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Pekka Tu i sku

Au l i s Kärk i

June 2010

Work ing Repor t 2010 -54

Metamorphic Petrology of Olkiluoto

June 2010

Base maps: ©National Land Survey, permission 41/MML/10

Working Reports contain information on work in progress

or pending completion.

The conclusions and viewpoints presented in the report

are those of author(s) and do not necessarily

coincide with those of Posiva.

Pekka Tu isku

Au l i s Kärk i

K iv i t i e to Oy

Work ing Report 2010 -54

Metamorphic Petrology of Olkiluoto

METAMORPHIC PETROLOGY OF OLKILUOTO ABSTRACT The bedrock of Olkiluoto consists mostly of pelitic migmatites, and lesser amounts of

biotite and hornblende gneisses, granite gneisses and granodiorite gneisses. The

metapelites underwent partial dehydration melting through decomposition of biotite,

sillimanite, plagioclase and quartz during the peak of regional metamorphism,

producing the migmatites. Simple model systems were used to estimate the conditions

of melting in metapelites and in quartzofeldspathic and granitic rocks. Pressure and

temperature conditions of metamorphism in Olkiluoto have also been calculated by

TWQ-thermobarometer and predicted from mineral assemblages of different rock types

in the area. Regional metamorphism of Olkiluoto culminated with a voluminous migmatization of

pelitic gneisses, in the temperature exceeding 660 ºC and relatively low pressure of

about 3.5 – 4 kbar. The temperature may have risen up to ~ 700 ºC producing granitic

melt, which later crystallized to leucosomes. The mineral assemblages produced during

the peak of regional metamorphism overprint the earlier S2 foliation. Granitic

leucosomes associated with the culmination of the metamorphism were deformed by D3

deformation. Accordingly, the peak of metamorphism took place between D2 and D3

deformation phases. Muscovite-bearing shear bands of some granite pegmatites indicate

that D3 deformation took place during subsequent cooling. Because granodiorite

gneisses, having an age of about 1.86 Ga, are deformed by D2 deformation, the peak of

regional metamorphism may be dated somewhere between 1.86 and 1.82 Ga, the latter

being the minimum age of granite deformed by D3 deformation. The culmination of

metamorphism is seen also in the mineral assemblages of more resistant rocks as biotite

gneisses, hornblende gneisses and granodiorite gneisses. The cooling phase took place

in low pressure, which is seen from some retrograde reaction products as andalusite,

chlorite and muscovite replacing cordierite.

Calculated pressures from some rocks are clearly higher than the average metamorphic

pressure in Olkiluoto. This might be caused by an earlier stage of metamorphism

connected with the extrusion and intrusion of granodioritic rocks. However, also this

stage may be classified as low-pressure type of regional metamorphism and it took

place shortly after 1.86 Ga. Because there is a pressure difference of two kbar between

the two stages of metamorphism, there has been an erosion phase between the

metamorphic phases, implying significant crustal uplift. Alternatively, a phase of

tectonic thinning of the crust took place before the main stage of migmatization. The

migmatization, however, was produced by heating, not decompression, as there is no

evidence of previous or relic garnet in the migmatites.

Keywords: Gneiss, Olkiluoto, metamorphism, mineral, temperature, pressure.

OLKILUODON METAMORFINEN PETROLOGIA TIIVISTELMÄ Olkiluodon kallioperä koostuu pääosin peliittisistä migmatiiteista. Myös biotiitti- ja

sarvivälkegneissejä sekä graniitti- ja granodioriittigneissejä esiintyy jonkin verran.

Alueellisen metamorfoosin aikana metapeliiteissä tapahtui biotiitin, sillimaniitin, pla-

gioklaasin ja kvartsin dyhydraatiosulaminen, jonka seurauksena migmatiitit syntyivät.

Olkiluodon alueen metamorfoosiolosuhteet on määritetty yksinkertaisten mallikoostu-

musten, sekä alueen kivilajien geotermobarometristen ja faasitasapainolaskelmien

avulla.

Olkiluodon alueellinen metamorfoosi kulminoitui peliittisten gneissien migmatisoi-

tuessa yli 660 ºC lämpötilassa ja alhaisessa 3.5 – 4 kbar paineessa. Lämpötila saattoi

nousta jopa noin 700 ºC:en, jolloin syntyi paljon graniittista silikaattisulaa, joka myö-

hemmin kivien jäähtyessä kiteytyi migmatiittien leukosomeiksi. Koska kulminaatio-

vaiheen mineraaliseurueet ovat syntyneet D2-deformaation jälkeen, ja sulamisen

seurauksena syntyneet graniitiset leukosomit ovat osittain deformoituneet D3:n aikana,

ajoittuu metamorfoosin päävaihe näiden deformaatiotapahtumien väliin. Myös joidenkin

graniittien muskoviittipitoiset D3-hiertosaumat osoittavat D3:n tapahtuneen jäähtymisen

aikana. Absoluuttiselta iältään Olkiluodon alueellisen metamorfoosin huippu sijoittuu

1,86 ja 1,82 Ga:n välille, koska 1,86 Ga vanhat granodioriitiset kivet ovat defor-

moituneet D2:n aikana ja huipun jälkeen kiteytyneiden graniittisten kivien minimi-ikä

on noin 1,82 Ga. Metamorfoosin kulminaatio ilmenee myös sulamista vastaan resis-

tanttien biotiitti- ja sarvivälkegneissien sekä granodioriittigneissien mineraaliseurueissa.

Jäähtyminen tapahtui alhaisessa paineessa, mitä osoittaa mm. kordieriittia syrjäyttävä

andalusiittipitoinen seurue.

Joistakin kivistä saadaan laskettua noin 2 kbar korkeampi metamorfoosipaine. Tähän

voi olla syynä aiempi metamorfoosivaihe, joka saattaa liittyä granodioriittisen sulan

purkautumiseen ja tunkeutumiseen alueen kallioperään. Tämäkin vaihe on luokitel-

tavissa alhaisen paineen metamorfoosiksi korkean T/P-suhteen vuoksi, ja sen ikä on

vähän alle 1,86 Ga. Paineen lasku tultaessa migmatiittiutumisvaiheeseen selittyy joko

intensiivisellä kohoamisella ja eroosiolla tai kuoren tektonisella ohenemisella näiden

kahden metamorfoosivaiheen välillä. Migmatiittiutumisen aiheutti kuitenkin lämpene-

minen, sillä migmatiiteissa pitäisi olla reliktejä granaatista, mikäli paineen lasku olisi

aiheuttanut dehydraatiosulamisen.

Avainsanat: Gneissi, Olkiluoto, metamorfoosi, mineraali, lämpötila, paine.

1

TABLE OF CONTENTS ABSTRACT

TIIVISTELMÄ

LIST OF ABBREVIATIONS ........................................................................................... 3 1 INTRODUCTION .................................................................................................... 5 2 GEOLOGICAL SETTING ........................................................................................ 7 3 SAMPLES ............................................................................................................... 9

3.1 Biotite Gneisses (a mica gneiss variant) ........................................................ 9 3.2 Hornblende Gneisses (a mafic gneiss variant)............................................. 10 3.3 Granite Gneisses (a TGG gneiss variant) .................................................... 11 3.4 Granodiorite Gneisses (a TGG gneiss variant) ............................................ 12 3.5 Migmatitic Gneisses .................................................................................... 13

4 ANALYTICAL METHODS ..................................................................................... 15 5 METAMORPHIC PETROLOGY ............................................................................ 17

5.1 General Considerations from Mineral Assemblages, Petrography and .......... Textures ...................................................................................................... 18

5.1.1 Mineral Assemblages ........................................................................... 18 5.1.2 Textures Related to Heating .................................................................. 25 5.1.3 Textures Related to Decompression ...................................................... 27 5.1.4 Textures Related to Cooling .................................................................. 28 5.1.5 Textures of the TGG Series ................................................................... 30

5.2 Phase Petrology and Petrogenetic Grids ..................................................... 30 5.2.1 Biotite Gneisses ..................................................................................... 30 5.2.2 Hornblende Gneisses ............................................................................ 33 5.2.3 Granite Gneisses ................................................................................... 34 5.2.4 Granodiorite Gneisses ........................................................................... 36 5.2.5 Migmatitic Gneisses ............................................................................... 38

5.3 Thermobarometry ........................................................................................ 47 5.3.1 Biotite Gneisses ..................................................................................... 47 5.3.2 Hornblende Gneisses ............................................................................ 54 5.3.3 Granite Gneisses ................................................................................... 57 5.3.4 Granodiorite Gneisses ........................................................................... 60 5.3.5 Migmatitic Gneisses ............................................................................... 64

6 DISCUSSION AND CONCLUSIONS .................................................................... 69 REFERENCES ........................................................................................................... 71 APPENDIX ................................................................................................................. 75

2

3

LIST OF ABBREVIATIONS Minerals and phases

Als aluminium silicate

And andalusite

Bt biotite

Chl chlorite

Crd cordierite

Cum cummingtonite

Cz clinozoisite

Czo clinozoisite

FSP feldspar

Grt garnet

H2O water fluid

Hbl hornblende

Ilm Ilmenite

Kfs potassium feldspar

LIQtc silicate liquid

LQ silicate liquid

m myrmekite

Mc microcline

Mic microcline

Ms muscovite

Or orthoclase

phng muscovite

Pl plagioclase

Py pyrite

q quartz

Qtz quartz

Ser sericite

Sil sillimanite

sill sillimanite

Sp spinel

Tr tremolite

Tur tourmaline

Urn uraninite

zo zoisite

Other

D deformation or deformation event / phase

Dn nth

deformation

PT pressure and temperature

S planar feature or structural element

Sn planar feature or structural element produced by nth

deformation

TGG tonalite-granodiorite-granite rock series or rock association

TWQ Thermobarometry With Estimation of Equilibration State (TWEEQU)

4

5

1 INTRODUCTION According to the Finnish Nuclear Energy Act, nuclear waste generated in Finland must

be treated, stored and disposed of within the Finnish borders. Posiva, a joint company of

Fortum Power & Heat Oy and Teollisuuden Voima Oy, has carried out petrological

studies on Olkiluoto area since 1988 as a part of bedrock investigations in Olkiluoto,

related to the studies on future final disposal of spent nuclear fuel.

The aim of the present study is the characterization of metamorphic evolution and

calculation of conditions of metamorphism of Olkiluoto by standard methods used in

petrology. The bedrock of Olkiluoto is composed of migmatitic rocks with minor

amount of resisters such as mafic gneisses and mica bearing gneisses, which form a

continuous transition from rather homogeneous gneisses to very heterogeneous

migmatites. Typical meso- and megascopic (mm scale – m scale) constituents of

migmatites are leocosome, mesosome and melanosome (Sederholm 1907, Mehnert

1968, Wimmenauer & Bryhni 2007) the types and proportions of which are directly

controlled by metamorphic processes and grade of metamorphism. Leucosome is the

lighest colored and typically igneous-looking part of the migmatite rock. Mesosome is

intermediate in color and it can be a more or less well preserved part of the protolith or

initial rock. Melanosome is the darkest part of the composite rock and it can be

composed purely of mafic minerals, eg. biotite and hornblende. The forms of

occurrence or shapes of individual migmatite components can vary widely but often the

final structure is layered or banded causing strong anisotropy in physical properties of

the rock material.

