low grade (greenschist facies) metamorphism in southern prins karls forland, svalbard

ALAN MORRIS:

LOW GRADE (GREENSCHIST FACIES) METAMORPHISM IN SOUTHERN PRINS KARLS FORLAND, SVALBARD

Abstract

Morns, Alan: Low grade (greenschist facies) metamorphism in southern Prim Karls Forland, Svalbard. Polar Research 2: 17-56.

The pre-Carboniferous miogeosynclinal sedimentary and volcanic sequences of southcentral Prins Karls Forland belong to a suite of rocks known as the Forland Complex. During early Palaeozoic diastrophism the Complex of Prins Karls Forland experienced predefomational, low grade metamorphism. A knowledge of the stratigraphy and structure of the region, together with an analysis of metamorphic mineral assemblages and reactions enable the conditions of metamorphism to be determined within certain limits. It is found that temperatures during metamorphism were between 4OO0C and 500°C, and pressures were about 6 kb. These conditions provide an estimate of palaeogeothermal gradient, consistent with observed mineralogies, of 18 to 2l0C.km-'. In addition, compositional variations within mineral phases are found to be largely dependent upon bulk-rock chemistry, and are of limited use in determining grade changes.

Alan Morris, Department of Geology, Wayne State University, Detroit, Mich- igan 48202, U.S.A.

I. Introduction

The area investigated in this study is south-central Prins Karls Forland (PKF), part of an island situated off the west coast of Spits- bergen, Svalbard (Fig. 1 a). Metamorphosed lithologies exposed in the study area are Precambiran to (?)Silurian in age and have been described elsewhere (Harland et al. 1979, Hjelle et al. 1979). Essentially they are miogeosynclinal in nature, consisting of well-laminated glaciogenic mix- tites, banded platform carbonates, shales (phyllites), well sorted arenites

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KE Y - rC?rtiory

Pre-Devc

Knivodden

uornn es

A/asdairhornet

Rossbukta

Gordon

Neukpiggen

Peferbukto

Hordio fieflet

Isochsen

!&J

Peochf/yo

Gei&ie

Ferrier

A

Fig. 1. - A. Location of Svalbard, Spitsbergen, and Prins Karls Forland. - B. Lithostratigraphic formations and geological map of south-central Prins

Karls Forland.

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and quartzarenites, with minor amounts of intermediate to mafic pyroclastics and high-level intrusives. These lithologies record a single, regional metamorphic event followed by three phases of deformation (Manby 1978a, Morris 1979). Ages of diastrophism are somewhat problematic, although it is generally accepted that the metamorphism and first deformation event are Caledonian (sensu lato) in age and the later events may be Tertiary (Atkinson 1960, Weiss 1953, Horsfield 1969, Hjelle et al. 1979). Fig. 1 b summarises the stratigraphy, structure, and areal distribution of lithologies of south-central PKF.

Previous work

Aspects of the metamorphic geology of PKF have been described, the principal contributors being Tyrrell (1924) and Atkinson ( 1 954) with some relevant remarks by Horsfield (1969). Tyrrell identified chloritoid ('ottrelite'), biotite, and muscovite in several localities but did not attempt to interpret his findings. Describing the rocks as generally 'chlorite or biotite zone', Atkinson (1956) also drew some conclusions about the origin of chloritoid. He described the chloritoid as a 'stress mineral', syn-defonnational in origin, and closely associated with thrusts. The fairly widespread occurrence of biotite was also noted by Hjelle et al. (1979). Finally, Horsfield in his study of the related rocks of Oscar I1 Land (OIIL) attributed the garnet, biotite, and hornblende assemblages t o temperatures of 3OOOC at the peak of metamorphism, and deduced a palaeogeothermal gradient of 3O0C.km-l.

The approach employed in this study is different from that of former works on PKF. It is consistent with more recent thinking on the development and significance of metamorphic rocks (Thompson 1955, Francis 1956, Albee 1965, Winkler 1976, Ohta 1979). Central to this method of study is the treatment of rocks as chemical systems and the comparisons made between observations in naturally occurring systems and their experimental analogues. Four compositional groups are recognised in the lithologies of south-central PKF:

*carbonates (>60% carbonate minerals) *pelites ( 1 5-50% quartz, >25% Mg-Al-Fe silicates, <60% carbonates) *psammites (>50% quartz) %nafics (<15% quartz, <lo% carbonates, >75% Mg-Al-Fe silicates).

These rock groups have been chosen and defined with due reference to general usage (eg. Wallis et al. 1968). Certain differences from other classi- fication schemes occur, for example: carbonate rocks are defined as above (and not 50% carbonate minerals) because rocks with psammitic characteristics frequently contain up to 50% carbonate material.

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Compositional data for the rocks of this study were obtained using standard x-ray diffraction goniometry and energy dispersive electron microprobe analysis. These facilities were available at the University of Cambridge. Bulk rock compositions have been estimated from modal mineralogy and probe analyses.

Timing of metamorphism

The earliest diastrophic event recorded in these rocks is a meta-

*quartz overgrowths and sl cleavage deflection around chloritoid laths;

:Vquartz overgrowths and sl cleavage deflection around pyrite crystals with chloritoid inclusions.

'%biotite porphyroblasts misaligned with the sl cleavage, and retrogressing t o chlorite that is concordant with the sl cleavage. Later events are essentially deformational: three fold phases are

recognised. The first of these is associated with the formation of the sl cleavage and post-dates the metamorphic peak. A second deformation event produced an inhomogeneously developed crenulation cleavage with no associated metamorphism, and a third event created some kinking and faulting (Morris 1979, 1981).

morphic peak. This is attested t o by abundant textural evidence:

These textural features are illustrated in Plate I.

11. Carbonate lithologies

Appendix 1 lists selected electron microprobe analyses of mineral phases occurring within the carbonate lithologies. Calcite, dolomite, and quartz are the dominant mineral phases, and muscovite, chlorite pyrite, and albite account for the minor constituents with an opaque, insoluble (in HCl and H2S04) 'dust' of carbonaceous material occurring in many of the rocks. In bulk terms, the chemical system of the carbonate lithologies can be characterised as:

CaO - MgO - Si02 - C 0 2 - H20.

Mineral chemistry and textures

CALCITE : The most widely distributed carbonate phase, calcite, is present in almost all the specimens collected. In purer carbonate lithologies, it occurs as a finely recrystallised matrix (grain size ca. 0.06mm).

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Variation in grain size is produced by the presence of impurities, notably carbonaceous material and phyllosilicates; these additional components usually increase the grain size (>0.6mm). This is because such impurities inhibit grain refinement by recrystallization through the promotion of pressure solution. The presence of impurities allows the strain to be taken up rapidly by pressure solution, and thus internal deformation and recrystallisation of individual grains do not occur. Recrystallisation in pure rocks seems to have continued throughout deformation. Analyses show the calcites to be very pure, with substitution of both Mg and Fe combined rarely exceeding 2 mol% (Appendix 1).

DOLOMITE: Less common than calcite, dolomite occurs in texturally very similar forms. Compositionally, the mineral is also very pure with only small amounts of Fe substitution. However, in four localities ankerite occurs.

TREMOLITE: Tremolite and tremolite-actinolite occur as the result of thermal metamorphism at two localities. They form lath-like crystals of between 0.05mm and lOmm in length, and they were formed prior t o regional metamorphism.

Fe-bearing epidote occurs in a thermally metamorphosed limestone/carbonate segregation within one of the eruptive horizons of the Alasdairhornet formation.

