Fredrik Sahlström Stable Isotope Systematics of Skarn ...706799/FULLTEXT01.pdf · Stable Isotope...

Examensarbete vid Institutionen för geovetenskaper ISSN 1650-6553 Nr 279

Stable Isotope Systematics of Skarn-hosted REE-silicate -

Magnetite Mineralisations in Central Bergslagen, Sweden

Stable Isotope Systematics of Skarn-hosted REE-silicate - Magnetite Mineralisations in Central Bergslagen, Sweden

Fredrik Sahlström

Fredrik Sahlström

The metasupracrustal-hosted, often polymetallic REE-Fe-deposits of Bastnäs-type are found along the “REE-line” in the Palaeoproterozoic Bergslagen ore province, south central Sweden. They essentially comprise REE silicate-bearing magnetite skarn mineralisations with variable contents of other metals. Even though these deposits have been important for mining and research for centuries, their origin still remains unclear. In this study, samples from 10 different deposits along the REE-line have been charactarised as to mineralogy, petrography and bulk geochemistry, in addition to their isotope systematics. Mineral separates of magnetite and, when present, co-existing quartz or carbonates have been analysed for their oxygen and (for carbonates) carbon isotope compositions, in order to put constraints on the sources for metals and fluids in these deposits. Magnetites have δ18O-values of -1.79 to 1.12 ‰, while quartzes lie between 7.19 and 8.28 ‰. Carbonates have δ18O-values between 5.77 and 7.15 ‰ and δ13C-values between -5.35 and -3.32 ‰. Thermometric calculations based on mineral pairs (magnetite-quartz, magnetite-calcite/dolomite), combined with available fluid inclusion data, indicate formation of primary magnetite assemblages between c. 650 to 400 °C. At these temperatures, magnetites from some of the deposits would have been in equilibrium with a magmatic fluid (δ18O = 6-8 ‰), while magnetites from other deposits would have been in equilibrium with fluids of lower δ18O (4-6 ‰). Oxygen and carbon isotope trends in carbonates can be explained by interaction between original host carbonates and a fluid of magmatic composition. The combined results indicate that the Bastnäs-type magnetite-REE mineralisations were deposited from an originally magmatic fluid at relatively high temperatures. At local scale, variable modification of the fluid isotopic composition can be explained by mixing with seawater-dominated fluids.

Uppsala University, Department of Earth SciencesMaster Thesis, 30 hp Solid Earth GeologyISSN 1650-6553 Nr 279Printed by Geotryckeriet, Uppsala University, Uppsala, 2014.

Supervisors: Erik Jonsson & Karin Högdahl

Examensarbete vid Institutionen för geovetenskaper ISSN 1650-6553 Nr 279

Stable Isotope Systematics of Skarn-hosted REE-silicate -

Magnetite Mineralisations in Central Bergslagen, Sweden

Fredrik Sahlström

1

Stable isotope systematics of skarn-hosted REE-silicate -

magnetite mineralisations in central Bergslagen, Sweden

Table of Contents

Abstract ................................................................................................................................................... 3

Sammanfattning på svenska ................................................................................................................... 4

1. Introduction ......................................................................................................................................... 5

1.1. Rare earth elements ..................................................................................................................... 5

1.2. Purpose of study and hypothesis ................................................................................................. 6

2. Geological Background ........................................................................................................................ 7

2.1. Regional geology and mineralisations of the Bergslagen ore province ....................................... 7

2.2. Skarn deposits .............................................................................................................................. 9

2.3. Local geology of the “REE-line” .................................................................................................. 10

2.3.1. Genesis of Bastnäs-type deposits ........................................................................................ 11

3. Materials and methods ..................................................................................................................... 12

3.1. Sample preparation .................................................................................................................... 13

3.1.1. Preparation of polished sections and thin sections ............................................................ 13

3.1.2. Preparation of samples for stable isotope analysis ............................................................. 14

3.2. Analytical methods ..................................................................................................................... 15

3.2.1. Powder x-ray diffraction ...................................................................................................... 15

3.2.2. Bulk geochemical analysis ................................................................................................... 15

3.2.3. Optical microscopy .............................................................................................................. 16

3.2.4. SEM-EDS .............................................................................................................................. 16

3.2.5. Electron probe microanalyser (EPMA) ................................................................................ 17

3.2.6. Stable isotope analysis ........................................................................................................ 19

4. Results ............................................................................................................................................... 22

4.1. Bulk geochemistry of REE-mineralised assemblages ................................................................. 22

4.2. Mineralogy and mineral chemistry ............................................................................................ 24

4.2.1. Östra Gyttorpsgruvan .......................................................................................................... 24

4.2.2. Johannagruvan .................................................................................................................... 32

4.2.3. Högforsfältet........................................................................................................................ 36

4.2.4. Danielsgruvan ...................................................................................................................... 37

4.2.5. Myrbacksfältet ..................................................................................................................... 38

2

4.2.6. Södra Hackspiksgruvan ........................................................................................................ 39

4.2.7. Östanmossagruvan .............................................................................................................. 41

4.2.8. Bastnäsfältet ........................................................................................................................ 44

4.2.9. Malmkärragruvan ................................................................................................................ 46

4.2.10. Rödbergsgruvan ................................................................................................................ 48

4.3. Chemical dating of uraninites ..................................................................................................... 49

4.4. Results stable isotopes ............................................................................................................... 50

5. Discussion .......................................................................................................................................... 52

5.1. Mineralogy, mineral chemistry and bulk geochemistry ............................................................. 52

5.2. Uraninite geochronology ............................................................................................................ 53

5.3. Stable isotope systematics ......................................................................................................... 54

5.3.1. Thermometry....................................................................................................................... 54

5.3.2. Fluid modeling ..................................................................................................................... 56

5.3.3. Effects of metamorphism .................................................................................................... 61

5.4. Fluid chemistry and ore mineralogy ........................................................................................... 61

6. Conclusions ........................................................................................................................................ 63

7. Acknowledgements ........................................................................................................................... 64

8. References ......................................................................................................................................... 65

9. Appendix ............................................................................................................................................ 72

3

Abstract

The metasupracrustal-hosted, often polymetallic REE-Fe-deposits of Bastnäs-type are found

along the “REE-line” in the Palaeoproterozoic Bergslagen ore province, south central Sweden.

They essentially comprise REE silicate-bearing magnetite skarn mineralisations with variable

contents of other metals. Even though these deposits have been important for mining and

research for centuries, their origin still remains unclear. In this study, samples from 10 different

deposits along the REE-line have been charactarised as to mineralogy, petrography and bulk

geochemistry, in addition to their isotope systematics. Mineral separates of magnetite and, when

present, co-existing quartz or carbonates have been analysed for their oxygen and (for

carbonates) carbon isotope compositions, in order to put constraints on the sources for metals

and fluids in these deposits. Magnetites have δ18O-values of -1.79 to 1.12 ‰, while quartzes lie

between 7.19 and 8.28 ‰. Carbonates have δ18O-values between 5.77 and 7.15 ‰ and δ13C-

values between -5.35 and -3.32 ‰. Thermometric calculations based on mineral pairs

(magnetite-quartz, magnetite-calcite/dolomite), combined with available fluid inclusion data,

indicate formation of primary magnetite assemblages between c. 650 to 400 °C. At these

temperatures, magnetites from some of the deposits would have been in equilibrium with a

magmatic fluid (δ18O = 6-8 ‰), while magnetites from other deposits would have been in

equilibrium with fluids of lower δ18O (4-6 ‰). Oxygen and carbon isotope trends in carbonates

can be explained by interaction between original host carbonates and a fluid of magmatic

composition. The combined results indicate that the Bastnäs-type magnetite-REE

mineralisations were deposited from an originally magmatic fluid at relatively high

temperatures. At local scale, variable modification of the fluid isotopic composition can be

explained by mixing with seawater-dominated fluids.

4

Sammanfattning på svenska

Polymetalliska REE-Fe-fyndigheter av Bastnästyp förekommer längs den så kallade ”REE-

linjen” i en metasuprakrustal enhet i den paleoproterozoiska malmprovinsen Bergslagen i

Mellansverige. De består av REE-förande silikater i magnetitskarnlager i marmor. Trots att

dessa fyndigheter har varit ekonomiskt betydande samt attraherat forskare i århundraden,

kvarstår många frågetecken beträffande deras bildningssätt. I denna studie har prov ifrån 10

olika fyndigheter längs REE-linjen karakteriserats med avseende på mineralogi, petrografi samt

geokemi. Mineralseparat av magnetit tillsammans med kvarts eller karbonat har analyserats för

syre- och (för karbonater) kolisotopsammansättning, för att få ytterligare information om

bildningssättet för denna malmtyp. Magnetiterna uppvisar δ18O-värden mellan -1,79 och 1,12

‰, medan kvarts ligger mellan 7,19 och 8,28 ‰. Karbonater har δ18O-värden mellan 5,77 och

7,15 ‰ samt δ13C-värden mellan -5,35 och -3,32 ‰. Termometriska beräkningar baserade på

mineralpar (magnetit-kvarts, magnetit-kalcit/dolomit), kombinerat med tillgänglig

vätskeinneslutningsdata, tyder på att primära magnetitassociationer bildats vid mellan ca. 650

och 400 °C. Vid dessa temperaturer är magnetit från några av fyndigheterna i jämvikt med

magmatiska fluider (δ18O = 6-8 ‰), medan magnetiter från andra fyndigheter är i jämvikt med

fluider med lägre δ18O (4-6 ‰). Trender för kol- och syreisotoper i karbonater kan förklaras

genom reaktioner mellan ursprungliga värdkarbonater och en fluid med en magmatisk

isotopsammansättning. Resultaten tyder på att malmer av Bastnäs-typ bildades från en primärt

magmatisk fluid vid relativt höga temperaturer. Lokalt har fluidens isotopsammansättning blivit

påverkad i varierande omfattning genom inblandning av havsvatten.

5

1. Introduction

1.1. Rare earth elements

In recent years, the rare earth elements (REE) have quickly become one of the more increasingly

sought-after natural resources in the world. The REEs comprise a group of elements in the

periodic table known as the lanthanides. Additionally, yttrium and sometimes scandium are

added to the group due to their chemical similarities (Fig 1). These elements are extensively

used in a number of different industries, including the booming “high tech” and “green tech”

sectors, leading to an increasing demand for these commodities (e.g. Haxel et al. 2002). Despite

their name, these metals are actually not very

rare. The least abundant REE, thulium and

lutetium, still exhibit average crustal

abundances of over 200 times that of gold

(Haxel et al. 2002). The main issue regarding

the REE is the fact that there are relatively few

geological processes that can concentrate them

into deposits of economic value. Moreover, the

heavy REE (HREE: Gd to Lu, and Y) and Eu

are even more scarce due to their more

compatible nature, and therefore deposits rich in these elements are especially valuable (Haxel

et al. 2002). Another important factor is the current global political climate. Since the closing

of the famous Mountain Pass mine in California in 1984, the REE-market has been dominated

by China, lately controlling over 95 % of the global REE-supply. Therefore, recent decreases

in the Chinese REE export quotas have made the future global supply of REE very uncertain

(Moffett & Palmer 2012). This has led to a boost in exploration and research on new REE

deposits worldwide, and the former so important Mountain Pass mine is being re-opened

(Wiens 2012, Molycorp). In Sweden, this newfound interest in REE and other strategic metals

is indeed very noticeable, with active exploration and planned mining of the Norra Kärr Zr-

REE deposit (Jonsson 2013, Nebocat 2009) as well as increasing research on other types of

potential ore deposits (e.g. this study, Högdahl et al. 2012, Sahlström 2012).

Fig 1. Periodic table with the rare earth elements

marked. Figure taken and modified from

geology.com.

6

1.2. Purpose of study and hypothesis

The Bastnäs-type Fe-REE (locally Fe-REE-Cu-(Co-Au-Bi-Mo)) skarn deposits all occur within

a belt of mostly significantly altered felsic metavolcanic rocks in the Nora-Riddarhyttan-

Norberg area, also called the “REE-line” (Jonsson & Högdahl 2013). This is located in the

western central part of the Bergslagen ore province, south central Sweden. These deposits,

among others, have been important for local mining of iron for centuries, and for REE

sporadically since the late 1800’s. They are also known for being the location of several

discoveries of both new minerals and elements (see Andersson et al. 2004, and references

therein). The most famous discovery is probably that of the element cerium, which was first

described by Wilhelm Hisinger and Jacob Berzelius in 1804. The element was found as a major

constituent of a reddish heavy mineral called “Bastnäs tungsten” (later formally named cerite).

After extensive chemical analyses, Hisinger and Berzelius were able to characterise and give a

detailed description of the new element. The name “cerium” was inspired by the discovery of

the asteroid Ceres three years earlier (Trofast 1996).

Despite their historic significance in both mining and science, little is still known about the

genesis of these deposits. Various theories have been proposed, and the consensus today is that

these ores are hosted by skarns formed during epigenetic reactions between a hydrothermal

mineralising silica and metal-bearing fluid and pre-existing carbonate rocks (Andersson et al.

2004, and references therein). However, the specific nature as well as origin of such a fluid are

yet to be determined. This project addresses the fundamental question of whether these ores

were formed from high, medium or low temperature fluids, and also try to shed light on the

possible source(s) of these fluids. This is primarily tested using oxygen isotope analyses of

magnetites together with co-existing quartz or carbonates, separated from mineralogically and

petrographically well-characterised REE-rich assemblages. The hypotheses tested are 1)

whether these magnetites could have formed in equilibrium with a water-dominated

hydrothermal fluid and if so at what temperature, and 2) whether this fluid had a magmatic,

meteoric, seawater or mixed origin. Additionally, carbon isotope analyses of associated

carbonate assemblages from a selection of deposits were performed in order to better constrain

the processes responsible for the formation of these deposits. The mineralogical and

petrographical characterisation of the assemblages selected for isotope analyses was conducted

with a special focus on deposits lacking modern studies. This study was done within the

framework of projects "Securing society's supply of hi-tech elements by recycling mine dumps

and probing the metallogenic setting of the host systems", on rare and high-tech metals in

7

central Swedish mines and dumps, financed by Vetenskapsrådet and the Geological Survey of

Sweden (SGU).

2. Geological Background

2.1. Regional geology and mineralisations of the Bergslagen ore province

Bergslagen is one of the largest and most important ore provinces in Sweden, and its rich

mining history dates back to medieval times and beyond (cf. Tegengren 1924). With over 8500

documented mines and prospects the Bergslagen province boasts a wide variety of different ore

types. These include stratiform and stratabound Zn-Pb-Cu-(Ag-Au) sulphide deposits, iron

mineralisations of banded iron formation (BIF)-, skarn- and Kiruna types as well as granite-

hosted Mo-deposits (Allen et al. 2008, Allen et al. 1996, Ripa & Kübler 2003). Iron has

historically been the most important metal in the region (Sidenvall 1942), with 32 major

deposits producing over 420 Mt of ore up until the 1990’s (Åkerman 1994). The nine major

sulphide deposits have accounted for over 70 Mt of ore, and additionally a number of smaller

scale operations are scattered all over the Bergslagen province, where everything from base

metals to precious metals, phosphorous and REE have been targeted (Tegengren 1924, Geijer

& Magnusson 1944, Åkerman 1994, Back 1981, Andersson et al. 2004).

The Bergslagen province is situated in the south-western part of the 1.9-1.8 Ga Svecofennian

domain in the Fennoscandian Shield (Fig 2; e.g Stephens et al. 2009). It is a part of a larger,

medium to high grade metamorphosed felsic magmatic region (Allen et al. 1996). The host-

rocks to most of the mineralisations comprise mainly of a metavolcano-sedimentary succession,

which was previously referred to as the “leptite-hälleflinta formation” (Magnusson 1970). The

volcanic and subvolcanic rocks are predominantly of rhyolitic to dacitic composition, and are

interpreted to have been deposited in a mostly shallow submarine, continental back-arc setting

(Allen et al. 1996, Lundström 1987, Ripa 2001, Vivallo & Rickard 1984). Additionally, at

various stratigraphic levels in the volcanic succession subordinate rocks include intermediate

to mafic volcanic rocks and chemical, epiclastic and biogenic sedimentary rocks (Ripa 2001).

On a larger scale, the metavolcanic succession and associated rocks are interlayered beween

two metasedimentary successions of clastic facies, of which the lower contact is exposed on the

island of Utö (Allen et al. 1996). The basement underlying the Svecofennian metasupracrustal

rocks is unknown. However, locally occurring quartzites containing 2.7-1.95 Ga old detrital

zircons as well as inherited zircons in metavolcanic rocks suggest a continental-type basement

8

which has subsequently been eroded and reworked (Lundqvist 1987, Allen et al. 1996,

Andersson et al. 2006). Deformation and metamorphism associated with the c. 1.87-1.80 Ga

Svecokarelian orogeny affected the older rocks to variable degrees, and extensive granitic

intrusions were emplaced pre-, syn- and post-tectonically (Wilson et al. 1984, Allen et al. 1996,

Stephens et al. 2009). Additionally, the western parts of the region have locally been

overprinted by the Sveconorwegian orogeny at 1.0 Ga (e.g. Stephens et al. 2009). The

Svecokarelian deformation is most obvious in the supracrustal rocks where it appears as steep,

tight to isoclinal folding (Allen et al. 1996).

Fig 2. Geological map (SGU Ba-58) over the Bergslagen province. The REE-line is marked with a black

rectangle. Figure modified from Stephens et al. (2009).

9

The Bergslagen province has gone through a magmatic-extensional-compressional cycle,

where an initial stage of intense magmatism, thermal doming and crustal extension was

followed by a stage of waning extension, thermal subsidence, reversal from extensional to

compressional deformation, metamorphism and structural inversion (Allen et al. 1996). This

evolution can be observed in the rock succession, where lower deep water sedimentary rocks

(e.g. turbidites) are overlain by shallow marine to subaerial volcanic rocks, representing the

first stage of the cycle. Continuing upwards in the successions, shallow marine rocks (e.g.

volcanic sand- to siltstones and carbonates) are found, followed by deep water sedimentary

rocks, representing the second stage of the cycle (e.g. Stephens et al. 2009 and references

therein). Additionally, second-order stratigraphic variations due to subregional differences in

uplift and subsidence together with variations in the evolution of individual volcanoes are

superimposed on this cycle (Allen et al. 1996). Extensive hydrothermal alteration of the

metavolcanic sequence is commonly seen, with metasomatic enrichment in K, Na and Mg

having occurred in partly successive pulses and areas (Trägårdh 1991, Stephens et al. 2009).

Most mineralisations are hosted by hydrothermally altered metavolcanic rocks and associated

metalimestones and/or skarns. They are, with the exception of the apatite-iron oxide ores, found

mainly in the upper parts of the volcanic succession in medial to distal facies, connecting them

to the second evolutionary stage with waning volcanism and deposition in a subaqueous

environment (Allen et al. 1996). Moreover, a genetic link between the ores and specific

magmatic-hydrothermal systems from individual volcanoes or volcanic complexes is possible

(Allen et al. 1996, Högdahl & Jonsson 2004, Jonsson 2004).

2.2. Skarn deposits

The term “skarn” originates from Bergslagen, and was first coined by Swedish geologist Alfred

Elis Törnebohm in 1875. Skarns can be divided into two main types – reaction skarns and

contact skarns. Reaction skarns are small scale (millimeter to meter) features, formed more or

less in-situ owing to metasomatic transport of components between adjacent lithologies during

high-grade metamorphism (e.g. Burt 1977 and references therein). Contact skarns on the other

hand are generally attributed to magmatic-hydrothermal activity related to dioritic to granitic

plutonism in orogenic belts, and are distinguished by having a very characteristic gangue

assemblage of coarse-grained mixtures of Ca-Mg-Fe-Al silicates (Einaudi & Burt, 1982). They

are often classified according to the rock type they replace, with endoskarns being replacements

10

of intrusive rocks and exoskarns being replacements of carbonate rocks. Contact skarns can be

further classified based on their dominant mineralogy (which is dependant on the composition

of the rock being replaced) and the dominant economic metal (Fe, W, Cu, Zn-Pb or Sn; Einaudi

& Burt 1982).

2.3. Local geology of the “REE-line”

The “REE-line” is situated in the western central part

of the Bergslagen region (Figs 2,3). The deposits along

this line are found in a narrow northeast-trending belt

of felsic metavolcanic rocks within the Svecofennian

succession (Holtstam & Andersson 2007, Jonsson &

Högdahl 2013; Fig 3). In this area a very prominent

Mg-alteration has overprinted the country rocks, which

were subsequently metamorphosed under amphibolite-

facies conditions (Trägårdh 1988, Geijer 1920). The

carbonate-bearing metavolcanic belt is surrounded by

granitoids of two generations, the older generation (1.9-

1.85 Ga) is an often deformed and metamorphosed

suite ranging from gabbro to tonalite-granodiorite-

granite, and a younger generation (1.85-1.75 Ga) of

mainly undeformed granites and associated migmatites

(Andersson et al. 2004, Holtstam & Andersson 2007).

The REE-silicate bearing mineralisations in the area,

known as Bastnäs-type deposits, generally occur as

seemingly epigenetic, massive to disseminated

magnetite-skarn replacements in mostly dolomitic

marbles (e.g. Geijer 1920, Geijer 1961, Holtstam

2004), hence there is a direct link between magnetite formation and that of REE mineralisation.