6

7

2 GEOLOGICAL SETTING Rauma map sheet 1132 is situated in the western end of the Svecofennian Accretionary

Arc Complex of Southern Finland having an age span between 1.82-1.90 Ga (Korsman

et al.1997) The lithology of the area has been described by Suominen et al. (1997).

Majority of supracrustal rocks consists of migmatized mica gneisses whereas

hornblende gneisses, amphibolites and metamorphosed uralite porphyrites are more

scattered The supracrustals are intruded by felsic, granitic – tonalitic igneous rocks.

Granites often crosscut other rocks. All the rock types, also granites seem to be more or

less foliated. Post-orogenic Laitila Rapakivi batholith is situated in the NE part of the

sheet area and numerous Post-Jotunic olivine diabases crosscut all other rock types

(Korsman et al. 1997). Literature based description of the geology of the region

surrounding Olkiluoto was compiled by Paulamäki (2007).

TIMS U-Pb dating of the zircons from a foliated tonalite body in the SE part of the area

gives an age of 1863 ± 7 Ma while the U-Pb age of the monazite from the same rock is

younger, 1813 ± 4 Ma. The zircon age was interpreted by Suominen et al. (1997) to

present the crystallization age of the intrusion, which means, according to them, that the

major regional deformation of the area is younger than the intrusion. The monazite age

was interpreted to represent post-intrusive diffusion. The Laitila batholith in the eastern

margin of the map sheet has been dated to have an age of 1574 ± 4 Ma and the age of

Sorkka olivine diabase is 1258 ± 13 Ma (Suominen 1991, Vaasjoki 1996). Olkiluoto is

situated in the NE part of the Rauma map sheet, just a few kilometres west of the

Eurajoki rapakivi stock, a satellite massif of the Laitila batholith (Figure 1).

Figure 1. Geological map of Olkiluoto. Modified from Kärki and Paulamäki (2006).

8

The rock types of Olkiluoto can be divided into five major classes 1) migmatites, 2)

homogeneous mica-bearing gneisses and quartzitic gneisses, 3) homogeneous tonalitic-

granodioritic-granitic gneisses (TGG gneisses), 4) amphibolites and other mafic

gneisses and 5) pegmatitic granites. The supracrustal rocks are deformed in four

deformation phases (D1 – D4) during collisional and post-collisional stages of

Svecofennian orogeny (Paulamäki and Koistinen 1991, Kärki and Paulamäki, 2006).

Tonalitic and granodioritic rocks were extruded and intruded before or during the phase

D2 and potassium granites for a part during the phase D3. Thus, the igneous rocks seem

to be syn-metamorphic with the progressive stage, because metamorphic segregation

defines the S1 foliation at least in some metapelitic rocks (Kärki and Paulamäki, 2006).

Mänttäri et al. (2006) dated zircons from a TGG gneiss in the NW part of Olkiluoto by

SIMS U-Pb. The zonal prismatic zircons form a homogeneous population and gave an

crystallization age of 1863 ± 6 Ma which is in accordance with the TIMS age from the

more southern part of the Rauma map sheet (Suominen et al., 1997, see above). This

would thus set the maximum age bracket for the deformation phase D2 (Kärki and

Paulamäki, 2006). Mänttäri et al. (2006) also obtained an 1822 ± 13 Ma age from one

homogeneous zircon grain which is in agreement with the monazite age of Suominen et

al. (1997) indicating post-intrusive thermal diffusion.

9

3 SAMPLES

The sample material considered in this work consists of about 450 thin sections.

Majority of the rock samples were collected earlier during drill core loggings (Gehör et

al. 1997 … 2007). About 35 sections were selected for more detailed investigation

based on coverage of the rock types and mineral assemblage(s), which could be useful

for petrologic research. In addition, an attempt was made to avoid rocks, which had

largely suffered late stage retrogressive alteration as chloritization of biotite,

sericitization of plagioclase and pinitization of cordierite, which are quite common in

Olkiluoto area. The classification presented below mainly follows the main lithological

subdivision of Olkiluoto presented by Kärki and Paulamäki (2006), who also estimated

volume percentages of each rock type in Olkiluoto drill cores. Description of sample

material origin and summary of results obtained by this study is in Appendix 1.

3.1 Biotite Gneisses (a mica gneiss variant)

Biotite gneisses are homogeneous, foliated and fine-grained. They are composed almost

solely of intermediate to anorthitic plagioclase, quartz and biotite (Figure 2). They may

contain small porphyroblasts of garnet and rarely minor amounts of potash feldspar.

Garnet is commonly replaced by plagioclase in variable amounts. Other phases include

accessory minerals as apatite, zircon, monazite, ilmenite, graphite and sulphides.

Alteration of plagioclase to sericite and biotite to chlorite is common. Samples OL103,

OL142, OL157, OL184, OL200 and OL206 are classified as biotite gneisses. OL184

exhibites banding, in which the rock grades to hornblende gneiss through an Fe-Mg-

amphibole bearing zone.

Figure 2. Micrograph of homogeneous biotite gneiss. Penetrative foliation is defined by

parallel biotite scales and elongation of quartz and plagioclase. Sample OL 103.

10

3.2 Hornblende Gneisses (a mafic gneiss variant)

Hornblende gneisses resemble biotite gneisses by fabric and grain size. They are

composed of anorthitic plagioclase, quartz and hornblende (Figure 3). Accessory phases

include ilmenite, sphene and apatite. Some diopside may be present as well as small

garnet porphyroblasts, which are partially replaced by plagioclase. Secondary

clinozoisite, calcite etc. may be present. Samples OL118 and OL121 are classified as

hornblende gneisses but the latter grades to biotite gneiss.

Figure 3. Micrograph of garnet-bearing hornblende gneiss. Foliation is less prominent

due to the lack of mica. Garnet poikiloblast is partly replaced by plagioclase most

probably due to the decompression during or slightly after the temperature peak of

regional metamorphism. Sample OL 121.

11

3.3 Granite Gneisses (a TGG gneiss variant)

Only one studied sample OL108 belongs to this rock type. The rock is coarse-grained

and has a blastomylonitic fabric. Large feldspar and quartz grains and porphyroclasts

are encircled by muscovite- and sillimanite-bearing shear zones (Figure 4). Garnet is

partially replaced by plagioclase and some biotite is found as inclusions in quartz.

Figure 4. Micrograph of pegmatitic granite. Larger porphyroclasts of microcline-

perthite, quartz and plagioclase (oligoclase) are enveloped by fine grained shear zones

of muscovite and some sillimanite. Sample OL 108.

12

3.4 Granodiorite Gneisses (a TGG gneiss variant)

Granodiorite gneisses are homogeneous, foliated and medium-grained. They are

composed of plagioclase (oligoclase), often with narrow albitic rim, perthitic

microcline, quartz and oriented biotite (Figure 5). Garnet is present in some samples and

at least the rims are grown over pre-existing S2 foliation. It is thus considered to be post-

or syn-kinematic in relation to D2 and has still later been partially replaced by

plagioclase. Muscovite may be present and microcline-replacing myrmekite is usual.

Accessory phases include apatite, sulphides, graphite and retrograde chlorite and

sericite. Samples OL145, OL147 and OL188 are granodiorite gneisses.

Figure 5. Micrograph of granodiorite gneiss. Sample OL 145.

13

3.5 Migmatitic Gneisses

Majority of the studied samples belong to this rock class. They are heterogeneous and

commonly composed of mesosomes, melanosomes and leucosomes. Mesosomes are

foliated, fine- to medium-grained and consist of oligoclase, quartz and biotite as well as

variable amounts of prismatic sillimanite, porphyroblastic cordierite and accessory

sulphides (mainly pyrrhotite), graphite, apatite, zircon and monazite (Figure 6).

Melanosomes are foliated, usually coarser than leucosomes and consist mainly of

biotite, sillimanite, sulphides, graphite, monazite, zircon, and occasionally some

cordierite (Figure 7). Leucosomes are mostly granitic, less foliated, medium- to coarse-

grained and myrmekitic. They consist of typical granite minerals microcline (-perthite)

or orthoclase (-perthite), oligoclase, quartz and biotite and may contain cordierite,

muscovite, graphite, sulphides, andalusite, zircon and monazite (Figure 7). Samples

OL105, OL112, OL114, OL122, OL123, OL133, OL134, OL135, OL136, OL137,

OL140, OL143, OL148, OL156, OL183, OL185, OL194, OL196 and OL197 are

migmatitic gneisses. Migmatitic gneisses include both veined and diatexitic gneiss types

of Kärki and Paulamäki (2006).

Figure 6. Micrograph of the mesosome in migmatitic gneiss Arrows from left to right

point to biotite relic, quartz inclusion and sillimanite relic in cordierite porphyroblast.

The texture and thermodynamic modelling indicate that post-tectonic cordierite

porphyroblast is most evidently were generated by dehydration melting of syn-tectonic

biotite, plagioclase and sillimanite. Sample OL 197.

14

Figure 7. Micrograph of migmatitic gneiss. Foliated biotite-sillimanite-pyrite

melanosome and granitic, microcline-rich leucosome, where myrmekite (m) and

muscovite were formed during cooling from potash feldspar and water released from

frozen leucosome melt. Sample OL 105.

15

4 ANALYTICAL METHODS

The whole rock analyses were made in the Laboratories of Geological Survey of

Finland and SGS Mineral Services laboratory in Canada. Analyses of major components

and some minor components were performed by X-ray fluorescence (XRF) while those

of minor and trace elements were undertaken in most cases by inductively coupled

plasma mass spectrometry (ICPMS). The analysis used in calculation in the present

study are those reported in Gehör et al. (1997 … 2007) and the summary of analytical

methods can be found there.

The mineral analyses were performed by Jeol JXA-8200 microprobe at the Institute for

Electron Microscopy at University of Oulu by P.Tuisku. Accelerating voltage of 15 kV

and beam diameter between 2-20 m was used. Synthetic oxide standards were used for

most elements, and KCl and CaF2 and albite for Cl, F and Na respectively.

16

17

5 METAMORPHIC PETROLOGY

Metamorphic petrology of Olkiluoto was already described by Kärki and Paulamäki

(2006). They discussed the probable source materials of the metamorphic rocks at

Olkiluoto, mostly based on the chemical composition of the rocks. We will discuss

source rocks here only when it is relevant for detection of metamorphic evolution and

pressure-temperature conditions during this evolution.

Kärki and Paulamäki (2006) also discussed the probable pressure and temperature (PT)

condition especially during the peak phase of regional metamorphism at Olkiluoto. The

discussion was mostly based on literature survey of similar rocks. Fewer attempts were

made to detect the PT-conditions during the progressive stage as well as retrogressive

stage of metamorphism. In the following, we try to quantify peak PT-conditions of

regional metamorphism of major rock types, but will also discuss the PT-evolution

during progressive stage as well as cooling stage. Of course, we will also discuss how

the evolution is visible in the mineralogy and texture in different rock types as this is

important for overall bedrock modelling of the area.

There are several methods for estimating conditions of formation of metamorphic rocks.

