu and Au in drill hole 170 

drilled across the Cu–Au rich part of the Laurinoja 

ore body shows that Cu–Au mineralisation chiefly


occurs within the ironstone, and that there is a positive 

correlation between Cu and Au (Fig. 8). Very little 

analytical data is available for rocks outside the ironstones, 

but the new data presented in this work show 

that elevated Cu and Au concentrations occur also in 

the altered wall rocks and skarns (Table 2). 

Hiltunen (1982) estimated that the size and grade of 

the Kuervitikko deposit is 1.2 Mt with Fe grades similar 

Fig. 8. Distribution of Cu and Au in selected cross sections from Laurinoja, Kuervitikko, and Cu-Rautuvaara. See Figs. 3 and 4 for drill hole 

locations. Data after Kinnunen (1980) and Keinänen (1995).

86 

Table 2 

Chemical composition of the wall rocks, skarns, and ores from the Laurinoja, Cu-Rautuvaara, and Kuervitikko 

Sample i.d. 1 2A 2B 2C 3A 3B 3C 3D 4A 4B 4C 4D 5 6A 6B 6C 7A 7B 8A 8B 

DDH 79 75 75 79 78 170 75 170 170 78 129 109 109 109 127 109 75 170 170 127 

Depth 37.50 33.60 43.75 148.30 212.45 113.40 136.25 99.00 65.85 147.80 90.40 237.30 231.70 219.30 65.20 274.00 108.70 48.45 71.65 42.80 

Rock 

Monzonite Diorite Diorite Diorite m-volc m-volc m-volc m-volc Skarn Skarn Skarn Skarn sca skarn Albitite Albitite Albitite+ 

Cu–Au 

Ironstone Ironstone Ironstone+ 

Cu–Au 

XRF (wt.%) 

SiO 2 58.78 56.69 55.79 51.23 48.44 54.50 49.20 33.31 48.64 41.61 27.14 43.22 44.70 67.08 54.77 45.15 2.73 26.48 17.56 18.36 

TiO 2 0.87 0.30 0.29 0.23 1.05 0.89 0.86 0.42 0.05 0.06 0.29 0.07 0.38 0.63 0.79 0.72 0.03 0.10 0.04 0.21 

Al 2 O 3 16.31 20.69 20.74 17.31 13.32 13.80 12.89 6.68 0.92 2.43 4.32 7.28 17.19 15.40 13.75 12.87 0.57 1.71 1.14 4.36 

Fe 2 O 3 (t) 6.83 4.85 8.23 13.42 15.21 13.79 13.41 32.99 12.22 29.03 14.28 12.06 12.79 4.49 10.35 21.51 86.49 59.39 59.33 63.46 

MnO 0.10 0.05 0.05 0.05 0.21 0.12 0.17 0.12 0.29 0.24 0.43 0.52 0.11 0.05 0.10 0.03 0.46 0.09 0.12 0.11 

MgO 2.67 2.11 1.24 1.33 6.49 5.11 7.06 5.43 11.06 5.20 6.27 12.29 0.62 1.00 3.25 1.65 4.65 7.38 4.31 2.66 

CaO 4.57 5.80 3.43 5.26 9.35 3.36 4.07 10.47 20.83 16.52 25.85 16.17 8.75 1.98 4.38 0.69 3.64 3.53 5.22 2.13 

Na 2 O 5.12 7.05 6.54 6.66 3.08 5.95 1.54 1.80 0.68 1.34 0.50 1.00 6.08 6.22 5.52 7.36 b0.07 0.59 0.14 1.69 

K 2 O 3.88 1.12 3.03 0.70 0.66 1.95 6.17 0.45 0.01 0.05 0.39 0.95 0.69 1.63 2.59 0.06 b0.01 0.26 0.12 0.24 

P 2 O 5 0.36 0.24 0.24 0.20 0.09 0.09 0.06 0.13 0.03 0.21 0.12 0.10 0.08 0.19 0.24 0.05 0.05 0.07 0.10 0.13 

C 0.01 0.02 b0.01 0.03 0.03 0.01 0.02 0.03 0.01 0.03 3.40 0.06 0.38 0.08 0.07 0.04 0.95 0.02 b0.01 0.02 

Cl 0.08 0.10 0.05 0.05 0.13 0.05 0.16 0.07 0.03 0.02 0.10 0.37 1.69 0.11 0.02 0.01 0.02 0.02 0.01 0.03 

S 0.01 0.01 0.05 4.60 0.12 0.27 0.02 6.07 0.25 0.04 2.51 0.01 4.85 0.08 3.11 6.43 0.20 0.06 7.25 5.66 

Total 99.61 99.03 99.67 101.07 98.15 99.89 95.63 97.97 95.03 96.76 85.60 94.10 98.30 98.94 98.95 96.58 99.80 99.69 95.32 99.06 

XRF (ppm) 

As b30 b30 b30 b30 b30 b30 b30 b30 b30 b30 b30 b30 b30 b30 b30 b30 147 b30 b30 b30 

Ba 924 405 1447 190 121 211 821 78 b20 25 34 78 156 1948 736 44 24 31 46 336 

Ce 115 45 62 42 b30 b30 50 68 b30 b30 54 55 51 365 81 b30 b30 854 b30 42 

Cr 31 b30 b30 b30 140 117 125 101 b30 b30 68 b30 b30 b30 189 100 30 b30 b30 54 

Cu 32 b20 28 955 140 59 b20 24,990 1744 29 5282 b20 2177 26 2133 51,360 995 44 65,130 27,910 

La 47 b30 b30 b30 b30 b30 b30 61 b30 b30 b30 b30 b30 185 32 b30 b30 729 b30 b30 

Mo b10 b10 b10 b10 b10 b10 b10 10 b10 b10 b10 b10 b10 b10 60 b10 b10 b10 b10 b10 

Nb 11 b10 b10 b10 b10 b10 b10 11 b10 b10 b10 46 50 b10 12 b10 b10 b10 b10 b10 

Ni b20 20 b20 194 85 70 79 64 21 67 50 b20 117 21 225 137 b20 137 66 b20 

Ironstone+ 

Cu–Au 

T. Niiranen et al. / Ore Geology Reviews 30 (2007) 75–105

Pb b30 b30 b30 b30 b30 b30 b30 b30 b30 b30 b30 154 b30 b30 b30 b30 b30 b30 b30 41 

Rb 148 42 67 18 38 73 128 19 b10 b10 b10 b10 11 48 43 b10 b10 b10 b10 20 

Sr 707 667 354 394 113 79 59 191 14 26 83 16 143 378 221 54 18 b10 12 61 

Th 11 b10 b10 b10 b10 b10 b10 b10 b10 b10 b10 180 40 19 b10 b10 b10 b10 b10 b10 

U b10 b10 b10 b10 b10 b10 b10 b10 b10 b10 b10 330 20 b10 b10 b10 b10 b10 b10 b10 

V 141 59 51 39 349 233 250 92 b3 95 95 b3 33 60 545 138 b3 156 b3 61 

Y 20 b10 b10 b10 31 27 29 18 b10 21 15 66 48 12 35 17 b10 b10 b10 b10 

Zn 69 b20 b20 b20 63 b20 58 156 44 34 105 70 24 26 25 54 36 22 123 33 

Zr 268 38 41 31 54 131 66 77 11 18 59 181 379 523 206 33 b10 17 13 44 

TiO 2 /Al 2 O 3 10 4 53.3 14.5 14.0 13.3 78.8 64.5 66.7 62.9 54.3 24.7 67.1 9.62 22.1 40.9 57.5 55.9 52.6 58.5 35.1 48.2 

Zr/Al 2 O 3 16.4 1.84 1.98 1.81 4.05 9.49 5.12 11.5 12.0 7.41 13.7 24.9 22.0 34.0 15.0 2.56 9.94 11.4 10.1 

Zr/TiO 2 308 127 141 135 51.4 147 76.7 183 220 300 203 2590 997 830 261 45.8 170 325 210 

ICP-AES (ppm) 

Ag b1 b1 b1 b1 b1 b1 b1 3 1 b1 n.a. n.a. n.a. n.a. n.a. n.a. 1 b1 6 n.a. 

Co 16 6 13 418 22 47 43 660 6 19 n.a. n.a. n.a. n.a. n.a. n.a. 383 24 154 n.a. 

GFAAS (ppb) 

Au b0.5 b0.5 b0.5 41 5 b0.5 3 818 191 b0.5 70 1 84 2 44 387 16 2 1020 3780 

Bi 5 b2 b2 172 8 8 17 389 73 b2 125 91 475 67 221 b20 2460 b2 396 853 

Sb 29 b25 b25 82 44 31 b25 b25 b25 b25 n.a. n.a. n.a. b25 n.a. n.a. 72 b25 b25 n.a. 