Some deposits are, however, hosted mainly by felsic metavolcanic rocks (e.g. Östra Gyttorp;

Nordenström 1890, Tegengren 1924). The ore mineral assemblages found include various

combinations of Fe-oxides, REE-silicates, REE-fluorocarbonates, sulphides of Cu, Co, Bi, Mo

and minor native Au, Ag (the latter specifically in the Bastnäs field; e.g. Tegengren 1924, Geijer

Fig 3. Geological map of the ”REE-

line”. Modified from Jonsson &

Högdahl (2013) and references

therein.

11

1920, Holtstam 2004). These mineral assemblages typically occur associated with skarn

minerals such as actinolite, tremolite, diopside and fluorite as well as often minor carbonates

(Geijer 1936). Based on slight local differences in chemistry and mineralogy of the deposits,

Holtstam & Andersson (2002) suggested a subdivision of Bastnäs-type deposits into two

subtypes. Subtype 1-deposits comprise the Bastnäs and Rödbergsgruvan areas where the

mineralisations are mainly enriched in LREE and Fe. In subtype 2-deposits, which are found in

the Norberg area (e.g. Östanmossa, Malmkärragruvan, Södra Hackspiksgruvan,

Johannagruvan), the ores are enriched in both LREE and HREE+Y together with Mg, Ca and

F (Holtstam & Andersson 2007). Mining in Nora-Riddarhyttan-Norberg has been on-going for

many centuries. Tens to hundreds of now abandoned mines, occurring scattered around the area,

have mainly been targeted for iron, base metals and, in limited extent, REE (Geijer 1936).

2.3.1. Genesis of Bastnäs-type deposits

The first theories regarding the genesis of the Bastnäs-type ores were born almost a century ago

by the Swedish geologist Per Geijer. He proposed that ores formed by metasomatic replacement

of pre-existing carbonate rocks by high temperature fluids (Geijer 1920), an idea that is widely

accepted today (cf. Andersson et al. 2004 and references therein). However, the exact nature of

the ore forming fluid is still uncertain as well as its sources. Additionally, the timing of

mineralisation has been debated. Geijer (1931) noticed the close spatial relation to the heavily

Mg-altered country rocks, as well as the high magnesium contents of minerals such as allanites

in some of the deposits. He concluded that widespread Mg-metasomatism occurred during

emplacement of the 1.9 Ga synorogenic granitoids, a common theory at that time, and that this

process was genetically related to ore formation (e.g. Geijer 1961). This type of metasomatism

has later been suggested to be primarily driven by seawater-dominated fluids (Trägårdh 1988,

Trägårdh 1991, Ripa 1994). More recent studies point towards a syn-volcanic magmatic-

hydrothermal origin of the ores (Holtstam & Andersson 2007, and references therein), thereby

discarding Geijer’s model. Fluid inclusion studies in bastnäsite indicate that this mineral was

deposited from a CO2-rich, highly to moderately saline fluid at minimum (not pressure

corrected) temperatures between 400 and 300 °C (Holtstam & Broman 2002, Andersson et al.

2013). However, bastnäsite mainly formed later than cerite-(Ce) and ferriallanite-(Ce),

suggesting that the initial temperatures of the ore forming fluid were probably higher

(Andersson et al. 2004). Fluorite were in turn deposited from a low to medium saline fluid,

during cooling, from 150 to 100 °C (Uncorrected; Holtstam & Broman 2002, Andersson et al.

12

2013). Furthermore, carbon and oxygen isotope studies of carbonate minerals associated with

REE mineralisation from type 2-deposits show values indicating a magmatic source of the ore

forming fluids (Andersson et al. 2013). This is also suggested by the Cu-Mo-Bi-W-Be-F

association in some of the deposits, which further indicates linkage to granitic magmatism

(Andersson et al. 2004). Re-Os geochronology on molybdenite, which is paragenetically related

to REE mineralisation, give ages of 1.90-1.84 Ga, with the youngest ages in the SW

(Rödbergsgruvan) and oldest ages in the NE (Norberg area) (Andersson et al. 2013). No

explanation has so far been given to explain this extended time interval fully. Textural and

structural interpretations of the extensively folded and recrystallised REE assemblages in the

Högfors BIF deposit (north of Bastnäs) suggest that the REE mineralisation was formed during

the Svecofennian syn-volcanic magmatic stage, thus pre-dating the later extensive polyphase

deformation and peak metamorphism of the Svecokarelian orogeny (Jonsson & Högdahl 2013).

3. Materials and methods

3.1. Sample preparation

Samples from 10 different localities in the Nora-Riddarhyttan-Norberg area were

petrographically and mineralogically characterised (Table 1). They were also analysed for

major and trace element geochemistry as well as for their oxygen and carbon isotope

composition. The samples are from Danielsgruvan, Östanmossagruvan, Malmkärragruvan,

Södra Hackspiksgruvan, Johannagruvan, Bastnäsfältet, Högforsfältet, Myrbacksfältet,

Rödbergsgruvan and Östra Gyttorpsgruvan (Fig 3). They were generously provided by the

Geological Survey of Sweden (SGU), the Swedish Museum of Natural History (NRM), Prof.

Erik Jonsson and Doc. Karin Högdahl.

13

Table 1. Description of samples used in this study including name of deposit, ore type, sample coding and analyses

performed on each sample. Abbreviations used are: BGC – bulk geochemical analyses, ISO – stable isotope

analyses, XRD – powder X-ray diffraction, OM – optical microscopy, SEM-EDS – scanning electron microscopy

with energy dispersive spectrometer, WDS – electron probe microanalyser with wavelength dispersive

spectrometer.

Deposit Type Sample ID Thin/thick section label Analyses

Danielsgruvan Banded mgt-qtz ore NRM-20020125 Danielsgr. Ö BGC, ISO, OM

Danielsgruvan Banded mgt-qtz ore NRM-20020124 Danielsgr. BGC, ISO, OM

Östanmossa Carbonate skarn mgt ore EJ-ÖM90-13-1 Östanmossa EJ-ÖM90-13-1 BGC, ISO, XRD, OM, SEM-EDS

Östanmossa Carbonate skarn mgt ore SGU-M4528 Östanmossa M4528 BGC, ISO, XRD, OM, SEM-EDS Östanmossa Carbonate skarn mgt ore EJ-ÖM90-13-2 Östanmossa EJ-ÖM90-13-2 BGC, ISO, XRD, OM

Östanmossa Carbonate skarn ore SGU-M4441 Östanmossa M4441 BGC, ISO, XRD, OM

Östanmossa Carbonate skarn ore SGU-M4529 - ISO, XRD

S. Hackspiksgruvan Fluorite skarn mgt ore SGU-M3563 S. Hackspiksgr. M3563 BGC, ISO, OM, SEM-EDS

Johannagruvan Skarn mgt ore Joha-KH-1 Johanna BGC, ISO, OM, SEM-EDS, WDS Johannagruvan Skarn mgt ore Joha-KH-2 Johannagruvan BGC, ISO, OM

Malmkärra Carbonate skarn mgt ore SGU-M4068 Malmkärragr. M4068 (*2) BGC, ISO, XRD, OM, SEM-EDS

Malmkärra Carbonate skarn mgt ore SGU-M4048 Malmkärragr. M4048 BGC, ISO, XRD, OM, SEM-EDS

Högfors BIF Högf-KH-1 Högfors BGC, ISO, OM

Högfors BIF Högf-KH-2 - ISO

Bastnäs Skarn mgt ore Bast-EJ-1 Ceritgruvan BGC, ISO, OM

Bastnäs Massive mgt ore SGU-M6777 Bastnäs M6777 ISO, OM

Bastnäs Skarn mgt ore SGU-M309 N. Bastnäs M309 BGC, ISO, OM, SEM-EDS

Myrbacksfältet Massive mgt/sulphide ore Myrb-EJ-1 Myrbacksfältet BGC, ISO, OM

Myrbacksfältet Massive mgt/sulphide ore Myrb-EJ-2 - ISO

Östra Gyttorp Massive mgt ore Gytt-EJ-1a Unmarked (polished rock piece) ISO, OM, SEM-EDS, WDS

Östra Gyttorp Massive mgt ore Gytt-EJ-1b Unmarked (polished rock piece) ISO, OM Östra Gyttorp Massive mgt ore Gytt-EJ-1c - ISO

Östra Gyttorp Skarn mgt ore Gytt-EJ-2 Gyttorp BGC, OM, SEM-EDS, WDS

Rödbergsgruvan Massive mgt/skarn ore NRM-880071 Rödbergsgr. 880071 BGC, ISO, OM, SEM-EDS Rödbergsgruvan Massive mgt/skarn ore NRM-19984100 - BGC, ISO

3.1. Sample preparation

3.1.1. Preparation of polished sections and thin sections

Polished (thick) sections were prepared at the Department of Earth Sciences, Uppsala

University (UU). First, hand specimens were cut into small pieces using a diamond saw,

thereafter the samples were ground in order to get a flat surface before polishing. For this a

silicon carbide powder was mixed with water onto a glass plate, on which the samples were

ground. This was done in four steps with finer powder used successively (80 µm-45 µm-18 µm-

12 µm). Finally, the samples were polished using a polishing machine in order to get a perfectly

flat surface for microscopy studies. On the machine a cloth was placed, onto which fine

diamond paste and water were mixed. The samples were polished in three successive steps,

14

with increasingly finer paste used at each one (6 µm-3 µm-1 µm). Between each step of grinding

and polishing, the samples were thoroughly cleaned using an ultrasonic bath in order to

minimise the risk of contamination of coarser polishing agents that could cause scratches during

the following polishing steps.

Additionally, a number of samples were prepared as thin sections. Samples were cut out into

thin pieces using the diamond rock saw. The samples were then sent for thin section preparation

at Minoprep, Sweden.

3.1.2. Preparation of samples for stable isotope analysis

The samples for stable isotope analysis were prepared at the Department of Earth Sciences, UU.

Based on the mineral paragenesis of the hand specimen, one or more of the following minerals

were separated: magnetite, quartz and carbonates. Small pieces of rock were cut out from the

hand specimens using a diamond saw. Care was taken to remove all weathered surfaces. The

pieces were then crushed using a hammer. Samples containing large amounts of gangue

minerals, such as silicates, needed occasionally to be ground in a small mortar to reduce the

grain size enough to get clean crystals for separation. The samples were then cleaned thoroughly

using an ultrasonic bath, in order to reduce dust and contaminants, and then dried in an oven at

100 °C for 24 hours. The final step was using a magnet to separate the magnetite from the

gangue material, which was done several times until no impurities were visible. For carbonates

and quartz, clean crystals or crystal fragments were handpicked using tweezers. The final

mineral concentrate was controlled under a stereo microscope, were any remaining contaminant

minerals were picked out using a pair of non-magnetic tweezers. For each sample, around 10

mg of crystal grains were separated for final analysis.

15

3.2. Analytical methods

3.2.1. Powder x-ray diffraction

In order to measure the proportions of calcite and dolomite in carbonates analysed for isotopic

composition, some of the mineral separates were sent for analysis with powder x-ray diffraction

(XRD) technique. The analyses were conducted using a PANalytical X’pert PRO automated

diffractometer at the Swedish Museum of Natural History in Stockholm. The samples were

ground to a fine powder to allow the crystals to orientate in all possible orientations, and then

placed in a silicon sample holder. The samples were then bombarded by x-ray radiation, which

will diffract only in directions with constructive interference, determined by Bragg’s law:

n*λ=2*d*sinθ. Using a goniometer and a detector, the intensity and angle of the diffracted

radiation can be measured. By applying Bragg’s law, it is possible to determine the lattice

spacings, d, and in turn the Miller indices of lattice planes of the crystals. Furthermore, the

intensity of the peaks are related to the position of a specific atom in the unit cell, as atoms with

dense electron clouds diffract the beam more efficiently than atoms with less dense electron

clouds. The result can then be matched with available x-ray diffractogram databases, and

thereby identifying the mineral(s) in the analysed sample (Putnis 1992, Borchardt-Ott 2011).

Analyses were performed using an acceleration voltage of 45 kV and a beam current of 40 mA.

2θ were measured in the interval 5-70° for 11 minutes. Mineral identification and quantification

were done using the Highscore Plus software.

3.2.2. Bulk geochemical analysis

Bulk rock samples from all localities were analysed for major and trace elements at the ALS

Labs, Canada. Major elements were analysed using X-Ray Fluorescence (XRF) technique,

while trace elements and REE were analysed with combined Inductively Coupled Plasma

Atomic Emission Spectroscopy (ICP-AES) and Inductively Coupled Plasma Mass

Spectrometry (ICP-MS) technique. In the XRF technique, the sample is excited by high energy

x-rays. The excited atoms emit secondary fluorescent x-rays as they return to their ground state,

with energies characteristic for each element. The concentrations of each element in the sample

can thereby be quantified (Norrish & Chappell 1977). The ICP-AES technique is very similar,

where a plasma source is used to dissociate the sample into charged atoms or ions, thereby

exciting them to a higher energy level. As they return to their ground state, x-rays characteristic

16

of each element will be emitted, allowing quantification of elemental concentrations in the

sample (Philips 2008). In ICP-MS analysis, the sample is ionised with an inductively coupled

plasma. The ions are separated and quantified using a mass spectrometer (Beauchemin 2008).

For both ICP-MS and ICP-AES analyses the sample must be dissolved, and for the studied

samples, an acid digestion method was used (ALS 2012).

3.2.3. Optical microscopy

Optical microscopy was done at the Department of Earth Sciences, UU, using a Nikon Eclipse

E600 POL microscope with an attached camera/computer system, which enables high quality

photomicrography. Thin sections were studied using transmitted and reflected light, whereas

the polished (thick) sections were studied using reflected light.

3.2.4. SEM-EDS

For reconnaissance investigations on the mineralogy of the samples, SEM-EDS analysis was

conducted at the Evolutionary Biology Centre, UU. The Scanning Electron Microscope (SEM)

used was a Carl Zeiss Supra 35-VP Field Emission SEM equipped with a Robinson backscatter

detector, a variable pressure secondary electron (VPSE) detector and an EDAX Apex-4 energy

dispersive spectrometer (EDS). In order to perform SEM-EDS analyses, the samples need to be

electrically conductive. This was done by coating the samples with a thin film of carbon by

means of sputtering at the Department of Earth Sciences, UU.

The SEM uses a Field Emission Source (FES) to generate a finely focused, high energy electron

beam, which is swept across the sample surface in a raster pattern. As the electron beam hits

the sample surface, it interacts with the sample to produce a variety of different signals, such as

secondary electrons, backscattered electrons, Auger electrons and x-rays. These signals depend

on various characteristic properties of the sample, e.g. chemical composition, crystallographic

properties and surface topography (Goldstein et al. 1975). In this study backscattered electrons

were used for imaging. As the electrons from the beam hits the sample surface, they will be

attracted by the positively charged nuclei of the elements in the sample. They will be pulled

towards the nuclei and circle it, emitting out of the sample again in the opposite direction where

they are collected and registered by a BSE detector. Minerals containing heavy elements will

backscatter electrons more effectively, as their nuclei have a higher positive charge and thereby

17

attracts more electrons. Heavy element-rich minerals will yield brighter images, making it

possible to distinguish between different mineral phases as well as compositional variations

(e.g. zoning) within a single mineral (Goldstein et al. 1981). The SEM also has a camera

function which can produce high resolution BSE or SE photomicrographs of the studied

samples. The analyses were performed using an acceleration voltage of 20 kV, a beam current

of 40 nA and a beam diameter of 3 µm.

Also present in the SEM is an EDS detector, allowing qualitative/semi-quantitative chemical

spot analysis of various mineral phases to be obtained e.g. for identification purposes. When

the electron beam strikes the sample, thereby exciting it, each element in the mineral will

produce characteristic x-rays of a specific energy. By measuring the energies of the x-rays from

the various elements simultaneously, a spectrum displaying the approximate elemental

abundances of the analysed mineral can be generated (Goldstein et al. 1975).

3.2.5. Electron probe microanalyser (EPMA)

For quantitative mineral analyses and chemical dating, EPMA analysis was done at the National

Electron Microprobe Laboratory, CEMPEG, Department of Earth Sciences, UU, using a Jeol

JXA-8530F Field Emission Electron Probe Microanalyser (FE-EPMA). The technique is very

similar to the SEM – an electron beam interacts with the sample to produce the signals and

generate an image. This instrument is also equipped with an EDS detector for rapid mineral

identification and a BSE detector was used for orientation and identification of the various

minerals. The EPMA normally has several wavelength dispersive spectrometers (WDS)

installed. The WDS spectrometers differs from the EDS in that they measure the wavelengths

of the generated x-rays, which - similarly to the energies - are characteristic for each element.

The spectrometer configuration requires each wavelength to be measured independently, and

by comparing the intensity to a known standard the exact concentration of each element in the

sample can be determined. This makes this technique much more precise and gives higher mass

resolution compared to EDS, and can therefore be used to yield good quantitative measurements

of the elemental compositions of different minerals (Goldstein et al. 1981). The EPMA also has

a camera function similar to the one in the SEM. The analyses were done using an acceleration

voltage of 20 kV, a beam current of 40 nA and a beam diameter of 1-10 µm. Counting times

(background + peak) were 30 s for Si, Al, Ca, P, Fe; 45 s for As, S, Y; 75 s for all La-Lu; 150

s for U, Th and 225 s for Pb. Standards used were Gallium arsenide (AsGa) for As; wollastonite

18

(CaSiO3) for Si, Ca; aluminum oxide (Al2O3) for Al; sphalerite (ZnS) for S; thorite (ThO2) for

Th; apatite (Ca5(PO4)3(F,Cl,OH)) for P; endmember synthetic phosphates (XPO4) for all REE;

elemental iron (Fe) for Fe and vanadinite (Pb5(VO4)3Cl) for Pb. X-ray lines used were AsLα,

SiKα, AlKα, SKα, UMβ, ThMα, CaKα, YLα, PKα, LaLα, CeLα, SmLβ, GdLβ, NdLβ, YbLα,

EuLα, PrLβ, FeKα, TbLβ, DyLβ, HoLβ, ErLβ, TmLα, LuLα and PbMβ.

3.2.5.1. Chemical dating of uraninite

Chemical dating can be done using WDS-data from certain U and Th-bearing minerals, such as

uraninite and monazite. Provided that all Pb present in the mineral is a product of the radioactive

decay of U and Th, precise measurement of these elements allows an age of the mineral to be

calculated (Bowles 1990). Programs for calculation of chemical ages (t) are based on the

composite decay equation:

Eq 1: 𝐶𝑃𝑏 = 𝐶𝑇ℎ[0.897(𝑒λ232𝑡 − 1)] + 𝐶𝑈[0.859(𝑒λ238𝑡 − 1) + 0.006(𝑒λ235𝑡 − 1)]

(Hurtado et al. 2007)

Where CPb, CTh and CU are concentrations of Pb, Th and U in ppm. λ 232, λ238 and λ235 are decay

constants for 232Th (4.9475·10-11 yr-1), 238U (1.55125·10-10 yr-1) and 235U (9.8485·10-10 yr-1)

(Steiger & Jäger 1977). The coefficient before the first exponential term is given by the mass

ratio of 208Pb to 232Th (208/232 = 0.897), while the coefficients before the second and third

exponential terms are given by the ratios of the abundance fractions of the respective U isotopes

to the mean atomic mass of U (0.9928/238.04 = 0.859 for 238U and 0.0072/238.04=0.006 for

235U).

In order for the chemical age to accurately reflect the true age of the mineral, the boundary

condition, that no Pb, U and Th has been lost or gained after mineral crystallisation, must be

fulfilled (Kempe 2003). This can be controlled by calculating the ThO2*-value:

Eq 2: 𝑇ℎ𝑂2∗ = 𝑇ℎ𝑂2 +

𝑈𝑂2𝑊𝑇ℎ

𝑊𝑈[𝑒𝑥𝑝(λ232𝑡) − 1]× [

exp(λ235𝑡) + 138 𝑒𝑥𝑝(λ238𝑡)

139− 1]

(Suzuki et al. 1991)

ThO2, UO2 and PbO are reported in wt %. W is the molecular weight of each oxide (WTh = 264,

WU = 270, WPb = 224). 238U/235U = 138. Other variables are the same as in the composite decay

equation above. If the U-Th-Pb system has remained closed, and the analysed minerals

19

crystallised at the same time, all measurements should plot along a single isochron line in a

PbO-ThO2* diagram (Suzuki et al. 1991, Kempe 2003).

3.2.6. Stable isotope analysis

The stable isotope (O and C) analyses were performed at the Department of Geological

Sciences, University of Cape Town, South Africa. For the magnetite oxygen isotope analysis,

the samples were prepared by laser fluorination. The samples were placed in a reaction

chamber, where they were reacted with BrF5 at 10 kPa pressure. The purified O2 was then

collected onto a 5 Å molecular sieve in a glass storage bottle (see Harris & Vogeli 2010, Weis

2013). For carbonates, CO2 was extracted by reacting the samples with 100 % phosphoric acid

at 50 °C (see McCrea 1950, Ganino et al. 2013). The material was initially assumed to be calcite,

in order to correct for the fractionation between CO2 and carbonate. Samples containing

dolomite were later recalculated based on calcite-dolomite fractionation. The oxygen and

carbon isotope ratios were measured using a Thermo DeltaXP mass spectrometer, using

Monastery garnet as standard for magnetites and Namaqualand Marble for carbonates.