Phyllosilicates are fairly common minor phases in these rocks. Although usually associated with quartz and other detrital material, they also occur in relatively pure carbonate lithologies. Muscovite and chlorite are the only layered silicates identified. In some cases pressure solution has produced concentrations of these minerals and other insoluble materials into pressure solution, tectonic cleavage planes (sl). All muscovites are phengitic and usually contain small amounts of Na, Mg, Ca, and Ti (<3 mol% in total). Chlorite is present in smaller amounts than muscovite.

EPIDOTE :

CHLORITE & MUSCOVITE :

Mineral assemblages and inferred reactions

The following are the mineral assemblages in the carbonate rocks of this study:

cc + qz 5 mv f ab . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . (A) cc + do + qz k mv . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . (B)

cc + qz f cl . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . (D) cc + qz + t r * do . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . (E) cc + qz + cl + ep f t r . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . (F)

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . do + qz k mv (0

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4

P kb

2

0

600

400

roc 200

I

I

I I

I I I

lkb

200 0 600 0 0.5 LO

r *c P f C 3 l

The cu rves were c a l c u l a t e d f o r t h e r e a c t i o n :

21qz + 3 c l + l O c c S 2 z o + 3 t r + loco2 + 8H20 ..... ( 3 )

us ing thermodynamic d a t a g iven by Robie 6 Waldbaum (1968) and Zen (1972) , and app ly ing t h e fo l lowing r e l a t i o n s h i p s :

- dP = - A S T = - AH y-1 i aT = + RT 7. - i) dT bV e AS apc02 A S P H20 ’C02

f o r r e a c t i o n ( 3 ) : 2 = 0.44 Y = 0.56 Ca lcu la t ed formulas f o r phases i n assemblage (F) u s ing probe a n a l y s e s w e r e :

1 .89 Fe0.48 *l2.54 si3.08 ‘12 (OH)

1 .93 Mg2.71 Fe2.07 *l0.28 si7.88 ‘22 ( O H ) 2

4.68 Fe 4.46 A15.06

ep = Ca

am = Ca

si5.61 ‘20 (OH)16 cl = Mg

On t h e b a s i s of t h e s e a n a l y s e s t h e fo l lowing s t o i c h i o m e t r i c r e a c t i o n is i n f e r r e d :

38qz + 19cc + 6 c l h 5 e p + 5 t r + lOC02 + 17H20

Because of t h e p re sence of i r o n i n t h e s e phases , s t a n d a r d thermodynamic d a t a a r e n o t a p p l i c a b l e . However, t h e r e a c t i o n is of t h e same form as (3) and shou ld have s imi la r c h a r a c t e r i s t i c s . P(CO2) t aken d i r e c t l y from t h i s r e a c t i o n i s c o n s i s t e n t w i t h t h e r ange g iven by Harte 6 Graham (19751, a l s o see t h e dashed l i n e on the curve .

P - p r e s s u r e T - e q u i l i b r i u m t empera tu re 4 - u n i v e r s a l gas c o n s t a n t

S - en t ropy H - en tha lpy

Fig. 2. Reaction (3), its significance for temperature and pressure conditions, and pro- bable stoichiometric form.

On the basis of these assemblages, and the textural relationships of their constituent minerals, certain reactions have been identified : 5do + 8qz + H20<--->tr + 3cc + 7 C 0 2 . . . . . . . . . . . . . . . . . . (1) 2do + ta + 4qz<--->tr + 4C02 . . . . . . . . . . . . . . . . . . . . . . . . . . (2)

(3) lOcc + 21qz + 3cl<--->2zo + 3tr + 8H20 + 1 0 C 0 2 . . . . . . .

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These reactions are prograde left to right (Metz and Trommsdorf 1968, Slaughter et al. 1975, Winkler 1976).

The coexistence of calcite, dolomite and quartz suggests that the reaction : 3do + 4qz + H20<--->ta + 3cc + 3C02 . . . . . . . . . . . . . . . , . . (4) (Metz and Trommsdorf 1968, Slaughter et al. 1975), did not take place, either because temperatures were not sufficiently high, or because P(C02) was too high. There are some indications, however, that quartz grains in a dolomitic matrix have suffered limited reaction with the carbonate: corroded quartz grains occur, although no products from the reaction can be detected. This suggests that temperatures may have approached those necessary for reaction (4) but that because of the dominance of carbonate over quartz, P(C02) values were too high to allow the reaction to proceed.

Muscovite is the prograde expression of the sheet silicate mineral impurities that were present in the original rock. It has not formed at the expense of any other phase present, and seems to have acted merely as a sink for excess elements A1 and K.

The assemblage (E) is only found in the thermal metamorphic aureole produced by a pre-regional-metamorphism sill: it indicates that reaction (1) has operated. Continued presence in the rock of quartz and dolomite might indicate that conditions were approximately at equilibrium, or that insufficient water was present for the reaction to reach completion.

Assemblage (F) can be attributed to reaction (3). The rock in which this assemblage occurs is a small carbonate xenolith, or segregation, within a basic volcanic flow. Thermal metamorphism during the extrusion of the flow resulted in the formation of epidote, with some tremolite, around these carbonate rocks. Subsequent partial retrogression has reversed this reaction producing assemblage (F).

Finally, assemblage (D) occurs in one rock in which the chlorite and quartz pseudomorph large (5mm long) bladed crystals of either tremolite or epidote are present. Only calcite, quartz, and chlorite are now detectable in the rock. Either reaction (2) or (3) has occurred. (2) is unlikely since no traces of either dolomite or talc are present, and a chlorite- producing reaction from these is very unlikely to occur. Reaction (3) having occurred from right to left (retrogression) is the most likely ex- planation.

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Implications of inferred reactions

Experimental and field data for reaction (4) are available from Metz and Trommsdorf (1968) and Slaughter et al. (1975); their data shows that if this reaction were on the point of occurring, as suggested above, then temperatures could have ranged from 4OO0C to 5OO0C at 1 kb pressure, and from 45OoC to 6OO0C at 5 kb pressure. The domi- nantly carbonate composition of the rocks in which reaction (4) may occur indicates high P(C02), say greater than 0.5. 5 km of low grade metamorphic rocks now overlie (stratigraphically) these carbonate horizons, therefore a minimum confining pressure of 1.5kb must have been attained during metamorphism. This gives temperatures of 450 - 6OO0C on the basis of Metz and Trommsdorf’s work. Slaughter et al. do not map reaction (4) above P(C02) = 0.5, but assemblages (A) to (D) must lie on the low temperature side of reaction (2); therefore temperature estimates from this scheme will be 50-1OO0C lower than those stated above. It should also be noted that reaction (4) taking place in a semi- closed system is self-inhibiting. When the fluid phase is not present to excess the production of C02 from the reaction will increase P(C02) and push the system along the line of reaction (4) with increasing T. This buffering of the fluid phase by the mineral reactions results in little mineralogical change until tremolite appears (Greenwood 1975).

The preservation of thermal metamorphic tremolite crystals in assemblage (6) is evidence that later regional metamorphism must have satisfied the following conditions: 1. Temperatures were not sufficient to prograde the assemblage to one

bearing diopside (Winkler 1976); 2. Temperatures below those necessary for reaction ( 1 ) to be in equi-

librium did not prevail for long enough for the assemblage to retrograde to a quartz and dolomite bearing assemblage. Thus, temperature conditions must have been fairly close to those

for the invariant point at the intersection of the lines of reactions ( I ) , (2), and (4) since reaction (1) has a limited range in T-P(CO2) space. Un- fortunately, the invariant point has a PT range of 0.8kb/490°C to 6.2 kb/610°C (Metz and Trommsdorf 1968) or 0.8 kb/410°C to 5 kb/460°C (Slaughter et al. 1976). However, this assemblage does provide information about the temperature of metamorphism during the initial intrusion of the sill. Mudge (1968) considers that a depth of 2.3 km may be the lower depth limit for small concordant igneous bodies. Therefore a maximum confining pressure of 0.8kb during intrusion is suggested; this gives thermal metamorphic temperatures of 490 - 5OO0C (Metz and Trommsdorf) or 410 - 45OoC (Slaughter et al.). These temperatures are consistent with the intrusion of a small body of basic magma into country rock at an ambient temperature of ca 50°C.