Se n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. 1960 b50 4140 b50 4130 470 n.a. n.a. n.a. 1560 

Te b2 2 3 812 127 43 11 1510 132 4 830 3 1710 10 1000 970 123 12 4070 892 

(1) Typical hanging wall monzonite, (2A) albite- and scapolite-altered diorite from upper distal alteration zone, (2B) biotite–K-feldspar-altered diorite from upper distal alteration zone, (2C) 

clinopyroxene–magnetite-altered diorite from upper proximal alteration zone, (3A) least altered footwall metavolcanic rock, (3B) albite–biotite-altered mafic metavolcanic rock from lower distal 

alteration zone, (3C) K-feldspar–biotite-altered mafic metavolcanic rock from lower proximal alteration zone, (3D) clinopyroxene–magnetite-altered and mineralised mafic metavolcanic rock, lower 

proximal alteration zone, (4A) clinopyroxene skarn, upper proximal alteration zone, (4B) magnetite-disseminated clinopyroxene dominated skarn, lower proximal alteration zone, (4C) weakly 

mineralised clinopyroxene skarn, upper proximal alteration zone, (4D) clinopyroxene–amphibole skarn, Cu-Rautuvaara, (5) clinopyroxene-bearing , magnetite-disseminated and weakly mineralised 

scapolite skarn, Cu-Rautuvaara, (6A) albitite, Cu-Rautuvaara. (6B) Albitite, Kuervitikko. Upper proximal alteration zone. (6C) Mineralised albitite, Cu-Rautuvaara. (7A) Barren, allanite–monatzitebearing 

ironstone, Laurinoja, (7B) barren ironstone, Laurinoja. (8A) Mineralised ironstone, Laurinoja, (8B) mineralised ironstone, Kuervitikko. DDH=diamond drill hole. The depth refers to the 

down-hole depth and is given as meters. n.a.=not analysed. Zr, TiO 2 , and Al 2 O 3 ratios calculated as Zr (ppm)/Al 2 O 3 (wt.%), and Zr (ppm)/TiO 2 (wt.%). 

T. Niiranen et al. / Ore Geology Reviews 30 (2007) 75–105 87

88 


to those at Laurinoja. Typical concentrations in the 

Kuervitikko deposit are 0.7 ppm and 0.5 wt.% for Au 

and Cu, respectively (Fig. 8; Keinänen, 1995). Besides 

ironstone, the Cu–Au mineralisation is hosted by a 

magnetite-disseminated albite–biotite-altered diorite 

that appears similar to the albitite that hosts the Cu– 

Au mineralisation at Cu-Rautuvaara (Fig. 8). The spatial 

correlation between Cu and Au is stronger at Kuervitikko 

than at Laurinoja. 

There is no previously published estimate of the size 

and grade of the Cu-Rautuvaara. Based on drilling the 

lens shaped magnetite-disseminated albitite unit hosting 

the bulk of the Cu–Au mineralisation (Fig. 8) is at least 

200 m in the longest dimension, up to 45 m thick, and 

is open at the depth of about 250 m (Fig. 4; Hiltunen, 

1982). According to our calculations, the size of the 

weakly to moderately mineralised albitite in the Cu- 

Rautuvaara deposit is 4.2 Mt. The Au concentration at 

Cu-Rautuvaara is generally 1 ppm or less, except for a 

few thin, V1 m, sections where it is up to 3 ppm. 

Copper concentrations up to 1.3 wt.% have been detected, 

but typically the Cu concentrations are 0.7 wt.% 

or less (Fig. 8; Keinänen, 1995). The correlation between 

Cu and Au is similar to that at Kuervitikko. 

5. Alteration 

Metasomatic alteration styles related to the ironstones 

and Cu–Au mineralisation consist of combinations of 

albite, amphiboles (actinolite–hornblende–anthophyllite), 

biotite, clinopyroxenes (diopside–hedenbergite), 

K-feldspar, magnetite, scapolite (marialite), calcite, and 

sulphides (pyrite–pyrrhotite–chalcopyrite) alteration 

(Fig. 9). At the deposit scale, zoning can be distinguished 

around the Laurinoja and Kuervitikko ironstones whereas 

clear zoning is difficult to distinguish around the Cu- 

Rautuvaara deposit. 

At Laurinoja, where the zoning is most distinct, the 

alteration halo can be divided into distal and proximal 

zones in both hanging wall and footwall host rocks 

(Fig. 9). The alteration assemblages vary somewhat 

depending on the primary rock type, but the general 

pattern is that albite Fscapolite, biotite, and K-feldspar 

are the dominant alteration minerals in the distal zones, 

and clinopyroxenes, amphiboles, and magnetiteF 

calcite are the dominant minerals in the proximal alteration 

zones. The change from distal to proximal alteration 

zones is defined by the appearance of 

clinopyroxene and the outer boundary of the distal 

alteration zone is defined by the appearance of albite. 

Distal alteration zones display a mineral variation with 

the amount of biotite and K-feldspar increasing toward 

Fig. 9. Schematic mineral paragenetic diagrams across the Laurinoja, 

Kuervitikko, and Cu-Rautuvaara deposits. Thickness of the lines 

corresponds to the abundance of the mineral. The proximal and distal 

zones cannot be distinguished at Cu-Rautuvaara due to poorly developed 

zonation.


proximal alteration zones. Next to the proximal zone, 

the distal alteration zone is biotite- and K-feldspar-dominated. 

This feature, and the change from distal to 

proximal alteration, are best developed in footwall 

mafic metavolcanic rocks, where the mineral assemblage 

changes from plagioclase–hornblende in the 

least altered rock (Fig. 6d), via biotite–albite (Fig. 6e) 

and K-feldspar–biotite–albite (Fig. 6f) in the distal zone, 

to K-feldspar–biotite–amphibole–clinopyroxene–magnetite 

in the proximal alteration zone (Fig. 6g). In the 

hanging wall distal alteration zone, the assemblages are 

similar, except that the albite alteration is more extensive 

and intense, and scapolite is more abundant than in the 

footwall. 

The dominant alteration mineral in proximal zones 

is clinopyroxene (Fig. 9). However, calcic amphibole 

commonly occurs together or instead of clinopyroxene 

as the main alteration mineral. Variable amounts 

of albite, biotite, K-feldspar, scapolite, calcite, and 

quartz occur in the proximal zones and locally their 

amount can exceed that of clinopyroxene and amphibole. 

Besides abundant magnetite, the alteration 

assemblages in ironstone are similar to the proximal 

alteration zones (Fig. 9). Scapolite skarn units (Fig. 

6i), dm to several m in thickness, occur locally within 

the proximal alteration zones, and in the ironstones 

scapolite is locally the main gangue mineral. Outside 

the ironstone, magnetite is most abundant in the 

proximal alteration zones, but can locally account 

for up to 15 vol.% in the distal alteration zones. In 

the Laurinoja ore body, most of the sulphides occur 

in the ironstone and in the proximal alteration zones, 

but both Fe- and Cu-sulphides are locally abundant (1 

to 10 vol.%) in all altered rocks independent of the 

alteration assemblages. 

In the Kuervitikko deposit, the alteration assemblages 

define a similar zoning as described for Laurinoja 

(Fig. 9). The most significant differences are that 

at Kuervitikko, K-feldspar and biotite are less common 

in the hanging wall distal alteration zone, and there is 

an albitite unit between the diorite and clinopyroxene– 

amphibole skarn. In addition, amphibole is a slightly 

more abundant proximal alteration mineral than at 

Laurinoja. Also, compared to Laurinoja, sulphides appear 

to be more abundant in the distal alteration zones, 

especially in the albitite (Fig. 9). 

At Cu-Rautuvaara, the alteration zoning is poorly 

developed, although the alteration assemblages are similar 

to Laurinoja and Kuervitikko (Fig. 9). A thin 

sequence of clinopyroxene skarns with alteration assemblage 

similar to the proximal alteration zones at 

Laurinoja and Kuervitikko occurs in the hanging wall. 

However, as stated above, the main Cu–Au mineralisation 

is within a magnetite-disseminated anthophylliteand 

biotite-bearing albitite (Figs. 6h and 9). 

6. Geochemistry 

6.1. Monzonite 

The chemical composition of the monzonite indicates 

that it is metaluminous, low in Nb and Y, and 

shows a calc-alkaline affinity (Table 2). These features 

Fig. 10. Chondrite-normalised REE patterns for the samples presented 

in Table 3. Normalising values for C1 chondrite after McDonough 

and Sun (1995).

90 


are typical for the 1900 to 1860 Ma Haparanda Suite 

intrusions abundant throughout northwestern Finland 

and northern Sweden (e.g., Mellquist et al., 2003). 

The chondrite-normalised REE pattern of the monzonite 

shows enrichment in LREE in respect to HREE and 

a weak negative Eu anomaly (Fig. 10a). 