The most general is the concept of metamorphic facies (Eskola 1915, 1921) which

involves a qualitative subdivision of metamorphic pressure and temperature regime

according to mineral assemblages found in rocks having certain chemical composition.

According Eskola´s concept the migmatites of Olkiluoto as well as the basic rocks

found as intercalations are clearly classified to upper amphibolite-facies. Eskola did not

know very much about mineral thermodynamics and the rocks compositions are not

strictly defined quantitatively in facies concept, so the facies-concept can only be used

as suggestive of relative high temperature and relatively low pressure as well as high

T/P-ratio at Olkiluoto.

However, metamorphic mineral assemblages and mineral reactions of certain chemical

bulk composition may be presented quite exactly in the PT-space, if we know enough

about thermodynamic properties of minerals and/or have enough experimental results

on mineral equilibria, further developing Eskola´s idea. This kind of phase diagram was

named “petrogenetic grid” by N.L. Bowen (1940). During the last half of the previous

century, large-scale experimental work has produced extensive thermodynamic data on

minerals, and further mathematical treatment of the data has made possible to compile

internally consistent data sets of thermodynamic properties of minerals (Berman 1988,

Holland and Powel 1985, 1998, Aranovich and Berman 1996, Spear et al. 2000). These

data sets allow the calculation of phase diagrams for almost all possible rock

composition by modern computation methods, because the stability of minerals and

mineral assemblages is governed by the laws of thermodynamics (de Capitani and

Brown 1987, Powell and Holland 2001, de Capitani 2005, de Capitani and Petrakakis

2008). Great number of these diagrams has been published for common rock types in

recent years (Vernon and Clarke 2008). But more importantly, if we know the chemical

bulk composition of our rock sample, we can always calculate a petrogenetic grid for

that specific composition of interest and compare the result phase diagram with the

natural mineral assemblage of the rock (de Capitani and Brown 1987, Powell and

Holland 2001, Spear 2001, de Capitani 2005, de Capitani and Petrakakis 2008).

18

Most common rock-forming minerals form solid solutions between different end-

members and the distribution of elements between these solution phases (minerals)

follow the general laws of thermodynamics. If we can measure the chemical

composition of co-existing minerals and know the thermodynamic properties of the

phases involved, it is possible to calculate precisely PT-conditions, where the chemical

equilibrium between the phases was reached. These kinds of methods are called

geological thermobarometry (Kretz 1959) and modern thermodynamic data sets and

computation methods allow the calculation with numerous mineral phases involved

(Powell and Holland 2001, Spear 2001, Berman 2007)

The above calculation methods suffer some pitfalls which, even when not making the

calculation impossible or the results meaningless, at least cause some error, variable

from case to case. Because the modern analytical methods and the thermodynamic data

for mineral end-members are relatively reliable, largest errors are produced by four

different facts. These are 1) incompetent application of analytical or calculation

methods, 2) impurities and chemical factors not taken into account in the datasets,

calculation or analytical methods, 3) missing data for certain end-members or

inadequate or missing solution models for the distribution of elements and 4)

unequilibria producer either by sluggishness of chemical reactions in solid rocks or

(partial) recrystallization in the course of PT-evolution of the rocks concerned.

However, the last one may also utilized, because all unequilibria, sluggish reactions etc.

are also recorded in the rocks and are visible as different mineral textures (replacement,

alteration, reaction rims etc.) or mineral zoning.

In the further discussion, we apply all these methods for Olkiluoto. The principal

software used in calculations is those described by de Capitani (2005), de Capitani and

Petrakakis (2008) and Berman (2007). Thermodynamic data sets used are referred in

text, figures and program manuals.

5.1 General Considerations from Mineral Assemblages, Petrography and Textures

5.1.1 Mineral Assemblages

The peak metamorphic mineral assemblages of all rock types found in Olkiluoto belong

to the upper amphibolite facies, which roughly constrict the conditions of

metamorphism between 550 – 770 oC and 2 – 12 kbar (e.g. Vernon and Clarke 2008).

The common occurrence of sillimanite limits the maximum pressure to about 8.5 kbar.

However, further general constraint based on mineral assemblages are difficult to set

because the typical biotite-sillimanite-cordierite assemblage of pelitic migmatites in

Olkiluoto has a relatively large stability field from 560 oC and 2 kbar to about 850

oC

and at least 8 kbar. Lower temperature limit is further confined to about 640 oC by the

occurrence of melt together with the typical solid phases (e.g. Spear, 1993, Tuisku & al

2006). So, without further consideration, the possible temperature interval of regional

metamorphism at Olkiluoto remains as large as about 200 oC between 560 – 850

oC,

and pressure interval as large as 6.5 kbar between 2 – 8.5 kbar.

19

Further constraints may be set by phase equilibria of two simple systems, which may

applied to metapelites, meta-arenites or silicic igneous rocks of Olkiluoto. Low- and

medium-grade metapelites commonly contain muscovite, oligoclase (plagioclase),

biotite and quartz and may thereby suffer migmatization due to dehydration melting of

muscovite or biotite. Dehydration melting of muscovite, oligoclase and quartz is the

first possible reaction to cause melt generation in metapelites. On the other hand

crystallization of water-bearing granitic magma often produce muscovite-, oligoclase-,

quartz- and K-feldspar-bearing granite or pegmatite, by the reverse reaction of

muscovite dehydration. It is thus useful to consider phase equilibria in a simple system

muscovite-plagioclase-quartz-water.

Phase diagrams calculated for system 1 Ms – 1 Pl (An20) – 4 Qtz – (zero – 10) H2O are

presented in Figures 8, 9a and 9b. The database used in calculation was from Holland

and Powell (1998) with liquid data of White et al. (2007). Figure 8 is a pressure-

temperature diagram of intermediate composition with two H2O. It illustrates first

melting reactions in muscovite-bearing metapelites as well as crystallization of water-

rich granitic magma. The diagram may be applied to the prograde heating of pelitic and

arenitic metasediments as well as crystallization or possible heating of granitic rocks of

Olkiluoto. Blue and blue gray fields are regions where liquid is absent and are thus

below the peak conditions in Olkiluoto. Red and orange fields are PT-regimes where

fluid exists in the system and are possible in Olkiluoto. The reaction curve between pink

and red is critical for Olkiluoto, especially pegmatitic granite sample OL 108 where

sillimanite probably did co-exist with melt and muscovite. The curve constricts

conditions where sillimanite first appears in the system during heating or tends to

disappear from the system during cooling to the pressure between 3.9 and 7.3 kbar and

the temperature of 650 – 680 oC. The curve also constricts the first possible appearance

of sillimanite in the metasediments, especially more arenitic members of the

metasedimentary sequence.

20

Figure 8. Pressure-temperature phase-diagram for model system 1 muscovite – 1

plagioclase (An20) – 4 quartz – 2 H2O. The diagram illustrates phase relations in

aluminous metapelites, arenites and granites in the presence of excess water. Blue

areas are stability fields of melt absent and orange and red melt-bearing assemblages.

The effect of water component is illustrated in Figure 9, which shows the phase

relations of the system in the pressure of four kbar. It is clear, especially from Fig 9b,

which is a detail of Fig 9a, that dehydration melting of Ms+Pl+Qtz takes place in a

small temperature interval from 652 to 658 oC. Decreased water content of the system

tends to stabilize K-feldspar in the expense of sillimanite, i.e., the first phase to appear

during heating in higher water activity is sillimanite and in lower potash feldspar. Fluid

influx cannot cause melting in this system at constant temperature and/or pressure.

Figure 8 shows the effect of pressure in melting temperature. It is seen that especially in

low-P near the reaction curve between sillimanite and andalusite decompression will

cause crystallization in this system.

21

Figure 9. Composition-temperature phase diagram for the model system 1 muscovite –

1 plagioclase (An20) – 4 quartz – (0-10) H2O in the pressure of 4 kbar. A) Phase

relations in 300 - 700 oC and H2O 2.5 – 0. B) Detail of A. Diagrams show how the

dehydration melting of muscovite takes place in limited T range from 652 to 658 oC.

22

Another simplified model system, which represents especially the pelitic migmatites of

Olkiluoto, is biotite – sillimanite – plagioclase (An20) – quartz – H2O. Phase relations in

the system are shown in Figs. 10 and 11. The databases used are as above and the

composition of the system is shown in Table 1. Biotite composition used in calculation

corresponds the average of biotite in metapelites in Olkiluoto. It is relatively magnesian,

having only slightly more iron than magnesium, and has 30 % siderophyllite-eastonite

series in solution. The model system illustrates most common evident migmatization

reaction in Olkiluoto i.e. dehydration-melting reaction of biotite and sillimanite.

Figure 10. P-T phase-diagram of the system biotite – sillimanite – plagioclase(An20) –

quartz – H2O illustrating the most prominent dehydration melting reaction taken place

in the migmatites of Olkiluoto in the presence of excess H2O-fluid. Blue fields illustrate

stability of melt absent and red to orange fields melt present assemblages. Red letters

show disappearance or appearance of some important minerals.

23

Figure 11. Binary phase-diagram of the system biotite – sillimanite – plagioclase (An20)

– quartz – H2O with variable water content. Diagram illustrates, that dehydration

melting of biotite begins at 665 oC and continues up to 825

oC with all compositions in

the binary system.

24

Table 1. Composition matrix of the model system used to illustrate phase relations in

migmatitic metapelites of Olkiluoto. This kind of rock contains approximately equal

volume of all solid phases biotite, sillimanite, oligoclase and quartz +/- water fluid.

Phase Si Al Fe n Mg Ca Na K O H

Bt 2.7 1.7 1.4 1 1.2 0 0 0.9 12 2

Sil 2 4 0 2 0 0 0 0 10 0

Pl(An20) 2.8 1.2 0 1 0 0.2 0.8 0 8 0

Qtz 4 0 0 4 0 0 0 0 8 0

H2O 0 0 0 2 0 0 0 0 2 4

Tot. 11.5 6.9 1.4 1.2 0.2 0.8 0.9 40 6

Phase Si Al n Ca Na K O H

Bt 2.7 1.7 1.4 1 1.2 0 0 0.9 12 2

Sil 2 4 0 2 0 0 0 0 10 0

Pl(An20) 2.8 1.2 0 1 0 0.2 0.8 0 8 0

Qtz 4 0 0 4 0 0 0 0 8 0

H2O 0 0 0 0 0 0 0 0 0 0

Tot. 11.5 6.9 1.4 1.2 0.2 0.8 0.9 38 2

Phase Si Al n Ca Na K O H

Bt 2.7 1.7 1.4 1 1.2 0 0 0.9 12 2

Sil 2 4 0 2 0 0 0 0 10 0

Pl(An20) 2.8 1.2 0 1 0 0.2 0.8 0 8 0

Qtz 4 0 0 4 0 0 0 0 8 0

H2O 0 0 0 10 0 0 0 0 10 20

Tot. 11.5 6.9 1.4 1.2 0.2 0.8 0.9 48 22

Figure 10 is a phase diagram of the system with nH2O = 2 (Table 1) in 500 – 7500 bar

and 500 – 900 oC. Blue and blue grey areas are melt absent and thus below the

temperatures of migmatization in Olkiluoto. Because biotite-plagioclase-sillimanite-

quartz-cordierite assemblage is relatively common in Olkiluoto migmatites, especially

in palaeosomes and rocks with lesser amount of migmatization, the dark blue region

well constricts the pre-migmatization pressure-temperature conditions in Olkiluoto. The

first melting in this system takes place in the temperature of 660 - 670 oC in the

presence of sillimanite, which is the case in almost every sample studied so far. It is thus

reasonable to state that everywhere in Olkiluoto the migmatization begun at 660 – 670 oC and 2.9 – 7.2 kbar. The lower pressure can be further delimited to 3.4 kbar, because

potash feldspar is rarely present in mesosomes in Olkiluoto. On the other hand, the

abundance of cordierite in mesosomes will confine the upper pressure during beginning

of melting to 5.9 kbar. In summary, the migmatization in Olkiluoto begun at 660 – 670 oC and 3.4 – 5.9 kbar. The absence of garnet in the migmatites would suggest still lower

upper pressure of 5.5 kbar but because garnet may be absent due to magnesian and/or

less aluminious bulk composition, this cannot be ascertained from this diagram alone.