3.2.6.1. Stable isotope geochemistry

The field of isotope geochemistry was developed after the discovery of the neutron in 1932 by

H. Urey, and the demonstration of variations in the isotopic composition of light elements by

A. Nier during the 1930’s and 1940’s (White 2011). Stable isotope geochemistry is based on

variations in isotopic composition owing to physicochemical processes (White 2011, Hoefs

1997), and the measurement of these variations has many different geological applications. In

this study, stable isotopes are used as geochemical tracers for identification of the sources for

ore forming fluids together with the possibility of performing geothermometric calculations

(Rollinson 1993). Here oxygen and carbon isotopes have been used, more specifically the pairs

18O/16O and 13C/12C. The relative abundance of the heavier isotope (i.e. 18O and 13C) in a rock,

mineral or fluid is usually determined by calculating the δ-value. This value is obtained, using

oxygen as example, by the equation:

20

Eq 3: δ 𝑂18 =

𝑂18

𝑂16 (𝑠𝑎𝑚𝑝𝑙𝑒) −𝑂18

𝑂16 (𝑠𝑡𝑎𝑛𝑑𝑎𝑟𝑑)

𝑂18

𝑂16 (𝑠𝑡𝑎𝑛𝑑𝑎𝑟𝑑)

× 1000 [‰]

(Hoefs 1997, Rollinson 1993, White 2011)

The standard used for oxygen is the Standard Mean Ocean Water (SMOW), which is

characterised as 𝑂18

𝑂16 = 2005.2 ± 0.43 ppm (Baertschi 1976). For carbon, the established

standard is the Pee Dee Belemnite (PDB), with 𝐶13

𝐶12 = 11237.2 ± 2.9 ppm (Craig 1957). A

positive δ-value means that the analysed substance is enriched in the heavier isotope, while a

negative δ-value means that the substance is depleted in the heavier isotope, relative to the

standard. Due to various geological processes, there can be a partitioning of isotopes between

two substances in equilibrium with one another, leaving one enriched in 18O or 13C and the other

one depleted. This is known as isotope fractionation (White 2011, Rollinson 1993).

3.2.6.1.2. Stable isotope fractionation

The stable isotope fractionation between two substances, 1 and 2, can be expressed by the

fractionation factor 𝛼1−2 obtained by the equation:

Eq 4: 𝛼1−2 =𝑅1

𝑅2=

(𝑂18

𝑂16 )1

(𝑂18

𝑂16 )2

(Javoy 1977, Hoefs 1997)

Furthermore, the isotope fractionation is also temperature dependent:

Eq 5: 1000𝑙𝑛(𝛼1−2) =𝐴 × 106

𝑇2+ 𝐵

(Bottinga & Javoy 1973, Hoefs 1997, Chiba et al. 1989)

T is the absolute temperature (Kelvin) while A and B are thermometric coefficients specific for

certain pairs of substances (Chacko et al. 2001 and references therein) Note that for large T,

𝛼1−2 will approach 1, meaning that there is practically no fractionation at sufficiently high

21

temperatures. In addition to temperature, other physical properties such as bond strength, ionic

potential and oxidation state will have influence on the fractionation of stable isotopes.

Generally the heavier isotope is attracted to the stronger bond (Hoefs 1997, Rollinson 1993).

Correlating the δ-values of two substances with the fractionation between them is possible by

the following approximations:

Eq 6: ∆1−2= (𝛿 𝑂18 )1

− (𝛿 𝑂18 )2

~1000𝑙𝑛(𝛼1−2) for δ 18O < 10 ‰

Eq 7: 𝛼1−2 =1000 + (𝛿 𝑂18 )

1

1000 + (𝛿 𝑂18 )2

for δ 18O > 10 ‰

(Hoefs 1997, Rollinson 1993, Friedman & O’Niel 1962)

In cases were fractionation factors for certain pairs of substances are not established, it is

possible to combine fractionation factors:

Eq 8: 1000𝑙𝑛(𝛼1−2) + 1000𝑙𝑛(𝛼2−3) =

= 1000𝑙𝑛(𝛽1) − 1000𝑙𝑛(𝛽2) + 1000𝑙𝑛(𝛽2) − 1000𝑙𝑛(𝛽3) =

= 1000𝑙𝑛(𝛽1) − 1000𝑙𝑛(𝛽3) =

= 1000𝑙𝑛(𝛼1−3)

(Schütze 1980, Zheng 1991, Zheng 1996)

Where β is a temperature-dependent thermodynamic isotope factor (see Zheng 1991, Zheng

1996).

By combining the above equations (Eqs 4-8), a variety of information can be extracted.

Knowing the δ-values of two mineral pairs and the fractionation between them enables one to

calculate at which temperature the fractionation took place. And by knowing the fractionation

at a specific temperature between a mineral and an equilibrium substance, the δ-value of said

equilibrium substance can be determined (cf. the methodology used in Jonsson et al. 2013,

Nyström et al. 2008, Weis 2013).

22

4. Results

4.1. Bulk geochemistry of REE-mineralised assemblages

REE-mineralised samples from all ten studied deposits were analysed for major and trace

elements using XRF, ICP-AES and ICP-MS technology. The REE composition of analysed

samples shows that samples generally plot in two groups - those with high REE enrichment and

those with slight to no REE enrichment, compared to average crustal values (Fig 4). The highly

enriched samples come from Bastnäs, Östra Gyttorp, Södra Hackspiksgruvan,

Malmkärragruvan and Johannagruvan, while the low-enriched samples are from

Danielsgruvan, Högfors and Myrbacksfältet. Samples from Östanmossa (three highly enriched,

one low enriched) and Rödbergsgruvan (one highly enriched, one low enriched) plot in both

groups. Bulk REE concentrations are of course highly dependant on the particular mineral

assemblage sampled from each deposit, explaining why samples from some deposits plot in

both groups. The LREE to HREE enrichment can be determined using the (Ce/Yb)cn-ratio (cf.

Belousova et al. 2002). The highly enriched group shows higher (Ce/Yb)cn-ratios, ranging from

48 to 637 with an average of 313. The low enriched group shows lower (Ce/Yb)cn-ratios,

ranging from 15 to 188 with an average of 77. Comparing subtype 1 and suptype 2 deposits,

subtype 1 deposits show LREE enrichment (average Ce/Yb = 424-608) compared to subtype 2

deposits (average Ce/Yb = 63-267). Deposits not covered in this classification (i.e. Östra

Gyttorp, Högfors, Myrbacksfältet) all lie in the same range as subtype 2 deposits, while

Danielsgruvan shows lower values (average Ce/Yb = 57).

23

Most samples show a pronounced Eu depletion, which can be quantified using the equation:

Eq 9: 𝐸𝑢

𝐸𝑢∗= [

𝐸𝑢𝑐𝑛

(𝑆𝑚𝑐𝑛×𝐺𝑑𝑐𝑛)0,5] − 1

(Taylor & McLennan 1985)

Most samples show Eu/Eu*-ratios between -0.5 and -0.8. However, three samples show slightly

positive Eu anomalies. These comprise of the two low enriched Danielsgruvan samples (0.32

and 0.38) and one of the highly enriched Bastnäs samples (0.26).

The samples show similar chondrite-normalised REE trends as the REE-rich apatite-iron oxide

ores from Grängesberg (Jonsson et al. 2010, Jonsson et al. 2013), Blötberget (Jiao 2011) and

Idkerberget (Sahlström 2012), all in northwestern Bergslagen. However the Bastnäs-type ores

are generally slightly more LREE-enriched compared to these deposits. The large Bayan Obo

deposit in China, which has been interpreted to have a magmatic-hydrothermal origin (from a

presumably carbonatitic source), shows a much higher LREE to HREE enrichment (> 4 times)

compared to the Bastnäs-type ores, and it lacks the negative Eu-anomaly (Zhongxin et al. 1992).

So called IOCG-type (iron-oxide-copper-gold) ores from the Olympic Dam deposit in

Australia, which are suggested to have formed hydrothermally by combination of fluids from

both deep-seated and meteoric sources (e.g. Haynes et al. 1995, Gow et al. 1994, Oreskes &

Einaudi 1992), show similar REE trends, yet featuring neutral to even positive Eu-anomalies

(Oreskes & Eunaudi 1990).

Fig 4. Chondrite-normalised spider diagram of whole rock samples from the studied Bastnäs-type ores.

Average crustal values were collected from Webelements and chondrite values from McDonough & Sun

(1995).

24

4.2. Mineralogy and mineral chemistry

4.2.1. Östra Gyttorpsgruvan

The Östra Gyttorpsgruvan samples consist of both massive magnetite and silicate dominated

assemblages, hosted in mainly metavolcanic rocks (Nordenström 1890). The REE-silicate-rich

assemblages occur as isolated, relatively small lensoidal bodies or pods and consist mainly of

anhedral amphiboles, fine and even-grained allanite and biotite (Fig 4a). The allanites often

show zoning, generally with BSE-brighter cores and BSE darker-rims. Uraninite is common

and mostly present within allanite grains (Fig 4b). A gadolinite group mineral showing

anhedral-subhedral crystals with complex zoning patterns occurs frequently, associated with

allanite. The zoning can primarily be attributed to variable Y and Ce contents in the different

parts of the crystal due to the antipathetic behavior of these elements in gadolinites, as described

by Holtstam & Andersson (2007). Furthermore, a few zircon crystals have been observed in

association with the allanite. The zircons exhibit a distinct morphology and zoning pattern, with

a euhedral core and an subhedral, altered rim (Fig 4c). Some uraninites have been partially to

wholly replaced by bastnäsite (Fig 4d). Bastnäsite is furthermore common in fractures, showing

that it is a paragenetically late mineral (Fig 4e). It often exhibits lamellar intergrowths with

parisite, which is related to the variability in stacking of (REEF),(CO3) and (Ca(CO3))-layers

in the bastnäsite group minerals (Yunxiang et al. 1993). The massive magnetite ore samples

consist of masses of medium-grained, homogenous and anhedral magnetite together with bands

of mica (Fig 4f).

25

Fig 4a-f. Photomicrographs (a,f) and BSE images (b-e) of samples from Östra Gyttorpsgruvan. Images a), b),

c), d) and e) are of sample Gytt-EJ-2 and image f) is of sample Gytt-EJ-1a.

a) Overview photograph (transmitted light) of silicate-dominated sample showing allanite-(Ce) (Aln), biotite

(Bt), amphibole (Am) and uraninite (Urn). b) BSE image showing the presence of gadolinite group minerals

(Gad) and uraninite in the allanite masses. The gadolinite minerals show complex zoning due to variable Y

and Ce contents. Allanites exhibit zoning with Ca, Fe3+-rich rims (BSE-light) and REE, Fe2+-rich cores (BSE-

dark). c) Zoned zircon (Zrn) with euhedral core and altered, subhedral rim. d) Bastnäsite-(Ce) (Bas) wholly

replacing uraninite inside an allanite grain. e) Bastnäsite/parisite-(Ce)(Bas/Par) in fractures. Notice the

lamellar intergrowth textures. f) Reflected light photomicrograph showing massive magnetite ore sample with

bands of mica.

26

Samples from Östra Gyttorpsgruvan were analysed using WDS technique. One of the aims was

to characterise the allanites, as they represent the main REE-bearing phase in the deposit. The

chemical compositions of 26 analysed points are shown in Table 5 (appx). Based on these

measurements, the average formula for Östra Gyttorp allanites can be expressed as:

A1(Ca0.96[]0.04)A2(Ce0.41Nd0.18La0.16Pr0.05Sm0.03Y0.02Gd0.01∑REE0.86Ca0.14)∑1.00

M3(Fe2+0.63Mg0.19

Fe3+0.18)∑1.00

M2(Al1.00)M1(Al0.79Fe3+

0.21)∑1.00T(Si3.00)(O11.99F0.01)∑12.00(OH)1.00 if calculated

according to the methodology described in Ercit (2002). From the REE concentration, it can be

concluded that the allanites belong to the species allanite-(Ce), with significant and equal

amounts of Nd and La (Fig 5). The compositional diagram (Fig 6) shows that there has been

some substitution in the allanite structure, where Fe3+ substitute for Al in M3 and Mg for Fe2+

in M1, indicating solid solution towards the end member CaCeMgAlFe3+Si3O12(OH). In the

chemical formula it is also apparent that some REE are replaced by Ca in A2, in order to balance

for Fe3+ substituting for Fe2+ in M1. This can be visualised by introducing epidote as another

end member of the solid solution (Fig 7). The zoning in allanite (Fig 4b) can be attributed to

these substitutions, with BSE-lighter cores being richer in REE+Fe2+ and BSE-darker rims

being richer in Ca+Fe3+.

Fig 5. Plot showing major REE (in atoms per formula unit - apfu) in allanite-group minerals.

27

Fig 6. Plot of Fe3+/(Fe3++VIAl) versus Fe2+/(Fe2++Mg) (in apfu) for Östra Gyttorpsgruvan and

Johannagruvan allanite group minerals, with compositional fields for allanites (from Holtstam & Andersson

2007) and ferriallanites (from Holtstam 2003) from the Bastnäs deposit. Modified after Holtstam & Andersson

(2007).

Fig 7. Plot of Mg/(Mg+Fe2+) versus Mg+Fe2+ (in apfu) for allanite group minerals from Östra

Gyttorpsgruvan and Johannagruvan.

28

The gadolinite group minerals were also analysed with WDS. The chemical compositions of 18

analysed points are shown in Table 6 (appx). The average chemical formula can be expressed

as: (Ca0.32Y1.27Dy0.11Gd0.09Yb0.04Er0.04Ho0.03Nd0.03Sm0.02Ce0.01∑REE1.64)∑1.96(Fe0.39Mg0.03)∑0.42

Si2.04(Be1.68B0.32)∑2.00O8.83(OH)1.17, based on 4 A+T cations (cf. Miyawaki 2007, DeMartin et

al. 2001). There has been significant substitution of Ca+B for REE+Be and (OH)2 for Fe (Fig

8), and all but three points should be classified as hingganites. However it cannot be excluded

that some of the addition of Ca for REE was balanced by replacing Fe2+ with Fe3+, i.e. solid

solution towards calciogadolinite, which was not possible to confirm with WDS data. This

could possibly explain the low totals (Table 6, appx), and the excess of Si and shortage of A-

site cations in the calculated mineral formula. These substitutions in Östra Gyttorp gadolinites

make them differ significantly from gadolinites at other localities along the REE-line (this

study, Holtstam & Andersson 2007), which exhibit close to ideal gadolinite composition. Fig

9 shows the dominant REE, and suggest that the gadolinite group minerals from Östra Gyttorp

should be classified as hingganite-(Y)/gadolinite-(Y). The Y/Ce-ratios vary between 13 and

271, which is remarkably contrasting to gadolinites from other REE-line occurrences, which

are much richer in Ce and Nd (this study, Holtstam & Andersson 2007). The large variability

in Y/Ce-ratios is reflected in the zoning described above.

Fig 8. Compositional plot for gadolinite group minerals from Östra Gyttorpsgruvan and Johannagruvan.

29

Bastnäsite replacing uraninite (Fig 4d) were also analysed by WDS. The crystals are generally

very small (a few µm) which hampers the analysis of volatile elements. The chemical

composition of four analysed points are presented in Table 7 (appx). The compositional diagram

(Fig 10) shows that the minerals plot along the bastnäsite-parisite line, indicating about 25 %

parisite. This could be due to solid solution, but more likely due to microscopic lamellar

intergrowths (as seen in Fig 4e) being included in the analysed points. Furthermore, there

appears to be significant substitution of OH for F (based on calculations assuming ideal

stoichiometry), suggesting hydration of the mineral (cf. Guastoni et al. 2009). However this

substitution might be slightly exaggerated due to difficulties of obtaining accurate fluorine

analyses of such small grains due to degassing under the electron beam. The average mineral

composition (based on 6 cations) can be expressed as: Ca1.01(Ce2.14La0.97Nd0.96Y0.29Pr0.25

Sm0.17Gd0.12Dy0.03Eu0.03U0.02)∑4.98(CO3)6.01(F2.64(OH)2.35)∑4.99. The dominance of Ce means that

the mineral should be classified as bastnäsite-(Ce) (Fig 11).

Fig 8. Plot of (REE+Be)/(REE+Be+Ca+B) versus (Fe+Mg)/(Fe+Mg+(OH)2) (in apfu) for gadolinite group

minerals from Östra Gyttorpsgruvan and Johannagruvan.

Fig 9. Dominant REE (in apfu) in gadolinite group minerals from Östra Gyttorpsgruvan and Johannagruvan.

http://en.wikipedia.org/wiki/Mu_(letter)

30

Fig 10. Plot of REE/CO3 versus Ca/(F+OH) (in apfu) of Östra Gyttorpsgruvan and Johannagruvan REE-

fluorocarbonates. Analysed points from Östra Gyttorp were in replacements of uraninite, while points from

Johannagruvan were in larger grains.

Fig 11. Plot of major REE (in apfu) in Östra Gyttorpsgruvan and Johannagruvan REE-fluorocarbonates.

31

Chondrite-normalised spider diagrams (Fig 12) show that allanites and bastnäsites are enriched

in LREE, similar to the whole rock, while gadolinite-group minerals favour HREE+Y have a

reverse LREE-trend. Bastnäsites are depleted in Er and enriched in Tm, whereas allanite and

gadolinite-group minerals show the opposite trend. All three minerals are depleted in Tb and

have slightly negative Eu-anomalies. Allanites and bastnäsites appear to fractionate REE in

proportions similar to that of the whole rock (Fig 13). This is especially valid for LREE, while

HREE show slight scatter (Fig 13). This indicates that crystallographic controls on REE

incorporation, particularly LREE, are limited in these minerals. Gadolinite group minerals

however show a very large scatter for both LREE and HREE, indicating that these minerals

fractionate REE rather independently of whole rock REE proportions and that there is a strong

crystallographic control on the incorporation of REE. This is indicated by the very high

concentrations of Y and certain HREE in these minerals despite relative depletion in the whole

rock (Fig 12), together with the zonation patterns that are explained by the antipathetic behavior

of Y and Ce (Fig 4b).

Fig 12. Chondrite normalised spider diagram of average REE concentrations in allanite (n=26), gadolinite

group minerals (n=18), bastnäsite (n=4) and whole rock from Östra Gyttorpsgruvan. Chondrite values from

McDonough & Sun (1995).

1

10

100

1000

10000

100000

1000000

10000000

La Ce Pr Nd Sm Eu Gd Tb Dy Y Ho Er Tm Yb Lu

Min

eral

/ch

on

dri

te (

pp

m)

Whole rock

Allanite

Gadolinite

Bastnäsite

32

4.2.2. Johannagruvan

Johannagruvan is a skarn iron ore deposit, which was mined up to 1909 (Geijer 1936). The

skarn consist of amphibole (mainly tremolite) with minor quartz. At deeper levels of the mine,

enclosed units of ophicalcite occur (Geijer 1936). The skarn minerals occur as masses of fine

grained, subhedral-anhedral grains (Fig 14a). Aggregates of brown dollaseite are also common,

while magnetite and minor chalcopyrite, pyrite and sphalerite occur as disseminations (Fig 14a).

Some samples exhibit banding of coarse-grained norbergite (Fig 14b). EDS analyses show that

the dollaseite-(Ce) masses contain a substantial amount of bastnäsite-(Ce) and “gadolinite-

(Nd)”, which occur as anhedral, irregular masses (Fig 14c). Some bastnäsite grains have

lamellar intergrowths of parisite-(Ce) (Fig 14d). Minor molybdenite is also present as isolated

grains in the dollaseite masses (Fig 14e).

Fig 13. Mineral REE versus whole rock REE content for minerals from Östra Gyttorpsgruvan. Elements that

were below detection limit in WDS analysis had to be omitted due to the logaritmisation. These elements are

Tb and Tm for allanite and gadolinite group minerals, and Tb and Er for bastnäsite.

33

Figure 14a-e. Photomicrographs (a-b) and BSE images (c-e) of ore samples from Johannagruvan. Images a),

c), d) and e) are of sample Joha-KH-1 and image b) is of sample Joha-KH-2.

a) Samples consist mainly of white amphibole (Am) and quartz (Qtz) together with anhedral aggregates of

dollaseite-(Ce) (Dla). Small amounts of magnetite (Mgt), pyrite (Py), chalcopyrite (Ccp) and sphalerite (Sp)

occur as disseminations. b) Banding of coarse-grained norbergite (Nrb) in XPL. c) Gadolinite-(Nd) (Gad)

and bastnäsite-(Ce) (Bas) occur as anhedral grains mostly concentrated in the dollaseite masses. d) Some

bastnäsite grains show lamellar intergrowths of parisite-(Ce) (Par). e) Small amounts of molybdenite (Mo)

are present in the samples.