6.2. Diorite and mafic metavolcanic rocks 

Due to the highly altered nature of the wall and 

host rocks of the Kolari deposits, the primary classification 

of the variably altered rocks are distinguished 

using TiO 2 –Al 2 O 3 , Zr–Al 2 O 3 , and Zr–TiO 2 ratios; 

elements that usually display immobile behaviour in 

most hydrothermal systems. The variably altered diorites 

display remarkably constant TiO 2 –Al 2 O 3 , Zr– 

Al 2 O 3 , and Zr–TiO 2 ratios (Fig. 11). The chondritenormalised 

REE pattern of the diorite is identical to 

monzonite, except for the slightly positive Eu anomaly 

and the lower general REE content (Fig. 10a). This, 

combined with the age data (see below), suggests that 

the monzonite and diorite were derived from the same 

magma source. 

The metavolcanic rocks form two groups based on 

their TiO 2 –Al 2 O 3 , Zr–Al 2 O 3 , and Zr–TiO 2 ratios (Fig. 

11). One group has a similar Zr–TiO 2 ratio to least 

altered mafic metavolcanic rock (average Zr/ 

TiO 2 =65F15 ppm/wt.%, n =6), while the other 

group shows a considerably higher (average Zr/ 

TiO 2 =174F50 ppm/wt.%, n = 6) ratio (Fig. 11). This 

indicates a primary geochemical variation. Based on the 

difference in Zr/TiO 2 ratios, the metavolcanic rocks are 

divided into type-1 (lower Zr–TiO 2 ) and type-2 (higher 

Zr–TiO 2 ). The Zr–TiO 2 ratio of the type-1 is typical for 

basalts while the ratio of the type-2 rocks straddles the 

border of basalts and andesites in the classification 

diagram of Winchester and Floyd (1977). Both types 

are present at Laurinoja and Kuervitikko whereas at Cu- 

Rautuvaara, only type-1 has been detected. However, 

the data from Cu-Rautuvaara are too limited to rule out 

the possibility that both types of mafic metavolcanic 

rocks occur there, too. 

At Laurinoja and Kuervitikko, the chemical composition 

of diorite and metavolcanic rocks displays a 

zoning that corresponds to zoning in alteration assemblages. 

The rocks in the distal alteration zones are rich 

in Na 2 OFK 2 O, Cl, Ba, and the proximally altered 

rocks are enriched in CaO FFe 2 O 3t , S, Ag, Au, Bi, C, 

Ce, Co, Cu, La, and Te compared to distal altered rocks 

(Table 2; Fig. 9). For the evaluation of the effect of 

alteration on the chemical composition of rocks, major 

element oxides and selected minor elements are plotted 

Fig. 11. Immobile element ratios in variably altered diorite, metavolcanic rocks, skarns, and ironstones from Laurinoja, Kuervitikko, and Cu- 

Rautuvaara. C8=slope for the lines defined by average TiO 2 –Al 2 O 3 , Zr–Al 2 O 3 , and Zr–TiO 2 ratios for diorite and type-1 and -2 metavolcanic rocks.


against the immobile Al 2 O 3 (Fig. 12). The variation in 

mobile elements shows apparent loss or gain of elements 

during the alteration. Correlation of mobile element 

with immobile Al 2 O 3 suggests that the change of 

the former was coupled with a net mass change of the 

rock during alteration. 

Fig. 12. Major element oxide variation against the immobile Al 2 O 3 in altered diorite, metavolcanic rock, clinopyroxene–amphibole-dominated 

skarns, and ironstones from the Laurinoja, Kuervitikko, and Cu-Rautuvaara deposits with interpreted alteration trends.

92 


Both the distal altered diorite and metavolcanic 

rocks show considerable variation in K 2 O and 

Na 2 O concentrations reflecting their proportions of 

albite, biotite, and K-feldspar. Barium displays very 

high (up to 1500 ppm; Table 2) local values in distal 

altered rocks. Since no barite has been detected, it is 

very likely that Ba is bound within biotite and/or K- 

feldspar in these rocks. The concentration of Si 2 O, 

Fe 2 O 3t , CaO and MgO show negative correlation 

with Al 2 O 3 in the proximal altered rocks indicating 

that proximal alteration was accompanied with significant 

net mass gain (Fig. 12). The increase in the 

CaO, Fe 2 O 3t , and MgO reflects the increase in 

amount of clinopyroxene, magnetite, and amphiboles 

in proximal altered rocks compared to distal altered 

rocks. In addition, high CaO concentrations coupled 

with elevated C concentrations in the proximal altered 

rocks reflect the presence of calcite. Sulphur 

shows elevated (N1.5 wt.%) values only in proximal 

altered diorites and metavolcanic rocks (Fig. 12). The 

elevated sulphur concentrations display positive correlation 

with concentrations of Ag, Au, Bi, Co, Cu, 

and Te (Table 2). 

6.3. Clinopyroxene–amphibole skarns and ironstones 

The skarns and ironstones, together with albitites, 

are the most intensely altered rocks in the Kolari deposits. 

The TiO 2 –Al 2 O 3 , Zr–Al 2 O 3 , and Zr–TiO 2 ratios in 

the clinopyroxene skarns and ironstones are similar for 

both rock types suggesting common origins for these 

rocks (Fig. 11). However, there is some variation in 

the immobile element ratios, implying that the protolith 

for these rocks was at least locally heterogeneous 

and/or consisted of several rock types. Nevertheless, 

the ratios in the skarns and ironstones show considerable 

similarities with other rock types, especially with 

type-2 metavolcanic rocks, and a continuum from 

altered metavolcanic rocks and diorites to skarns and 

ironstones can be envisioned in a number of the major 

element oxide bivariant plots (Figs. 11 and 12). The 

elements that display high concentrations in proximal 

altered diorites and metavolcanic rocks also have high 

concentrations in the skarns and ironstones (Fig. 12; 

Table 2). The main difference between the chemical 

composition in the skarns and ironstones is the drastically 

higher Fe 2 O 3t concentrations in the latter (Fig. 

12). In addition, concentration of CaO is significantly 

higher in skarns than ironstones, yet this can be 

explained with the higher amount of clinopyroxene 

and calcic amphibole in the former compared to the 

latter. The very high (N20 wt.%) CaO concentrations 

in the skarns are coupled with elevated C values and 

this is due to significant amount of calcite in the rock 

(Fig. 12; Table 2). The behaviour of Ag, Au, Bi, Co, 

Cu, S, and Te in the skarns and ironstones is similar to 

the proximal altered diorites and metavolcanic rocks, 

but the concentration of these elements is typically 

higher in the former rock types indicating higher 

sulphide contents (Table 2). Concentrations of As, 

Pb, and Zn are typically low in all rocks. The 180 

ppm Th and 330 ppm U concentration in the uraninite-bearing 

clinopyroxene–amphibole skarn suggest 

that these elements have also been mobile during the 

skarn alteration. 

The P REE in the skarns and ironstones is typically 

low, although locally the ironstones display considerably 

high REE concentrations (Table 3). The main 

REE-bearing minerals in these ironstones are allanite 

and monazite. The chondrite-normalised REE patterns 

for both skarns and ironstones are relatively flat with 

only slight enrichment in LREE in respect to HREE 

(Fig. 10b, c). However, the allanite- and monazite-rich 

ironstone displays a strong LREE enrichment (Fig. 

10c). This suggests that at least LREE, and possibly 

also some of the HREE, have been mobile during the 

alteration, and further explains the locally elevated La 

and Ce values in all altered rocks (Table 2). Nevertheless, 

the REE patterns of the skarns and ironstones 

are relatively similar to distal altered type-1 metavolcanic 

rock (Fig. 10). This further supports the idea that 

the metavolcanic rocks are the precursors for the 

skarns and ironstones. 

6.4. Albitites and scapolite skarns 

Geochemically the most peculiar rock types at Kolari 

are the scapolite skarns and albitites. Typical for both 

rock types is the high Na 2 O concentration (up to 8.6 

wt.%) and the highly variable concentrations in Al 2 O 3 , 

TiO 2 ,Zr(Table 2). These rocks can be found throughout 

the ore-bearing sequence around all the deposits. Albitites 

are most commonly found in the hanging and 

footwalls whereas the scapolite skarns are typically 

located in the proximal alteration zones. Thus, it is 

likely that the protolith for these rocks includes several 

rock types. In places, especially at Cu-Rautuvaara, these 

rocks are mineralised with a metal association that is 

similar to mineralised ironstones and skarns (Table 2). 