Because biotite is always present in Olkiluoto, the upper temperature limit of the

stability of biotite in this diagram, ranging from 812 to 855 oC with increasing pressure,

is also the maximum possible peak metamorphic temperature in the area.

Figure 11 is a binary phase diagram that illustrates the effect of water in the system at

the pressure of 4000 bar. The migmatization (melting) begins at about 665 ºC and is

almost independent of water content of the system. The melting temperature is raised to

25

about 680 ºC only at very low water content (XH2O = 0, i.e. XBulk(2) = 1). Thus, a

dehydration melting could be accomplished by influx of fluid at 665 – 680 oC in this

system in dry starting conditions. It is also seen that K-feldspar is favored in the expense

of sillimanite in the dehydration melting of biotite and sillimanite by relatively dry

conditions, similarly as to muscovite dehydration melting (compare Figs. 9b and 11).

On the basis of mineral assemblages, we can summarize that migmatization in Olkiluoto

was most probably caused by dehydration melting of biotite, plagioclase, sillimanite and

quartz due to heating in the temperature above 660 oC and pressure of 3.4 – 5.9 kbar.

5.1.2 Textures Related to Heating

Reaction textures are quite useful in relative timing of metamorphic evolution in

Olkiluoto. The growth of cordierite, which is produced by dehydration melting of biotite

etc., is clearly post-tectonic in relation to the major penetrative foliation S2 of the

palaeosomes in Olkiluoto. Growing cordierite porphyroblasts have replaced oriented

matrix mica and sillimanite as well as plagioclase and the foliation is seen as a relic

helisitic texture of the biotite and sillimanite inclusions in cordierite. Moreover,

cordierite is practically undeformed contrary to all matrix minerals. These relations are

visible in Figs. 6 and 12. The penetrative foliation preceding the peak of metamorphism

is considered to represent the S2 composite foliation in Olkiluoto. The post-S2 nature of

migmatization is also visible macroscopically in outcrops (Kärki and Paulamäki 2006,

Figure 3-7).

Figure 12. BE-image of a cordierite poikiloblast in migmatitic metapelite. Oriented

biotite and sillimanite inclusions in cordierite are parallel to S2 foliation, and the

cordierite is thus post-tectonic in relation to D2 deformation. Cordierite is undeformed

in this sample (OL 197).

26

In granodiorite gneisses, at least the rims of garnet have internal relic foliation (Figure

13a). This is best explained assuming that the garnet was produced by post-tectonic

heating and the penetrative foliation in granodiorite gneisses is composite S2 similarly to

metapelites. Also, in biotite gneisses and hornblende gneisses at least the rims of garnets

have often internal foliation, indicating that they are post-tectonic relative to S2

foliation (Figure 13b), but in these rocks the texture is often obliterated due to the later

decompression and cooling textures (see below).

Figure 13. Micrographs of garnet poikiloblasts showing internal foliation S2. The

poikiloblasts are post-tectonic in relation to D2 deformation. A) Granodiorite gneiss OL

1xx. B) Biotite gneiss OL 1xx.

27

5.1.3 Textures Related to Decompression

Replacement of garnet by plagioclase is the most common decompression texture in

Olkiluoto. It is visible in every garnet-bearing sample in biotite gneisses, hornblende

gneisses, and in tonalitic and granodioritic gneisses in lesser amount, and even in

granitic sample OL108 (Figs. 14 and 15). The texture is a result of decreased stability of

dense minerals (garnet) relatively to less dense in reduced pressure, as already stated by

Eskola (1914, 1921). Decompression textures seem always to be post-tectonic relative

to S2 foliation (Figure 14).

Figure 14. Relics of a garnet poikiloblast in biotite gneiss (white, some pointed by

arrows). BE-image of Sample OL 142.

28

Figure 15. BE-image of a garnet grain partly replaced by plagioclase in pegmatitic

granite gneiss OL 108.

5.1.4 Textures Related to Cooling

As already noted, low temperature hydration or carbonatization reaction/alteration

products like sericite, carbonate, zoisite etc. are quite common in Olkiluoto. These are

relatively late and often associated with fractures, which were generated after the ductile

stage of deformation in Olkiluoto. The late minerals are already discussed in several

Posiva reports (see references in Kärki and Paulamäki 2006) and, thus, will not be

discussed more in the present paper.

Some of the hydration products, however, were apparently generated in the ductile or

semi-ductile cooling stage and deserve special attention here. The first of these is seen

in sample OL 108, where muscovite and sillimanite are found in shear bands enveloping

the felsic porphyroclasts. The texture seems to be largely a result of replacement of

sillimanite by muscovite during shearing and is in accordance with the reaction 13 in the

model system muscovite-plagioclase-quartz, which denotes transition from K-feldspar-

plagioclase-quartz-muscovite-sillimanite (+LQ) to K-feldspar-plagioclase-quartz-

muscovite (+LQ) assemblage. According to model system, this would take place in the

temperature of 655 oC (at four kbar) in a low water activity. The model system would

thus set relatively limited T-aH2O conditions in pegmatite during this shearing-

deformation. According to Kärki and Paulamäki (2006), the shearing in the pegmatites

is caused by D3 deformation, which thus took place in relatively high temperature.

29

Another interesting hydration reaction is seen in sample OL 183, where cordierite

porphyroblasts are replaced by andalusite, chlorite and muscovite (Figure 16). This,

almost isochemical cooling reaction is post-tectonic relative to D2, and will be

considered below in section which describes phase petrology of Olkiluoto ( 5.2) to set

further PT constraints for the cooling stage.

Replacement of garnet by biotite is observed in most garnet-bearing potassic rocks, both

in biotite gneisses and granodiorite gneisses. This is often accompanied with minor

generation of muscovite and replacement of potash feldspar by plagioclase or

myrmekite, especially in granodiorite gneisses.

Cooling has also produced some common exsolution textures. Exsolution of albite from

K-feldspar is a common feature in granite gneisses, granodiorite gneisses and

migmatites. We have analysed the pre-exsolution composition of some microperthitic

feldspar using average of several analyses with large microprobe beam diameter. The

primary albite content exceeds 20 %, which denotes in ternary feldspar system pre-

exsolution temperatures well exceeding 500 oC. In cummingtonite- and hornblende-

bearing sample OL 184 both amphiboles exist as exsolution lamellas in each other.

Similar exsolution lamellas were estimated to form during cooling from 640 to 580 ºC

by Klein et al. (1996).

Myrmekite commonly replaces potash feldspar in Olkiluoto. It was formed by enhanced

stability of micas in the expense of feldspar during cooling.

Figure 16. Micrograph of a cordierite porphyroblast replaced by andalusite, chlorite

and muscovite. Sample OL 183.

30

5.1.5 Textures of the TGG Series

Although the textures of the TGG series rocks largely are result of crystallization from

magma (Figure 5) the rocks commonly also exhibit penetrative foliation S2 (Kärki and

Paulamäki 2006, Fig 3-6) and show evidence of annealing like grain boundary

migration and subgrain development. TGG rocks are considered thus to be pre-tectonic

or latest syn-tectonic with D2 deformation.

5.2 Phase Petrology and Petrogenetic Grids

Chemical composition of c. 250 rock samples in Olkiluoto have been previously

determined by Gehör et al. (1997 …2007). In this section, we present selected phase

diagrams calculated for specific rock compositions in Olkiluoto. The results may be

used to set PT-constraints for the metamorphic evolution of Olkiluoto as the modal

mineral composition of the rock samples has also been quantified. Furthermore, from

textural evidence it is sometimes possible to deduce how the assemblages changed

during the metamorphic evolution and thus make conclusions on the PT evolution. The

chemical analyses were recalculated to ionic basis and the amount of ferric iron was

either calculated from Fe3+

= Fetot/5.5 or assumed to be zero. Because some rocks

contain substantial amounts of P2O5 or sulphur, sufficient number of calcium was

allocated to phosphorus to form apatite and iron to sulphur to form pyrrhotite and

removed from the total amount of these elements. The calculation conditions are given

in each phase diagram. The data set holland2 is a slightly modified version of the set by

Holland and Powell (1998) and White et al. (2007).

5.2.1 Biotite Gneisses

Figure 17 illustrates phase relations in sample OL 103. Stability field of the assemblage

observed in the rock is coloured in blue. A small amount of interstitial potash feldspar is

found in the rock. It is most probably generated during cooling from the blue field in the

direction of red arrows, in the temperature below 635 oC. Red labels (Crd in, Ms in, LQ

in, Kfs in) illustrate where cordierite, muscovite, silicate melt or potash feldspar would

be generated in the rock in changing the PT-conditions.

Phase relations of the sample OL 142 are illustrated in Figure 18. Again, the stability

field of the observed assemblage is coloured in blue. In this rock, the observed

assemblage plagioclase-biotite-quartz-garnet is stable in a relative wide PT-region. Red

labels indicate, where melt (LQ), cordierite, muscovite and K-feldspar should first

appear in the rock with the changing P and T. There are no signs of partial melting in

this sample, so the maximum temperature during metamorphism did not exceed 700 oC.

Also in this sample, there are tiny late, interstitial microcline grains. These were

probably generated in low pressure during cooling underneath 535 oC.

Sample OL 184 is banded, contains relatively complex exsolved amphiboles and

represents thus hardly an equilibrium assemblage in the hand specimen scale.

Nevertheless, calculated phase relations for the bulk composition are illustrated in

31

Figure 19. Because the sample contains substantial hornblende, it is expected that the

tremolite field would be suggestive for PT-conditions of metamorphism. Blue field is

tremolite stability field calculated by Holland and Powell (1998) data set and turquoise,

for comparison, the stability field of tremolite (not existing with carbonate) calculated

with the set of Berman (1988). The diagram suggests approximately the same PT

conditions as previous diagrams (Figs. 17 and 18).

Figure 17. Phase diagram calculated with the composition of biotite gneiss sample OL

103. Blue field shows the stability field of the mineral assemblage observed in the rock.

Red arrows indicate the PT evolution which leads to the generation of small amounts of

potash-feldspar in the rock.