34

Additional mineral chemical analyses by WDS was performed on dollaseites, gadolinites and

bastnäsites. The dollaseite chemistry is shown in Table 8 (appx); these are rather pure

dollaseites (Fig 6), with only minor substitution of Fe2+ for Mg. Similarly to the Östra Gyttorp

allanite-group minerals, they are Ce-dominant (Fig 5). An average empirical formula can be

expressed as: A1(Ca0.88Y0.06Gd0.01Sm0.02[]0.03)∑1.00A2(Ce0.49La0.30Nd0.17Pr0.05)∑1.01

M3(Mg1.02)

M2(Al0.95Fe3+0.04)∑0.99

M1(Mg0.68Fe2+0.30Mn0.02)∑1.00

T(Si2.97Al0.03)∑3.00(O11.23F0.77)∑12.00(OH1.00),

calculated using the methodology described in Ercit (2002). The stoichiometry indicates that

there is an excess of 0.1 apfu REE that cannot fit in the A2-site. Therefore 0.09 apfu of the three

REE with the largest ionic radius (Y, Gd and Sm) was used to fill up the A1-site, which had a

deficit of 0.12 apfu. This excess in REE can most likely be attributed to charge balancing of the

deficit in F, which in ideal dollaseite should be 1 apfu.

Gadolinites from Johannagruvan exhibit some distinct differences from those from Östra

Gyttorp. All but one of the analysed points are Nd-dominant, with Ce and Y in approximately

equal proportions (Table 9 (appx) and Fig 9). It is also evident that only very minor hingganite

substitution has occurred here (Fig 8). An empirical mineral formula can be expressed as:

(Ca0.02Nd0.57Y0.40Ce0.40Sm0.18Gd0.15La0.09Pr0.09Dy0.04As0.01)∑1.95(Fe2+0.76Mg0.10)∑0.86Si2.06(Be1.98

B0.02)∑2.00O9.72(OH)0.28, based on 4 A+T cations (cf. Miyawaki 2007, DeMartin et al. 2001).

WDS data of bastnäsites is shown in Table 10 (appx). The influence of parisite substitution is

very limited, evident by the very low Ca contents (Fig 10). This was expected due to the

intergrowths of parisite (Fig 14d). Based on analytical totals, there is significant OH-

substitution, albeit not as large as in Östra Gyttorp bastnäsites. The dominant REE is Ce (Fig

11), with proportions very similar to bastnäsites from Östra Gyttorp. The average formula

(based on 6 cations) was calculated as: (Ca0.02K0.02)∑0.04(Ce2.55La1.68Nd1.26Pr0.29Sm0.10Y0.06

Gd0.02Tm0.01As0.01Si0.01)∑5.99(CO3)6.00(F4.61(OH)1.37)∑5.98.

Chondrite-normalised spider diagrams (Fig 15) show that dollaseites and bastnäsites are LREE-

enriched, similar to the whole rock. Gadolinite has high HREE+Y and a reverse LREE-trend,

albeit not to the same extent as in Östra Gyttorp samples. All minerals show strongly negative

Eu-anomalies. It is evident that dollaseites and bastnäsites from Johannagruvan exhibit REE in

proportions similar to those of the whole rock (Fig 16). Dollaseites show an almost perfect

linear trend, while bastnäsites show slight scatter. In contrast to gadolinite group minerals from

Östra Gyttorp, the Johannagruvan gadolinites do not scatter as strongly, indicating that there

has not been the same crystallographic control on REE incorporation in these minerals.

35

Fig 15. Chondrite-normalised spider diagram of average REE concentrations in dollaseite (n=9), gadolinite

(n=13), bastnäsite (n=8) and whole rock from Johannagruvan. Chondrite values from McDonough & Sun

(1995).

1

10

100

1000

10000

100000

1000000

10000000

La Ce Pr Nd Sm Eu Gd Tb Dy Y Ho Er Tm Yb Lu

Min

eral

/ch

on

dri

te (

pp

m)

Whole rock

Dollaseite

Gadolinite

Bastnäsite

Fig 16. Mineral REE plotted against whole rock REE for Johannagruvan samples. The following elements

had to be omitted since they were below detection limit and therefore could not be logaritmised: Gadolinite -

Eu, Tb, Tm, Yb. Dollaseite - Eu, Tb, Tm. Bastnäsite - Eu, Tb, Ho, Er, Yb.

36

4.2.3. Högforsfältet

The Högfors deposit is banded iron formation (BIF) with minor skarn and small amounts of

REE (Jonsson et al. 2013, Geijer 1961), and represents the only observed occurrence of this

type in the area. Layers of massive magnetite and hematite are mixed with bands of quartz and

thin bands of skarn assemblages with mainly tremolite and REE silicates. The magnetite is

coarse grained and occur as anhedral-subhedral masses in bands (Fig 17a). Within the quartz

bands numerous small grains of recrystallised magnetite and hematite occurring as

disseminations (Fig 17b). Furthermore, the banded assemblages show a very distinct folding,

visible in the hand specimens. In the ores minor amounts of REE-bearing minerals such as

cerite-(Ce), gadolinite-(Ce), allanite-(Ce), bastnäsite-(Ce), monazite-(Ce), an Fe-analogue of

västmanlandite-(Ce) and an Mg-analogue of rowlandite-(Y) have been identified (Jonsson &

Högdahl 2013, Jonsson et al. 2014). The coarse-grained, primary magnetite and quartz are

interpreted to be coeval and to be in textural equilibrium.

Figure 17a-b. Reflected light photomicrographs of ore sample Högf-KH-1 from Högfors.

a) Band of coarse magnetite (Mgt) interbedded between quartz (Qtz) bands. b) Small dissiminations of

hematite (Hem) and magnetite in silicate bands.

37

4.2.4. Danielsgruvan

Danielsgruvan is a quartz-banded iron deposit, similar to Högfors, albeit not with the same

prominent layering. It occurs together with a carbonate rock rich in tremolite and serpentinte

skarn, containing REE mineralisation (Geijer 1961). Coarse grained massive magnetite bands

are interbedded with bands of hematite and bands of quartz and minor skarn amphibole (Fig

18a). The quartz bands contain numerous inclusions of recrystallised magnetite and hematite.

The magnetite inclusions are rather large and randomly oriented, while the smaller hematite

grains are elongate and aligned with the banding (Fig 18b). Pyrite is rather common, and occurs

mainly as fracture fillings (Fig 18d). Previous studies suggest that minor amounts of dollaseite-

(Ce) and allanite-(Ce) occur as isolated grains or small masses. Additionally, rare törnebohmite-

(Ce) has been identified within in the dollaseite (Geijer 1961). Similarly to the Högfors samples,

coarse-grained and primary magnetite is interpreted to be coeval and in textural equilibrium

with the quartz.

Figure 18a-d. Reflected light photomicrographs

of Danielsgruvan ore samples. Images a) and b)

are of sample NRM-20020125 and image c) is

from sample NRM-20020124.

a) Overview showing the coarse magnetite (Mgt)

and bands of hematite (Hem) and Quartz (Qtz). b)

Quartz band with inclusions of magnetite and

hematite. Notice the elongate, aligned hematite

grains. c) Pyrite (Py) is present mainly as fracture

fillings.

38

4.2.5. Myrbacksfältet

Myrbacksfältet contain a sulfide rich massive magnetite ore hosted in metavolcanic rocks,

sometimes also with large amounts of quartz, with mining of mainly iron having occurred in

over 30 local operations of variable size (Geijer 1923, Geijer & Magnusson 1944). REE

enrichment in the area is mainly visible as microscopic allanite and REE-epidote in quartzite

(Geijer 1923). The sulphides consist of pyrite and chalcopyrite that occur as coarse grained

disseminations in the anhedral massive magnetite together with minor fine grained silicates

(mainly quartz and amphiboles) (Fig 19a). The pyrites often show signs of brittle deformation

(Fig 19b), and in contact with silicates it often display remobilisation, forming networks of thin

veinlets (Fig 19c). In the silicates small disseminations of magnetite, pyrite and chalcopyrite

are common (Fig 19d).

39

4.2.6. Södra Hackspiksgruvan

Södra Hackspiksgruvan is a fluorite-rich skarn iron ore deposit, representing a, for the region,

rather unusual mineral paragenesis (Geijer 1961, Geijer 1936). The fluorite make up masses of

subhedral grains, which coexist with amphibole (mainly tremolite) and masses of dissakisite-

(Ce), cerite-(Ce), fluorbritholite-(Ce) and bastnäsite-(La) - bastnäsite-(Ce) (Fig 20a-d).

Additionally, törnebohmite-(Ce) have been found in the deposit (Geijer 1936). Minor phases,

mainly magnetite but also pyrite, scheelite and an orange humite group mineral, occur as

disseminations. The REE-mineralogy is rather complex. Large coherent masses of dissakisite

dominate in the samples, with the other minerals occurring as both irregular masses and isolated

grains (Fig 20b). Cerite is commonly altered into bastnäsite, giving it a patchy texture (Fig 20c).

Figure 19a-d. Reflected light photomicrographs of ore sample Myrb-EJ-1 from Myrbacksfältet.

a) Overview showing coarse-grained anhedral pyrite (Py), chalcopyrite (Ccp) and silicates in the massive

magnetite (Mgt) ore. b) Brittle deformation in pyrite. c) Pyrite occurring as networks in silicates.

d) Disseminations of magnetite, pyrite and chalcopyrite in silicate.

40

Minor amounts of fluorbritholite are found, and these are often partially altered into cerite (Fig

20d). Fluorbritholite is considered to be the paragenetically oldest REE-mineral in the Bastnäs-

type deposits (Holtstam & Andersson 2007), the likely paragenetic sequence is thus

fluorbritholite cerite bastnäsite.

Figure 20a-d. Photomicrograph (a) and BSE images (b-d) of ore sample SGU-M3563 from Södra

Hackspiksgruvan.

a) Masses of dissakisite-(Ce) (Dis), cerite-(Ce) (Cer), bastnäsite-(La) (Bas) and fluorbritholite-(Ce) (Flb)

occur in a matrix of coarse, subhedral fluorite (Fl). Magnetite (Mgt) occur as disseminations. b) Dissakisite

make up large coherent masses, while the other REE-bearing minerals occur in irregular masses and as

isolated grains. c) Cerite (BSE-lighter) is often altered into bastnäsite (BSE-darker), creating a patchy texture.

Minor scheelite (Sch) is also found in the samples. d) Cerite also appears as an alteration product of

fluorbritholite.

41

4.2.7. Östanmossagruvan

The Östanmossa deposit is hosted in mainly dolomitic carbonates, which have been partially

replaced by an association of magnetite and skarn silicates (Geijer 1927), and was mined until

1931 (Geijer 1936). The samples from Östanmossa represent the most carbonate-rich samples

in the study. The matrix consists of masses of fine to coarse-grained carbonates with minor

amphibole skarn minerals (actinolite and minor tremolite), which co-exist with REE-bearing

minerals, magnetite (occurring as bands and disseminations), fluorite and minor amounts of an

orange humite-group mineral. Fluorbritholite-(Y) is very common and form large coherent

masses (Fig 21a-b). Along contacts to carbonates as well as along fractures, the fluorbritholite

is partially altered to parisite-(Ce), giving it a prominent darker color both in BSE and

transmitted light (Fig 21b-d). Within fluorbritholite masses small bastnäsite-(Ce) grains are

common, which often show lamellar intergrowths with parisite-(Ce) (Fig 21c-d). A few

individual grains of monazite-(Ce) were found in the altered rims of the fluorbritholite masses

(Fig 21d). Additionally, some uraninite is present as inclusions in the fluorbritholite (Fig 21e).

The uraninites are partly replaced by bastnäsite-(Ce) - parasite-(Ce), similar to the replacement

seen in Östra Gyttorp samples. Dollaseite-(Ce) is also common and occur together with

fluorbritholite-(Y) as large irregular masses in the carbonate matrix (Fig 21a,f). Bastnäsite-(Ce)

also occurs with fluorbritholite, where the former may have formed at the expense of the latter

(Fig 21g). Additionally, some minor isolated grains of scheelite were found associated with

cerite (Fig 21g).

43

Fig 21a-g. Photomicrographs (a-b) and BSE images (c-g) of Östanmossa samples. Image a) is from sample

SGU-M4441, images b), c), d) and e) are from sample SGU-M4528 and images f) and g) are from sample EJ-

ÖM90-13-1.

a) Dollaseite-(Ce) (Dla) and fluorbritholite-(Y) (Flb) form anhedral aggregates in a carbonate (Cb) matrix.

b) Fluorbritholite masses adjecent to carbonates. Note the dark rim of parisite-(Ce) (Par) alteration. c) BSE

image again showing the parisite alteration along rims. Also present in the fluorbritholite masses are small

bastnäsite-(Ce) (Bas) grains, often showing lamellar intergrowths of parisite-(Ce) (BSE dark). d) Small

isolated grains of monazite-(Ce) (Mnz) found at parisite rims. e) Uraninite (Urn) with bastnäsite-(Ce)

replacement along edges. f) Dollaseite and fluorbritholite masses in carbonate matrix (cf. 21a).

g) Bastnäsite co-existing with fluorbritholite. Also seen is a small, isolated grain of scheelite (Sch).

44

4.2.8. Bastnäsfältet

The Bastnäs area consist of a variety of polymetallic deposits, and mining copper and iron

occurred for centuries before REE mineralisation were discovered in the middle of the

eighteenth century (Geijer 1921). At the Nya Bastnäs deposit northeast of Riddarhyttan village,

160 tons of REE ore (mainly cerite) were extracted in the two local mines Ceritgruvan and

Sankt Göransgruvan up until their closure in 1919 (Andersson et al. 2004), with later extraction

of REE from the old mine dumps. Studied samples from Bastnäs both represent massive

magnetite (Fig 22a) and skarn-dominated assemblages (Fig 22b-f). The massive magnetite

samples contain additional minor skarn minerals, allanite and are also rather rich in chalcopyrite

and malachite (on oxidised surfaces). The skarn samples are dominated by fine-grained

amphibole in addition to masses of ferriallanite. The amphiboles and the ferriallanite show a

pronounced foliation occasionally folded in some parts of the samples (Fig 22b). Opaque phases

appear to be more diverse in the skarn-dominated samples, containing rather large amounts of

pyrite and sphalerite in addition to the magnetite and chalcopyrite (Fig 22c). SEM-EDS analysis

shows the presence of small amounts of molybdenite (Fig 22d). Additionally, other minerals

such as bismuthinite, linnæite, cuprite as well as native gold has been found in the area (Geijer

1927). The ferriallanites appear to have a rather ideal composition, with very little Mg

substitution, and the dominant REEs are Ce and La in approximately equal proportions. The

ferriallanites also display slight zonation with BSE-darker and BSE-lighter parts (Fig 22e), but

any major compositional variability between the different parts cannot be identified with EDS.

Small accumulations of tiny monazite-(Ce) grains occur in the amphibole masses (Fig 22f).

45

Fig 22a-f. Photomicrographs (a-c) and BSE images (d-f) of Bastnäs ore samples. Image a) is from sample

SGU-M6777, image b) is from sample Bast-EJ-1 and images c), d), e) and f) are from sample SGU-M309.

a) Massive magnetite (Mgt) ore with minor amphibole (Am) and chalcopyrite (Ccp). b) Skarn-dominated ore

with amphibole and ferriallanite-(Ce) (Fln). Note the pronounced foliation in the sample. c) Skarn-dominated

ore sample with sphalerite (Sp) and pyrite (Py) in addition to the chalcopyrite. d) Overview image showing

the ferriallanite masses and minor chalcopyrite and molybdenite (Mo). e) A slight zonation in the ferriallanite

is visible. f) Minor amounts of monazite-(Ce) (Mnz) was found in the ferriallanite.

46

4.2.9. Malmkärragruvan

Malmkärragruvan is located in a syncline, and the mineralisation is hosted by a dolomitic

limestone layer (Geijer 1936). Studied samples are skarn-dominated assemblages rich in

carbonates. Amphiboles (mainly actinolite), carbonates and fluorite make up a rather coarse-

grained matrix, in which masses of mainly västmanlandite-(Ce) and fluorbritholite-(Ce) are

found (Fig 23a). Magnetite and minor amounts of chalcopyrite and pyrite are found as

disseminations mainly in carbonates. Västmanlandite makes up large aggregate masses, and

also occur associated with fluorbritholite (Fig 23b). The vätmanlandite masses exhibit fractures,

which are often filled with parisite-(Ce). Fluorbritholite occurs both as isolated anhedral grains

in västmanlandite masses and as large masses in the carbonate matrix (Fig 23b-c). A Cl-rich,

not fully characterised, REE-silicate occurs as intergrowths in fluorbritholite (Fig 23d). This

phase most likely corresponds to “Mineral E” in Holtstam & Andersson (2007). Preliminary

EDS-analysis indicate that this mineral contains Ca, REE, Cl and minor Mg, Al and F. The

REE are dominated by Ce and Nd in equal proportions with significant amounts of La, Pr, Sm

and Gd. Sparse fine-grained uraninites are present in the fluorbritholite masses, and similarly

to other REE-line deposits they are partly altered to bastnäsite-(Ce) (Fig 23e). Fluorite appears

to be intimately related to fluorbritholite, often growing along grain boundaries of the

fluorbritholite masses (Fig 23f). Bastnäsite occurring as individual grains in the matrix often

shows skeletal textures. The grains are partially to fully replaced by fluorite (Fig 23g).

47

Figure 23a-g. Photomicrograph (a) and BSE images (b-g) of ore samples from Malmkärragruvan. Images a),

b), c) and e) are from sample SGU-M4048 and images d), f) and g) are from sample SGU-M4068.

a) Masses of västmanlandite-(Ce) (Väs) and fluorbritholite-(Ce) (Flb) are found in a matrix of carbonate (Cb),

amphibole (Am) and fluorite (Fl). Magnetite (Mgt) occur as disseminations. b) Fluorbritholite in

västmanlandite masses. Note the fracture fillings of parisite-(Ce) in the västmanlandite. c) Parisite in fractures

in fluorbritholite. d) Not fully characterised Cl-bearing REE-silicate occuring as intergrowths in

fluorbritholite masses. e) Uraninites (Urn) in fluorbritholite. Bastnäsite-(Ce) (Bas) replacements are found in

the halos. f) Fluorite occuring along the rims of fluorbritholite masses. g) Isolated bastnäsite crystals in the

carbonate/amphibole matrix often show skeletal textures, and they are being replaced by fluorite to variable

degrees.

48

4.2.10. Rödbergsgruvan

Rödbergsfältet comprise of a skarn iron ore hosted in metavolcanic rocks, which has been

mined in a collection of small mines called Rödbersgruvorna (Geijer & Magnusson 1944).

Studied samples are rich in both massive magnetite and skarn (mainly actinolite), with minor

disseminated chalcopyrite and molybdenite (Fig 24a). The samples are rich in various REE-

bearing minerals. Allanite-(Ce), bastnäsite-(Ce) and cerite-(Ce) occur intergrown as anhedral

grains in rather complex aggregates (Fig 24b). Some cerite grains appear to have been altered

into bastnäsite, resulting in a patchy texture (Fig 24c). The allanites are exhibit only minor Mg

substitution (as indicated by EDS-analyses), similarly to Bastnäs and Östra Gyttorp samples.

Fig 24a-c. Photomicrograph (a) and BSE images

(b-c) of Rödbergsgruvan sample NRM-880071.

a) Amphibole (Am), magnetite (Mgt) and minor

amounts of molybdenite (Mo). b) Allanite-(Ce)

(Aln), bastnäsite-(Ce) (Bas) and cerite-(Ce) (Cer)

are found in intergrown aggregates in the skarn.

c) Some cerite grains show alteration to

bastnäsite.

49

4.3. Chemical dating of uraninites

Chemical dating was conducted on 24 uraninites from Östra Gyttorpsgruvan, and ages

calculated using the data obtained from WDS analyses (Table 11, appx) are shown in Fig 25.

The ages display a large span, ranging from 1898 to 1623 Ma with most ages lying between

1675 to 1775 Ma (2σ = ± 3 Ma). Only three points lie in the range of Re-Os ages (1.90-1.84

Ga) from Andersson et al. (2013) (1843, 1869 and 1898 Ma), and only two of these (1869 and

1898 Ma) show ages corresponding to the syn-volcanic stage of the Svecofennian (1.90 to 1.87

Ga; e.g. Stephens et al. 2009). This can also be seen in the PbO vs. ThO2* diagram (Fig 26),

where all but two points plot below the Svecofennian syn-volcanic stage isochrons. From the

WDS analyses (Table 11, appx), it is clear that the REE contents are high. An average empirical

formula (based on one cation) for the uraninites from Östra Gyttorp can be expressed as:

(U0.56Pb0.19Y0.09Ce0.02Nd0.02Dy0.02Fe0.02Gd0.01Sm0.01Yb0.01Ca0.01)∑1.00O2.

Fig 25. Calculated chemical ages of Östra Gyttorp uraninites, based on WDS-data (Table 11, appx). 2σ errors

for all points are less than ± 3 Ma. Inserted are fields for Re-Os ages of Bastnäs-type deposits (1.9-1.84 Ga;

Andersson et al. (2013) and the Svecofennian syn-volcanic stage (SSVS - 1.9 to 1.87 Ga; e.g. Stephens et al.

2009).