Like the other altered rocks, the albitites and scapolite 

skarns contain locally elevated concentrations of Ba 

(1950 ppm), Ce (365 ppm), and La (185 ppm). Generally 

the concentrations of Cr, Mo, Nb, Ni, Th, U, and V 

are in places higher in these rocks than anywhere in


Table 3 

Rare earth element composition of unaltered and altered rocks and ironstones 

Sample i.d. 1 2A 3E 4E 4F 6C 7A 7B 7C 7D 

DDH 79 75 75 75 75 109 75 170 170 130 

Depth (m) 37.50 33.60 144.15 53.80 71.50 274.00 108.70 48.45 107.15 107.40 

La 37.0 9.00 7.32 4.32 2.68 10.0 2.12 698 6.50 1.93 

Ce 75.1 18.5 17.0 12.0 5.25 16.8 4.39 841 6.16 2.92 

Pr 8.36 2.29 2.23 1.91 0.77 1.74 0.57 58.4 0.63 0.32 

Nd 33.4 9.43 10.7 8.95 3.94 6.92 2.26 145 2.49 1.57 

Sm 5.35 1.77 2.79 2.14 1.14 1.20 0.52 7.75 0.69 0.45 

Eu 1.32 0.66 0.94 0.66 0.53 0.39 0.10 1.83 0.47 0.18 

Gd 5.16 1.44 3.77 1.82 0.96 1.67 0.69 7.71 1.13 0.41 

Tb 0.71 0.20 0.61 0.32 0.16 0.26 0.10 0.86 0.22 b.d. 

Dy 3.28 1.04 3.35 1.87 0.98 1.76 0.44 1.25 1.24 0.23 

Ho 0.6 0.15 0.74 0.34 0.17 0.38 b.d. 0.18 0.25 b.d. 

Er 1.70 0.50 2.33 0.97 0.56 0.98 0.30 0.45 0.72 b.d. 

Yb 1.69 0.44 1.85 1.17 1.00 1.15 0.21 0.54 0.68 b.d. 

Lu 

P 

0.23 b.d. 0.29 0.16 0.19 0.16 b.d. 0.12 0.11 b.d. 

REE 173.90 45.42 53.92 36.63 18.33 43.41 11.70 1763.09 21.29 8.01 

Sample numbering as in Table 2. 1=unaltered monzonite, Laurinoja, 2A=albitised diorite, Laurinoja, 3E=weakly albitised mafic metavolcanic 

rock, Laurinoja, 4E–F=clinopyroxene skarn from upper proximal alteration zone, Laurinoja, 6C=Cu–Au mineralised albitite, Cu-Rautuvaara, 

7A=weakly mineralised ironstone, Laurinoja, 7B=barren, allanite and monazite bearing ironstone, Laurinoja, 7C=Cu–Au mineralised ironstone, 

Laurinoja, 7D=barren ironstone, Kuervitikko. All values given in ppm. b.d.=below detection limit. DDH=diamond drill hole. Depth as down-hole 

depth (meters). 

other altered rock types (Table 2). The chondrite-normalised 

REE pattern of the Cu–Au-mineralised, magnetite-disseminated 

albitite at Cu-Rautuvaara shows 

LREE enrichment relative to HREE, but the pattern 

does not seem to correlate to any other rock type 

(Fig. 10). 

6.5. Mass balance calculations 

There probably are several protoliths for skarns and 

ironstones at Kolari, but the immobile-element ratios 

suggest that the dominant protolith is the type-2 metavolcanic 

rock (Fig. 11). It is evident that significant net 

mass changes did occur during the proximal alteration 

(Figs. 11 and 12). This hinders the estimation of which 

rocks were protoliths to which parts of skarn and 

ironstone units. However, the effect of the net mass 

changes is eliminated by plotting the data presented in 

Figs. 11 and 12 into the TiO 2 /Al 2 O 3 vs. Zr/Al 2 O 3 

binary plot (Fig. 13). For comparison, the albitites, 

scapolite skarn and monzonite presented in Table 2 

are also plotted on the diagram. In the ratio to ratio 

plot the rocks form distinct groups of their own. Variably 

altered diorites form group I, type-1 metavolcanic 

rocks, group II, and type-2 metavolcanic rocks, group 

III (Fig. 13). In addition, three skarn, and three ironstone 

samples plot near to each other, forming a distinct 

group (IV). The majority of skarn and ironstone sam- 

Fig. 13. TiO 2 –Al 2 O 3 vs. Zr–Al 2 O 3 plot of the samples shown in Figs. 11 and 12. Monzonite, albitites, and scapolite skarn presented in Table 2 

added for comparison. Encircled areas defined as follows: (I) distal and proximal altered diorites, (II) distal and proximal altered type-1 

metavolcanic rocks, (III) distal and proximal altered type-2 metavolcanic rocks and overlapping skarns and ironstones, (IV) skarn and ironstone 

samples that appear to correlate with each other but not with any other rock types.

94 


ples overlap the area defined by group III indicating 

that the protolith for these is the type-2 metavolcanic 

rock. The skarn and ironstone samples forming group 

IV as well as other samples that scatter around the 

diagram do not seem to correlate with any known 

rock type, and thus the protolith for these remains 

unknown. Possible candidates are the sedimentary 

rocks in the sequence (Figs. 3 and 4). 

All samples within the group III (Fig. 13) were 

selected for the mass balance calculations. Due to 

heterogeneity of the immobile element ratios in the 

type-2 metavolcanic rocks average concentrations 

were used for calculations instead of individual samples. 

The average concentrations were calculated for 

distal and proximal altered type-2 metavolcanic rocks, 

skarns, and ironstones representing moderate proximal 

alteration, intense proximal alteration, and intense Fealteration, 

respectively (Table 4). The mass balance 

calculations were carried out using isocon method described 

by Grant (1986). Since there is no unaltered 

type-2 metavolcanic rock available, the calculated 

averages of the distal altered varieties were used as 

reference samples for the calculations instead of unaltered 

rocks. Aluminium, titanium and zirconium were 

used as immobile elements in the calculations. Calculations 

are based on constant volume assumption that is 

probably not true for the most intense alteration. Nevertheless, 

the results (Fig. 14) are in accordance with 

the changes in mineral assemblage, and we thus feel 

that they reflect the true chemical changes that occurred 

during the alteration. 

The isocons indicate constant and considerable net 

mass gains between both proximally altered samples 

(moderately and intensely altered) and intensely Fealtered 

samples (Fig. 14). The Zr, TiO 2 , and Al 2 O 3 

composition plot very close to the isocon in all diagrams 

indicating that they were conserved during the 

alteration. In addition, SiO 2 , MgO, MnO, and Cr 

appear to have been conserved, for the most part, in 

moderate proximal alteration. All other major elements, 

except alkalis, show gains in all three groups. 

Of these the gains in Fe 2 O 3t , CaO, and S are highest 

in all three groups, and MgO display significant gain 

in intensely altered groups. Of the trace elements the 

highest gains in all groups are in light rare earth 

elements and chalcophile elements. Only V shows 

constant depletion in all groups. Since the isocons 

were constructed using distal altered rocks as reference 

samples (i.e., Na–K altered rocks) instead of 

truly unaltered rocks, it is possible that some Na 2 O 

and/or K 2 O was also added into the skarns and ironstones 

during the alteration event though the isocon 

Table 4 

Average chemical composition values in the rocks used in mass 

balance calculations 

A B C D 

n 3 3 5 7 

SiO 2 54.75 42.51 32.32 23.58 

TiO 2 0.83 0.56 0.27 0.18 

Al 2 O 3 14.19 10.09 4.31 3.28 

Fe 2 O 3 (t) 11.97 22.22 12.15 54.76 

MnO 0.08 0.10 0.43 0.20 

MgO 5.00 4.24 7.48 5.01 

CaO 2.59 8.47 25.34 6.67 

Na 2 O 4.53 4.10 0.46 0.97 

K 2 O 3.90 0.61 1.21 0.37 

P 2 O 5 0.11 0.12 0.17 0.10 

C 0.01 0.02 2.74 0.08 

Cl 0.08 0.03 0.06 0.05 

S 0.95 5.95 1.86 3.71 

(ppm) 

Ba 390 176 182 90 

Ce 39 68 49 164 

Cr 130 88 55 46 

Cu 122 10,584 1979 20,550 

La 22 30 27 152 

Mo 29 57 9 44 

Ni 119 122 63 80 

Sr 63 153 67 45 

V 347 110 139 66 

Zn 22 62 47 91 

Zr 134 100 47 36 

(ppb) 

Au 7 303 61 849 

Bi 74 279 105 333 

Te 221 1803 519 828 

TiO 2 /Al 2 O 3 0.0584 0.0550 0.0630 0.0542 

Zr/Al 2 O 3 * 9.46 9.94 10.96 10.89 

Zr/TiO 2 * 162 181 174 201 

A=distal altered type 2 metavolcanic rocks, B=proximal altered type 

2 metavolcanic rocks, C=skarns, D=ironstones, barren and Cu–Au 

mineralised. n =number of samples. *(ppm/wt.%). 0.5detection 

limit is used for values below detection limits in calculations. See 

Table 2 for the detection limits. 

diagrams suggest depletion of these elements in some 

of the groups. 