32

Figure 18. Phase diagram calculated with the composition of biotite gneiss sample OL

142. Blue field shows the stability field of the mineral assemblage observed in the rock.

33

Figure 19. Phase diagram calculated with the composition of biotite-hornblende gneiss

sample OL 184. Blue and turquoise fields show the stability field of tremolite. Blue field

calculated by the database of Holland and Powell (1998) and turquoise by that of

Berman (1988).

5.2.2 Hornblende Gneisses

Phase relations of sample OL 118 were calculated using Berman (1988) data set and are

shown in Figure 20. The field of the assemblage observed in the rock is coloured darker

blue. The reason for lower temperature than expected from above discussion is most

probably inadequate solution data for amphiboles. If hornblende was neglected, the field

would be enlarged to higher temperatures, colored with lighter blue.

34

Figure 20. Phase diagram calculated with the composition of hornblende gneiss sample

OL 118. Blue field shows the stability field of the mineral assemblage observed in the

rock. Diagram calculated by the database of Berman (1988). Lighter blue field shows

the stability field of the rock assemblage when hornblende is neglected.

5.2.3 Granite Gneisses

Phase relations for sample OL 108 were calculated by Holland and Powell (1998) data

set and are shown in Figure 21. The dark red narrow area shows the PT-field where all

the phases observed in the rock occur in equilibrium with each other. Univariant curves

where sillimanite, biotite and melt would disappear are shown by symbols “Sil out”, “Bt

out” and “LQ out”, respectively. On textural basis, it is not very probable that large

amount of melt were present when sillimanite was replaced by muscovite during D3-

shearing. Thus it is expected that sillimanite out and liquid out univariants should lie

near each other, which is the case, if the rock suffered cooling in as low pressure as

possible allowed by the stability fields presented in Figure 21. Thus, a probable cooling

path for this rock is drawn by red arrow in the lower part of the stability fields. The path

35

is also supported by the low amount of biotite, found only as inclusions in quartz in the

sample as biotite should disappear from the assemblage slightly before sillimanite. Error

ellipses derived from thermobarometric calculation are also shown in the figure and will

be discussed below.

Figure 21. Phase diagram calculated with the composition of granite gneiss sample OL

108. Red field is the stability field of peak temperature assemblage of the rock (in

partially molten stage). Blue field shows the field of subsolidus assemblage (with

muscovite).

36

5.2.4 Granodiorite Gneisses

The composition of sample OL145 was used to calculate phase equilibria representing

granodiorite gneisses. The phase diagram is shown in Figure 22 where the large blue

area represents the equilibrium assemblage of the rock. This diagram seems not to help

a lot in first sight. However, the amount of muscovite should increase in the assemblage

during cooling until finally K-feldspar would totally be replaced at about 420 oC.

Because this is not observed in the rock, it is likely that the mineral assemblage reached

equilibrium in the higher temperature side of the blue field. The red line represents the

beginning of melting of the mineral assemblage of the sample, or reversely, a reaction

curve where melt would disappear during cooling from silicate liquid having the

composition of the sample and the observed assemblage would become stable. In other

words, the observed assemblage of the rock represents solidus assemblage on the red

curve and the curve gives crystallization conditions of the sample. On the structural and

textural basis discussed before, granodioritic rocks were deformed by D2 deformation

phase, which is a pre-metamorphic deformation stage. It thus seems probable, that the

granodioritic sample cooled below 650 ºC before or during D2 phase and was not heated

above (or much above) this temperature during subsequent regional metamorphism. It is

important to observe, that the pressure during the crystallization of the granodiorite

magma was above 3.5 kbar, as there is no sillimanite or cordierite or any evidence of

their previous existence in the rock. Figure 23 illustrates the amount of garnet in the

rock or crystal-melt mush in the bulk composition of sample OL 145. This was done

trace the development of garnet in changing PT-conditions. It is clear, that largest

amount of garnet (80 cm3/~3800 cm

3 of mush) would occur in the field of partially

molten stage.

37

Figure 22. Phase diagram calculated with the composition of granodiorite gneiss

sample OL 145. Blue field is the stability region of the observed subsolidus assemblage

of the rock and red curve is the solidus. The area to the right of the red curve is the

region where melting of the rock assemblage should begin.

38

Figure 23. Diagram illustrating the amount of garnet in the composition of

granodiorite gneiss sample OL 145. Comparison to figure 22 shows that largest amount

of garnet would exist in this rock in hypersolidus conditions (i.e. in partially molten

stage).

5.2.5 Migmatitic Gneisses

Because the bedrock of Olkiluoto consists mostly of migmatitic gneisses, we calculated

several phase diagrams of them, to illustrate the general metamorphic evolution and

especially migmatization processes of the area. Most of the samples were selected from

drill core OL-KR1 from a depthinterval of about 350 – 740 m. Therefore, this section

represents the central area in the Olkiluoto study site.

Figure 24 is a phase diagram for the composition of sample OL 133. Red area is the

stability field of assemblage observed in the rock and light red the stability field when

sillimanite is removed from the rock assemblage. Small green ellipse gives the region

where first melting most probably took place, as above this pressure garnet should be a

solidus phase and below the area sillimanite should not be a solidus phase, and, there

were no signs of relic garnet in the rock but sillimanite relics are common. Therefore,

the PT-conditions for the beginning of melting in this rock may be defined quite

accurately to have been 2.9 – 3.2 kbar and 660 oC.

39

Figure 24. Phase diagram calculated with the composition of migmatitic gneiss sample

OL 133. Red field is the stability region of the observed assemblage of the rock. Lighter

red field is the stability region if sillimanite would be removed from the assemblage.

Green ellipse shows PT conditions where the migmatization of this rock began.

40

Figure 25 is a phase diagram for the bulk composition of sample OL 134. It is almost

identical to Figure 24 and similar conclusions can be drawn for the evolution of this

sample.

Figure 25. Phase diagram calculated with the composition of migmatitic gneiss sample

OL 134. Red field is the stability region of the observed assemblage of the rock. Lighter

red field is the stability region if sillimanite would be removed from the assemblage.

Green ellipse shows PT conditions where the migmatization of this rock began.

41

Figure 26 similarly shows the phase relations for the bulk composition of sample

140. This is also nearly similar to the previous two diagrams, although melting would

take place in this composition after the disappearance of sillimanite due to the lower

content of water in the system. However, there is certain uncertainty on the water

content of the system during metamorphism, as the estimate of volatile content of the

rock sample is based on the loss on ignition. Small excess of water would change this

diagram identical to previous ones.

Figure 26. Phase diagram calculated with the composition of migmatitic gneiss sample

OL 140. Red letters show, that in this composition sillimanite would disappear from the

assemblage before the melting begins. Small increase in water content would shift the

melt in (LQ in) curve into the stability field of sillimanite.

42

Figure 27 is similar diagram for the bulk composition of sample OL 183. This sample is

different from the previous ones because it does not contain potash feldspar. The

equilibrium assemblage of the partially molten rocks is shown in red color and the

lighter red field shows an area where sillimanite would not exist with the other phases

involved. Again, the PT-conditions are quite strictly defined by the assemblage (dark

red field) and the beginning of melting is still more strictly defined as shown by a small

green ellipse (665 – 670 oC and 3.7 – 4.3 kbar). It is possible that melting took place in a

little lower pressure from 3 to 3.7 kbar, but then sillimanite would have tried to react out

of the assemblage before the beginning of melting at 3 – 3.5 kbar or simultaneously

with melting at 3.5 -3.7 kbar.

Figure 27. Phase diagram calculated with the composition of migmatitic gneiss sample

OL 183. Red field is the stability region of the observed assemblage of the rock. Lighter

red field is the stability region if sillimanite would be removed from the assemblage.

Green ellipse shows PT conditions where the migmatization of this rock began.

Diagram is identical to those presented in Figs. 24 and 25.

43

In addition, the PT-conditions during the cooling phase may be considered by phase

equilibria. Sample OL 183 is outstandingly useful for this purpose as it contains almost

isochemical replacement of cordierite by andalusite, chlorite and muscovite (Figure 16).

Figure 28 is a phase diagram calculated for the composition of cordierite in the sample.

Light blue area shows the region, where cordierite is stable, dark blue narrow field is the

region where cordierite is stable together with the alteration products and yellow field is

the stability region of alteration product assemblage alone. The cooling path of the rock

has necessarily passed the dark blue region, which quite strictly defines the passing

point on the line between 575 ºC and 4.0 kbar, and 470 ºC and 0.5 kbar.

Figure 28. Phase diagram calculated with the composition of cordierite in migmatitic

gneiss sample OL 183. Light blue field is the stability region of cordierite, blue is a

region, where cordierite exists in equilibrium with its alteration products andalusite,

chlorite and muscovite and yellow field is the stability region of alteration products

alone. Compare to micrograph taken from the thin section of the rock (Figure 16).

44

Phase equilibria calculation may also be used to illustrate metamorphic evolution of

Olkiluoto by following the progression of metamorphism and especially volume

changes of solid and liquid phases. This approach is efficient especially with migmatitic

rocks, where volume ratios of melt and solid may be macroscopically estimated already

from the outcrops in the field, tunnel walls, or from drill cores during logging. Volume

change may be estimated or calculated from leucosome/ (meso+melanosome) ratio

assuming no loss of melt from the system, which was also the assumption of Kärki and

Paulamäki (2006). This kind of calculation is illustrated in Figure 29 where volume

changes in the composition of sample OL134 are shown in constant pressure of four

kbar, which is reasonable in Olkiluoto, and temperatures from 460 – 850 oC, also well

covering the temperature range of Olkiluoto. Logarithmic scale is used to better

illustrate the amounts of minor phases, which might be useful in microscopic

petrographic studies. In the Figure, typical cordierite-bearing migmatite in Olkiluoto

falls between 710 and 750 oC. Slightly lower pressure would expand the field to 660 –

760 oC, which is seen by comparison to Figure 25.

45

Figure 29. Diagram demonstrating changes of volume proportions of minerals, total

volume of solids and volume of silicate liquid in the of bulk composition of migmatite

gneiss sample OL 134 during progressive heating from 550 to 850 ºC in the pressure of

four kbar.

46

Figure 30 is a phase diagram for the composition of sample OL 135. The red colour

covers the stability field of the assemblage observed in the rock and light red shows the

expansion of the field if sillimanite would be removed from the assemblage. This

sample contains only small amount of K-feldspar. The reason for this can be seen by

comparing the Figure to Figs. 24 – 27. Potash feldspar should become stable in this

composition in relative low pressure or with profound melting in larger pressure. Again,

the calculated phase diagram is consistent with the observed assemblage.

Figure 30. Phase diagram calculated with the composition of migmatitic gneiss sample

OL 135. Red field is the stability region of the observed assemblage of the rock. Lighter

red field is the stability region if sillimanite would be removed from the assemblage.

Green ellipse shows PT conditions where the migmatization of this rock began.

Diagram is identical to those presented in Figs. 24, 25 and 27.

47

5.3 Thermobarometry

Pressure-temperature conditions were calculated by winTWQ program written by

Berman (2007). The number of ions in the mineral formula was calculated by the

program winCMP of the package, including ferric iron calculation and 3 % ferric iron

consideration for garnet. The error propagation is based on the reaction intersections

with exclusion of those points, which were outside one of average.