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Assemblages (D) and (F) are indicative of reaction (3). (F) is found in a thermal metamorphic carbonate rock which is retrogressing through reaction (3). This reaction has been noted by Greenwood (1967), Harte and Graham (1975), Watts (19731, and Barron and Barron (1976). However, none of these sources provides precise data for the reaction since it is fairly restricted in extent by P(C02) conditions. Fig. 3 shows the results of elementary thermodynamic calculations based on data given by Robie and Waldbaum (1968) and Zen (1972). The composition of the mineral phases in assemblage (F) suggests a P(C02) of about 0.5 which corresponds to temperatures during retrogression of between 3OO0C and 55OoC depending upon confining pressure (Fig. 2). Textural evidence from these rocks suggests that the prograde assemblage tr + ep + cc was pre-deformational and formed in a low differential stress environment. it is likely that these occurrences are limestone zenoliths in a basic flow. Initial temperatures of metamorphism would probably have been between 2OO0C and 300OC. The natural system is complicated by the presence of Fe introduced by contamination from the lava. Assemblage (D) is interpreted as being the result of reaction (3) retrogressing to completion. This rock is almost completely Fe-free thus indicating temperatures of between 3OO0C and 55OoC for retrogression. The metamorphic history of the rock seems t o include a thermal event sufficient to operate reaction (3) in the prograde sense, followed by a prolonged cooling during which (3) was reversed. This is consistent with the structural position of the rock: under the lower limb of a large overturned nappe, and hence subject to a thermal 'blanket' effect.

Summary of carbonate data

The very widespread occurrence of carbonate- and quartz-bearing assemblages suggests low grade metamorphism occurring below any talc- o r tremolite-forming reactions. Preservation of pre-regional-metamorphism contact aureole assemblages does not contradict this conclusion. A single occurrence of a retrograde reaction involving tremolite and epidote indicates that in certain zones temperatures were higher and cooling may have been prolonged. The lack of knowledge of the confining pressure during regional metamorphism limits the interpretation of the carbonate data.

111. Pelites

Appendix 1 lists selected mineral analyses from pelitic horizons. The term pelite is used in a strictly compositional sense, although these rocks are generally finer grained than the psammites discussed later.

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6C

SO

W

4c

30 30 40 50 60 70

mV

C/

d

40 i 0

30 40 so 60 n, mV

o carbonate pelire

0 ahminous pelile

Fig. 3 . Cation distribution coefficient plots of Mg/(Mg t Fe)% for coexisting mineral pairs in pelitic rocks. Cation ratios are derived from micro-probe analyses.

Since the approach in this present work is that of defining chemical systems, the compositional groups have to a large extent been chosen from a consideration of common assemblages. This method leads to a sub-division of pelitic rocks as follows :

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1. Aluminous pelites 2. Carbonate-bearing pelites. In terms of Fe, Mg content these two sub-divisions do not run

absolutely parallel: the carbonate group tends to include more mafic rocks than the aluminous group. The two subdivisions represent different chemical systems, these are:

1 A1203 - Si02 - K20 - MgO - FeO - H 2 0 2 A1203 - Si02 - K 2 0 - CaO - MgO - FeO - H 2 0 - C02

Studies on similar systems have been made by Brown (1 975), Thompson (1976 a & b), Barron (19761, and Leitch (1976). The carbonate group rocks are broadly equivalent to those described as mark by Winkler ( 1 976).

Mineral chemistry and texture

Phyllosilicates are the dominant non-silica minerals in these pelites with varying amounts of carbonates and feldspar (albite), epidote and tourmaline occur in accessory amounts. Chloritoid is common in those rocks with a high Al, and low Fe:Mg content.

Throughout the area muscovites have a phengitic composition, they constitute the most widely distributed phyllosilicate phase, and show recrystallisation after the peak metamorphism, and during deformation (Plate 1). Variation in the composition of muscovite is chiefly the result of bulk rock composition, for example, muscovites associated with chloritoid have low phengite content (Si4+ = 3.05; Velde 1967) although grains immediately adjacent to chloritoid laths have Si4+ = 3.2 1. The distribution coefficient plots of Fig. 3 show that equilibrium was at least approximated between coexisting muscovites and biotites, a less well-defined plot between the muscovites and chlorites indicates some disequilibrium. The porphyroblast size of muscovite increases with stratigraphic depth from about 0.01 mm t o 0.05 mm or larger. There are essentially two textural modes of occurrence of muscovites, these are:

MUSCOVITE:

1 Relict pre-deformational banding - often bedding, 2 Pressure solution planes.

In the former, especially in the older rocks, it is often intergrown with biotite, showing a coarse crystalline form, and frequently kinked. Musco- vite occupying pressure solution planes shows evidence not only of reorientation, but also of recrystallisation, and has a finer crystalline texture than the pre-deformational muscovite. There is no detectable consistent chemical variation between the two generations of muscovite.

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CHLORITE: Second in abundance t o muscovite in the phyllosilicate component is chlorite. Almost identical to muscovite in its textural occurrence, chlorite is slightly more restricted with respect t o rock composition. It is not a major constituent of the high-A1 pelites, and its presence increases with increasing Mg and Fe, particularly the former. Chlorite coexisting with chloritoid has an Fe2+ content of about 15 mol%: chlorite with Fe2+ of 17.5 mol% usually shows some reaction with the chloritoid while Fe2+ contents greater than this are found in most other chlorite-bearing rocks. Compositional variation is closely linked with bulk rock chemistry, there is a close correlation between the Fe content of the chlorites and the overall Fe content of the rock. Distribution coefficients for chlorites are shown in Fig. 3, as with muscovite, equilibrium with coexisting biotite seems t o have been reached.

Biotite occurs in many of the pelitic rocks in the area. Its highest structural-stratigraphic occurrence is in the mafic pelites of the Alasdairhornet formation. Whenever it occurs it is pre-deformational, strongly crystalline and associated with muscovite or chlorite, or both. This association is due to the syn-deformational recrystallisation and retrogression of biotite to chlorite and/or muscovite, and this in turn accounts for the good equilibrium between biotite and muscovite, and biotite and chlorite (Fig. 3). Bulk rock composition is also a controlling factor in the chemistry of biotite and the Mg, Fe contents are similar to those of other phyllosilicates in any one rock. In composition, the biotite is generally annitic. There is a noticeable increase in both the coarseness of crystal growth and in the proportions of biotite from the younger t o older rocks of approximately the same composition.