No suitable sample pairs were found for mass balance 

calculations in distal alteration. However, since the 

distal alteration (albite–scapolite–biotite–K-feldspar alteration) 

caused no significant mass changes (Figs. 11 

and 12), the change in chemistry between least altered 

type-1 metavolcanic rocks and distal altered varieties of 

them reflects the absolute gains or losses that occurred 

during the alteration (Table 2). According to this, the 

main gains in the distal altered rocks are in Na 2 O, and 

K 2 O accompanied with loss in CaO. The Ba concentrations 

in biotite–K-feldspar altered rocks are typically


Fig. 14. Isocon diagrams for type-2 metavolcanic rocks showing elemental changes during alteration. Average values of the groups plotting within 

field III in Fig. 13 are used. The calculated averages for all three groups and reference samples are shown in Table 4. Major oxides, C, Cl, and S 

plotted in wt.%, trace elements in ppm. C8=slope of the isocon, R 2 =correlation coefficient between Zr, TiO 2 , and Al 2 O 3 .

96 


high, thus it is evident that the distal altered rocks also 

gained Ba during the alteration. 

7. Fluid inclusion data 

7.1. Sample description 

Fluid inclusion work was carried out on five samples, 

three from the Laurinoja deposit (DDH78-148.80, 

DDH170-51.40, 1-LAU-01), and two from the Kuervitikko 

deposit (DDH129-135.40, DDH130-115.10). The 

DDH78-148.80 is a quartz–magnetite vein from clinopyroxene 

skarn next to the massive ironstone and 

DDH170-51.40 is a barren ironstone with quartz, K- 

feldspar, actinolite, and albite gangue. Sample 1-LAU- 

01 was taken from sulphide-bearing quartz–scapolite 

vein in ironstone (Fig. 5a, b). Sample DDH129- 

135.40 is a pyrrhotite–chalcopyrite–pyrite–magnetitebearing 

quartz vein from footwall proximal altered 

metavolcanic rock and DDH130-115.10 is a chalcopy- 

Fig. 15. (a, b) Typical samples from which fluid inclusion work were carried out. (a) Pyrrhotite–pyrite–chalcopyrite–magnetite-bearing quartz vein 

from Kuervitikko. (b). Chalcopyrite- and carbonate-bearing quartz vein from the Kuervitikko. Pyrite and chlorite filled fractures brecciate the vein. 

(c) Photomicrograph of a pyrrhotite grain from sample (a) with a dense, sphere-like cloud of type-1 and -3 inclusions around the grain. (d) Close up 

of the inclusions in quartz around the pyrrhotite grain shown in (c). Secondary, lower salinity inclusions occur in a thin fracture that crosscut the 

type-1 and -3 inclusions. (e) Type-1a inclusions in calcite of the sample shown in (b). The type-1a inclusions in this sample contain ubiquitously an 

opaque solid. (f) A typical type-1a inclusion that contains up to five different solids in addition to halite, Kuervitikko. (g) Co-existing type-1b and -3 

inclusions, Laurinoja. (h) Type-1b inclusions from the brittle scapolite–quartz–carbonate vein shown in Fig. 5a–b. Halite is the only solid phase in 

this case. (i). Type-2 inclusion in quartz.


rite-bearing quartz–calcite vein, brecciated by pyrite 

and chlorite filled fractures, from hanging wall proximal 

altered diorite (Fig. 15). 

7.2. Fluid inclusion types and microthermometry 

Three main types of fluid inclusions were recognised 

in the Kolari samples. These are identified here using a 

combination of textural, phase proportional, and microthermometric 

observations: (1) H 2 OFCO 2 +halite F 

solids, (2) H 2 OFCO 2 , and (3) CO 2 (Table 5; Fig. 15). 

Type-1 inclusions are divided into two subgroups 

based on their mode of occurrence and solid content. 

Type-1a inclusions are complex aqueous- or aqueouscarbonic 

liquid-vapour multisolid (up to 6 phases) 

inclusions that were found as solitary inclusions or 

primary clusters in quartz paragenetically associated 

with sulphides (Fig. 15). Type-1b inclusions were 

found as a network of brecciating inclusion trails in 

quartz and calcite associated with sulphides and as 

solitary inclusions together with type-2 inclusions 

(Fig. 15). The type-1a inclusions occur between the 

type-1b inclusion trails. The longest dimensions of 

both inclusion types range from about 10 to 50 Am 

and the vapour bubble occupies only a small volume of 

them (less than 30 vol.%). Vapour bubbles of the 

inclusions contain enough CO 2 to form clathrate but 

not a visible phase of liquid CO 2 during cooling. Salinity 

estimates based on halite dissolution temperatures 

are in the range of 45 to 48 and 32 to 56 wt.% 

NaCl equiv. for the type-1a and -1b inclusions, respectively. 

In type-1a inclusions, halite disappears prior to 

vapour bubble and in the type-1b inclusions vapour 

bubble disappears simultaneously or more typically 

prior to halite dissolution. Some daughter minerals 

remained undissolved up to halite dissolution temperatures. 

Density of the type-1a and -1b inclusions varies 

from 1.08 to 1.09 g/cm 3 and from 1.14 to 1.28 g/cm 3 , 

respectively. 

Type-2 inclusions are simple aqueous Fcarbonic liquid-vapour 

inclusions in which the vapour bubbles do 

not contain enough CO 2 to form either clathrate or a 

visible phase of liquid CO 2 during cooling. This suggests 

that the CO 2 content of the bubble is less than 

0.85 mol.% (Hedenquist and Henley, 1985). The vapour 

bubble occupies less than 15 vol.% and the longest 

dimensions of the inclusions range from about 10 to 

50 Am. They occur as inter- and intragranular trails in 

quartz and calcite associated with sulphides (Fig. 15). 

In calcite, some inclusions contain a single daughter 

mineral that dissolves at temperatures around 120 8C 

during heating. The inclusions have homogenisation 

temperatures to the liquid phase ranging from 126 to 

315 8C. Salinity estimates based on final ice melting 

temperatures are in the range from 9 to 22 wt.% NaCl 

equivalent. First melting temperatures of about 35 8C 

Table 5 

Summary of fluid inclusion microthermometric data (8C) 

Deposit Min FI type ThCO 2 (L) TmCO 2 TmH 2 O ThH 2 OFCO 2 (L) TdNaCl Salinity 

Kuervitikko Qtz Type 1a 386 to 463 375 to 402 45 to 48 

Qtz Type 1b 132 to 266 200 to 447 32 to 56 

Qtz Type 2 7.3 to 5.1 155 to 174 8 to 11 

Qtz Type 3 20.3 to 25.7 56.7 to 56.8 Distinct 

co-occurring 

type-1b and type-3 

inclusion trails in 

inclusion bcloudsQ 

or halos surrounding 

sulphides (Fig. 15) 

10.2 to 24.5 57.4 to 58.4 

0.6 to 24.6 61.1 to 60.2 Distinct 

intragranular trails 

Cc Type 1b 146 to 209 198 to 244 32 to 35 

Cc Type 2 5.6 to 4.5 142 to 157 7 to 9 

Laurinoja Qtz Type 1b 109 to 216 209 to 251 32 to 35 

Qtz Type 2 19 to 6 126 to 315 9 to 22 

Qtz Type 3 20 to 2.7 56.6 Inter- and 

intragranular trails 

Min=mineral, FI=fluid inclusion type-1a=H 2 OFCO 2 -halite-solids, type-1b=H 2 OFCO 2 -halite, type-2=H 2 OFCO 2 , type-3=CO 2 FCH 4 , 

ThCO 2 =temperature of homogenisation of CO 2 to liquid (L), TmCO 2 =final melting temperature of CO 2 , TmH 2 O=temperature of melting of 

ice, ThH 2 O/CO 2 =temperature of homogenisation of H 2 O to liquid (L) and temperature of total homogenisation of H 2 O and CO 2 to liquid (L), 

TdNaCl=dissolution temperature of halite, Salinity=equivalent weight % NaCl (TmH 2 O) or weight % NaCl (TdNaCl).

98 


for the inclusions match the eutectic for the H 2 O– 

MgCl 2 –NaCl system (Crawford, 1981). Their density 

varies from 0.92 to 1.00 g/cm 3 . 