5.3.1 Biotite Gneisses

PT-conditions calculated for the sample OL 103 are given in Figure 31. Garnet in this

sample has different core (c) and rim (r) composition (Table 2). Because the amount of

garnet is small in this sample relative to that of plagioclase and biotite, the composition

of the latter probably have not changed much, due to the mass balance, during to

retrograde Fe/Mg exchange between garnet and biotite. The core of garnet may thus be

used with matrix biotite and plagioclase to calculate conditions during the peak of

regional metamorphism. Figures 31A and 31B are intersections calculated by garnet

core, matrix plagioclase and two different matrix biotite analyses. Three different

analysis point combinations were used to calculate the error ellipse, which thus shows

the effect of biotite heterogeneity in the sample. Biotite2 evidently gives the most

reliable result (3.3 kbar, 647 oC) because the two other analysed biotite grains most

probably are alteration products of garnet. The result is in good agreement with the

calculated phase diagram of the sample (Figure 17).

Figure 31. TWEEQU thermobarometry for biotite gneiss sample OL 103. A)

composition of biotite probably replacing garnet used in calculation; B) composition of

matrix biotite parallel to S2 foliation used in calculation. Error ellipse (two σ, 95 %

confidence) in both figures is calculated from intersections of three different mineral

analysis-point groups and illustrates the effect of biotite heterogeneity in the sample.

48

Table 2. Composition of minerals in sample OL 103.

Point SiO2 TiO2 Al2O3 FeO MgO CaO MnO Cr2O3 V2O3 Na2O K2O Total

OL103Grt1r 38.75 0.00 20.38 31.32 2.31 2.28 5.93 0.04 0.02 0.03 0.00 101.06

OL103Grt2c 38.61 0.00 20.27 32.46 2.95 2.24 3.76 0.05 0.02 0.01 0.01 100.36

OL103Grt3r 38.49 0.00 21.04 30.22 2.21 2.26 6.25 0.00 0.04 0.00 0.00 100.51

OL103Pl1 59.01 0.00 25.18 0.08 0.00 8.56 0.01 0.00 0.00 6.65 0.08 99.56

OL103Pl2 58.30 0.00 24.47 0.02 0.00 8.82 0.00 0.02 0.02 6.45 0.15 98.23

OL103Bt2 36.68 3.26 15.90 20.71 8.90 0.03 0.16 0.19 0.13 0.08 9.36 95.41

OL103Bt1 36.89 3.01 16.51 20.54 9.25 0.00 0.18 0.20 0.09 0.07 9.23 95.96

OL103Bt3 36.50 2.96 17.06 20.23 9.44 0.00 0.18 0.14 0.04 0.09 9.49 96.14

OL103Ilm1 0.04 53.33 0.00 44.20 0.15 0.02 2.22 0.08 0.21 0.00 0.02 100.26

Figure 32 shows PT-conditions calculated for the sample OL 142. The amount of garnet

is small and due to the mass balance effect, the composition of the core of garnet can be

used together with that of matrix minerals to calculate the conditions during the peak of

regional metamorphism. The result, 4.3 kbar and 642 oC, is again in good agreement

with the equilibrium assemblage calculated for this rock in these conditions (blue field

in Figure 18). The compositions of minerals in the sample are given in Table 3.

49

Figure 32. TWEEQU thermobarometry for biotite gneiss sample OL 142.

Table 3. Composition of minerals in sample OL 142.

Point SiO2 TiO2 Al2O3 FeO MgO CaO MnO Na2O K2O Total

OL142Grt2c 38.88 0.00 21.79 30.23 3.62 2.08 5.44 0.05 0.00 102.08

OL142Grt1r 38.57 0.00 21.43 28.58 2.64 2.20 8.13 0.00 0.00 101.53

OL142Bt1 36.82 2.84 17.33 18.17 10.34 0.00 0.18 0.10 9.21 94.99

OL142Pl1 58.48 0.02 24.67 0.00 0.00 8.31 0.03 6.71 0.17 98.38

50

Figure 33 shows PT-conditions calculated for sample OL 184. Composition of garnet

core is used. Because the amount of biotite is relatively small in this sample, and the

composition of the rim of garnet has evidently changed during cooling due to Fe-Mg

exchange with biotite, the composition of biotite is corrected by 10% for this exchange.

Figure 33. TWEEQU thermobarometry for biotite gneiss sample OL 184.

Figure 34 shows the error ellipses calculated, when matrix biotite composition is used

without correction and when the correction is made. The average P and T are

correspondingly 613 ºC, 3.596 kbar and 654 ºC, 3.937 kbar. The values are within the

stability field of tremolite bearing assemblage of the bulk composition, calculated

earlier with Berman (1988) and Holland and Powell (1998) thermodynamic data sets

(Figure 19). Compositions of minerals are given in Table 4.

51

Figure 34. TWEEQU thermobarometry for biotite gneiss sample OL 184. Error ellipses

illustrating the effect of retrograde Fe/Mg-exchange in thermobarometry when mass

balance is less effective (i.e. relatively small amounts of matrix mineral biotite present).

Table 4. Composition of minerals in sample OL 184.

Point SiO2 TiO2 Al2O3 FeO MgO CaO MnO Cr2O3 V2O3 ZnO Na2O K2O Total

Ol184Ilm1 0.03 52.68 0.01 41.10 0.09 0.00 5.22 0.01 0.26 0.01 0.00 0.02 99.41

Ol184Bt1 37.69 3.08 16.19 16.45 12.46 0.00 0.35 0.21 0.09 0.02 0.10 9.22 95.85

Ol184Pl1 46.31 0.00 34.59 0.02 0.00 18.54 0.01 0.00 0.00 0.00 1.14 0.01 100.63

Ol184Grt1.2 38.66 0.00 21.89 25.97 3.34 3.59 8.36 0.09 0.03 0.03 0.00 0.02 101.99

Ol184Grt1.1 38.18 0.01 22.07 25.62 2.46 3.61 9.32 0.05 0.04 0.04 0.01 0.00 101.41

Ol184Cum2 54.01 0.03 1.72 21.84 16.50 1.58 2.38 0.07 0.04 0.01 0.06 0.03 98.27

Ol184Hbl3 avg 50.56 0.71 6.62 14.23 14.60 9.50 1.19 0.07 0.08 0.00 0.51 0.24 98.32

Ol184Hbl3 lamel 49.86 0.79 7.55 12.80 13.79 11.03 0.94 0.07 0.10 0.04 0.59 0.33 97.87

Ol184Hbl3 lamel 52.65 0.43 4.15 17.72 15.76 5.78 1.80 0.07 0.05 0.07 0.34 0.12 98.94

OL184Cum3 rim 55.20 0.09 0.80 21.25 17.27 1.05 2.55 0.00 0.02 0.08 0.08 0.02 98.39

Ol184Cum2 pure 54.96 0.06 1.20 21.80 17.25 1.01 2.65 0.03 0.02 0.10 0.12 0.00 99.19

Ol184Hbl2 lamel 48.93 0.62 8.55 15.19 12.91 9.63 1.09 0.32 0.28 0.08 0.80 0.36 98.75

52

Figure 35 shows PT-conditions calculated for sample OL 200. The calculated PT-values

(4.16 kbar, 697 ºC) are similar to those of other biotite gneiss samples.

Figure 35. TWEEQU thermobarometry for biotite gneiss sample OL 200.

Figure 36 shows PT-conditions calculated for sample OL 206. The obtained pressure is

higher than in other samples in Olkiluoto, except the granodioritic sample discussed

below. It could be possible that the core of garnet in this case already crystallized during

the intrusion of the granodiorite body. This could have been caused by the heat of the

granodiorite magma, and thus there could be two metamorphic stages in Olkiluoto. The

first phase would then have a close connection to the magmatic processes.

53

Figure 36. TWEEQU thermobarometry for biotite gneiss sample OL 206.

Table 5. Composition of minerals in samples OL 200 and OL 206.

Point SiO2 TiO2 Al2O3 FeO MgO CaO MnO Cr2O3 NiO Na2O K2O Total

Ol200Grt1.2 38.60 0.04 21.03 30.29 3.22 2.32 5.97 0.06 0.00 0.00 0.01 101.53

Ol200Grt1.1 38.13 0.00 20.50 29.31 2.53 2.38 8.12 0.00 0.01 0.00 0.00 100.97

Ol200Bt1 36.68 2.78 16.84 18.89 9.91 0.00 0.24 0.24 0.10 0.15 9.10 94.92

OL200Pl1 57.17 0.00 25.51 0.04 0.00 9.45 0.02 0.00 0.04 5.97 0.14 98.33

Ol200Chl 34.91 0.09 14.08 14.51 22.13 0.27 0.25 0.07 0.00 0.13 0.05 86.49

Ol206Grt1c 38.49 0.04 20.96 27.78 2.96 4.39 6.16 0.09 0.02 0.00 0.00 100.87

Ol206Grt2r 38.88 0.00 21.47 26.26 2.35 4.81 7.65 0.01 0.08 0.00 0.03 101.54

Ol206Ilm3 0.00 53.63 0.02 40.41 0.08 0.00 6.62 0.01 0.00 0.05 0.01 100.83

Ol206Bt4 35.93 3.86 15.89 19.38 9.42 0.00 0.24 0.04 0.00 0.08 8.98 93.83

Ol206Pl4 56.61 0.00 25.41 0.01 0.00 10.07 0.02 0.02 0.01 5.79 0.07 97.99

54

5.3.2 Hornblende Gneisses

Figure 37 shows PT-conditions calculated for sample OL118. The result of calculation

is about 50 ºC higher than the stability field of the assemblage (Figure 20). As already

noted, there are great uncertainties in hornblende solution model and the stability field

could actually expand to higher temperature. So the calculated PT-conditions 5.4 kbar

and 644 ºC may be considered reliable, as far as the values are similar to those

calculated for different samples above. The slightly higher pressure could reflect

slightly earlier stage than recorded in other samples. Compositions of minerals in the

sample are given below in Table 6.

Figure 37. TWEEQU thermobarometry for hornblende gneiss sample OL 118.

55

Table 6. Composition of minerals in sample OL 118.

Point SiO2 TiO2 Al2O3 FeO MgO CaO MnO Na2O K2O Total

OL118Grt2 38.66 0.00 20.64 28.24 1.16 9.74 2.83 0.00 0.00 101.27

OL118Grt1 39.16 0.01 21.13 26.27 1.27 11.11 2.56 0.05 0.00 101.55

OL118Hbl1 47.15 0.59 8.04 24.80 5.66 11.61 0.37 0.57 0.34 99.13

OL118Di1 52.28 0.03 0.68 18.50 7.10 22.65 0.46 0.07 0.00 101.76

OL118Ilm1 0.02 53.14 0.01 44.48 0.09 0.04 2.51 0.01 0.03 100.32

OL118Di2 51.99 0.06 0.60 18.66 7.23 22.84 0.48 0.08 0.01 101.95

OL118Hbl3 53.02 0.02 2.68 20.73 9.74 12.74 0.28 0.18 0.09 99.48

OL118Grt3 39.31 0.00 20.93 26.47 1.27 10.83 2.56 0.00 0.00 101.36

OL118Pl1 45.37 0.00 33.77 0.08 0.02 19.41 0.01 0.51 0.00 99.17

OL118Czo1 40.11 0.00 30.27 3.31 0.01 24.41 0.01 0.00 0.05 98.17

OL118Pl2 45.01 0.02 32.88 0.06 0.00 19.20 0.00 0.58 0.03 97.78

56

Figure 38 shows PT-conditions calculated for sample OL121. The data set used was that

of Berman (1988) because it contains thermodynamic parameters and solution models

for several end-members in the hornblende series. Calculated PT-conditions (626 ± 8.2 ºC and 3260 ± 386 bar) are similar, although with somewhat lower average values, to

those obtained above. Table 7 gives the composition of minerals in the sample.