CHLORITOID: Chloritoid occurs in a number of highly aluminous phyllites as pre-deformational laths and rosettes up to 2 or 3 mm in length. Density of growth and porphyroblast size is related to bulk rock chemistry and grain size. Coarse grain size and relative paucity of alumina result in large, widely distributed grains, coarse grain size and abundance of alumina produce rosette growths, and in fine grained rocks the laths are small and well disseminated. This variation is due to the interrelation- ship between the number of nucleation sites and the amount of chloritoid-forming material present: the larger the grain size, the fewer the number of nucleation sites, therefore, the porphyroblasts are more widely spaced in coarser grained rocks than in finer rocks. In addition, in coarser rocks, porphyroblast size is limited by amount of material available, thus variable sizes, and in Al-rich rocks rosettes form, whereas in finer lithologies porphyroblast size is limited by the closer spacing of nucleation sites. Muscovite and quartz are the principal accompanying minerals, and pyrite may also be associated as a late (post-chloritoid) metamorphic growth. Chlorite does coexist with chloritoid but not in large amounts,

BIOTITE:

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and in some instances chloritoid is being chloritised. Chloritoid occurrence is restricted by bulk rock composition, and the limited areal extent of its presence renders it unsuitable as a useful grade indicator.

Occurring as small grains, albite is a fairly common minor constituent. It appears as small (0.05 - 0.1 5 mm) grains, rounded in outline, untwinned, and often with calcite inclusions. Size of grains and frequency of occurrence increase with increasing stratigraphic depth.

Epidote, tourmaline, and apatite are all quite common accessory minerals; tourmaline and apatite are restricted to the Ferrier Group of metasediments, while epidote occurs throughout. Epidote is not common in the carbonate group, probably because the increased P(C02) inhibited its growth, excess Ca forming carbonate minerals.

CARBONATE MINERALS: Usually occupying interstitial, ce- menting positions, calcite, dolomite, and ankerite are found throughout the pelitic rocks. Although carbonate minerals do occur in aluminous lithologies they are best developed in more mafic systems.

ALBITE:

OTHERS:


A number of typical assemblages and possible reactions can be inferred from the various sets of data. These are:

Aluminous assemblages qz + mv + cd f op . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . G ) qz + mv + cd + cl f o p = (H) qz + mv + cl f ab f ep f op .(I) qz + mv + bi f cl f ep f t o f op (J)

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . . . . . . . . . . . . . . . . .

Carbonate assemblages qz + mv + cl f cc f do f ab . . . . . . . . . . . . . . . . . . . . . . . (K) qz + mv + cl + bi * cc f do/ak f a b . . . . . . . . . . . . . . . . (L) qz + mv + bi f cl k cc k ab . . . . . . . . . . . . . . . . . . . . . . . ( M )

These assemblages when found to be in equilibrium are interpreted as being steps of the following continuous reactions:

Al-mica + F e d <- --> cd + qz + H20. ( 5 ) (Seidel & Okrusch 1975)

K-mv + c l < - - -> mv + bi + H 2 0 . . (6) (Brown 1975)

mv + cl + do<-- ->bi + cc + qz + H 2 0 (7)

. . . . . . . . . . . . . . .

. . . . . . . . . . . . . . . . .

. . . . . . . . . . .

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Aluminous pelites

Two sets of reactions are of interest in the aluminous pelites, the first concerns the occurrence and stability of chloritoid, and the second, the initial production of biotite during prograde metamorphism. Reaction (5) can be seen occurring from right t o left in a number of horizons. It is suggested that a similar reaction is responsible for the original formation of chloritoid within the high alumina lithologies. Chloritoid does not form in rocks with high (Mg + Fe)/Al ratios, in addition, it seems from textural evidence and bulk mineralogies that the stability field of chloritoid can be reduced by increasing (Mg + Fe)/Al. Hoschek ( 1 969) noted that elevated P( H20) could raise the equilibrium temperature for reaction ( 5 ) . The occurrence of chloritoid retrograding to chlorite in certain more mafic lithologies than those adjacent in which chloritoid remains stable can be explained if (Mg + Fe)/A1 ratios have a similar effect to P(H20) (Plate I) . La Tour et al. (1980) suggest that the Mg/(Mg t Fe) ratio is more important than (Mg + Fe)/Al in determining chloritoid stability. However, in this case the only significant difference between lithologies containing stable chloritoid and retrogressing chloritoid is in the (Mg + Fe)/Al ratio.

Reaction (6) is proposed on the basis that the appearance of biotite in the aluminous pelites is associated with a reduction in the amount of chlorite, and chlorite may be absent from biotite-bearing rocks.

A number of assemblages showing a sequence from ( I ) to ( K ) indicates that the muscovites become slightly depleted in K , as do the chlorites which also disappear entirely (Fig. 4).

Carbonate-bearing pelites

In the carbonate-bearing pelites, in addition t o reaction (6) another biotite-producing reaction is thought to operate, reaction (7). The occurrence of (7) is suspected because of the antipathetic association between dolomite and biotite. Rocks in which biotite is well formed contain little or no dolomite and the carbonate phases are usually calcite or ankerite. In many rocks, coarse grained biotite is associated with Fe-calcite (Fig. 4). Possible alternative reactions t o (7) are:

* F e d o + qz + ph + hm <---> ep + bi + mc + C02 (Thompson 1973). However, none of these rocks contains microcline and there is no evidence for any reactions involving hematite and epidote.

This seems unlikely as biotite has formed at the expense of chlorite.

* do + mv < ---> bi + cl + cc + C 0 2 (Bethune 1976).

Rea

ctio

n

(6)

Rea

ctio

n

(7)

B

El

Mg

A1

Si

K Fe

Mg

A1

Si

Fe

Mi3

A1

S

i

K Fe

Mg

Ca

Fe

Mg

Ca

Fe

Mg

Ca

Fe

~g~

A1

35.3

0 :.2.1 33.36 2.78 S

i 45

.75

48.1

8 K

13.1

5 12

.64

Fe

3.27

2.

65

Mg

23.2

7 21

.85

A1

24.9

0 23

.76

Si

30.0

8 28

.47

'* 5.36

32

.21

47.6

8 11

.70

2.59

27

.31

25.7

0 29

.88

11.9

3

37.1

8 15

.28

3.01

0.

59

96.6

7 0.

68

A1

I I

9.38

-

14.6

5 21

.97

38.4

0 11

.57

29.0

4

2.18

93

.78

2.35

36.3

0 11

.25

Fe

16.4

1

Min

- m

iner

al

El

- el

emen

t

34.0

6 46

.90

12.2

5

20.0

2 36

.14

16.4

2

Min

i M

....

spec

imen

num

ber

Fig

ure

s a

re i

n o

xid

e w

eig

ht

%.

mv

- m

usc

ov

ite

bi -

bio

tite

* -

M27

84,

M29

65,

M29

70

cl -

cc -

Fig.

4 c

hlo

rite

d

o -

do

lom

ite

ca

lcit

e

ak -

an

ke

rite

11.7

2121

.301

25..2

6 20

.161

18.1

7115

.00

10.9

1 16

.50

19.0

8 38

.79

50.4

2

Two

biot

ite-f

orm

ing

reac

tions

in p

eliti

c ro

cks,

(6) a

nd (7

), an

d re

late

d ch

ange

s in

min

eral

ass

embl

ages

and

com

posi

tions

. B.

Qua

litat

ive g

rade

cha

nges

from

pel

itic

min

eral

ass

embl

ages

and

min

eral

com

posi

tions

.

-3 2-


Reaction (5 ) is the only cited reaction which approaches discontinuity, and for which experimental data are available (Hoschek 1969, Seidel 1975). The occurrence of this reaction indicates temperatures of approximately 42OoC.

Continuous reactions such as (6) and ( 7 ) are very difficult to pin- point in terms of pressure and temperature conditions at equilibrium. A relative grade indication can be obtained by comparing the proportions of reacting phases present.