Type-3 inclusions are single phase liquid CO 2 -rich 

inclusions which were found in quartz (not in calcite) as 

intra- and intergranular trails, as halos around sulphide 

grains, or they occur together with the type-1b inclusions 

in short intragranular trails (Fig. 15). Their longest 

dimensions vary between 5 and 25 Am. Some 

inclusions contain a single solid phase. Based on 

CO 2 -freezing point depression, in intergranular and 

intragranular trails with the type-1b inclusions they 

contain pure CO 2 , whereas in halos and in intragranular 

trails they contain 7 to 10 and 15 to 20 mol.% CH 4 /N 2 , 

respectively. Density of the type-3 inclusions range 

from 0.72 to 1.03 g/cm 3 in inter- and intragranular 

trails, 0.70 to 0.77 g/cm 3 in trails together with the 

type-1b inclusions, and from 0.72 to 0.98 g/cm 3 in 

halos surrounding sulphides. 

8. U–Pb geochronology 

8.1. Previous age determinations 

Previous U–Pb age data from the Kolari deposits 

have been published by Hiltunen (1982) and were 

obtained from a monzonite sample from the Hannukainen 

deposit (A840), albitite in the hanging wall of the 

Rautuvaara deposit (A958), a coarse-grained bmafic 

pegmatoidQ (A959) from the Rautuvaara deposit, and 

clinopyroxene skarn from the Hannukainen deposit 

(A963). The ages for samples A840, A958, A959, 

and A963 given below are from Väänänen and Lehtonen 

(2001) in which Hiltunen’s (1982) recalculated U– 

Pb data are presented. All these previously dated samples 

were re-checked from the GSF archives, and brief 

updated sample descriptions are given below. 

Sample A840 is a typical monzonite that occurs in 

the hanging wall for the Kuervitikko and Cu-Rautuvaara 

and Hannukainen deposits. In macroscopic investigation, 

it appears as slightly altered, possibly albitised. 

It displays a well developed foliation and lineation, 

probably S 3 and L 3 . The 1862F4 Ma U–Pb zircon 

age of the monzonite is based on four fractions that 

are slightly discordant. The age for titanite is 1784 Ma. 

Hiltunen (1982) proposed that A958 is an albitic 

variety of the monzonite derived via magmatic fractionation. 

In contrast, our opinion is that albitites are metasomatic 

rather than orthomagmatic rocks. Macroscopic 

investigation of A958 revealed that it is similar to the 

distal altered diorite at Laurinoja and Kuervitikko. The 

A958 albitite is strongly foliated and the lineation is 

well developed. The ages for zircon and titanite are 

1849F19 Ma and 1783 Ma, respectively. 

According to Hiltunen (1982), sample A959 was 

taken from a narrow, coarse-grained mafic pegmatoid 

vein that crosscuts the skarn and skarn hosted ore in the 

Rautuvaara deposit. The macroscopic review of A959 

showed that it is unoriented, coarse-grained (V15 mm), 

and consists of plagioclase (albite?), amphibole, clinopyroxene, 

quartz, and possibly scapolite. The origin of 

the vein is ambiguous but it could be linked to the late 

(D 4 ) brittle deformation (e.g., Fig. 5a, b). The 1748F7 

Ma zircon age of the vein from six highly discordant 

zircon fractions is the youngest age obtained from the 

whole of northern Finland (Hanski et al., 2001). 

Sample A963 is a typical clinopyroxene skarn that 

occurs as the main host rock for the ironstones for the 

Kuervitikko and Hannukainen deposits. It consists almost 

solely of clinopyroxene and is in very pristine 

condition showing no signs of retrograde alteration. 

The 1797 F5 Ma age for the skarn is constructed by 

three slightly discordant zircon fractions and titanite is 

coeval. 

8.2. U–Pb age determinations of A1697 granite and 

A1698 diorite 

Two new rock samples (A1697 and A1698) were 

taken for U–Pb dating. U–Pb data on zircon and titanite 

are shown in Table 6 and the concordia plots presented 

in Fig. 16. All U–Pb ages from the Kolari deposits are 

summarised in Fig. 17. 

The A1697 granite was taken from the waste rock 

piles at Hannukainen. It is even grained, unoriented and 

brecciates and crosscuts Cu–Au-mineralised ironstone 

(Fig. 5c–d). Four fractions of prismatic, magmatic zircon 

and one fraction of dark brown, translucent titanite 

were dated (Table 6). On the concordia diagram, the 

prismatic zircons give an upper intercept age of 

1766F5 Ma(Fig. 16). The concordia age for titanite 

is 1759F3 Ma. Accordingly, the ages of the magmatic 

zircon and titanite are coeval within the error limits 

(Fig. 17). 

The granite also contains a small amount of bright, 

colourless and rounded zircon without crystal faces. 

Consequently, these grains have characteristics typical 

of metamorphic zircon. The Pb–U ratios of these zircon 

grains plot quite close to the concordia curve (Fig. 16). 

The age of the only slightly discordant data can be 

approximated using 207 Pb/ 206 Pb ratio to be ca. 

1800 Ma. As the magmatic age of the Hannukainen 

granite is significantly younger, the 1800 Ma zircon 

must be inherited in origin. Interestingly, this age is

Table 6 

Conventional U–Pb age data on zircon and titanite, samples A1697 granite, and A1698 diorite 

Sample information, analysed 

mineral and fraction 

Sample 

(weight/mg) 

U 

(ppm) 

Pb 

(ppm) 

206 Pb/ 204 Pb 

(measured) 

208 Pb/ 206 Pb 

(radiogenic) 

Isotopic ratios a 

Apparent ages/MaF2r 

206 Pb/ 238 U F2r 207 Pb/ 235 U F2r 207 Pb/ 206 Pb F2r Rho 206 Pb/ 238 U 207 Pb/ 235 U 207 Pb/ 206 Pb 

A1697 granite 

(A) Zircon d: N4.2 g/cm 3 , N75 Am, 

bright, colourless, rounded, 

abraded 4 h 

(B) Zircon d: N4.2 g/cm 3 , b75 Am, 

long prismatic, translucent, 

abraded 10 h 

(C) Zircon d: N4.2 g/cm 3 , b75 Am, 

long prismatic, translucent, 

abraded 1 h 

(D) Zircon d: 4.2–4.0 g/cm 3 , 

long prismatic, 

translucent, brownish, abraded 10 h 

(E) Zircon d: 4.2–4.0 g/cm 3 , 

long prismatic, 

translucent, brownish, abraded 1 h 

(F) Titanite, dark brown, transparent, 

abraded 1/3 h 

A1698 diorite 

(A) Zircon d: N4.2 g/cm 3 , N75 Am, 

prismatic, transparent, colourless, 

abraded 4 h 

(B) Zircon d: N4.2 g/cm 3 , b75 Am, 


abraded 5 h 

(C) Zircon d: N4.2 g/cm 3 , b75 Am, 


abraded 3 h 

(D) Zircon d: N4.2 g/cm 3 , b75 Am, 

prismatic, transparent, colourless 

(E) Titanite, dark brown, translucent, 

abraded 1/2 h 

0.59 152 52 2078 0.16 0.3058 0.65 4.619 0.65 0.1096 0.15 0.97 1720 1753 1792F2 

0.32 1425 361 3244 0.14 0.2318 0.65 3.373 0.65 0.1055 0.15 0.97 1344 1498 1724F2 

0.26 1120 229 3061 0.13 0.1895 0.65 2.710 0.65 0.1037 0.15 0.97 1119 1331 1691F2 

0.26 1100 297 2978 0.14 0.2472 0.65 3.623 0.65 0.1063 0.15 0.97 1424 1555 1737F2 

0.48 1317 317 2554 0.13 0.2212 0.65 3.212 0.65 0.1053 0.15 0.97 1288 1460 1720F2 

1.97 173 96 654 0.83 0.3143 0.40 4.662 0.42 0.1076 0.20 0.88 1762 1761 1759F4 

0.24 579 221 1659 0.25 0.3149 0.40 4.939 0.42 0.1137 0.13 0.95 1765 1809 1860F3 

0.50 456 183 1828 0.30 0.3205 0.40 5.027 0.42 0.1137 0.07 0.98 1792 1824 1860 

0.31 432 170 966 0.28 0.3132 0.40 4.908 0.42 0.1136 0.10 0.96 1757 1804 1858F2 

0.16 537 202 1014 0.23 0.3115 0.40 4.872 0.42 0.1135 0.17 0.89 1748 1798 1855F3 

1.85 127 63 759 0.56 0.3241 0.40 4.924 0.60 0.1102 0.46 0.62 1810 1806 1803F9 

a Isotopic ratios corrected for fractionation, blank (50 pg), and age related common lead (Stacey and Kramers, 1975; 206 Pb/ 204 PbF0.2, 207 Pb/ 204 PbF0.1, 208 Pb/ 204 PbF0.2). 

T. Niiranen et al. / Ore Geology Reviews 30 (2007) 75–105 99

100 


Fig. 16. Concordia diagrams showing the conventional U–Pb age data on zircons and titanites, samples A1697 granite (a) and A1698 diorite (b) 

from the Laurinoja deposit. See Table 5 for the U–Pb data. Error ellipses are 2r. Concordia ages are calculated at 2r confidence level and decayconstant 

errors ignored. 

equal with the age of zircon and titanite from the A963 

clinopyroxene skarn (Fig. 17). Furthermore, the age 

from the titanite in A840 monzonite is also ca. 1800 

Ma. 