50

4.4. Results stable isotopes

In total 21 magnetite samples, three quartz samples and seven carbonate samples were analysed

from 10 different deposits with respect to their oxygen (and carbon) isotopic compositions. The

obtained δ18O and δ13C-values, together with dolomite proportions of carbonates, is shown in

Table 2. δ18O in magnetite vary between -2.3 and 1.1 ‰, relative to SMOW, while quartz and

carbonates show ranges of 7.2 to 8.3 ‰ and 5.8 to 7.2 ‰, respectively. δ13C in carbonates range

from -5.4 to -3.3 ‰, relative to PDB. In general, magnetites from the same deposit tend to have

rather similar δ18O-values, while the δ18O can vary quite substantially between magnetites from

different deposits.

Fig 26. PbO vs. ThO2* diagram of Östra Gyttorp uraninites. Assuming mineralisation during the 1.90-1.87

Ga Svecofennian syn-volcanic stage (Andersson et al. 2013, Jonsson & Högdahl 2013), any points plotting

below these isochrons have most likely experienced either lead loss or a later re-equilibration event. Note that

the reason for the artificial ThO2*-values being >100 wt % is due to the much slower decay rate of Th (t1/2 =

1.41∙1010 years) compared to U (t1/2 = 4.47∙109 years for 238U and t1/2 = 7.04∙108 years for 235U).

51

Table 2. Analytical results from mass spectrometry on magnetite, quartz and carbonates. Also included are

dolomite proportions of carbonates obtained from powder XRD analysis.

Deposit Type Sample ID δ18Omgt (‰) δ18Oqtz (‰) δ18Ocarb (‰) δ13Ccarb (‰) 2σ % dol

Danielsgruvan Banded mgt-qtz ore NRM-20020125 -0.7 7.5 ±0.2

Danielsgruvan Banded mgt-qtz ore NRM-20020124 -1.9 ±0.2

Östanmossa Carbonate skarn mgt ore EJ-ÖM90-13-1 -1.2 7.2 -5.4 ±0.2 0

Östanmossa Carbonate skarn mgt ore SGU-M4528 -1.4 5.8 -3.6 ±0.2 100

Östanmossa Carbonate skarn mgt ore EJ-ÖM90-13-2 6.4 -4.8 ±0.2 10

Östanmossa Carbonate skarn ore SGU-M4441 6.8 -5.3 ±0.2 20

Östanmossa Carbonate skarn ore SGU-M4529 5.9 -5.1 ±0.2 85

S. Hackspiksgruvan Fluorite skarn mgt ore SGU-M3563 0.5 ±0.2

Johannagruvan Skarn mgt ore Joha-KH-1 -0.5 ±0.2

Johannagruvan Skarn mgt ore Joha-KH-2 0.5 ±0.2

Malmkärra Carbonate skarn mgt ore SGU-M4068 -0.6 6.8 -3.3 ±0.2 80 Malmkärra Carbonate skarn mgt ore SGU-M4048 0.2 7.0 -3.7 ±0.2 70

Högfors BIF Högf-KH-1 -1.7 8.3 ±0.2

Högfors BIF Högf-KH-2 -0.4 7.2 ±0.2

Bastnäs Skarn mgt ore Bast-EJ-1 -0.6 ±0.2

Bastnäs Massive mgt ore SGU-M6777 -1.2 ±0.2

Bastnäs Skarn mgt ore SGU-M309 0.0 ±0.2

Myrbacksfältet Massive mgt/sulphide ore Myrb-EJ-1 0.7 ±0.2

Myrbacksfältet Massive mgt/sulphide ore Myrb-EJ-2 1.1 ±0.2

Östra Gyttorp Massive mgt ore Gytt-EJ-1a -1.8 ±0.2

Östra Gyttorp Massive mgt ore Gytt-EJ-1b -2.3 ±0.2

Östra Gyttorp Massive mgt ore Gytt-EJ-1c -2.1 ±0.2

Rödbergsgruvan Massive mgt/skarn ore NRM-880071 1.1 ±0.2

Rödbergsgruvan Massive mgt/skarn ore NRM-19984100 0.6 ±0.2

52

5. Discussion

5.1. Mineralogy, mineral chemistry and bulk geochemistry

Table 3. Summary of analyses performed on studied samples, as well as their dominant mineralogy. Abbreviations used for

analyses are: BGC – bulk geochemical analyses, ISO – stable isotope analyses, XRD – powder X-ray diffraction, OM – optical

microscopy, SEM-EDS – scanning electron microscopy with energy dispersive spectrometer, WDS – electron probe

microanalyser with wavelength dispersive spectrometer. Mineral abbreviations used are: Qtz - quartz, Mgt - magnetite, Hem

- hematite, Cb - carbonates, Am - amphibole, Flb - fluorbritholite, Dla - Dollaseite, Bas - bastnäsite, Hu - humite group

minerals, Urn - uraninite, Fl - fluorite, Dis - dissakisite Gad - gadolinite, Väs - västmanlandite, FC - fluorocarbonates, Fln –

ferriallanite, Bt - biotite.

Deposit Sample ID Analyses Dominant mineralogy

Danielsgruvan NRM-20020125 BGC, ISO, OM Qtz, Mgt, Hem. Minor Am-skarn, LREE-silicates, Sulfides Danielsgruvan NRM-20020124 BGC, ISO, OM Qtz, Mgt, Hem. Minor Am-skarn, LREE-silicates, Sulfides

Östanmossa EJ-ÖM90-13-1 BGC, ISO, XRD, OM, SEM-EDS Cb, Am-Fl-skarn, Mgt, Flb-(Y), Dla-(Ce), Bas-(Ce). Minor Hu Östanmossa SGU-M4528 BGC, ISO, XRD, OM, SEM-EDS Cb, Am-Fl-skarn, Mgt, Flb-(Y), Dla-(Ce), Bas-(Ce). Minor Hu, Urn

Östanmossa EJ-ÖM90-13-2 BGC, ISO, XRD, OM Cb, Am-Fl-skarn, Mgt, Flb-(Y), Dla-(Ce), Bas-(Ce). Minor Hu

Östanmossa SGU-M4441 BGC, ISO, XRD, OM Cb, Am-Fl-skarn, Flb-(Y), Dla-(Ce), Bas-(Ce). Minor Hu Östanmossa SGU-M4529 ISO, XRD Cb, Am-Fl-skarn, Flb-(Y), Dla-(Ce), Bas-(Ce). Minor Hu

S. Hackspiksgr. SGU-M3563 BGC, ISO, OM, SEM-EDS Fl-Am-skarn, Mgt, Dis-(Ce), Cer-(Ce), Bas-(La)/(Ce), Flb-(Ce), Sulfides

Johannagruvan Joha-KH-1 BGC, ISO, OM, SEM-EDS, WDS Am-Qtz-skarn, Dla-(Ce), Mgt, Bas-(Ce), Gad-(Nd), Sulfides Johannagruvan Joha-KH-2 BGC, ISO, OM Am-Qtz-skarn, Nrb, Dla-(Ce), Mgt, Bas-(Ce), Gad-(Nd), Sulfides

Malmkärra SGU-M4068 BGC, ISO, XRD, OM, SEM-EDS Cb, Fl-Am-skarn, Mgt, Väs-(Ce), Flb-(Ce). Minor Sulfides, LREE-FC, Urn Malmkärra SGU-M4048 BGC, ISO, XRD, OM, SEM-EDS Cb, Fl-Am-skarn, Mgt, Väs-(Ce), Flb-(Ce). Minor Sulfides, LREE-FC, Urn

Högfors Högf-KH-1 BGC, ISO, OM Mgt, Hem, Qtz. Minor Am-skarn, LREE- silicates/FC, Y-HREE-silicates Högfors Högf-KH-2 ISO Mgt, Hem, Qtz. Minor Am-skarn, LREE- silicates/FC, Y-HREE-silicates

Bastnäs Bast-EJ-1 BGC, ISO, OM Am-skarn, Mgt, Fln-(Ce), Sulfides Bastnäs SGU-M6777 ISO, OM Mgt, Am-skarn, Fln-(Ce), Sulfides

Bastnäs SGU-M309 BGC, ISO, OM, SEM-EDS Am-skarn, Mgt, Fln-(Ce), Sulfides

Myrbacksfältet Myrb-EJ-1 BGC, ISO, OM Mgt, Sulfides, Am-Qtz-skarn. Minor LREE-silicates Myrbacksfältet Myrb-EJ-2 ISO Mgt, Sulfides, Am-Qtz-skarn. Minor LREE-silicates

Östra Gyttorp Gytt-EJ-1a ISO, OM, SEM-EDS, WDS Mgt, Bt-Am-skarn. Minor LREE-FC, Urn Östra Gyttorp Gytt-EJ-1b ISO, OM Mgt, Bt-Am-skarn. Minor LREE-FC, Urn

Östra Gyttorp Gytt-EJ-1c ISO Mgt, Bt-Am-skarn. Minor LREE-FC, Urn Östra Gyttorp Gytt-EJ-2 BGC, OM, SEM-EDS, WDS Am-Bt-skarn, Aln-(Ce), Mgt, Gad-(Y). Minor LREE-FC, Zrn, Urn

Rödbergsgruvan NRM-880071 BGC, ISO, OM, SEM-EDS Mgt, Am-skarn, Aln-(Ce), Bas-(Ce), Cer-(Ce). Minor Sulfides Rödbergsgruvan NRM-19984100 BGC, ISO Mgt, Am-skarn, Aln-(Ce), Bas-(Ce), Cer-(Ce). Minor Sulfides

The studied mineral parageneses in this work show variable primary ore mineral assemblages

of REE-silicates, REE-fluorocarbonates, iron oxides and sulphides occurring together with

carbonates and skarn minerals (Table 3). The deposits have been variably affected by later

processes, visible mainly as recrystallisation textures as well as mineral replacements by e.g.

fluorocarbonates and fluorite. Similar results as Holtstam & Andersson (2007) were found in

samples from the deposits they covered, with assemblages rich in LREE and Fe found in

subtype 1 deposits and assemblages rich in Mg, Ca, F and HREE+Y in subtype 2 deposits.

Additionally, whole rock geochemical data on REE-silicate bearing assemblages verifies their

53

suggestion that subtype 2 deposits are richer in HREE, indicated by average Ce/Yb-ratios that

are roughly two to three times lower compared to subtype 1 deposits.

The previously relatively unknown Östra Gyttorp deposit displays a mix of features, partly

represented by both the subtype 1 deposits and the subtype 2 deposits (Holtstam & Andersson

2002). Allanites are Mg-poor, typical of subtype 1. Gadolinites/hingganites are very rich in

Y+HREE, even more so than subtype 2-deposits. Whole rock REE indicate similar Ce/Yb-

ratios as subtype 2 deposits, but remnants of any possible host carbonates and larger amounts

of coherent skarn are absent (Nordenström 1890). Finally, the assemblages are generally F-

poor, with little incorporation into e.g. allanites and only minor amounts of fluorocarbonates,

fluorite or humite group minerals.

Important for this study was to investigate and characterise parageneses and textural

relationships of the minerals analysed for isotopic compositions. It can be concluded that

magnetite and quartz used are primary minerals crystallised during ore formation, and that they

occur in textural equilibrium with each other. The carbonates however could either have a

sedimentary, exhalative precipitation or even biological origin. They are also prone to

recrystallise under various post-depositional conditions. Therefore they cannot be concluded to

be in isotopic equilibrium with magnetite soley based on textural relationships. The magnetite,

quartz and carbonate analysed for isotopic composition were mainly collected from massive

units.

5.2. Uraninite geochronology

Provided that mineralisation is synchronous with the 1.90-1.87 Ga Svecofennian syn-volcanic

stage (Jonsson & Högdahl 2013), only two of the calculated chemical ages are valid (1898 Ma

and 1869 Ma). A previous attempt to date uraninite from Bastnäs by U-Pb mass spectrometry

was done by Welin (1963). The isotopic analysis yielded a discordant result with a 207Pb/206Pb

date of 1785 Ma (later corrected to 1760 Ma using modern decay constants; Welin 1980). The

sample size was large (5-10 g) and the uraninites were reported to be altered. The uraninites

used for chemical dating were often partly to wholly replaced by bastnäsite-(Ce) (Figs 4d, 21e

and 23d ), which is also indicated by the high REE contents of the uraninites (Table 11, appx).

It is therefore likely that the uraninites crystallised at an early stage of mineralisation (around

1.90 Ga), and were later affected by fluids during regional metamorphism and/or other

54

subsequent tectonothermal events (e.g. TIB magmatism). These later events lead to, among

others, replacement by fluorocarbonates and likely a loss of lead in the analysed grains (Fig

26). Another possibility is that during these later events the radiometric clock in the uraninites

was reset due to e.g. recrystallisation (Fig 26), however this is not supported by textural

observations. It can be concluded that chemical dating of uraninites in these ores is not very

reliable due to the extensive alteration as an affect of later events, and that dating of other

minerals and/or using isotopic analytical methods is needed. However, despite this, some

support is provided for mineralisation during the early Svecofennian, since it is not very likely

that uraninites have gained U and Th after crystallisation, thus indicating that the c. 1.90 Ga

date should be an approximate minimum age for these deposits.

5.3. Stable isotope systematics

5.3.1. Thermometry

Using the δ18O-data, it is possible to calculate equilibrium temperatures for mineral pairs of

magnetite-quartz and magnetite-calcite/dolomite (Eqs 5-6). However, the use of magnetite-

calcite/dolomite pairs requires some caution, due to the problems stated in section 5.1. For

further discussion, see section 5.3.2. Results indicate that crystallisation of these mineral pairs

took place between c. 650 to 520 °C (Table 4). Published fluid inclusion data suggest minimum

temperatures of crystallisation of bastnäsite at 400-300 °C and fluorite at 150-100 °C (Holstam

& Broman 2002, Andersson et al. 2013). These minerals are texturally interpreted to be, at least

in part, secondary in relation to other REE-minerals (Andersson et al. 2004), and the data is not

pressure-corrected. It is therefore likely that formation of primary assemblages occurred at

higher temperatures. Thus primary magnetite assemblages can, by combining stable isotope-

and fluid inclusion data, be inferred to have crystallised between c. 650 and 400 °C.

55

Table 4. Calculated equilibrium temperatures for mineral pairs and modelled equilibrium fluids.

Deposit Sample ID δ18Omgt (‰) δ18Oqtz (‰) δ18Ocarb (‰) T (°C) δ18O fluid( ‰)

Danielsgruvan NRM-20020125 -0.7 7.5 602 ± 211 5.64-5.63

Danielsgruvan NRM-20020124 -1.9 400-650 4.13-5.73

Östanmossa EJ-ÖM90-13-1 -1.2 7.2 571 ± 202 5.45-5.43

Östanmossa SGU-M4528 -1.4 5.8 624 ± 242 4.55-4.83

Östanmossa EJ-ÖM90-13-2 6.4 400-650 4.35-5.55

Östanmossa SGU-M4441 6.8 400-650 4.75-5.95

Östanmossa SGU-M4529 5.9 400-650 3.35-4.95

S. Hackspiksgruvan SGU-M3563 0.5 400-650 6.53-8.13

Johannagruvan Joha-KH-1 -0.5 400-650 5.63-7.23

Johannagruvan Joha-KH-2 0.5 400-650 6.53-8.13

Malmkärra SGU-M4068 -0.6 6.8 615 ± 232 5.55-5.73

Malmkärra SGU-M4048 0.2 7.0 649 ± 252 6.05-6.23

Högfors Högf-KH-1 -1.7 8.3 522 ± 161 5.23-5.24

Högfors Högf-KH-2 -0.4 7.2 641 ± 231 5.83-5.84

Bastnäs Bast-EJ-1 -0.6 400-650 5.53-7.03

Bastnäs SGU-M6777 -1.2 400-650 4.93-6.53

Bastnäs SGU-M309 0.0 400-650 6.03-7.63

Myrbacksfältet Myrb-EJ-1 0.7 400-650 6.73-8.33

Myrbacksfältet Myrb-EJ-2 1.1 400-650 7.23-8.83

Östra Gyttorp Gytt-EJ-1a -1.8 400-650 4.23-5.83

Östra Gyttorp Gytt-EJ-1b -2.3 400-650 3.73-5.33

Östra Gyttorp Gytt-EJ-1c -2.1 400-650 3.93-5.53

Rödbergsgruvan NRM-880071 1.1 400-650 7.13-8.73

Rödbergsgruvan NRM-19984100 0.6 400-650 6.63-8.23

1Calculated based on mgt-qtz fractionation (Clayton et al. 1989 and Chiba et al. 1989) 2Calculated based on mgt-cc/dol fractionation (Chiba et al. 1989 and Sheppard & Schwarz 1970) 3Calculated based on mgt-water fractionation (Zheng 1991) 4Calculated based on qtz-water fractionation (Chiba et al. 1989, Clayton et al. 1989 and Zheng 1991) 5Calculated based on cc/dol-water fractionation (Chiba et al. 1989, Zheng 1991, Sheppard & Schwarz 1970 and Clayton et al. 1975)

56

5.3.2. Fluid modeling

If it is assumed that the mineralising fluid was magmatic in origin, it should have had a δ18O-

value between approximately 6 and 8 ‰ (Hoefs 1997). Equilibrium fluids were calculated (Eqs

6, 8) for mineral pairs (from Högfors, Malmkärra, Danielsgruvan and Östanmossa samples),

and the ranges are presented in Table 4. Two equilibrium fluids were calculated based on

magnetite-water fractionation and quartz-water or calcite/dolomite-water fractionation,

respectively; the latter depending on which mineral was used as mineral pair. Fluid modelling

was done for the specific temperatures calculated for each mineral pair. The results indicate that

only one sample (SGU-M4048 from Malmkärra) is in equilibrium with a magmatic fluid, while

the others are in equilibrium with fluids of slightly lower δ18O (4.5-6 ‰). Equilibrium fluids

for both minerals in all mineral pairs lie within standard analytical error, indicating that all

mineral pairs are in or very near isotopic equilibrium. This is especially important for the

carbonates, since a textural equilibrium between carbonates and magnetite cannot be easily

established. Another indication that carbonates obtained isotopic equilibrium with magnetite is

that the calculated temperatures for magnetite-carbonate mineral pairs (571-649 °C) essentially

overlap with temperatures from magnetite-quartz pairs (522-641 °C), where the latter are in

textural equilibrium. Therefore the magnetite-carbonate pairs are interpreted as if they obtained

isotopic equilibrium, despite the difficulties of proving this texturally.

Singular analysed magnetites (and two carbonates) were modeled for equilibrium fluids in the

temperature range 650-400 °C (Table 4, Fig 27). It is shown that some deposits

(Rödbergsgruvan, Myrbacksfältet, Södra Hackspiksgruvan, Johannagruvan (one sample)) are

in equilibrium with magmatic fluids (or fluids of slightly higher δ18O) for the entire temperature

range. All samples from Östra Gyttorp, Danielsgruvan and Östanmossa are in turn in

equilibrium with lower-δ18O fluids (3-6 ‰) for the entire temperature range. The remaining

deposits (Bastnäs, Johannagruvan (one sample)) could be in equilibrium with both magmatic

and lower-δ18O fluids, this depends on the temperature, with low temperatures (i.e. 400 °C)

giving higher δ18O in the mineralising fluid and vice versa.

57

Fig 27a-b. Diagram over modelled equilibrium fluids from Table 4. Equilibrium fluids for minerals in mineral

pairs were calculated for the specific equilibrium temperatures for each mineral pair. Singular analysed

magnetites and carbonates were modelled for the temperature interval 650 °C (Fig 27a) to 400 °C (Fig 27b).

Inserted are also fields for most magmatic fluids and unmineralised host carbonate rocks, as well as vectors

for the expected fluid isotopic shift due to mixing with low-δ18O fluids and fluid-host carbonate interaction.

58

Assuming that both the remnant host carbonates (Andersson et al. 2013) and carbonates in REE

mineralisation were originally of typical composition of unmineralised carbonate rocks in

Bergslagen (NRSC; De Groot & Sheppard 1988), it is clear that they have experienced a

significant shift in isotopic compositions towards lighter δ18O and δ13C values (Fig 28). Some

carbonates plot in the field of carbonates from Fe-oxide skarn deposits (CFSD; De Groot &

Sheppard 1988) while others plot in the field of carbonates from granite-related W-Mo-bearing

skarn deposits (CGSD; De Groot & Sheppard 1988). Modeling of equilibrium decarbonation

(Rayleigh decarbonation to the “calc-silicate limit”, lines a and b in Fig 28; Valley 1986) show

that pure thermal breakdown of original host carbonate rocks is not enough to explain the

observed shifts in isotopic composition of carbonates from the REE-line. Lines 1 and 2 (Fig 28)

represent mixing between original host carbonate rocks and a hydrothermal fluid with a

magmatic isotopic composition. It is shown that all analysed carbonates from the REE-line are

enveloped by these mixing curves, implying that reactions between host carbonates and a fluid

of a pre-dominantely magmatic origin are responsible for formation of the skarn units.

59

The results indicate that the mineralising fluid was originally magmatic-dominated.