The predeformational nature of biotite porphyroblasts (Plate I) indicates that the metamorphic peak was reached prior to deformation. Retrogression of the biotite took place in many lithologies during deformation, producing chlorite. Quite large amounts of chlorite occur in some rocks with small relict inclusions of biotite and muscovite. This retrograde reaction seems unlikely to be a reversal of either reaction (6) or (7), since it seems to result only in the production of chlorite. It may be that this provides the reason why the distribution coefficients for coexisting mv + bi and cl + bi give evidence of equilibrium, while those for mv + cl show more scatter (Fig. 3): muscovite coexisting with biotite is prograde, but chlorite coexisting with muscovite and/or biotite is retrograde, forming from the hydration of biotite, not muscovite.

This also helps to explain the poor agreement of chlorite analyses with the reactions inferred from assemblage variations (Fig. 4). Although the reactions which have been deduced are continuous and difficult to use as precise grade indicators, certain trends are visible (Fig. 4). It is noticeable that the frequency and abundance of occurrence of biotite increases with stratigraphic depth, and that bulk rock composition is an important mineral-phase controlling factor.

Summary of pelitic data

The rocks as described exhibit 'greenschist facies' mineral assemblages, or in the terminology employed here, low grade metamorphic mineral assemblages (Winkler 1976). Prograde mineral reactions within these rocks, as deduced from variations in mineral assemblages, are generally continuous and do not provide useful grade indicators. Very little temperature and pressure information about conditions of metamorphism can be gained from this compositional rock group.

-33-

IV. Psammites

Mineral assemblage data show that the major difference between psammites and pelites is the overwhelming presence of quartz in the former. This being so, there are, nevertheless, certain variations in the metamorphic mineralogy of the psammites that require separate dis- cussion.


Quartz is the major mineral phase present, phyllosilicates and carbonates occur in varying amounts and small percentages of feldspar (usualIy albite), epidote, tourmaline, pyrite, and ilmenite can also be found. Mineral composition data are given in Appendix 1.

MUSCOVITE: Without exception the muscovites are phengitic. Fe content vanes from about 0.3 mol% to 3.4 mol%. There are, as in the pelites, two textural modes of muscovite occurrence :

1. pre-deformational banding - possibly bedding, 2. pressuresolution cleavage planes.

Where muscovite appears in the purer, quartzitic lithologies, it almost invariably occupies type 2 sites, and is usually associated with chlorite.

The modes of occurrence of chlorite are similar to those of muscovite. In addition, it frequently exhibits syn-deformational growth as a product of retrograde breakdown of biotite.

BIOTITE : Invariably pre-deformational, biotite occurs as laths and variably-sized porphyroblasts frequently associated with chlorite and ankerite. Biotite composition is generally annitic.

CALCITE: Calcite occupies interstitial sites and is very common throughout the psammites. Fe-bearing calcites (Fe = 8 mol%) are associated with biotites.

DOLOMITE: Similar in its textural features to calcite, dolomite is very variable in composition. Within a single lithology at one location it may vary from a standard Mg-Fe dolomite (Mg: Fe > 4) to the Fe-rich ankerite (Mg: Fe < 4, sometimes as low as 0.05). Ankeritic dolomites tend to be associated with biotite.

OTHERS: Albite is a common minor constituent, as in pelitic rocks grains are generally rounded, untwinned, and with very small calcitic inclusions. Such grains may represent metamorphically equilibrated detrital plagioclase grains. At one locality plagioclase grains with anort hite content of an (27) occur. Epidote is also quite common in less carbonate- rich lithologies. Tourmaline, apatite, and pyrite complete the minor mineral phases represented.

CHLORITE:

-34-


The following equilibrium assemblages have been distinguished in the psammites:

qz f mv f cl f ep f ab f op . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ( N )

qz + mv + cl + bi k ep ( Q ) (J) q z + m v f c c k d o f a b f e p (PI qz + mv + cl + cc f do f ab . . . . . . . . . . . . . . . . . . . . . . . . . (3 ( K ) qz + mv + cl + bi + cc f do/ak f ab . . . . . . . . . . . . . . . . . . (T) (L) qz + mv + cl + bi + cc f do/ak f an . . . . . . . . . . . . . . . . . . . . . (U) qz + mv + bi + cc f ab . ( V ) There are several obvious parallels between these assemblages and

those found in the pelites. Differences are best explained by the more variable proportions of minor components in psammites. In addition, because of the banded nature of some of these rocks, potential reacting phases may not be in mutual contact and are therefore unable to react. Similarly the overwhelming presence of quartz acting as a physical, dispersive barrier may also serve to inhibit reactions. In general, the assemblages observed define essentially the same reactions as for the pelites. There are, however, two exceptions: the definition of psammites employed here precludes high alumina content, therefore no chloritoid forms, and low Fe-epidotes occurring in the psammites appear to react with the carbonate phase. These are the proposed reactions:

K-mv + cl <-- ->mv + bi + H 2 0 . . . . . . . . . . . . . . . . . . (6)(Pelites) mv + cl + do<--->bi + cc + qz + H 2 0 . . . . . . . . . . . . (7)(pelites) zo + C02<- -->an + cc + H 2 0 . . . . . . .(8)(Storre & Nitsch 1972). Biotite-producing reactions in pelitic rocks have been discussed

already and to a large extent the same arguments apply here. There are, however, other alternative reactions which have been proposed, and which may be relevant to psammites:

q z + m v f e p f a b f o p . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . (0) qz + mv + cl f ep f ab f o p . . . . . . . . . . . . . . . . . . . . . . . . . .(P) ( I )

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

a) do + mv <- - - >bi + cl + cc + C 0 2 (Bethune 1976) This reaction is a possible substitute for reaction (7) and may well

these rocks, there is , however, some evidence to suggest operate in that biotite formed at the expense of chlorite. b) cc + mv + cl<- - - >bi + an + C02 (Bethune 1976)

Anorthite only occurs in one known locality and this is best explained by reaction (8). A similar reaction t o a) involving epidote production is also proposed by Bkthune, however, there is no evidence that biotite and epidote are closely associated. Reaction (8) as proposed by Storre & Nitsch (1972) is as follows:

-35-

220 + C02<- - - >cc + 3an + H 2 0 . . . . . . . . . . . . . . . . . . . . . . (8) It is possible that anorthitic feldspar is more widespread in occur-

rence than is indicated by the assemblage lists. Neither epidote nor anorthite is a major constituent of psammites and those grains which have been identified are of the order of 1 t o 10 micrometres in diameter only. It seems likely that reaction (8) can be used to explain the relative paucity of epidote in the carbonate-bearing psammites. Storre & Nitsch (1 972 j concluded that elevated P(C02) would restrict the field of epidote stability, allowing a considerable increase in the anorthite content of plagioclase if reaction (8) proceeds from left to right. An interesting conse- quence of this line of reasoning is that since reaction (7) is a water- producing reaction and elevated P(H20) should increase the epidote stability field, then in order to invoke reaction (8), it may be that ( 7 ) is not valid, and that reaction a) (Bethune 1976) explains the situation best. However, this would mean that two anorthite-producing reactions were operating and that anorthite should be more abundant than it is. Using the relatively high proportions of carbonate minerals in the carbonate-bearing psammites as an indication of high P(C02) it is suggested that water produced by (7) is not sufficient t o inhibit reaction (8).

The arguments, although somewhat tortuous, lead t o the conclusion that reaction (6) is the dominant biotite-producing reaction in the non-carbonate-bearing psammites, and reaction (7 j is responsible for biotite in the carbonate-bearing psammites. Reaction (8) is proposed in order to explain the limited occurrence of anorthitic feldspar and the reduced abundance of epidote in carbonate-bearing psammites and psammites in general.

Implications of inferred reuctions

As with the pelitic rocks the reactions proposed here are continuous and ill-defined with respect to P and T conditions during metamorphism. Experimental data is available for reaction (8) but this gives little or no indication as t o possible P and T conditions.