The A1698 diorite was also taken from the waste 

rock piles at Hannukainen. It is a typical distal altered 

diorite that occurs as a wall rock for all studied deposits 

displaying well developed S 3 foliation and L 3 lineation 

(e.g., Fig. 6b). Four fractions consisting of prismatic, 

colourless, transparent zircon were dated from the sample 

(Table 6). The data plot on a discordia line which 

intercepts the concordia curve at 1864F5 Ma and 

170F160 Ma (Fig. 16). Within the error limits, the 

A1698 diorite is contemporaneous with the A840 monzonite 

and A958 albitite (Fig. 17). 

A titanite fraction from the diorite gives a concordia 

age of 1805 F5 Ma(Table 6 and Fig. 16). The significant 

age difference with the zircon and titanite in the 

A1698 diorite is consistent with the age difference 

between these minerals in the A840 monzonite and 

A958 albitite (Fig. 17). The titanite ages thus must 

reflect the time of a later metamorphism. 

9. Discussion 

The data presented in this work suggest that the 

skarns and ironstones in the Kolari deposits overprint 

several different rock types, chiefly type-2 metavolcanic 

rocks (Fig. 13). This strongly suggests that the 

deposits are of metasomatic replacement type rather 

than metamorphosed syngenetic iron formations as 

was previously suggested (Mäkelä and Tammenmaa, 

1978; Frietsch et al., 1995). Although the calcic alteration 

assemblages in the Kolari deposits are quite similar 

to the ones described in association with calcic-iron 

skarns (sensu stricto), the zoning at Laurinoja and 

Kuervitikko is atypical for a classic skarn deposit. In 

the skarn deposits, the zoning is typically characterised 

Fig. 17. Summary of U–Pb ages from the Kolari deposits. A840, A958, A959, and A963 after Hiltunen (1982) and Väänänen and Lehtonen (2001).


by variations in proportions and compositions of pyroxenes 

and garnet (e.g., Einaudi et al., 1981). In the 

Kolari deposits, garnet-rich skarns are practically absent, 

and no clear variation can be seen in the colour or 

optical properties of clinopyroxene that would indicate 

significant variation in its composition. Instead, the 

zoning in Kolari deposits consists of moderately to 

intensely alkali-altered distal, and calcic-iron-altered 

proximal zones. The zoning is roughly symmetrical 

around the ironstones, although the distal alteration 

appears to be more extensive and intense in the hanging 

wall compared to footwall. The characteristics of the 

zoning, together with the intimate relationship of the 

deposits with thrust and shear zones suggest a hydrothermal 

system where the ironstones represent the structural 

base of the shear system and the zoning is 

controlled by fluid flux, fluid evolution, and fluidwall 

rock reactions in and around the fault and thrust 

zones. The lack of suitable structures that would focus 

the fluid flow could explain the poorly developed zoning 

in the Cu-Rautuvaara deposit. 

9.1. Temporal relationship between ironstones, Cu–Au 

mineralisation, and deformation stages 

Two generations of sulphides exist in the Kolari 

deposits, and both of these contain chalcopyrite. The 

majority of sulphides pre-date the brittle deformation. 

In the magnetite-disseminated and Cu–Au-mineralised 

albitites the textural features suggest that magnetite and 

sulphides were contemporaneous (Fig. 7). The gangue 

mineralogy is similar between the barren and sulphiderich 

parts of the ironstones. No systematic distinction 

can be made in chemical composition between the Cu– 

Au mineralised and unmineralised ironstone units, except 

for concentrations of sulphur and elements correlated 

with that. These facts suggest that precipitation of 

the magnetite and sulphides was contemporaneous. The 

S 3 foliation in the ironstones and the orientation of the 

ironstone lenses parallel to L 3 lineation at Laurinoja and 

Kuervitikko (Figs. 3 and 4), and the U–Pb age data (see 

below) suggests that the formation of ironstones and 

Cu–Au mineralisation were contemporaneous to the D 3 

thrusting event. The good correlation between Cu and 

Au in all deposits suggests that they most likely were 

precipitated in the same event. 

The youngest generation are related to brittle fractures 

that crosscut the S 3 foliation and therefore postdate 

D 3 being probably linked to D 4 stage (Figs. 5 and 

15b). However, it is unknown whether the sulphides in 

this stage are just remobilised varieties of the pre-existing 

ones, or if they are derived from a more distant 

source. Nevertheless, considering the low abundance of 

late fractures in the Kolari deposit, their contribution to 

the Cu–Au budget is low. 

9.2. Implications of the U–Pb age data 

The 1864F5 Ma U–Pb zircon age of the Hannukainen 

diorite is, within the error limits, the same as the U– 

Pb zircon age of the Hannukainen monzonite suggesting 

that they are coeval (Fig. 17). The age of the diorite 

also defines a maximum age for the ores. Based on the 

contact relations between the intrusives and the supracrustal 

rocks, Hiltunen (1982) proposed that monzonite 

and diorite intrusions intruded conformably into the 

supracrustal rocks during the early stages of the main 

folding. This would mean that they intruded during the 

D 2 stage of Väisänen (2002) and Hölttä et al. (in press). 

Thus, the D 2 stage in the Kolari region is 1865 Ma or 

older. The minimum age for the ores is constrained by 

the undeformed, 1766F5 Ma granite that brecciates the 

ore. Coeval undeformed granites occur throughout 

northern Finland (e.g., Hanski et al., 2001), and they 

also constrain the minimum age for the D 3 thrusting 

event (e.g., Väisänen, 2002). 

The 1797 F5 Ma U–Pb zircon age of the clinopyroxene 

skarn, which is ca. 65 Ma younger than the age 

of the hanging wall intrusions, makes the genetic model 

presented by Hiltunen (1982) dubious. A number of 

analyses of titanite display metamorphic ages that are 

similar to the zircon age of the skarn (Fig. 17). Furthermore, 

the inherited zircon from the granite that brecciates 

the ore is ca. 1800 Ma in age. Hiltunen (1982) does 

not give any explanation to age difference between the 

skarn and the monzonite intrusions, which is clearly too 

large if the clinopyroxene skarn is related to the hanging 

wall intrusions. The U–Pb age of 1797F5 Ma for 

the clinopyroxene skarn is constructed by only three 

bulk zircon fractions which actually form a two point 

discordia line (Hiltunen, 1982). The titanite is apparently 

slightly younger than the upper intercept age for 

the zircon fractions, and the U–Pb data for A840 monzonite 

plot on the bolder sideQ of the A963 data. Thus, 

although the skarn zircon can also contain some inherited 

zircon components, the transparent, brown, euhedral 

zircon most likely crystallized at the time of skarn 

formation. The age of the skarn cannot be younger than 

the 1780 Ma, titanite age and the age for the new zircon 

must be younger than the age of the monzonite especially 

when inherited, altered zircon also are mixed with 

the dated fractions. However, further ion microprobe 

dating would be required to thoroughly understand the 

ages and systematics of the skarn zircon. Regardless of

102 


all the problems dealing with the skarn zircon age, we 

believe that the skarn alteration in the Kolari region 

took place at ca. 1800 Ma, and that the titanite ages of 

the altered monzonite, diorite, and albitite reflect the 

age of the distal alteration that was contemporaneous 

with the proximal, skarn alteration. Hence, the formation 

of the ironstones and Cu–Au occurrences at Kolari 

most likely took place ca. 1800 Ma. This is coincident 

with the thermal event that led to the formation of ca. 

1800 to 1770 Ma granites that are abundant throughout 

northern Finland and northern Sweden. Furthermore, 

the intimate relation of the Kolari ores to the D 3 thrust 

zones suggests that the D 3 stage in the Kolari region 

took place around 1800 Ma, if our interpretation of the 

zircon age of the skarn is correct. This interpretation is 

supported by the U–Pb data from the Pajala Shear Zone 

in northern Sweden, which forms the southern extension 

to the Kolari Shear Zone, where ages between 

1810 and 1774 Ma were defined from metamorphic 

zircon in mica gneiss (Bergman and Skiöld, 1998). This 

also indicates that a thrust event corresponding to D 3 in 

Kolari region took place between 1810 and 1774 Ma. 

The 1748F7 Ma age of the bmafic pegmatoidQ is 

ambiguous. No intrusives with a similar age have been 

dated from northern Finland (cf. Hanski et al., 2001). If 

the pegmatoid vein is related to brittle fractures, the age 

of the D 4 deformation in the northern Finland is indeed 

significantly younger than previously considered as 

proposed by Väisänen (2002). 