Additionally, the overall geochemical character of the fluid (see section 5.4) is also compatible

with a magmatic source. However, it is clear that there have been variable modification of the

fluid on a local scale (Fig 27). There are mainly two processes that can be considered to have

been involved in shifting the isotopic signature of the mineralising fluid and the precipitating

minerals. These are 1) mixing with low-δ18O fluids, such as seawater or meteoric water

(generally δ18O ≤ 0; White 1961) and 2) fluid-rock interactions. Mixing with low-δ18O fluids

would lead to a decrease in the δ18O of the mineralising fluid (Fig 27). As calculated equilibrium

fluids generally trend from magmatic δ18O-values towards lower values, it can be suggested

that fluid mixing has been involved to variable degrees in the different deposits, leading to the

Fig 28. Plot of δ18O versus δ13C in carbonates from REE mineralisation (diamonds; this study) and remnant

host carbonate rocks of the “REE-line” (triangles; Andersson et al. 2013). Added are also carbonates from

Bergslagen (dots; De Groot & Sheppard 1988). These include NRSC (stratiform carbonates not related to ore

deposits), RSC (carbonates related to stratiform oxide deposits), CFSD (carbonates from Fe-oxide skarn

deposits) and CGSD (carbonates from granite-related W-Mo-bearing skarn deposits). Also included are fields

for most Proterozoic calcites and most Proterozoic dolomites (MPC and MPD; De Groot & Sheppard 1988

and references therein).

Inserted are mixing lines 1-2 showing binary mixing in 5 % increments between original host carbonate rocks

(δ18O = 12.5 to 19.6 ‰ and δ13C = -1.8 to 1.7 ‰; De Groot & Sheppard 1988) and a magmatic source (δ18O

= 5.3 to 7.3 ‰ and δ13C = -8 to -4 ‰; Taylor et al. 1986). In 1) both are of average composition and in 2)

both are of end-member compositions. Also inserted are Rayleigh decarbonation lines a-b, showing pure

thermal decarbonation to the “calc-silicate limit” of original host carbonates of average composition using

fractionation factors α13C(CO2-rock) = 1.0022 and α18O(CO2-rock) = 1.006 to 1.012 (Valley 1986). All

calculations were done using the methodology of Jolis et al. (2013), and references therein. Figure modified

after De Groot & Sheppard (1988), Jolis et al. (2013) and Andersson et al. (2013).

60

negative shifts in δ18O of some of the minerals and their calculated equilibrium fluids. The

possible fluid-rock interactions are harder to quantify. It is possible that some interaction

occurred during fluid transport, from the source to the area of mineralisation. Furthermore, it is

shown that there has been extensive interaction between the mineralising fluid and the host

carbonate rocks (Fig 28), which should lead to an increase in both δ18O and δ13C of the fluid

(Fig 27). However, since this interaction occurs essentially during mineralisation, it is possible

that this process only marginally affected mineral δ18O-compositions. If the analysed minerals

were removed from the fluid at an early stage during mineralisation, they would not show any

significant increase in the calculated equilibrium fluid δ18O-composition. Additionally, any

possible increase in δ18O is masked by the negative shifts due to mixing with low-δ18O fluids.

However, assuming that the analysed carbonates were in equilibrium with the mineralising

fluid, the slight enrichment in δ13C in some carbonates compared to magmatic values (Fig 28)

could be a sign of host rock input, since low-δ18O fluids would not have any significant impact

on the carbon isotope composition (J. Hoefs, personal communication 2013).

5.3.2.1. The nature of the low-δ18O fluid

With the data obtained in this study it is impossible to distinguish between different types of

low-δ18O fluids, e.g. seawater or meteoric water, since both lead to a lowering of the δ18O and

they have insignificant impact on the δ13C. One way of testing which type of water that could

have shifted δ18O to lower values is by using hydrogen isotopes on cogenetic phases. In a δD

vs. δ18O-plot these different waters have distinct signatures (see Hoefs 1997), and suitable

minerals for this could be e.g. skarn amphiboles.

However, it has been established that most Svecofennian, volcanic-associated ores in

Bergslagen were deposited mainly in a subaqueous setting in a back arc tectonic environment

(Allen et al. 1996), making it very likely that the ores of the REE-line too were deposited in a

marine basin with (essentially) unlimited access to marine waters. Furthermore, the very

prominent Mg-metasomatism in the area has been attributed to seawater-dominated fluids

(Trägårdh 1988, Trägårdh 1991, Ripa 1994). Therefore, seawater appears to be a probable

candidate in the ore forming process.

61

5.3.3. Effects of metamorphism

The temperatures derived from thermometric calculations based on mineral pairs lie in the same

range as regional amphibolite facies metamorphism, and it is therefore necessary to address the

possible effects of metamorphism on the isotopic signatures. A metamorphic overprint is visible

texturally to variable extents in the different deposits. However, reaction of minerals with later

metamorphic fluids is unlikely to completely re-equilibrate mineral oxygen isotopic signatures,

since oxygen rich minerals such as magnetite, quartz and (at least when occurring in massive

form) carbonates should be expected to act as δ18O-buffers (cf. Jonsson et al. 2012, Zheng et

al. 1999), and very large amounts of fluid would be required. Additionally, particularly

magnetite is a refractory and chemically robust phase. The buffering ability of these minerals

is especially great when they occur in massive units, as the majority of samples used in this

study. Overprinting by a metamorphic fluid would more likely only partially alter the isotopic

composition of minerals, thereby removing the mineral pairs out of isotopic equilibrium (cf.

Zheng et al. 1999). The mineral pairs in this study are interpreted to be in isotopic equilibrium.

Therefore any metamorphic overprinting has most likely insignificantly affected the isotopic

compositions, and the results are interpreted to reflect the conditions during original mineral

formation.

5.4. Fluid chemistry and ore mineralogy

As pointed out by Holtstam & Andersson (2007), mineral assemblages indicate that the

mineralising fluid contained a variety of ligands, including F, Cl, CO2, S, Si and OH, which all

could have played a role in transporting the metals (mainly REE and Fe together with minor

Al, Cu, Mo, Zn, Co, Bi…). Ca and Mg were most likely supplied by the host carbonate rocks,

as magmatic waters are normally depleted in these elements (White 1957). High F-

concentrations in the fluid require depletion in Ca owing to the very low solubility of CaF2

(Plimer 1984). The relative importance of the ligands would depend on their relative

concentrations, but also on pH and temperature, with Cl forming stronger complexes in strongly

acidic fluids, F in slightly acidic to neutral fluids and OH in basic fluids (Haas et al. 1995).

The slightly contrasting assemblages between the two subtypes are interpreted by Holtstam &

Andersson (2007) to be related to variability in both concentrations of F in the mineralising

fluid and in fluid:host carbonate ratios, with subtype 2 deposits experiencing initial fluids with

62

higher F-concentrations and lower fluid:host carbonate ratios. The lower fluid:host carbonate

ratios would lead to high concentrations of Mg and Ca from fluid interaction with the host

carbonates (Holtstam & Andersson 2007). Ca in turn would destabilise F-complexes leading to

deposition of large amounts of HREE+Y, REE-fluorocarbonates, fluorite, and F-bearing

silicates (cf. Williams-Jones et al. 2000), while Mg could be incorporated into e.g. allanite

group minerals. This suggests that fluid:carbonate interaction during mineralisation was a very

important factor controlling the final mineral paragenesis.

As suggested by Holstam & Andersson (2007), variable fluid composition was also important

for mineral parageneses in the deposits. Not only could there have been variability in

concentrations of certain crucial elements (such as F), but variable mixing with external fluids

(e.g. seawater, as indicated by this study) and a variability of the hydrothermal facies could also

be relevant for the mineralogical differences in the deposits along the REE-line. These

additional processes could possibly explain the seemingly contradictory features of the Östra

Gyttorp deposit (i.e. high Y+HREE but low F, Ca, Mg), since oxygen isotope data suggest that

mixing with external waters has been extensive in this mineralisation relative to other deposits.

Additonally, in Östra Gyttorp the REE-rich assemblages occur within felsic metavolcanic

rocks, and show rather small amounts of coherent skarn and an absence of remnant host

carbonates (Nordenström 1890). This suggests a relatively small, if any, input from host

carbonates on the mineral assemblages. The stability and relative importance of different

complexing species for REE is highly dependent on fluid chemistry (mainly pH and T; Haas et

al. 1995), it could therefore be possible that the external fluids changed the chemistry of the

mineralising fluid enough to destabilise F-complexes carrying preferentially HREE+Y.

Seawater should have near neutral pH for temperatures up to 250 °C (Bischoff & Seyfried 1978,

Plimer 1978). Assuming a slightly acidic magmatic fluid mixing with seawater, a rise of the pH

should be expected, together with a decrease in temperature (Plimer 1984). This would likely

decrease the stability of REE-F complexes, and especially HREE-F complexes are much more

sensitive to increases in pH than LREE-F complexes (Haas et al. 1995). The host carbonates (if

any) and the seawater probably provided an insignificant increase in fluid Ca, Mg-concentration

compared to for the other deposits, thus the concentrations of Ca, Mg in the fluid were too low

for formation of any larger amounts of fluorite, fluorocarbonates and Mg-F-bearing silicates.

Therefore F could remain (largely) in the fluid, forming more stable complexes with other

elements, while HREE+Y could be incorporated into crystallising minerals.

63

Secondary re-distribution of REE and other elements during later events has surely also been

important. This lead to formation of e.g. REE-fluorocarbonates in fractures and as mineral

replacements. It can likely also explain the heterogeneous composition and zonation in minerals

like allanite (cf. Fig 4b, Figs 6-7) as well as the observed lead loss in uraninites.

6. Conclusions

The results of this study suggest that the Bastnäs-type deposits formed through reactions

between pre-existing carbonate rocks and originally magmatic-dominated fluids at relatively

high temperatures (c. 650-400 °C) at approximately 1.9 Ga. Different deposits experienced

variable shifts of the mineralising fluids isotopic signature due to mixing with available

seawater-dominated fluids.

The overall mineralogy of studied REE-mineralised assemblages from the various deposits

indicate that metals were transported in a fluid containing a variety of complexing ligands.

Differences in mineral paragenesis between different deposits is suggested to be connected to

variability in: ligand concentrations in mineralising fluids, extent of fluid-rock interaction,

hydrothermal facies and degree of mixing with seawater. This lead to variable concentrations

and stabilities of crucial mineral forming elements in the fluid, which is reflected in the

somewhat heterogeneous mineralogy of the deposits within the REE-line. Secondary processes

during later events have also affected the mineralogy to different extents.

64

7. Acknowledgements

First and foremost I would like to give a big thank you to my supervisors, Prof. Erik Jonsson

and Doc. Karin Högdahl, for letting me work on this interesting project and for all the help and

inspiration. I would also like to thank Prof. Chris Harris for analysing my samples and for

answering many of my questions throughout the work. Dr. Jaroslaw Majka, Doc. Abigail

Barker and Gary Wife all helped me tremendously during my analytical work – thank you! I

owe Franz Weis, Prof. Valentin Troll, Dr. Ester Jolis and Prof. Jochen Hoefs my gratitude for

kindly helping me with various problems I encountered during my project. Finally, I would like

to thank my family for always supporting me.

Samples for the study were generously supplied by the Geological Survey of Sweden, the

Swedish Museum of Natural History, Prof. Erik Jonsson and Doc. Karin Högdahl.

The study was financially supported by Uppsala University, the Geological Survey of Sweden

and Vetenskapsrådet.

65

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72

9. Appendix

Table 5. WDS data for allanites from the Östra Gyttorp deposit. All analysed points are from sample Gytt-

EJ-2. Calculations are based on 6 M+T cations (see Ercit 2002).

Point 1 Point 2 Point 3 Point 4 Point 5 Point 6 Point 7 Point 8 Point 9 Point 10

SiO2 (wt %) 31.38 31.64 31.64 32.13 32.26 31.49 31.62 32.17 31.51 31.75

Al2O3 15.15 15.54 15.34 16.59 16.64 14.94 15.44 16.67 15.62 16.01

MgO 1.33 1.39 1.41 1.25 1.31 1.43 1.35 1.31 1.42 1.24 Fe2O3 4.40 3.24 2.87 5.18 6.17 2.35 3.84 6.30 5.36 3.95

FeO 9.55 10.11 10.51 8.11 7.15 11.10 9.65 7.22 8.17 9.57

MnO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 UO2 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.01 0.00 0.00

ThO2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

CaO 10.52 10.64 10.62 11.70 11.31 10.13 10.70 11.23 10.29 11.00 La2O3 4.81 4.89 5.03 4.08 4.52 5.34 4.46 4.72 4.95 4.48

Ce2O3 12.50 13.36 12.85 10.80 11.36 13.01 12.54 11.41 12.40 11.30

Pr2O3 1.39 1.51 1.40 1.35 1.41 1.50 1.58 1.36 1.47 1.42 Nd2O3 5.72 5.51 5.65 5.43 4.90 5.56 5.95 4.93 5.69 5.38

Sm2O3 0.94 0.83 0.92 1.08 0.77 0.86 0.94 0.79 0.99 1.15

Eu2O3 0.14 0.12 0.12 0.11 0.12 0.15 0.12 0.12 0.12 0.12 Gd2O3 0.47 0.25 0.37 0.66 0.35 0.46 0.39 0.33 0.41 0.68

Tb2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Dy2O3 0.13 0.05 0.06 0.12 0.07 0.06 0.03 0.06 0.07 0.11 Ho2O3 0.03 0.00 0.00 0.00 0.01 0.04 0.00 0.02 0.00 0.00

Er2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00 Tm2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Yb2O3 0.01 0.00 0.02 0.01 0.00 0.02 0.00 0.01 0.00 0.00

Lu2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.01 Y2O3 0.33 0.17 0.26 0.43 0.38 0.27 0.23 0.34 0.20 0.45

H2O* 1.56 1.57 1.57 1.61 1.61 1.56 1.57 1.61 1.57 1.59

F 0.00 0.05 0.08 0.00 0.07 0.16 0.02 0.09 0.00 0.17 O=-F* 0.00 -0.02 -0.03 0.00 -0.03 -0.07 -0.01 -0.04 0.00 -0.07

Total 100.35 100.80 100.61 100.63 100.34 100.22 100.42 100.61 100.23 100.13

Si (apfu) 3.01 3.02 3.02 3.00 3.00 3.03 3.02 2.99 3.01 3.00

Al 1.71 1.75 1.73 1.83 1.83 1.70 1.74 1.83 1.76 1.78 Mg 0.19 0.20 0.20 0.17 0.18 0.21 0.19 0.18 0.20 0.18

Fe3+ 0.32 0.23 0.21 0.36 0.43 0.17 0.28 0.44 0.38 0.28

Fe2+ 0.77 0.81 0.84 0.63 0.56 0.89 0.77 0.56 0.65 0.76 Mn 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

U 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Th 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Ca 1.08 1.09 1.09 1.17 1.13 1.05 1.10 1.12 1.05 1.11

La 0.17 0.17 0.18 0.14 0.16 0.19 0.16 0.16 0.17 0.16

Ce 0.44 0.47 0.45 0.37 0.39 0.46 0.44 0.39 0.43 0.39 Pr 0.05 0.05 0.05 0.05 0.05 0.05 0.06 0.05 0.05 0.05

Nd 0.20 0.19 0.19 0.18 0.16 0.19 0.20 0.16 0.19 0.18

Sm 0.03 0.03 0.03 0.03 0.02 0.03 0.03 0.03 0.03 0.04 Eu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Gd 0.01 0.01 0.01 0.02 0.01 0.01 0.01 0.01 0.01 0.02

Tb 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Dy 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Ho 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Er 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Tm 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Yb 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Lu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Y 0.02 0.01 0.01 0.02 0.02 0.01 0.01 0.02 0.01 0.02

OH* 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00

F 0.00 0.02 0.02 0.00 0.02 0.05 0.01 0.03 0.00 0.05 ∑REE 0.93 0.93 0.93 0.82 0.81 0.96 0.91 0.82 0.91 0.87

∑cations 8.01 8.01 8.02 7.99 7.94 8.00 8.01 7.94 7.97 7.98

73

Table 5 continued.


SiO2 (wt %) 31.61 32.35 32.46 31.63 31.26 32.15 32.13 31.12 31.07 32.04

Al2O3 15.58 16.66 16.89 15.77 16.17 16.67 16.24 15.24 15.13 16.09 MgO 1.44 1.42 1.29 1.34 1.48 1.21 1.46 1.44 1.42 1.32

Fe2O3 4.75 8.41 8.13 4.85 6.78 5.58 8.15 3.94 4.50 6.48

FeO 8.76 5.41 5.52 8.52 6.89 7.64 5.63 9.50 9.06 7.34 MnO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

UO2 0.00 0.03 0.00 0.02 0.00 0.00 0.01 0.00 0.00 0.00

ThO2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 CaO 10.54 11.22 11.34 10.51 10.10 11.64 11.14 9.75 9.77 10.81

La2O3 4.90 4.59 4.71 4.84 5.05 5.43 4.57 4.22 4.04 4.60

Ce2O3 12.41 11.18 11.00 12.37 13.00 11.53 11.20 14.14 13.46 11.28 Pr2O3 1.45 1.31 1.21 1.50 1.51 1.16 1.32 1.72 1.64 1.36

Nd2O3 5.82 4.98 4.66 5.46 5.43 4.31 4.88 6.21 6.48 5.13

Sm2O3 0.97 0.79 0.76 0.94 0.85 0.55 0.74 0.77 0.93 1.05 Eu2O3 0.11 0.13 0.13 0.10 0.12 0.14 0.11 0.13 0.10 0.13

Gd2O3 0.47 0.33 0.36 0.46 0.38 0.30 0.32 0.24 0.28 0.60

Tb2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Dy2O3 0.04 0.11 0.08 0.02 0.05 0.00 0.06 0.02 0.01 0.12

Ho2O3 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.03 0.00 0.03

Er2O3 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Tm2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Yb2O3 0.01 0.01 0.00 0.00 0.00 0.01 0.03 0.00 0.00 0.00

Lu2O3 0.00 0.00 0.01 0.00 0.00 0.01 0.00 0.01 0.00 0.00 Y2O3 0.19 0.30 0.41 0.29 0.17 0.25 0.34 0.15 0.16 0.55

H2O* 1.58 1.62 1.63 1.57 1.58 1.60 1.61 1.55 1.55 1.60 F 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.06 0.07 0.04

O=-F* 0.00 0.00 0.00 0.00 0.00 0.00 0.00 -0.03 -0.03 -0.02

Total 100.62 100.86 100.60 100.18 100.83 100.18 99.93 100.16 99.58 100.52

Si (apfu) 3.01 2.99 3.00 3.01 2.96 3.01 3.00 3.00 3.01 3.00 Al 1.75 1.81 1.84 1.77 1.80 1.84 1.79 1.73 1.73 1.78

Mg 0.20 0.20 0.18 0.19 0.21 0.17 0.20 0.21 0.21 0.18

Fe3+ 0.34 0.58 0.56 0.35 0.48 0.39 0.57 0.29 0.33 0.46

Fe2+ 0.70 0.42 0.43 0.68 0.55 0.60 0.44 0.77 0.73 0.58

Mn 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

U 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Th 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Ca 1.08 1.11 1.12 1.07 1.02 1.17 1.11 1.01 1.01 1.09

La 0.17 0.16 0.16 0.17 0.18 0.19 0.16 0.15 0.14 0.16 Ce 0.43 0.38 0.37 0.43 0.45 0.39 0.38 0.50 0.48 0.39

Pr 0.05 0.04 0.04 0.05 0.05 0.04 0.04 0.06 0.06 0.05

Nd 0.20 0.16 0.15 0.19 0.18 0.14 0.16 0.21 0.22 0.17 Sm 0.03 0.03 0.02 0.03 0.03 0.02 0.02 0.03 0.03 0.03

Eu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Gd 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.02 Tb 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Dy 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Ho 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Er 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Tm 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Yb 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Lu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Y 0.01 0.01 0.02 0.01 0.01 0.01 0.02 0.01 0.01 0.03

OH* 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 1.00 F 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.02 0.01

∑REE 0.91 0.80 0.79 0.90 0.92 0.81 0.80 0.97 0.95 0.85

∑cations 7.99 7.91 7.91 7.98 7.94 7.98 7.92 7.98 7.97 7.94

74

Table 5 continued.