Summary of psummitic data

Justification for separating psammites from pelites is based largely on lithological grounds, since the chemical system is essentially the same for both groups. However, certain reactions are different and similar reactions are less well-defined in the psammites. conditions of metamorphism, as inferred from the psammitic lithologies, were sufficient to produce biotite, implying a minimum temperature of 3OO0C (Winkler 1976).

-36-

V. Mafics

Metavolcanic flows, tuffs, mixed sediments and pyroclastics, and minor intrusives are all included in the term 'mafic' as employed in this study. There are no remnants of primary igneous mineralogy and the matrix textures are generally metamorphic. Occurrences of calcite as deformed blebs may represent original vesicles with calcitic infill. Pre- deformational, recrystallised apatites may be of metasomatic origin.


The most common phases in the mafic rocks are phyllosilicates and amphiboles, albitic feldspars and K-feldspars are also quite common. Epi- dote, sphene, and ilmenite are almost universal as minor constituents. Carbonate minerals, although fairly common, usually occupy interstitial sites or, in some cases, form cavity infillings in amygdales. Mineral analyses are given in Appendix l .

MUSCOVITE: Muscovite occurs in some of the mafics. In the mixed sediment/pyroclastic lithologies it has a phengitic composition and Mg: Fe ratios of 0.5, similar to muscovites found in quartz-rich pelites and psammites. Paragonitic muscovite occurs at one locality in a basic intrusive lithology, in this case the muscovite is probably due to contamination of the igneous material by the country rock.

BIOTITE : Compositions of biotite are usually annitic with between 1 and 2 mol% Ti. As with the psammites and pelites, occasional high Mg, phlogopitic biotites occur. Unlike the occurrences in the other rock groups, however, the mafic biotite tends to be a retrograde product, and replaces amphibole in many cases. There is textural evidence t o suggest that biotite formation took place continuously during deformation (Plate I). Biotite and chlorite are very closely associated, sometimes apparently forming parts of the same grain. The distribution coefficients for Fe and Mg in coexisting biotite and chlorite show excellent agreement with equilibrium (Fig. 5 ) .

CHLORITE: In general, chlorites have the same textural modes of occurrence as the biotites. Their compositions are variable and this seems to be a reflection of the bulk rock composition. Optically, these chlorites are very green and form quite large crystalline masses. Equilibrium with biotite is good since both are retrograde products.

Tremolite is the most common amphibole, with two exceptional occurrences of actinolite. These amphiboles are rarely found in equilibrium, they are usually altering to biotite, chlorite and epidote.

AMPHIBOLE:

-37-

EPIDOTE : The epidotes are generally of clinozoisite composition with A12.Fe% between 20 and 35. This contrasts with epidotes found in the calcareous xenoliths (see section 11) which have A12.Fe% of about 15. As with chlorite and biotite, epidote seems to have formed throughout defomration as a retrograde production of tremolite.

ILMENITE and SPHENE : The occurrence of these two Ti-bearing phases is related to a pro-grade reaction involving the breakdown of biotite and calcite. Both are pro-grade products and they tend to occur together. Presence or absence of sphene seems to depend upon P(C02), or the abuncance of calcite: if calcite is a fairly common constituent ( > I 0%) then sphene does not usually form.

FELDSPAR: All the feldspars analysed using the microprobe are albitic, however, some x-ray diffraction peaks indicate the presence of K- feldspars in small amounts. Most feldspar grains have the appearance of being retrograde minerals.

OTHERS: Calcite and dolomite occur as interstitial minerals in some rocks and as inflills of relict vesicles. Apatite, presumably of igneous metasomatic origin, is found at one locality.

(see Fig. 6, next page): -+

The graphs are drawn from the reactions published in Metz & Trommsdorff (1968), Carmichael (1970), Storre & Nitsch (1973), and Winkler (1976). Thermodynamic data were taken from Robie & Waldbaum (1968) and Barron & Barron (1976). Experi- mental data for the invariant points I, 11, 111, and IV has been extrapolated to 7 kb where :

using graphs published by Winkler (1976). From these points the equilibrium curves were estimated by comparing with experimental data, and by graphical analysis using the relationship:

7 kb is the maximum pressure estimate (section VII) and the temperature ranges at this pressure are regarded as the maximum possible. Ti-poor systems may involve analogous reactions but without the formation of ilmenite and sphene. nS - change in entropy, Te - equilibrium temperature, R - universal gas constant P - ambient pressure.

----t

-3 8-

Ix’

70 .

50

I 30 50 70

C l

Fig. 5 . Cation distribution coefficient plots of Mg/ (Mg + Fe)% for coexisting biotite and chlorite in mafic rocks. Cation ratios are derived from microprobe analyses.

r %

- 2to+lCO2 LI Icc + 3an + Iffn 0 300 -

d

0 0.5 I. 0

500 -

0 0. o/ 0.02 0.03 0.04

Reaction sequence (9) t o (12) . Fig. 6 . - (see text p. 37)

-39-


The following are the identified, simplified assemblages : qz + cl + bi rf: mv _+ cc f do rf: fp . . . . . . . . . . . . . . . . . . . . . . . . . (W) qz + cl + bi + ep + am + il k sp k cc k mv . . . . . . . . . . . . . . . . (X) qz + c l + cc + d o + f p + a m + i l + e p . . . . . . . . . . . . . . . . . . . . ( Y ) c l + a m + f p f b i . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . (Z)

Assemblage (W) is representative of the pelitic element in many of the pyroclastic deposits, such lithologies usually contain lithic fragments with a more mafic assemblage, commonly (Z).

The textural evidence suggests that assemblages (X) to (Z) inclusive can be explained by a complex series of prograde/retrograde reactions. These reactions are principally controlled by the P(C02) : P(H20) ratio. They are: prograde :

. . . . . . . . qz + c l + bi +cc<--->tr + i l +K-fp + H 2 0 + C 0 2 (9) q z + c l + b i + c c < - - - > t r + s p + K - f p + H 2 0 + C 0 2 . . . . . . (10)

retrograde : tr + K-fp + H 2 0 + C02<--->cl t bi + cc + ab . . . . . . . . . . (1 1) tr + K-fp + H20 + sp<--->cl+ bi + ep + ab (12)

(see Fig. 6) Rocks with a high ambient P(C02) during the reaction sequence

now contain assemblage (Y), those with a lower P(C02) contain assemblage (X). The reactions proposed are not reversible and are probably quite slow, this accounts for the continued presence of tremolite and sphene in the retrograde assemblages. Lithologies containing little tita- nium probably conform more closely with the reaction proposed by Carmichael ( 1970) :

. . . . . . . . . . . .

bi + cc + qz < - - - >Ca-am + K-fp + C02 + H20, subsequent retrogression after degassing results in assemblage (Z).


Fig. 6 shows that the possible range of temperatures for the reactions (9) to (1 2) is large and depends upon confining pressure, and where the reactions occur in P(C02)/T space in relation t o the limits of tremolite stability. Assuming, since reactions (9) to (1 2) represent the lower limits of tremolite stability, that the reaction sequence occurs close to the curve :

3tr + 6C02 + 2H20<--->5ta + 6cc + 4qz,

-40-

then the temperature ranges are: P(f) = 5 kb .... T = 380 - 45OOC P(f) = 7 kb .... T = 460 - 52OOC

Why does a reaction sequence similar t o (9) to (1 2) not occur in the pelitic and psammitic rocks which contain both biotite and calcite? The precise reasons for this are not clear but it can be seen that an increase in P(C02) raises the equilibrium temperature for tremolite-forming reactions. Although it is not always obvious that P(C02) was higher in the pelites and psammites than in the mafics it is genenrally true that the carbonate phases, when present in the former two, dominate over the phyllosilicates, The implication of this is that the ratio of carbonate to phyllosilicate is a truer representation of P(C02): P(H20) than simple carbonate proportion of the whole rock.