9.3. Pressure and temperature considerations, 

evolution of the fluid 

The early type-1a H 2 OFCO 2 –haliteFsolids inclusions 

in quartz with abundant sulphides indicate conditions 

under which the enclosing host were formed. The 

fluids were trapped homogeneously (constant phase 

ratios) in a single-phase region of the H 2 O–CO 2 –halite 

system (Fig. 18). However, the original P–T conditions 

of entrapment cannot be reconstructed from fluid inclusion 

evidence alone. Nevertheless, the total homogenisation 

temperature (386 to 463 8C) and the total 

homogenisation pressure (ca. 1.5 kbar) serve as minimum 

constraints on the entrapment conditions. The 

pressure during the peak of the regional metamorphism 

in the Kolari region was between 3.3 and 3.8 kbar 

(Hölttä et al., in press). Since the D 3 stage post-dates 

the metamorphic peak, these pressures can be considered 

to represent the upper limits during the alteration 

and the Fe and Cu–Au mineralisation. Based on this, the 

pressure during the entrapment of the type-1a inclusions 

was between 1.5 and 3.5 kbar and the temperature was 

Fig. 18. Pressure–temperature diagram showing the location of the 

isochors for the relevant fluid inclusion types detected at Kolari. 

Solvus for the H 2 O–40 wt.% NaCl system with 5 mol.% CO 2 is 

from Schmidt and Bodnar (1994) and liquidus for the 30 and 40 wt.% 

NaCl are from Cline and Bodnar (1994) and Bodnar (1994). Pressure 

during the peak of the regional metamorphism after Hölttä et al. (in 

press). 

between 450 and 550 8C (Fig. 18). Subsequent cooling 

and brittle fracturing resulted in trapping of the type-1b 

inclusions. The co-existing type-1b and -3 inclusions 

suggest that during the evolution of the hydrothermal 

system at Laurinoja and Kuervitikko, a phase separation 

occurred between the aqueous and carbonic components 

of the fluid. This could have occurred either due to 

cooling, or due to pressure fluctuations or both as 

indicated by the considerable variations between the 

densities of the type-3 inclusions. 

9.4. IOCG mineralisation at Kolari? 

The ca. 65 Ma age difference between the monzonite 

and diorite intrusions and skarn indicates that the genesis 

of the Kolari deposits cannot be related to the 

intrusions as suggested by Hiltunen (1982), and therefore 

an alternative origin should be considered. The 

general features of the Kolari deposits display characteristics 

typical for IOCG occurrences that form a controversial 

spectrum of deposits into which several 

diverse FeFCu, Au, REE, U deposits have been included 

(Hitzman et al., 1992; Barton and Johnson, 

1996, 2000; Hitzman, 2000; Williams and Pollard, 

2003). Common characteristics for both the Kolari 

and the IOCG deposits elsewhere include (1) high 

bulk Fe–S ratio, (2) metal association Fe–Cu– 

AuFAg, Co, Mo, REE, U, (3) relatively low Cu and 

Au concentrations, (4) relation to highly saline aqueous


to carbonic brines, (5) distal Na FK, Cl and proximal 

Ca–FeFCO 2 , S alteration (e.g., Hitzman et al., 1992; 

Barton and Johnson, 2000; Perring et al., 2000; Pollard, 

2001; Oliver et al., 2004). Generally in IOCG occurrences, 

the proximal alteration is dominated by K- and 

Fe-rich phases, i.e., biotite, K-feldspar, sericite, and Feoxides 

(e.g., Ernest Henry; Mark et al., 2000). Nevertheless, 

in a number of occurrences (e.g., Candelaria, 

Mt Elliott; cf. Pollard, 2000) the Cu–Au mineralisations 

are hosted by magnetite-rich dskarnsT. Of the known 

IOCG occurrences, the Kolari deposits are perhaps 

most similar to the Mt Elliott deposit (Cloncurry district, 

Australia), where the Cu–Au mineralisation is 

hosted by clinopyroxene–magnetite skarn comparable 

to the one at Laurinoja and Kuervitikko (Wang and 

Williams, 2001). 

We can only speculate about the ultimate source of 

the metals enriched in the Kolari deposits. Since the 

general features of the Kolari deposits are so apparently 

similar with deposits classified into the IOCG type, we 

feel that it is reasonable to try to apply the IOCG 

genetic model. The IOCG models emphasise the role 

of distal sodic alteration for the source of the elements 

enriched in the ironstones, barren or Cu–Au mineralised 

(e.g., Williams, 1994; Barton and Johnson, 1996; 

Hitzman, 2000; Oliver et al., 2004). The true extent of 

the sodic alteration around the Kolari deposits is not 

known, but it probably covers several km 2 of the country 

rocks along the strike of the thrust zones. Also, the 

alteration can be vertically extensive considering that 

the deposits are located in major thrust structures. Thus, 

similar to IOCG models, the regionally albite-altered 

rocks could be the source for the bulk of the elements 

enriched in the Kolari deposits. It is, however, possible 

that part of the CaO and CO 2 in the proximal alteration 

zones are derived from a local source, i.e., from a 

carbonaceous unit, since such rocks are known to 

occur in the supracrustal sequence in which the deposits 

are located (Hiltunen, 1982). 

Recent work by Oliver et al. (2004) shows that Cu is 

not systematically stripped from rocks subjected to 

albite alteration, in contrast to Fe, K, Ba, Rb FCa, Sr, 

Co, V, Mn, Pb, and Zn that are more consistently 

leached out of the albitised rocks. They propose a 

model where Cu in the IOCG deposits is derived 

from the source of the brines, and precipitation of the 

Cu was a result of mixing of Cu-bearing brine that is 

evolved extensively via albitisation with sulphur bearing 

fluid, or reaction of the brine with sulphur-bearing 

country rocks. In their model the ultimate source for the 

albitising brines and Cu is a deep seated granitoid. In 

the cases where the fluid lacks Cu, or sulphur source is 

absent, or both, brines enriched in elements released by 

albitisation produce barren ironstones. Applying this 

model to the Kolari deposits could explain the large 

variation in the sulphur isotope compositions in them 

(Mäkelä and Tammenmaa, 1978; Hiltunen, 1982), i.e., 

that the sulphur in the Kolari deposits was derived from 

the Savukoski Group metavolcanic and sedimentary 

rocks, and also possibly from the abundant intrusives. 

Similarly, the availability of Cu and/or S could explain 

why some of the ironstones in the Kolari region are Cu 

mineralised and others are not. 

10. Conclusions 

Current data on the Kolari Fe–Cu–Au deposits show 

that they are hydrothermal replacement-type deposits 

and they overprint the ca. 1860 Ma Haparanda-type 

intrusives and the metavolcanic rocks and metasediments 

of the Savukoski Group. U–Pb age data presented 

in this work and by previous authors indicate 

that the ores were formed between 1864F5 Ma and 

1766F5 Ma, most likely at ca. 1800 Ma as suggested 

by the age of the clinopyroxene skarn at Hannukainen. 

The temperature during the main alteration and mineralisation 

event was between 450 and 550 8C and the 

pressure between 1.5 and 3.5 kbar. The key features of 

the Kolari deposits that should be considered in any 

discussion are: 

1. Metal association Fe–Cu–Au FAg, Bi, Ba, Co, Mo, 

Sb, Se, Te, Th, U, LREE. 

2. Alteration styles that, in deposits where structural 

control is most evident, define a zoning where the 

distal zones are characterised with NaFCl, K-rich 

assemblages, proximal zones with Ca–FeFCO 2 - 

rich assemblages, and the core that is probably related 

to a structural base, is iron rich. In deposits 

where the zoning is well developed, the Cu–Au 

mineralisation is located in the ironstone core whereas 

in deposits with no clear zoning Cu–Au mineralisation 

is in the magnetite-disseminated albitite. 

3. Highly saline aqueous-carbonic fluids that circulated 

in the system during the main alteration and mineralisation 

event and, subsequently, during a later 

brittle stage. 

4. The deposits are controlled by the structures of a 

major shear zone system rather than being bound to 

certain stratigraphical position. 

Based on the data presented in this work, we suggest 

that an IOCG mineralisation model would best explain 

the features of the Kolari Fe–Cu–Au deposits.

104 


Acknowledgements 

We are grateful to Dr. Pasi Eilu, Dr. Juhani Ojala, Dr. 

Raimo Lahtinen for the general assistance and constructive 

comments on the manuscript. We also wish to 

express our gratitude to the staff of the GSF isotope 

laboratory. Reviews by Dr. Hamid Mumin and Dr. Ping 

Dong are highly appreciated. Senior author wishes to 

thank the management of the Geological Survey of 

Finland in Espoo and Rovaniemi offices for responsiveness 

that allowed the finishing of this work. Outokumpu 

Oyj Foundation and Finnish Graduate School in Geology 

provided the funding for the senior author’s work. 

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