Figure 38. TWEEQU thermobarometry for biotite-hornblende gneiss sample OL 121.

57

Table 7. Composition of minerals in sample OL 121.

Point SiO2 TiO2 Al2O3 FeO MgO CaO MnO Cr2O3 V2O3 Na2O K2O F Total

Ol121Grt1 38.19 0.03 22.10 29.60 2.18 6.47 2.79 0.03 0.03 0.00 0.00 0.00 101.42

Ol121Grt2 38.44 0.01 22.14 29.08 2.13 7.44 2.72 0.03 0.00 0.01 0.02 0.00 102.01

Ol121Hbl1 45.86 0.65 9.27 19.90 8.42 10.92 0.38 0.07 0.01 0.69 0.36 0.00 96.62

Ol121Ilm1 0.00 52.20 0.02 44.89 0.08 0.02 1.49 0.02 0.20 0.02 0.00 0.00 98.92

Ol121Pl1d 44.83 0.00 35.13 0.02 0.00 18.86 0.00 0.06 0.03 0.75 0.02 0.00 99.20

Ol121Hbl1d 45.99 0.66 9.45 20.10 8.40 10.99 0.36 0.11 0.06 0.71 0.27 0.00 97.21

Ol121Grt1d 38.74 0.00 22.31 28.43 2.04 8.17 2.70 0.07 0.04 0.03 0.00 0.00 102.53

Ol121Grt2d 38.65 0.00 22.92 29.97 2.18 6.33 2.91 0.00 0.04 0.00 0.02 0.00 103.00

Ol121Grt1.1 38.04 0.03 22.64 30.14 2.32 4.46 3.99 0.00 0.00 0.00 0.00 0.00 101.61

Ol121Ilm1.1 0.05 52.29 0.04 41.69 0.08 0.00 4.32 0.05 0.21 0.02 0.00 0.00 98.74

Ol121Bt1.1 36.59 2.59 17.39 18.34 10.10 0.00 0.11 0.15 0.23 0.02 9.43 0.10 95.11

Ol121Pl1.1 45.36 0.00 35.10 0.03 0.01 18.00 0.02 0.03 0.02 1.18 0.01 0.00 99.75

Ol121Grt1.2 37.91 0.00 22.54 30.95 2.67 3.81 3.62 0.00 0.00 0.02 0.01 0.00 101.53

5.3.3 Granite Gneisses

Figure 39 shows PT-conditions calculated for sample OL108. In this calculation, we

omitted some equilibria causing intersections outside one outside average. The error

ellipse derived in this calculation is shown in Figure 21 by dark blue colour. It lays

slightly below the calculated phase relations. However, it is still in relatively good

agreement with the phase relations as only small amounts of cordierite should exist in

addition of the observed phases in the field occupied by the ellipse, and cordierite would

surely be the first phase to disappear from the rock during cooling and shearing. In

addition, some discrepancy is produced due to the different data sets used in phase

equilibria and thermobarometric calculation.

A larger open error ellipse in Figure 21 involves all the calculated equilibria, except the

parallel ones, and most probably indicates local unequilibria produced during cooling

phase. This ellipse fits quite well with the cooling path deciphered from the phase

equilibria and rock texture i.e. replacement of sillimanite by muscovite in the shear

zones and relic nature of biotite and garnet.

58

Figure 39. TWEEQU thermobarometry for granite gneiss sample OL 108. Aranovich

and Berman (1996) and Berman (2007) data set used.

Table 8. Composition of minerals in sample OL 108.

Point SiO2 TiO2 Al2O3 FeO MgO CaO MnO Na2O K2O Total

OL108Grt2 38.48 0.02 21.27 32.91 2.69 0.59 6.28 0.06 0.00 102.30

OL108Grt1 38.12 0.01 21.03 32.30 2.40 0.59 6.75 0.03 0.02 101.26

OL108Pl1 64.93 0.00 21.12 0.04 0.00 3.50 0.03 9.63 0.11 99.37

OL108Pl2 65.42 0.00 21.24 0.00 0.00 3.41 0.00 10.02 0.08 100.16

OL108Kfs 65.45 0.00 17.37 0.00 0.00 0.00 0.00 0.58 15.69 99.09

OL108Chl1 25.24 0.10 20.55 30.14 10.76 0.00 0.71 0.03 0.02 87.54

OL108Chl1b 25.18 0.14 20.24 30.11 10.74 0.00 0.74 0.02 0.03 87.19

OL108Bt3 35.23 3.21 16.84 21.72 7.55 0.00 0.24 0.08 9.29 94.16

59

Another calculation was made with Berman (1988) data set. The results of this

calculation are shown in Figure 40. The error ellipse from this calculation is shown in

green colour in Figure 21 and is even in better agreement with the phase equilibria

calculation and proposed PT-path for the cooling stage of metamorphism. Compositions

of minerals in the sample OL 108 are given below in Table 8.

Figure 40. TWEEQU thermobarometry for granite gneiss sample OL 108. Berman

(1988) data set used.

60

5.3.4 Granodiorite Gneisses

Figure 41 shows PT-conditions calculated for the granodioritic gneiss sample OL145.

The core of garnet was used in the calculation in order to find out whether the garnet

composition used represent igneous crystallization conditions or conditions of regional

metamorphism, which seems to be somewhat later event. The obtained conditions, as

compared to the phase diagram (Figure 22), seem to be well in hypersolidus conditions

(i.e. where melt exists in the rock). Therefore, the garnet core could be of igneous rather

than metamorphic origin. It should be noted that biotite is an important phase used in

the calculation and it will change its composition readily due to changing conditions.

However, the great amount of biotite compared to that of garnet serves efficiently as

mass balance buffer, at least what comes to Fe/Mg exchange reaction, which has the

greatest effect on temperature estimate.

Figure 41. TWEEQU thermobarometry for granodiorite gneiss sample OL 145.

Composition of the core of garnet grain used in calculation.

61

Figure 42 shows the results of calculation, when the composition of the rim of garnet is

used. The obtained pressure and temperature are dramatically lower than those obtained

from the core and do in no means correspond the peak conditions of regional

metamorphism.

When these results are compared to Figure 23, which illustrates the amount of garnet in

the system, it is obvious that the conditions calculated from the rim composition

represent the stage where garnet was replaced by matrix plagioclase. This is in

accordance with the textural evidence. The development of garnet, that is grown in post-

or syn-D2 stage of deformation remains unresolved with the present data alone. Detailed

garnet composition mapping would be necessary to decipher the crystallization history

of garnet more precisely.

Figure 42. TWEEQU thermobarometry for granodiorite gneiss sample OL 145.

Composition of the rim of garnet grain used in calculation.

62

Table 9. Composition of minerals in sample OL 145.

Point SiO2 TiO2 Al2O3 FeO MgO CaO MnO Na2O K2O Total

OL145Grt1 37.89 0.00 20.97 34.65 1.78 0.47 4.93 0.00 0.01 100.68

OL145Grt2 38.26 0.02 21.24 35.00 2.65 0.48 3.80 0.02 0.00 101.46

OL145Grt3 38.51 0.00 21.24 34.35 3.49 0.47 2.97 0.03 0.00 101.04

OL145Bt2 35.44 0.06 18.73 21.91 7.90 0.00 0.05 0.15 8.94 93.18

OL145Bt1 35.67 1.24 19.65 21.32 7.30 0.00 0.19 0.12 9.45 94.93

OL145Kfs1 65.21 0.08 17.58 0.00 0.00 0.04 0.00 0.68 15.46 99.05

OL145Pl1 65.57 0.00 20.58 0.03 0.00 2.74 0.00 9.72 0.21 98.84

OL145Pl2 65.91 0.03 20.10 0.02 0.02 2.77 0.01 9.82 0.10 98.78

OL145Ms1 46.26 0.92 33.36 0.93 0.52 0.04 0.00 0.62 10.37 93.02

Figure 43 shows PT-conditions calculated for the granodiorite sample OL188. The

results (648 ± 3 ºC and 3231 ± 61 bar) obtained from this granodiorite are similar to

those of other rock types in Olkiluoto.

63

Figure 43. TWEEQU thermobarometry for granodiorite gneiss sample OL 188.

Table 10. Composition of minerals in sample OL 188.

Point SiO2 TiO2 Al2O3 FeO MgO CaO MnO Na2O K2O ZnO Total

Ol188Grt1.2 38.14 0.00 20.46 35.58 2.36 0.90 3.70 0.00 0.01 0.07 101.22

Ol188Grt1.1 37.96 0.00 20.62 33.34 1.58 0.69 6.54 0.00 0.02 0.00 100.75

Ol188Bt1 35.01 1.55 19.59 22.26 6.70 0.00 0.22 0.14 9.56 0.09 95.32

Ol188Pl1 63.75 0.00 21.89 0.02 0.00 4.37 0.00 8.95 0.10 0.00 99.08

Ol188Bt2 35.65 3.43 18.27 22.29 5.92 0.00 0.18 0.16 9.36 0.15 95.42

Ol188Kfs2 65.07 0.00 18.22 0.04 0.00 0.00 0.00 0.92 14.85 0.03 99.13

Ol188Pl3.2 63.30 0.00 21.59 0.00 0.01 4.66 0.02 8.93 0.32 0.00 98.82

Ol188Pl3.1 62.74 0.01 21.60 0.00 0.01 4.35 0.01 8.56 0.16 0.00 97.44

Ol188Pl3.1b 62.86 0.00 21.71 0.00 0.00 4.72 0.02 8.30 0.31 0.00 97.92

64

5.3.5 Migmatitic Gneisses

The results obtained by TWQ calculation on the sample OL 133 are shown in Figure 44.

The temperature obtained is much lower than would be expected from the phase

diagram (Figure 24). The obvious reason for this must be retrogressive re-equilibration

of the phases due to the water fluid released from the melt fraction of migmatites, when

they crystallized during cooling. It seems that migmatites are not so resistant against

retrogression as non-migmatitic rocks. This is obvious also from the calculation results

for samples OL 135 (Figure 45) and OL 183 (Figure 46), although the latter seems to

give results, which are more consistent, even somewhat dispersed, with other criteria

derived from Olkiluoto. It should also be kept in mind that migmatitic gneisses often

contain relic phases (i.e. sillimanite) which do not necessarily belong to the equilibrium

assemblage. Compositions of minerals in the migmatites of Olkiluoto are given in Table

11.

Figure 44. TWEEQU thermobarometry for migmatitic cordierite gneiss sample OL 133.