Summary of mufic data

Information about the mafic rocks of the study area indicates the importance of degassing during prograde metamorphism as a factor in determining retrograde reactions, and hence observed mineral assemblages. Comparisons with existing experimental data do not provide well defined temperature estimates: the minimum likely is 38OoC (5kb) and the maximum is 52OoC (7kb).

Degassing - the removal of C02 during prograde metamorphism - has resulted in the operation of non-reversible reactions to form the observed assemblages. Many of the volcanic lithologies are characterised by 1 t o 2 mm thick calcitic partings along tectonic cleavage planes. This could be one of the sites for the precipitation of C02 released by metamorphism, if Ca were also available.

VI. Mineralogy

Compositions of mineral phases in the individual lithological groups have already been described. This section is not a detailed descrip- tion of the mineralogies encountered but a comparison of the salient features which show some variation.

Biotite

Metamorphic biotites from similar metamorphic terrains to that studied here are described by Cooper (1972) and Ramsay (1973). Fig. 7

mofics

40

- K

Fig. 7 . denotes pelites.

Compositional plots of analysed biotites. Ps denotes psammites, and Pe

-42-

shows the chemical variation of all the biotites analysed. The range of compositions is greater than those of either Cooper or Ramsay. Certain features are apparent from the plots:

I ) Rock composition has a strong influence on mineral chemistry. 2) Some variations within rock compositional groups can be dis-

tinguished. 3 ) The AFM diagram illustrates significant changes more clearly

than the AKF plot.

Two general trends can be deduced from the AFM plot. In both pelites and psammites increasing grade is marked by an increase in Mg/Fe ratios, and in the latter, there is a concomitant increase in the proportion of A13. These two separate trends within the psammites may be explained by the paucity of material to form biotite. Initially, reactions 6 and 7 form biotite at the expense of chlorite, hence the gradual increase in Mg/Fe ratio of the biotite. However, with the depletion of Chlorite, reaction a (section IV) could operate which would then increase the A1 content and slightly depress the Mg/Fe ratio in the biotite. The assess- ment of grade is discussed later (see section VII). Mafic biotites form a separate group, principally because of their different mode of origin, and the composition of their host rocks.

Muscovites

Bulk rock Composition also seems to have a strong control over muscovite chemistry. Two well defined regions are visible on the AKF plot of Fig. 8: aluminous pelites are distinct from psammites and pelites in general. Certain other aspects are also apparent on closer inspection. Psammite and pelite assemblages containing both biotite and dolomite have bimodal muscovite populations. This is taken as confirmation that muscovite is an active participant in biotite formation. Rocks of a similar composition with no chlorite contain muscovites which form part of the variation sequence A-B-C (Fig. 8) representing reactions (6) and (7). Certain authors (Guidotti and Sassi 1976, Hudson 1980) have suggested a correlation between celadonite content and grade of metamorphism. Decreasing celadonite in muscovites is thought to be representative of increasing grade. The muscovites of this study when plotted on an Si- Al-(Mg + Fe) diagram (Fig. 8), show a larger spread of compositions than that shown by the above authors, but do not show consistent trends. Variability in composition is best explained in terms of bulk rock chemistry and does not appear to be grade dependent.

-43-

40 30 20 I0 - K'

30 20 I0 0 - f No

S

A QF

inc.

content celodonite

pelites

0 psommites

0 mofics

0 curbonotes

\ 0 . \ \

50 45 40 35 30 1

Fig. 8. Compositional plots of analysed muscovites.

-44-

Chlorite

In a similar way to both biotite and muscovite, chlorite composition reflects the chemistry of the rock in which it occurs. Two conclusions can be drawn from the spread of chlorite points (Fig. 9): 1 ) Pelitic and psammitic rocks show approximately the same vari-

ation, while mafic lithologies contain chlorites with a higher Mg content. There is a slight trend for Mg and A1 to increase with respect to Fe with increasing grade, whitin the psammite and pelite lithologies.

2)

A mp h iboles

All analysed amphiboles fall in the tremolite-actinolite compositional field with limited A1 substitution towards the hornblendic amphiboles. Those from meta-carbonate horizons approach the ideal tremolite composition, while others show a range towards the actinolite end of the series.

There are a number of workers who propose the presence of a miscibility gap between coexisting amphiboles in the tremolite-actinolite and hornblende ranges (e.g. Cooper and Lovering 1970). This gap is thought to be temperature dependent (Misch and Rice 1975). If such a chemical discontinuity exists, it furnishes some indication of temperatures of metamorphism. Fig. 10 illustrates the possible size of the gap analysed in coexisting amphiboles of this study. Three specimens yield useful information and the temperature range indicated is 50O-58O0C (Pc = 5 kb).

Epidote

Epidote compositions are intermediate between piemontite and clinozoisite. No consistent patterns are apparent, but those formed in pelites and carbonate-bearing rocks are more Al-rich than others.

Carbonate minerals

Carbonate minerals calcite, dolomite, and ankerite are common accessory minerals in most rocks, and dominate the carbonate horizons. Calcite and dolomite, with compositions close to the ideal, occur in all the carbonate lithologies. Variations in carbonate chemistry are common in the other lithological groups. This variation in chemistry is closely linked with biotite-producing reactions. Increase in the proportions of biotite is associated with Fe-bearing calcites (Fe mo1%>3) and an increase in the amount of ankeritic dolomite.

-45-

Table 1

Table to show the relationship between geothermal gradient (dT), overburden pressure (P) , and overburden thickness (o.t.1 for the pre-Carboniferous rocks o f Prim Karls

Forland, assuming an average density of rock of 2.65 g r n J ~ r n - ~

dT (OC) p (kb) 0.t. (km)

9 10.3 12 7.7 15 6.2 18 5.2 21 4.4 25 3.7 28 3.3

38.9 29.2 23.3 19.4 16.7 14.0 12.5

0 100 80 70 60 50 40 30 20 10 0

1

Fe'+ Mn'

pe/ites o psammites mafics

Fig. 9. Compositional plot of analysed chlorites.

-46-

VII. Inferred conditions of metamorphism

Grade

There is no variation in the metamorphic 'grade' within the study area either as defined in the original sense of the term 'facies' or using the up-dated definitions of Winkler (1976). Explicitly, this means that the mineral assemblage data do not provide evidence for a wide variation in the metamorphic conditions over the current areal extent of the rocks studied. However, certain observations allow estimates of grade change to be made: 1 ) Schistosity development: rocks of the Ferrier Group are more

obviously schistose than those of the Peachflya and Geikie Groups, micaceous flakes are easily visible to the naked eye ( 1 - 2 nim diameter), whereas in the younger rocks, although a cleavage is apparent, individual phyllosilicate mineral grains are not obvious.

2) Biotite development: porphyroblasts of biotite in lithologies of similar composition are larger and more numerous in the older rocks.

45 - figs. in formula ions

curve from

Misch b Rice I975 data from

400

0.5 40 0.5 LO L5 2.0

40 -

A/

0.5 -

A I ~ A/

Fig. 10. Miscibility gap in coexisting amphboles.

- 47-

3) Mineral composition: data from muscovites, biotites, and chlorites, when combined with that from points 1 and 2 above, indicate certain broad trends with changing grade (Figs. 7, 8 , 9 ) . Although points 1 and 2 above do not allow a quantitative grade

variation to be characterised, they do indicate that, in general terms, grade does vary with inferred stratigraphic depth.