Point 21 Point 22 Point 23 Point 24 Point 25 Point 26

SiO2 (wt %) 31.53 32.48 32.18 32.89 31.81 32.43

Al2O3 15.80 17.42 16.80 17.56 15.74 17.17

MgO 1.40 1.15 1.37 1.14 1.31 1.33 Fe2O3 4.36 5.67 7.42 7.71 4.27 8.15

FeO 9.00 7.27 6.00 5.94 9.25 5.36

MnO 0.00 0.00 0.00 0.00 0.00 0.00 UO2 0.00 0.01 0.00 0.05 0.00 0.03

ThO2 0.00 0.00 0.00 0.00 0.00 0.00

CaO 10.16 12.25 11.37 12.15 10.68 11.75 La2O3 4.76 4.25 5.18 4.82 3.12 5.23

Ce2O3 13.35 10.86 11.15 10.64 11.83 10.54

Pr2O3 1.59 1.27 1.16 1.17 1.77 1.18 Nd2O3 5.72 4.64 4.46 4.28 7.17 4.18

Sm2O3 0.85 0.76 0.65 0.69 1.25 0.65

Eu2O3 0.14 0.10 0.16 0.14 0.13 0.13 Gd2O3 0.40 0.38 0.30 0.37 0.59 0.30

Tb2O3 0.00 0.00 0.00 0.00 0.00 0.00

Dy2O3 0.03 0.06 0.08 0.08 0.10 0.07 Ho2O3 0.01 0.00 0.08 0.00 0.00 0.00

Er2O3 0.00 0.00 0.01 0.00 0.00 0.00

Tm2O3 0.00 0.00 0.00 0.00 0.00 0.00 Yb2O3 0.01 0.02 0.02 0.00 0.02 0.00

Lu2O3 0.00 0.02 0.00 0.00 0.00 0.01 Y2O3 0.30 0.40 0.37 0.35 0.38 0.41

H2O* 1.58 1.63 1.62 1.65 1.58 1.63

F 0.02 0.00 0.00 0.00 0.00 0.00 O=-F* -0.01 0.00 0.00 0.00 0.00 0.00

Total 100.99 100.62 100.38 101.63 101.00 100.55

Si (apfu) 3.00 2.99 2.99 2.99 3.02 2.98

Al 1.77 1.89 1.84 1.88 1.76 1.86 Mg 0.20 0.16 0.19 0.15 0.18 0.18

Fe3+ 0.31 0.39 0.52 0.53 0.30 0.56

Fe2+ 0.72 0.56 0.47 0.45 0.73 0.41

Mn 0.00 0.00 0.00 0.00 0.00 0.00

U 0.00 0.00 0.00 0.00 0.00 0.00

Th 0.00 0.00 0.00 0.00 0.00 0.00 Ca 1.04 1.21 1.13 1.18 1.09 1.16

La 0.17 0.14 0.18 0.16 0.11 0.18

Ce 0.47 0.37 0.38 0.35 0.41 0.35 Pr 0.06 0.04 0.04 0.04 0.06 0.04

Nd 0.19 0.15 0.15 0.14 0.24 0.14

Sm 0.03 0.02 0.02 0.02 0.04 0.02 Eu 0.00 0.00 0.01 0.00 0.00 0.00

Gd 0.01 0.01 0.01 0.01 0.02 0.01

Tb 0.00 0.00 0.00 0.00 0.00 0.00 Dy 0.00 0.00 0.00 0.00 0.00 0.00

Ho 0.00 0.00 0.00 0.00 0.00 0.00

Er 0.00 0.00 0.00 0.00 0.00 0.00 Tm 0.00 0.00 0.00 0.00 0.00 0.00

Yb 0.00 0.00 0.00 0.00 0.00 0.00

Lu 0.00 0.00 0.00 0.00 0.00 0.00 Y 0.02 0.02 0.02 0.02 0.02 0.02

OH* 1.00 1.00 1.00 1.00 1.00 1.00

F 0.01 0.00 0.00 0.00 0.00 0.00 ∑REE 0.94 0.77 0.80 0.75 0.91 0.77

∑cations 7.98 7.98 7.93 7.93 8.00 7.92

75

Table 6. WDS data for Östra Gyttorp gadolinite group minerals. All analysed points are from sample Gytt-

EJ-2. Calculations are based on 4 A+T cations (see Miyawaki 2007, DeMartin et al. 2001).


SiO2 (wt %) 25.96 25.67 26.07 26.22 26.27 26.29 25.65 26.63 26.53 26.74

Al2O3 0.01 0.00 0.00 0.00 0.00 0.00 0.19 0.00 0.00 0.00 MgO 0.08 0.12 0.07 0.22 0.12 0.42 0.41 0.32 0.35 0.41

FeO 4.37 6.47 3.82 5.11 4.31 6.29 6.36 6.62 6.74 6.50

MnO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 UO2 0.00 0.02 0.00 0.02 0.00 0.01 0.12 0.00 0.00 0.00

ThO2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

CaO 2.61 2.30 2.15 3.55 3.29 5.35 4.99 5.32 5.31 6.15 La2O3 0.11 0.13 0.00 0.01 0.06 0.06 0.80 0.03 0.02 0.03

Ce2O3 0.42 0.62 0.33 0.19 0.20 0.31 2.76 0.23 0.16 0.23

Pr2O3 0.10 0.14 0.08 0.08 0.10 0.14 0.33 0.00 0.09 0.00 Nd2O3 0.80 1.56 1.13 0.54 0.98 0.88 1.98 0.70 0.62 0.59

Sm2O3 1.10 1.51 1.24 0.53 1.11 0.83 0.99 0.77 0.65 0.60

Eu2O3 0.13 0.15 0.08 0.03 0.11 0.06 0.10 0.08 0.06 0.08 Gd2O3 4.97 4.62 4.46 2.60 4.34 3.25 3.04 3.49 3.21 2.97

Tb2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Dy2O3 5.06 4.91 4.85 3.96 4.81 4.01 3.45 3.99 3.90 3.75 Ho2O3 0.84 0.97 1.12 1.28 1.10 1.15 0.98 1.09 1.03 0.99

Er2O3 0.73 1.05 1.17 1.75 1.14 1.58 1.49 1.56 1.68 1.63

Tm2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Yb2O3 0.78 1.19 1.00 1.80 1.02 2.01 2.25 1.89 2.08 2.21

Lu2O3 0.00 0.00 0.00 0.01 0.00 0.00 0.05 0.01 0.02 0.04

Y2O3 33.20 31.15 33.43 32.99 32.83 28.22 24.23 27.84 28.11 27.67 BeO* 9.65 9.66 9.89 9.33 9.47 8.56 8.55 8.71 8.68 8.39

B2O3* 1.62 1.43 1.33 2.20 2.04 3.32 3.10 3.30 3.30 3.82

H2O* 2.72 2.09 2.85 2.50 2.77 2.11 1.99 2.07 2.02 2.11 F 0.00 0.00 0.00 0.00 0.00 0.00 0.08 0.00 0.00 0.00

O=-F* 0.00 0.00 0.00 0.00 0.00 0.00 -0.03 0.00 0.00 0.00

Total 95.27 95.77 95.07 94.92 96.09 94.84 93.86 94.67 94.55 94.91

Si (apfu) 2.02 2.04 2.04 2.03 2.01 2.03 2.04 2.06 2.06 2.05 Al 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.00

Mg 0.01 0.01 0.01 0.02 0.01 0.05 0.05 0.04 0.04 0.05

Fe 0.28 0.43 0.25 0.33 0.28 0.41 0.42 0.43 0.44 0.42 Mn 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

U 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Th 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Ca 0.22 0.20 0.18 0.29 0.27 0.44 0.43 0.44 0.44 0.50

La 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.00

Ce 0.01 0.02 0.01 0.01 0.01 0.01 0.08 0.01 0.00 0.01 Pr 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00

Nd 0.02 0.04 0.03 0.01 0.03 0.02 0.06 0.02 0.02 0.02

Sm 0.03 0.04 0.03 0.01 0.03 0.02 0.03 0.02 0.02 0.02 Eu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Gd 0.13 0.12 0.12 0.07 0.11 0.08 0.08 0.09 0.08 0.08

Tb 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Dy 0.13 0.13 0.12 0.10 0.12 0.10 0.09 0.10 0.10 0.09

Ho 0.02 0.02 0.03 0.03 0.03 0.03 0.02 0.03 0.03 0.02 Er 0.02 0.03 0.03 0.04 0.03 0.04 0.04 0.04 0.04 0.04

Tm 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Yb 0.02 0.03 0.02 0.04 0.02 0.05 0.05 0.04 0.05 0.05 Lu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Y 1.38 1.32 1.39 1.36 1.34 1.16 1.03 1.15 1.16 1.13

Be* 1.78 1.80 1.82 1.71 1.73 1.56 1.57 1.56 1.56 1.50 B* 0.22 0.20 0.18 0.29 0.27 0.44 0.43 0.44 0.44 0.50

OH* 1.41 1.11 1.48 1.29 1.42 1.09 1.06 1.07 1.05 1.08

F 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.00 ∑cations 6.32 6.49 6.29 6.38 6.30 6.49 6.53 6.53 6.54 6.51

76

Table 6 continued.

Point 11 Point 12 Point 13 Point 14 Point 15 Point 16 Point 17 Point 18

SiO2 (wt %) 26.71 25.44 26.12 26.87 26.48 26.26 25.64 26.17

Al2O3 0.00 0.00 0.73 0.00 0.00 0.00 0.00 0.00

MgO 0.42 0.10 0.17 0.43 0.13 0.20 0.09 0.13 FeO 6.54 7.67 8.04 6.03 4.64 5.76 6.08 5.19

MnO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

UO2 0.00 0.01 0.04 0.02 0.01 0.03 0.04 0.02 ThO2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

CaO 5.70 1.84 2.92 6.38 2.97 3.67 1.79 3.02

La2O3 0.01 0.06 0.55 0.04 0.03 0.03 0.13 0.05 Ce2O3 0.20 0.28 1.57 0.29 0.18 0.22 0.47 0.18

Pr2O3 0.11 0.03 0.27 0.08 0.02 0.00 0.11 0.09

Nd2O3 0.56 0.75 1.40 0.82 0.50 0.52 1.08 0.50 Sm2O3 0.57 0.69 0.82 0.80 0.53 0.44 0.89 0.37

Eu2O3 0.04 0.07 0.04 0.06 0.05 0.05 0.06 0.02

Gd2O3 2.96 3.43 3.33 3.20 3.14 2.80 3.82 2.56 Tb2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Dy2O3 3.77 4.89 4.07 3.72 4.15 3.98 4.89 3.69

Ho2O3 1.07 1.17 1.08 0.93 1.17 1.14 1.18 1.15 Er2O3 1.66 1.55 1.39 1.43 1.62 1.66 1.45 1.83

Tm2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Yb2O3 2.10 1.59 1.61 1.80 1.66 1.56 1.41 2.60 Lu2O3 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.05

Y2O3 28.41 32.63 29.49 27.51 33.48 32.25 32.88 32.77 BeO* 8.58 9.77 9.93 8.34 9.70 9.29 9.87 9.55

B2O3* 3.54 1.14 1.81 3.96 1.84 2.28 1.11 1.87

H2O* 2.08 1.74 1.74 2.24 2.66 2.32 2.19 2.47 F 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

O=-F* 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Total 95.03 94.86 97.13 94.97 94.96 94.46 95.20 94.29

Si (apfu) 2.05 2.06 2.04 2.04 2.05 2.04 2.05 2.05 Al 0.00 0.00 0.07 0.00 0.00 0.00 0.00 0.00

Mg 0.05 0.01 0.02 0.05 0.01 0.02 0.01 0.02

Fe 0.42 0.52 0.53 0.38 0.30 0.37 0.41 0.34

Mn 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

U 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Th 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Ca 0.47 0.16 0.24 0.52 0.25 0.31 0.15 0.25

La 0.00 0.00 0.02 0.00 0.00 0.00 0.00 0.00

Ce 0.01 0.01 0.04 0.01 0.01 0.01 0.01 0.01 Pr 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00

Nd 0.02 0.02 0.04 0.02 0.01 0.01 0.03 0.01

Sm 0.01 0.02 0.02 0.02 0.01 0.01 0.02 0.01 Eu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Gd 0.08 0.09 0.09 0.08 0.08 0.07 0.10 0.07

Tb 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Dy 0.09 0.13 0.10 0.09 0.10 0.10 0.13 0.09

Ho 0.03 0.03 0.03 0.02 0.03 0.03 0.03 0.03

Er 0.04 0.04 0.03 0.03 0.04 0.04 0.04 0.04 Tm 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Yb 0.05 0.04 0.04 0.04 0.04 0.04 0.03 0.06

Lu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Y 1.16 1.40 1.23 1.11 1.38 1.34 1.40 1.37

Be* 1.53 1.84 1.76 1.48 1.75 1.69 1.85 1.75

B* 0.47 0.16 0.24 0.52 0.25 0.31 0.15 0.25 OH* 1.06 0.94 0.91 1.14 1.37 1.20 1.17 1.29

F 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

∑cations 6.52 6.59 6.65 6.47 6.36 6.44 6.46 6.40

77

Table 7. WDS data for Bastnäsite group minerals from Östra

Gyttorp sample Gytt-EJ-1a. Calculations are based on 6 cations.

Point 1 Point 2 Point 3 Point 4

SiO2 (wt %) 0.01 0.00 0.02 0.00

Al2O3 0.00 0.00 0.00 0.00 MgO 0.00 0.00 0.01 0.00

FeO 0.00 0.00 0.00 0.00

MnO 0.00 0.00 0.00 0.00 UO2 0.50 0.15 0.50 0.48

ThO2 0.00 0.00 0.00 0.00

CaO 5.04 5.17 5.53 3.57 La2O3 13.33 13.13 11.60 15.91

Ce2O3 30.01 29.99 27.80 32.17

Pr2O3 3.55 3.47 3.31 3.80 Nd2O3 13.92 13.33 13.04 14.94

Sm2O3 2.62 2.68 2.64 2.01

Eu2O3 0.41 0.40 0.38 0.34 Gd2O3 2.05 2.10 2.05 1.18

Tb2O3 0.00 0.00 0.00 0.00

Dy2O3 0.58 0.75 0.64 0.05 Ho2O3 0.05 0.05 0.01 0.00

Er2O3 0.00 0.00 0.00 0.00

Tm2O3 0.12 0.08 0.00 0.00 Yb2O3 0.02 0.00 0.00 0.00

Lu2O3 0.01 0.00 0.00 0.00

Y2O3 3.49 3.48 3.30 1.13 H2O* 2.19 2.38 2.12 2.39

CO2* 23.12 22.95 22.05 22.10

F 4.66 4.17 4.25 4.01 O=-F -1.96 -1.76 -1.79 -1.69

Total 103.75 102.60 97.64 102.41

Si (apfu) 0.00 0.00 0.00 0.00

Al 0.00 0.00 0.00 0.00 Mg 0.00 0.00 0.00 0.00

Fe 0.00 0.00 0.00 0.00

Mn 0.00 0.00 0.00 0.00 U 0.02 0.01 0.02 0.02

Th 0.00 0.00 0.00 0.00

Ca 1.03 1.06 1.18 0.76 La 0.94 0.93 0.85 1.17

Ce 2.09 2.10 2.03 2.35

Pr 0.25 0.24 0.24 0.28 Nd 0.95 0.91 0.93 1.06

Sm 0.17 0.18 0.18 0.14

Eu 0.03 0.03 0.03 0.02 Gd 0.13 0.13 0.14 0.08

Tb 0.00 0.00 0.00 0.00

Dy 0.04 0.05 0.04 0.00 Ho 0.00 0.00 0.00 0.00

Er 0.00 0.00 0.00 0.00 Tm 0.01 0.00 0.00 0.00

Yb 0.00 0.00 0.00 0.00

Lu 0.00 0.00 0.00 0.00 Y 0.35 0.35 0.35 0.12

OH* 2.17 2.41 2.13 2.71

CO3* 6.01 6.00 6.01 6.01 F 2.81 2.53 2.68 2.53

78

Table 8. WDS data for dollaseites in sample Joha-KH-1 from the Johannagruvan

deposit. Calculations are based on 6 M+T cations (see Ercit 2002).

Point 1 Point 2 Point 3 Point 4 Point 5 Point 6 Point 7 Point 8 Point 9

SiO2 (wt %) 31,26 31,16 31,19 31,36 31,65 31,62 31,56 31,32 31,56

Al2O3 8,82 8,68 8,44 8,31 9,61 8,25 9,08 8,79 8,94

MgO 11,74 11,84 12,05 12,78 11,16 13,63 12,03 11,36 11,59 FeO 4,06 4,37 4,06 3,00 4,52 1,91 2,95 5,05 3,91

Fe2O3 0,54 0,19 0,46 0,60 0,38 0,39 1,19 0,05 1,01

MnO 0,17 0,18 0,14 0,27 0,19 0,27 0,20 0,15 0,18 TiO2 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00

UO2 0,00 0,00 0,00 0,00 0,00 0,00 0,01 0,00 0,00

ThO2 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 PbO 0,00 0,00 0,00 0,01 0,01 0,00 0,00 0,00 0,00

CaO 8,53 8,44 8,58 8,57 9,10 8,80 8,83 8,44 8,52

Na2O 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 K2O 0,00 0,01 0,01 0,00 0,00 0,01 0,01 0,01 0,00

P2O5 0,01 0,02 0,02 0,03 0,02 0,00 0,02 0,00 0,01

As2O5 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00

La2O3 8,02 8,18 10,20 8,27 8,35 9,11 8,02 8,17 8,18

Ce2O3 13,87 14,02 14,45 14,56 14,15 15,00 14,00 13,79 13,92

Pr2O3 1,51 1,43 1,24 1,51 1,46 1,40 1,49 1,46 1,49 Nd2O3 5,36 5,41 3,70 4,97 4,85 4,26 5,23 5,21 5,26

Sm2O3 0,69 0,63 0,31 0,47 0,35 0,32 0,47 0,71 0,62

Eu2O3 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 Gd2O3 0,28 0,29 0,14 0,22 0,03 0,12 0,15 0,40 0,27

Tb2O3 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00

Dy2O3 0,17 0,20 0,06 0,11 0,00 0,17 0,12 0,18 0,12 Ho2O3 0,00 0,05 0,03 0,07 0,05 0,01 0,02 0,01 0,00

Er2O3 0,04 0,02 0,00 0,05 0,00 0,08 0,05 0,04 0,03

Tm2O3 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 Yb2O3 0,04 0,05 0,02 0,07 0,02 0,05 0,01 0,01 0,03

Lu2O3 0,00 0,03 0,04 0,02 0,00 0,03 0,00 0,02 0,00

Y2O3 1,13 1,26 1,05 1,36 0,70 1,27 0,87 1,32 1,06 H2O* 1,58 1,57 1,57 1,58 1,60 1,59 1,59 1,57 1,59

F 2,49 2,51 2,61 2,83 2,33 3,11 2,59 2,34 2,38

O=-F* -1,05 -1,06 -1,10 -1,19 -0,98 -1,31 -1,09 -0,99 -1,00

Total 99,27 99,49 99,27 99,83 99,55 100,08 99,42 99,41 99,68

Si (apfu) 2,97 2,97 2,97 2,97 2,97 2,97 2,97 2,98 2,98 Al 0,99 0,97 0,95 0,93 1,06 0,91 1,01 0,99 0,99

Mg 1,66 1,68 1,71 1,80 1,56 1,91 1,69 1,61 1,63

Fe2+ 0,32 0,35 0,32 0,24 0,35 0,15 0,23 0,40 0,31 Fe3+ 0,04 0,01 0,03 0,04 0,03 0,03 0,08 0,00 0,07

Mn 0,01 0,01 0,01 0,02 0,01 0,02 0,02 0,01 0,01

Ti 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 U 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00

Th 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00

Pb 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 Ca 0,87 0,86 0,88 0,87 0,92 0,89 0,89 0,86 0,86

Na 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 K 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00

P 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00

As 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 La 0,28 0,29 0,36 0,29 0,29 0,32 0,28 0,29 0,28

Ce 0,48 0,49 0,50 0,50 0,49 0,52 0,48 0,48 0,48

Pr 0,05 0,05 0,04 0,05 0,05 0,05 0,05 0,05 0,05 Nd 0,18 0,18 0,13 0,17 0,16 0,14 0,18 0,18 0,18

Sm 0,02 0,02 0,01 0,02 0,01 0,01 0,02 0,02 0,02

Eu 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 Gd 0,01 0,01 0,00 0,01 0,00 0,00 0,00 0,01 0,01

Tb 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00

Dy 0,01 0,01 0,00 0,00 0,00 0,01 0,00 0,01 0,00 Ho 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00

Er 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00

Tm 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 Yb 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00

Lu 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00 0,00

Y 0,06 0,06 0,05 0,07 0,04 0,06 0,04 0,07 0,05 OH* 1,00 1,00 1,00 1,00 1,00 1,00 1,00 1,00 1,00

F 0,75 0,76 0,79 0,85 0,69 0,93 0,77 0,70 0,71

∑REE 1,09 1,11 1,10 1,11 1,04 1,11 1,06 1,11 1,08 ∑cations 7,96 7,98 7,98 7,99 7,96 8,00 7,95 7,97 7,94

79

Table 9. WDS data for Johannagruvan gadolinite group minerals. All analysed points are from sample Joha-KH-1.

Calculations are based on 4 A+T cations (see Miyawaki 2007, DeMartin 2001).