65

Figure 45. TWEEQU thermobarometry for migmatitic cordierite gneiss sample OL 135.

66

Figure 46. TWEEQU thermobarometry for migmatitic cordierite gneiss sample OL 183.

67

Table 11. Composition of minerals in migmatite samples OL 133, 134, 135 and 183.

Point SiO2 TiO2 Al2O3 FeO MgO CaO MnO Cr2O3 NiO ZnO Na2O K2O Total

Ol133Crd1a 49.49 0.01 33.00 9.69 7.18 0.00 0.20 0.01 0.04 0.04 0.24 0.01 99.89

OL133Sil1 37.44 0.00 63.82 0.43 0.00 0.00 0.01 0.11 0.00 0.00 0.02 0.01 101.84

Ol133Bt1 36.15 3.02 21.01 19.03 7.34 0.00 0.07 0.11 0.00 0.11 0.03 9.53 96.38

133Pl1 64.16 0.00 23.66 0.06 0.00 4.24 0.00 0.00 0.03 0.04 8.94 0.10 101.22

OL133Kfs1a 65.85 0.01 19.16 0.01 0.00 0.05 0.00 0.00 0.00 0.00 1.29 14.63 101.01

OL133Kfs1b 65.53 0.00 18.97 0.02 0.00 0.05 0.00 0.00 0.10 0.05 1.88 13.43 100.01

Ol133Kfs1c 66.64 0.00 19.03 0.00 0.00 0.08 0.03 0.00 0.00 0.00 2.32 13.18 101.28

Ol133Kfs1d 65.87 0.00 19.04 0.00 0.00 0.07 0.00 0.05 0.00 0.01 2.03 13.61 100.68

Ol134Tur1 39.37 0.87 33.65 6.73 5.53 0.53 0.00 0.04 0.00 0.02 1.93 0.08 88.75

ol134Pl1 63.30 0.00 21.47 0.06 0.00 4.78 0.01 0.03 0.00 0.05 9.33 0.25 99.27

ol134Kfs1 65.30 0.00 17.44 0.05 0.00 0.03 0.00 0.00 0.00 0.00 1.60 14.33 98.75

OL135Kfs1.2 65.05 0.00 17.20 0.05 0.00 0.02 0.01 0.00 0.01 0.00 1.74 13.78 97.87

OL134Kfs1.3 65.03 0.00 17.25 0.00 0.00 0.07 0.01 0.00 0.10 0.00 2.33 13.51 98.29

OL134Bt1 35.67 3.25 18.54 20.18 7.41 0.00 0.03 0.11 0.01 0.05 0.13 9.39 94.77

OL134Bt1.2 35.75 2.91 19.26 19.97 7.36 0.00 0.07 0.12 0.06 0.03 0.16 9.49 95.18

Ol134Crd1.2c 48.78 0.01 29.88 9.94 6.97 0.00 0.19 0.00 0.00 0.00 0.24 0.00 96.01

OL134Crd1.3b 48.80 0.00 29.62 9.80 7.03 0.02 0.23 0.01 0.01 0.01 0.23 0.00 95.73

OL135Pl1 62.43 0.00 21.73 0.00 0.00 5.69 0.00 0.01 0.00 0.03 8.85 0.19 98.92

OL135Bt1 35.75 3.03 18.23 19.44 7.77 0.00 0.08 0.08 0.08 0.03 0.11 9.57 94.17

OL135Kfs1 65.15 0.00 16.83 0.01 0.00 0.03 0.03 0.00 0.03 0.04 2.26 13.37 97.75

OL135Crd1.2c 48.95 0.01 30.78 9.31 7.35 0.00 0.19 0.03 0.07 0.00 0.19 0.01 96.87

Ol135Crd1.1r 48.69 0.03 30.19 9.13 7.61 0.05 0.19 0.00 0.03 0.00 0.17 0.02 96.10

Ol183Pl1 64.03 0.00 22.91 0.00 0.00 4.23 0.00 0.03 0.03 0.04 9.78 0.46 101.50

Ol183Kfs1 66.37 0.05 18.52 0.00 0.00 0.01 0.01 0.00 0.03 0.00 1.32 14.73 101.04

Ol183And2 38.09 0.02 62.78 0.23 0.01 0.04 0.02 0.01 0.00 0.03 0.00 0.01 101.25

Ol183Crd2 48.92 0.00 32.64 8.80 7.36 0.00 0.30 0.00 0.07 0.06 0.47 0.02 98.62

Ol183Ms2 49.50 0.00 34.07 0.80 0.96 0.01 0.04 0.01 0.00 0.00 0.66 9.31 95.36

Ol183Chl2 25.95 0.02 23.15 22.02 16.24 0.00 0.40 0.00 0.00 0.10 0.05 0.02 87.94

Ol183Bt3 36.24 2.62 19.98 18.67 9.00 0.00 0.11 0.06 0.00 0.04 0.36 9.41 96.49

Ol183Pl3 64.39 0.00 22.94 0.00 0.00 4.37 0.01 0.05 0.00 0.00 9.58 0.32 101.67

Ol183And3 37.76 0.04 62.46 0.26 0.06 0.01 0.00 0.05 0.00 0.00 0.02 0.00 100.65

68

69

6 DISCUSSION AND CONCLUSIONS

Regional metamorphism of Olkiluoto culminated with a voluminous migmatization of

pelitic gneisses due to the dehydration melting of biotite, sillimanite, plagioclase and

quartz in the temperature exceeding 660 ºC and relatively low pressure of about 3.5 – 4

kbar. The temperature may have risen up to ~ 700 ºC producing granitic melt, which

later crystallized to leucosomes. The mineral assemblages produced during the peak of

regional metamorphism (e.g. cordierite) clearly overprint the earlier S2 foliation. In

addition, granitic leucosomes associated with the culmination of the metamorphism

were deformed by D3 deformation, so it is clear that the peak took place between D2 and

D3 deformation phases. D3 deformation took place during subsequent cooling which is

seen for example from the muscovite-bearing shear zones of some granite pegmatites

and granitic gneisses. Because TGG gneisses, having an age of about 1862 Ma, are

deformed by D2, the peak of regional metamorphism may be dated somewhere between

1.86 and 1.82 Ga, the latter being the minimum age of granite deformed by stage D3.

The culmination of metamorphism is seen also in the mineral assemblages of more

resistant rocks as biotite gneisses, hornblende gneisses and granodiorites.

The cooling phase took place also in low pressure as seen from some retrograde reaction

products as andalusite, chlorite and muscovite replacing cordierite.

Calculated pressures from some rocks are clearly higher than the average metamorphic

pressure in Olkiluoto. This might be caused by an earlier stage of metamorphism

connected with magmatic processes and emplacement of protolith of TGG-rocks.

However, also this stage may be classified as low-pressure type of regional

metamorphism and it took place soon after 1.86 Ga. Because there is a pressure

difference of circa two kbar between the two stages, there must have been an exquisite

erosion phase between the metamorphic phases, implying significant crustal uplift, or

alternatively a phase of tectonic thinning of the crust before the main stage of

migmatization. The migmatization, however, was produced by heating, not

decompression, as there is no evidence of previous or relic garnet in the migmatites.

As a summary, pressure and temperature conditions during the culmination of regional

metamorphism in Olkiluoto have been calculated and predicted from mineral

assemblages. The conditions were circa 660 – 700 ºC and 3.7 – 4.2 kbar. There was

probably an earlier metamorphic stage in the pressure of ~ six kbar (Figure 47). The

major phase of metamorphism took place between deformation phases D2 and D3, and

1860 – 1820 Ma, which are the ages of pre-metamorphic TGG-rocks and cooling stage

granitic leucosomes, respectively.

There is clearly need for further dating of migmatites to set time constraints for the peak

of metamorphism and further constrict the age of second deformation phase D2.

70

Figure 47. Metamorphic pressure and temperature values detected from the migmatites

and gneisses of Olkiluoto.

71

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74

75

APPENDIX

Origin of sample material and reference to previous studies and obtained results.

Abbreviations: pd = phase diagram, mc = chemical composition of minerals, TWQ =

result of TWEEGU thermobarometry, T = temp, P = pressure

Sample

Drillcore/

length

Petro-

graphy,

Geo-

chemistry

Phase

petrology

Thermo-

barometry

Additional data

PT results

Biotite gneisses

OL 103 OL-KR15

94,15 m

Fig 2

Gehör et al.

2007b

pd Fig 17 TWQ Fig 31 mc Table 2

T=647 oC; P=3,3 kbar

OL 142 OL-KR1

831,35 m

Gehör et al.

1996

pd Fig 18 TWQ Fig 32 mc Table 3

T=642 oC; P=4,3 kbar

OL 184 OL-KR8

497,06 m

Gehör et al.

2007a

pd Fig 19 TWQ Fig 33

and Fig 34

mc Table 4

T=613 oC; P=3,6 kbar

T=654 oC; P=3,9 kbar

(corrected for

retrogression)

OL 200 OL-KR19

428,45 m

Gehör et al.

2007c

TWQ Fig 35 mc Table 5

T=697 oC; P=4,2 kbar

OL 206 OL-KR19

18,90 m

Gehör et al.

2007c

TWQ Fig 36 mc Table 5

T=6 oC; P=6,7 kbar

Hornblende gneisses

OL 118 OL-KR18

52,79 m

pd Fig 20 TWQ Fig 37 mc Table 6

T=644 oC; P=5,4 kbar

OL 121 OL-KR18

98,73 m

Fig 3 TWQ Fig 38 mc Table 7

T=626 oC; P=3,3 kbar

Granite gneiss (from TGG series)

OL 108 OL-KR16

109,39 m

Fig 4 pd Fig 21 TWQ Fig 39

TWQ Fig 40

mc Table 8

T=670 oC; P=4,0 kbar

(Berman 2007 data set)

T=676 oC; P=4,2 kbar

(Berman 1988 data set)

Granodioritic gneisses (from TGG series?)

OL 145 OL-KR1

943,65 m

Fig 5

Gehör et al.

1996

pd Fig 22

garnet

Fig 23

TWQ Fig 41

TWQ Fig 42

mc Table 9

T=680 oC; P=5,5 kbar

(garnet core)

T=511 oC; P=2,3 kbar

(garnet rim)

OL 188 OL-KR15

387,98 m

Gehör et al.

2007b

TWQ Fig 43 mc Table 10

T=648 oC; P=3,2 kbar

76

Migmatitic gneisses

OL 133 OL-KR1

349,50 m

Gehör et al.

1996

pd Fig 24 TWQ Fig 44 mc Table 11

T=585 oC; P=3,2 kbar

OL 134 OL-KR1

368,13 m

Gehör et al.

1996

pd Fig 25,

Fig 29

mc Table 11

OL 135 OL-KR1

406,43 m

Gehör et al.

1996

pd Fig 30 TWQ Fig 45 mc Table 11

T=649 oC; P=3,8 kbar

OL 140 OL-KR1

738,90 m

Gehör et al.

1996

pd Fig 26

OL 183 OL-KR8

369,28 m

Fig 16

Gehör et al.

2007a

pd Fig 27,

Fig 28

TWQ Fig 46 mc Table 11

T=651 oC; P=4,4 kbar