Point 1 Point 2 Point 3 Point 4 Point 5 Point 6 Point 7 Point 8 Point 9 Point 10 Point 11 Point 12 Point 13

SiO2 (wt %) 22.89 22.94 22.95 22.75 22.84 22.41 22.35 22.33 22.88 22.93 22.92 23.18 22.92

Al2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 MgO 0.91 0.85 0.77 0.70 0.78 0.65 0.51 0.72 0.70 0.80 0.75 0.81 0.91

FeO 9.64 10.14 10.00 10.16 10.26 9.91 9.82 10.04 10.14 10.13 10.03 9.99 10.01

MnO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 TiO2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

UO2 0.03 0.04 0.02 0.02 0.01 0.00 0.00 0.03 0.02 0.04 0.02 0.04 0.04

ThO2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 PbO 0.00 0.03 0.01 0.00 0.02 0.01 0.00 0.00 0.01 0.02 0.00 0.00 0.00

CaO 0.28 0.14 0.17 0.18 0.15 0.12 0.20 0.14 0.18 0.23 0.15 0.19 0.18

Na2O 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 K2O 0.01 0.00 0.00 0.01 0.00 0.00 0.01 0.02 0.00 0.00 0.00 0.01 0.00

P2O5 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.02 0.02

As2O5 0.18 0.18 0.21 0.19 0.13 0.13 0.16 0.12 0.22 0.17 0.18 0.19 0.20 La2O3 1.42 2.53 2.63 2.63 3.35 3.66 1.81 3.92 1.73 2.65 3.51 2.13 2.22

Ce2O3 7.37 11.76 11.73 12.07 15.31 13.87 10.09 15.47 9.96 12.36 14.69 10.23 10.60

Pr2O3 1.92 2.74 2.70 2.83 3.16 2.75 2.65 3.02 2.67 2.77 2.90 2.46 2.49 Nd2O3 15.18 16.92 17.09 18.40 18.34 16.44 19.27 17.26 18.78 17.84 17.46 17.08 17.29

Sm2O3 7.35 5.54 5.51 5.70 4.42 4.97 6.45 4.53 6.49 5.32 4.85 5.93 6.24

Eu2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Gd2O3 8.20 4.67 4.68 4.51 3.38 4.44 5.20 3.62 5.48 4.46 3.94 5.32 5.80

Tb2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Dy2O3 2.39 1.52 1.58 1.31 1.12 1.40 1.48 1.20 1.60 1.33 1.34 1.81 1.75 Ho2O3 0.10 0.04 0.10 0.12 0.08 0.07 0.08 0.02 0.05 0.04 0.06 0.13 0.14

Er2O3 0.14 0.19 0.13 0.04 0.03 0.06 0.00 0.11 0.13 0.12 0.11 0.20 0.06

Tm2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Yb2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Lu2O3 0.09 0.10 0.08 0.09 0.07 0.11 0.10 0.09 0.08 0.13 0.08 0.08 0.09

Y2O3 10.79 8.97 9.12 7.54 6.46 7.48 7.79 6.34 8.40 8.11 7.26 10.13 8.91 BeO* 9.14 9.16 9.19 9.08 9.12 9.02 8.96 9.00 9.13 9.12 9.19 9.27 9.18

B2O3* 0.17 0.09 0.11 0.11 0.09 0.08 0.13 0.09 0.11 0.14 0.09 0.12 0.11

H2O* 0.51 0.40 0.48 0.44 0.38 0.49 0.57 0.42 0.46 0.43 0.49 0.51 0.42 F 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

O=-F* 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Total 98.69 98.95 99.26 98.89 99.50 98.08 97.63 98.50 99.19 99.13 100.03 99.80 99.56

Si (apfu) 2.06 2.07 2.06 2.07 2.07 2.06 2.06 2.05 2.07 2.07 2.06 2.06 2.06 Al 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Mg 0.12 0.11 0.10 0.10 0.11 0.09 0.07 0.10 0.09 0.11 0.10 0.11 0.12

Fe 0.72 0.77 0.75 0.77 0.78 0.76 0.76 0.77 0.77 0.76 0.75 0.74 0.75 Mn 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Ti 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

U 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Th 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Pb 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Ca 0.03 0.01 0.02 0.02 0.01 0.01 0.02 0.01 0.02 0.02 0.01 0.02 0.02 Na 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

K 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 P 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

As 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.01

La 0.05 0.08 0.09 0.09 0.11 0.12 0.06 0.13 0.06 0.09 0.12 0.07 0.07 Ce 0.24 0.39 0.39 0.40 0.51 0.47 0.34 0.52 0.33 0.41 0.48 0.33 0.35

Pr 0.06 0.09 0.09 0.09 0.10 0.09 0.09 0.10 0.09 0.09 0.09 0.08 0.08

Nd 0.49 0.55 0.55 0.60 0.59 0.54 0.63 0.57 0.61 0.58 0.56 0.54 0.56 Sm 0.23 0.17 0.17 0.18 0.14 0.16 0.20 0.14 0.20 0.17 0.15 0.18 0.19

Eu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Gd 0.24 0.14 0.14 0.14 0.10 0.14 0.16 0.11 0.16 0.13 0.12 0.16 0.17 Tb 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Dy 0.07 0.04 0.05 0.04 0.03 0.04 0.04 0.04 0.05 0.04 0.04 0.05 0.05

Ho 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Er 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.00

Tm 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Yb 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Lu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Y 0.52 0.43 0.44 0.36 0.31 0.37 0.38 0.31 0.40 0.39 0.35 0.48 0.43

Be* 1.97 1.99 1.98 1.98 1.99 1.99 1.98 1.99 1.98 1.98 1.99 1.98 1.98 B* 0.03 0.01 0.02 0.02 0.01 0.01 0.02 0.01 0.02 0.02 0.01 0.02 0.02

OH* 0.31 0.24 0.29 0.27 0.23 0.30 0.35 0.26 0.28 0.26 0.29 0.30 0.25

F 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 ∑cations 6.85 6.88 6.86 6.87 6.88 6.85 6.82 6.87 6.86 6.87 6.85 6.85 6.87

80

Table 10. WDS data for bastnäsite group minerals from Johannaguvan sample Joha-

KH-1. Calculations are based on 6 cations.

Point 1 Point 2 Point 3 Point 4

SiO2 (wt %) 0.01 0.01 0.06 0.06 Al2O3 0.00 0.00 0.00 0.00

MgO 0.01 0.02 0.00 0.00

FeO 0.00 0.00 0.02 0.00 MnO 0.00 0.00 0.00 0.00

TiO2 0.00 0.00 0.00 0.00

UO2 0.00 0.00 0.00 0.00 ThO2 0.00 0.00 0.00 0.00

PbO 0.00 0.00 0.00 0.01

CaO 0.09 0.04 0.15 0.02 Na2O 0.00 0.00 0.00 0.00

K2O 0.15 0.17 0.02 0.00

P2O5 0.02 0.03 0.01 0.01 As2O5 0.12 0.08 0.06 0.12

La2O3 20.25 18.41 21.04 20.66

Ce2O3 31.23 30.31 31.48 30.19 Pr2O3 3.59 3.60 3.36 3.37

Nd2O3 16.99 16.11 14.18 15.14

Sm2O3 1.35 1.31 1.00 1.31 Eu2O3 0.00 0.00 0.00 0.00

Gd2O3 0.38 0.27 0.09 0.26

Tb2O3 0.00 0.00 0.00 0.00 Dy2O3 0.05 0.00 0.00 0.05

Ho2O3 0.00 0.00 0.00 0.00

Er2O3 0.00 0.00 0.00 0.00 Tm2O3 0.16 0.17 0.10 0.15

Yb2O3 0.00 0.00 0.00 0.00

Lu2O3 0.00 0.02 0.00 0.01 H2O* 1.01 0.80 0.88 0.95

CO2* 20.10 19.05 19.31 19.22

F 6.53 6.53 6.42 6.27 O=-F -2.75 -2.75 -2.70 -2.64

Total 99.56 94.42 95.59 95.37

Si (apfu) 0.00 0.00 0.01 0.01

Al 0.00 0.00 0.00 0.00

Mg 0.00 0.01 0.00 0.00 Fe 0.00 0.00 0.00 0.00

Mn 0.00 0.00 0.00 0.00

Ti 0.00 0.00 0.00 0.00 U 0.00 0.00 0.00 0.00

Th 0.00 0.00 0.00 0.00

Pb 0.00 0.00 0.00 0.00 Ca 0.02 0.01 0.04 0.01

Na 0.00 0.00 0.00 0.00 K 0.04 0.05 0.01 0.00

P 0.00 0.01 0.00 0.00

As 0.01 0.01 0.01 0.01 La 1.63 1.56 1.77 1.75

Ce 2.50 2.56 2.62 2.53

Pr 0.29 0.30 0.28 0.28 Nd 1.33 1.32 1.15 1.24

Sm 0.10 0.10 0.08 0.10

Eu 0.00 0.00 0.00 0.00 Gd 0.03 0.02 0.01 0.02

Tb 0.00 0.00 0.00 0.00

Dy 0.00 0.00 0.00 0.00 Ho 0.00 0.00 0.00 0.00

Er 0.00 0.00 0.00 0.00

Tm 0.01 0.01 0.01 0.01 Yb 0.00 0.00 0.00 0.00

Lu 0.00 0.00 0.00 0.00

OH* 1.47 1.23 1.34 1.45 CO3 5.99 5.99 6.00 6.01

F 4.51 4.76 4.62 4.54

81

Table 11. WDS data for uraninites from Östra Gyttorp sample Gytt-EJ-2. Calculations are based on 1 cation.

Point 1 Point 2 Point 3 Point 4 Point 5 Point 6 Point 7 Point 8 Point 9 Point 10 Point 11 Point 12

UO2 (wt %) 67.15 67.59 65.97 67.10 65.97 67.10 63.78 64.98 65.85 66.55 66.55 65.28

ThO2 0.00 0.02 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.02 PbO 18.67 18.66 18.87 20.34 18.87 20.34 18.32 17.88 17.75 18.16 18.16 18.15

SiO2 0.04 0.04 0.03 0.03 0.03 0.03 0.05 0.04 0.03 0.06 0.06 0.01

Al2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 MgO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

CaO 0.16 0.17 0.32 0.08 0.32 0.08 0.06 0.18 0.22 0.20 0.20 0.34

FeO 0.78 0.81 0.68 0.38 0.68 0.38 0.97 0.77 0.88 1.40 1.40 0.53 MnO 0.00 0.00 0.00 0.01 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.00

TiO2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

K2O 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 P2O5 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.02 0.00 0.00 0.00 0.00

Na2O 0.05 0.04 0.06 0.07 0.06 0.07 0.00 0.04 0.07 0.06 0.06 0.04

As2O5 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 La2O3 0.06 0.04 0.05 0.03 0.05 0.03 0.10 0.11 0.03 0.09 0.09 0.02

Ce2O3 0.30 0.22 0.32 0.80 0.32 0.80 3.66 1.49 0.40 0.47 0.47 0.39

Pr2O3 0.04 0.03 0.08 0.12 0.08 0.12 0.54 0.19 0.06 0.07 0.07 0.08 Nd2O3 0.56 0.40 0.63 1.37 0.63 1.37 3.07 1.44 0.66 1.05 1.05 0.55

Sm2O3 0.45 0.42 0.49 0.46 0.49 0.46 1.30 0.79 0.49 0.45 0.45 0.44

Eu2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Gd2O3 0.85 0.71 0.98 0.86 0.98 0.86 1.26 1.01 1.01 0.84 0.84 0.89

Tb2O3 0.27 0.26 0.25 0.18 0.25 0.18 0.11 0.25 0.29 0.25 0.25 0.27

Dy2O3 1.44 1.33 1.34 1.18 1.34 1.18 1.18 1.43 1.72 1.35 1.35 1.48 Ho2O3 0.26 0.26 0.28 0.23 0.28 0.23 0.15 0.27 0.25 0.37 0.37 0.28

Er2O3 0.87 0.87 0.89 0.60 0.89 0.60 0.41 0.70 0.98 0.76 0.76 0.81

Tm2O3 0.13 0.06 0.11 0.11 0.11 0.11 0.08 0.06 0.08 0.00 0.00 0.10 Yb2O3 0.85 0.90 0.79 0.45 0.79 0.45 0.26 0.55 0.89 0.75 0.75 0.79

Lu2O3 0.17 0.18 0.23 0.16 0.23 0.16 0.18 0.17 0.26 0.17 0.17 0.20

Y2O3 5.21 5.05 5.60 3.62 5.60 3.62 3.10 5.47 6.22 5.40 5.40 5.77 Cl 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Total 98.29 98.05 97.96 98.19 97.96 98.19 98.58 97.83 98.16 98.44 98.44 96.43

U (apfu) 0.58 0.59 0.57 0.59 0.57 0.59 0.55 0.56 0.56 0.56 0.56 0.57

Th 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Pb 0.20 0.20 0.20 0.22 0.20 0.22 0.19 0.19 0.18 0.19 0.19 0.19

Si 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Al 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Mg 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Ca 0.01 0.01 0.01 0.00 0.01 0.00 0.00 0.01 0.01 0.01 0.01 0.01

Fe 0.03 0.03 0.02 0.01 0.02 0.01 0.03 0.02 0.03 0.04 0.04 0.02 Mn 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Ti 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

K 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 P 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Na 0.00 0.00 0.00 0.01 0.00 0.01 0.00 0.00 0.01 0.00 0.00 0.00

As2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 La 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Ce 0.00 0.00 0.00 0.01 0.00 0.01 0.05 0.02 0.01 0.01 0.01 0.01 Pr 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00

Nd 0.01 0.01 0.01 0.02 0.01 0.02 0.04 0.02 0.01 0.01 0.01 0.01

Sm 0.01 0.01 0.01 0.01 0.01 0.01 0.02 0.01 0.01 0.01 0.01 0.01 Eu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Gd 0.01 0.01 0.01 0.01 0.01 0.01 0.02 0.01 0.01 0.01 0.01 0.01

Tb 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Dy 0.02 0.02 0.02 0.02 0.02 0.02 0.01 0.02 0.02 0.02 0.02 0.02

Ho 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Er 0.01 0.01 0.01 0.01 0.01 0.01 0.00 0.01 0.01 0.01 0.01 0.01 Tm 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Yb 0.01 0.01 0.01 0.01 0.01 0.01 0.00 0.01 0.01 0.01 0.01 0.01

Lu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Y 0.11 0.11 0.12 0.08 0.12 0.08 0.06 0.11 0.13 0.11 0.11 0.12

Cl 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

82

Table 11 continued.

Point 13 Point 14 Point 15 Point 16 Point 17 Point 18 Point 19 Point 20 Point 21 Point 22 Point 23 Point 24

UO2 (wt %) 64.36 64.05 62.21 63.48 62.14 63.43 66.07 62.52 67.55 63.54 66.78 60.74

ThO2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 PbO 18.56 18.41 17.66 17.78 18.29 19.58 18.77 16.42 17.73 17.32 17.32 19.12

SiO2 0.02 0.08 0.08 0.06 0.07 0.07 0.01 0.21 0.14 0.11 0.07 0.12

Al2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 MgO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

CaO 0.10 0.12 0.04 0.07 0.07 0.08 0.31 0.22 0.14 0.11 0.26 0.14

FeO 0.50 1.61 0.34 0.65 0.69 0.36 0.68 1.01 1.18 1.84 1.03 1.04 MnO 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

TiO2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

K2O 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 P2O5 0.00 0.01 0.01 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.02 0.00

Na2O 0.02 0.01 0.00 0.00 0.09 0.06 0.01 0.00 0.00 0.00 0.00 0.00

As2O5 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 La2O3 0.06 0.18 0.13 0.20 0.12 0.16 0.00 0.24 0.01 0.28 0.02 0.24

Ce2O3 1.42 2.01 3.82 3.88 2.60 2.84 0.34 4.06 1.41 4.32 0.81 3.43

Pr2O3 0.20 0.32 0.56 0.53 0.33 0.46 0.07 0.60 0.23 0.50 0.16 0.52 Nd2O3 1.48 3.76 3.85 2.95 2.12 3.42 0.61 3.86 1.45 3.64 0.95 3.52

Sm2O3 0.76 0.92 1.66 1.49 1.41 1.31 0.56 1.77 0.67 1.46 0.66 1.59

Eu2O3 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.13 0.06 0.14 0.05 0.11 Gd2O3 0.99 0.85 1.54 1.43 1.57 1.28 0.98 1.75 1.11 1.35 1.28 1.68

Tb2O3 0.25 0.25 0.31 0.21 0.29 0.25 0.25 0.00 0.00 0.00 0.00 0.00

Dy2O3 1.40 1.02 1.29 1.29 1.69 1.30 1.42 0.84 0.51 0.62 0.81 0.77 Ho2O3 0.23 0.39 0.20 0.12 0.33 0.21 0.24 0.09 0.09 0.05 0.12 0.05

Er2O3 0.74 0.67 0.35 0.34 0.63 0.35 0.89 0.12 0.18 0.09 0.28 0.10

Tm2O3 0.10 0.02 0.16 0.11 0.12 0.14 0.12 0.03 0.00 0.00 0.02 0.05 Yb2O3 0.64 0.36 0.16 0.35 0.53 0.21 0.87 0.25 0.46 0.20 0.78 0.21

Lu2O3 0.19 0.12 0.10 0.14 0.21 0.11 0.24 0.00 0.00 0.00 0.00 0.00

Y2O3 4.67 2.90 2.90 2.58 3.73 3.65 5.91 3.21 3.62 2.30 5.22 2.81 Cl 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Total 96.67 98.06 97.35 97.65 97.00 99.28 98.34 97.44 96.67 98.06 96.76 96.35

U (apfu) 0.57 0.55 0.54 0.55 0.54 0.54 0.57 0.53 0.60 0.54 0.58 0.53

Th 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Pb 0.20 0.19 0.19 0.19 0.19 0.20 0.19 0.17 0.19 0.18 0.18 0.20

Si 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.01 0.00 0.00 0.00

Al 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Mg 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Ca 0.00 0.00 0.00 0.00 0.00 0.00 0.01 0.01 0.01 0.00 0.01 0.01

Fe 0.02 0.05 0.01 0.02 0.02 0.01 0.02 0.03 0.04 0.06 0.03 0.03 Mn 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Ti 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

K 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 P 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Na 0.00 0.00 0.00 0.00 0.01 0.00 0.00 0.00 0.00 0.00 0.00 0.00

As2 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 La 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Ce 0.02 0.03 0.05 0.06 0.04 0.04 0.00 0.06 0.02 0.06 0.01 0.05 Pr 0.00 0.00 0.01 0.01 0.00 0.01 0.00 0.01 0.00 0.01 0.00 0.01

Nd 0.02 0.05 0.05 0.04 0.03 0.05 0.01 0.05 0.02 0.05 0.01 0.05

Sm 0.01 0.01 0.02 0.02 0.02 0.02 0.01 0.02 0.01 0.02 0.01 0.02 Eu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Gd 0.01 0.01 0.02 0.02 0.02 0.02 0.01 0.02 0.01 0.02 0.02 0.02

Tb 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Dy 0.02 0.01 0.02 0.02 0.02 0.02 0.02 0.01 0.01 0.01 0.01 0.01

Ho 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Er 0.01 0.01 0.00 0.00 0.01 0.00 0.01 0.00 0.00 0.00 0.00 0.00 Tm 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Yb 0.01 0.00 0.00 0.00 0.01 0.00 0.01 0.00 0.01 0.00 0.01 0.00

Lu 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 Y 0.10 0.06 0.06 0.05 0.08 0.07 0.12 0.07 0.08 0.05 0.11 0.06

Cl 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00

Tidigare utgivna publikationer i serien ISSN 1650-6553 Nr 1 Geomorphological mapping and hazard assessment of alpine areas in Vorarlberg, Austria, Marcus Gustavsson Nr 2 Verification of the Turbulence Index used at SMHI, Stefan Bergman Nr 3 Forecasting the next day’s maximum and minimum temperature in Vancouver, Canada by using artificial neural network models, Magnus Nilsson

Nr 4 The tectonic history of the Skyttorp-Vattholma fault zone, south-central Sweden, Anna Victoria Engström Nr 5 Investigation on Surface energy fluxes and their relationship to synoptic weather patterns on Storglaciären, northern Sweden, Yvonne Kramer Nr 269 Changes in Arsenic Levels in the Precambrian Oceans in Relation to the Upcome of Free Oxygen. Emma H.M. Arvestål, November 2013

Nr 270 Environmental and Climate Change During Holocenein Hjaltadalur, Skagafjördur,North Iceland -Peat core analysis and pollen identification. Jenny N. Johansson, November 2013

Nr 271 Reprocessing of 2D Reflection Seismic Marine Data and Investigation into the AVO behavior of Cambrian Sandstones, Southern Baltic Sea, Sweden Mahboubeh Montazeri, December 2013 Nr 272 Afrikas klimat med fokus på Västafrika, Eric Sönnert, January 2014 Nr 273 Shear-Wave Splitting Observed in Local Earthquake Data on the Reykjanes Peninsula, SW Iceland. Darina Buhcheva, February 2014 Nr 274 Unravelling Temporal Geochemical Changes in the Miocene Mogan and Fataga group ignimbrite succession on Gran Canaria, Canary Islands, Spain. Peter A. Nicholls, February 2014 Nr 275 Time Lapse VSP Monitoring of Small scale Injected CO2 in the Frio Formation, Texas, USA. Ershad Gholamrezaie, February 2014 Nr 276 A 2D Electrical Resistivity Survey of Palsas in Tavvavuoma, sub-arctic Sweden. Per Marklund, Mars 2014

Nr 277 Metamorphic Evolution of the Middle Seve Nappe in the Snasahögarna area, Scandinavian Caledonides. Åke Rosén, Mars 2014

Nr 278 Analogue Modelling of Ductile Deformation at Competent Lenses in Grängesberg, Bergslagen, Sweden. Sara Eklöf, Mars 2014

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