These considerations lead to the conlcusion that while grade varies only slightly, it does increase from top to base of the stratigraphic sequence. In addition there seems (from mineral compositions) to be a grade increase from south to north within the Ferrier schist belt. This is a result of differential movement on the Southern Grampian Slcde, the tectonic break that separates the Ferrier Group from the younger Geikie and Peachflya Groups (Morris 1979).

Pressure and temperature conditions

It is clear that temperature estimates depend upon a knowledge of pressure during metamorphism. Although there is not yet full agreement on the stratigraphy of Prim Karls Forland (see for example Harland et al. 1979 and Hjelle et al. 1979), the author is in general agreement with the interpretations of Harland et al. At present a total of some 6 km of metasediments are exposed on Prins Karls Forland (Harland et al. 1979) and throughout this sequence grade variation is roughly parallel to the variation with stratigraphic depth. All Forland Complex rocks contain evidence for a single thermal event and similar post-metamorphic histories (Manby 1978a, Morris 1979), so that it seems reasonable to assume that the main metamorphism occurred simultaneously throughout the Com- plex. Thus, the rocks suffered 'burial' metamorphism to low grade. The youngest rocks on the island - the Barents Group (Harland et a1 1979, Manby 1978a) - show evidence of incipient biotite formation in certain lithologies indicating an absolute minimum temperature of about 350OC. At the base of the stratigraphic sequence exposed in south-central PKF are the Ferrier schists. These contain assemblages which are consistent with a range of temperatures from 400-5OO0C. This gives a clue as to the possible range of geothermal gradient ambient during metamorphism, i.e. 9-2SoC/ km' . Since the temperature at the top of the column must be about 35OoC, an estimate of the overlying thickness of rock during metamorphism can now be obtained. By assuming an average rock density of 2.65 g m / ~ m - ~ (e.g. Ramberg 1967) a pressure estimate can be made giving consistency with the inferred metamorphic reactions. Table 1 shows the possible values of overburden thickness (OT), pressures (P), and geothermal gradients (dT).

-48-

100 200 300 400 500 600

doto from: Wink& l976;

Sloughfer &? 2.1975; &i&/ 8 OATusch, 1975.

2

4 k, kb

6

8

The graph shows the PT conditions for Some of the critical reactions which are inferred to have been important in the metamorphism of PKF. Each heavy line, labelled with a palaeogeothermal gradient represents the PT space occupied by the rocks of PKF at the indicated geothermal gradient during metamorphism. The top left end of each line represents the Grampian Group and the bottom right end of each line, the Ferrier Group. Geothermal gradients of 25 and 28'C.km-' give ranges which are consistent for most assemblages, however, a large part of the tremolite field (as defined by Slaughter et a1 1975) is covered. Tremolite might therefore be expected to be abundant in carbonate rocks, but it is not. The 12 and 15°C.km-1 ranges, although fitting for the lower biotite reaction, and for chloritoid, indicate that the higher biotite reactions (c) might not occur. Gradients of 18 and 2l0C.km-' give the best fit to all the data: allowing biotite to form by two sets of reactions (hence biotite in the Grampian, Peachflya and Ferrier Groups), chloritoid is marginally stable, and only a small overlap into the tremolite field occurs.

a) st + ph + ac bi + cl + ep + H20 b) Al-mica + Fe-cl cd + qz + H20

c) K-mv + cl mv + bi + H20 mv + do + cl bi + cc + qz + H20 d)+f) do + qz + H20 ta + cc + C02

e) ep + qz gt + an + H20 g) cd + qz sr + gt + H20

Fig. 1 1 . Critical reactions, palaeo-geothermal gradients, and Prins Karls Forland's pre-Carboniferous rocks represented in pressure/temperature space.

-49-

Pressures in excess of 7 kb do not give consistent results for the reactions inferred, for carbonate and psammitic rocks temperatures would need to have been above 7OO0C which would radically alter the pelitic assemblages. The minimum temperature of 35OoC is an absolute minimum and given pressures of 4 kb is more likely to be about 43OoC (Winkler 1976) and is therefore not consistent with a higher dT value of 2 l0C/ km- ' or higher. It would seem, therefore, that the minimum temperatures (those at the top of the presently exposed Forland Complex) were probably of the order of 40O-45O0C and the maximum temperatures (at the base) were perhaps about 500°C or slightly higher. These temperatures give dT = ca 18°C/km-' which leads to pressure estimates of about 6 kb. These pressure and temperature limits are consistent with all the data from the various lithological groups (Fig. 1 I ).

Conclusions

Southern PKF forms part of a low grade metamorphic belt which shows characteristics of low geothermal gradients during metamorphism. This, together with the predeformational timing of the thermal event has resulted in widely separated 'isograds' which approximately parallel the stratigraphic boundaries.

Stratigraphic links with the main island have been discussed elsewhere (Harland et al. 1979). There is also much evidence to suggest that the metamorphism suffered by PKF may be coeval with that recorded in Oscar 11 Land (Horsfield 1969, Manby 1978b, Hjelle et al. 1979).

Although it would be rash to attempt regional interpretations solely on the basis of this study, a number of points do emerge:

1) PKF, and probably other parts of western Spitsbergen, form part of a large metamorphic/orogenic region.

2) Geothermal gradients obtaining at the time of metamorphism were low, close to the 'normal' geothermal gradiet at the present time.

Acknowledgements

The author is grateful to W.B. Harland for providing the oppor- tunity and logistical support for work in Spitsbergen. The Natural En- vironment Research Council supplied funding for field and laboratory research. Drs. Norman Charnley, Jim Long, and Geoff Manby gave in- valuable advice and assistance during the analytical stages of the work, and Louise Speed produced two typewritten drafts from handwritten notes.

-50-

References

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Atkinson, D.J., 1954: The geology of Prince Charles Foreland and adjacent parts of northwestern Spitsbergen. Ph.D. Thesis. London University.

- 1956: The occurrence of chloritoid in the Hecla Hoek formation of Prince Charles Foreland, Spitsbergen. Geological Magazine 93: 63 -71.

- 1960: Caledonian tectonics of Prins Karls Forland. International Geological Congress 21st Session, Copenhagen 1960. Section 19: 17-27.

Barron, B.J. and L.M. Barron, 1976: A model for greenschist facies equilibria in altered mafic volcanic rocks at Sofala, New South Wales. American Journal o f Science 276: 637-670.

Bdthune, S., 1976: Formation of metamorphic biotite by decarbonation. Lithos

Brown, E.H., 1975: A petrogenetic grid for reactions producing biotite and other Al-Fe-Mg silicates in the greenschist facies. Journal of Petrology 16: 258-27 1.

Carmichael, D.M., 1970: Intersecting isograds in the Whetstone Lake Area, Ontario. Journal ofPetrology 11: 147-181.

Cooper, A.F. and J.F. Lovering, 1970: Greenschist amphiboles from Haast River, New Zealand. Contributions to Mineralogy and Petrology 27: 11-24.

Ernst, W.G., 1963 : Significance of phengitic micas from low-grade schists. Ameri- can Mineralogist 48: 1357-1373.

Francis, G.H., 1956: Facies boundaries in pelites at the middle grades of regional metamorphism. Geological Magazine 93 : 253-268.

Greenwood, H.J., 1967: Mineral equilibria in the system MgO-Si02-H20-C02. In: Researches in Geochemistry (Ed. Abelson, P.H.) 2: 542-567. Wiley, N.Y.

Guidotti, C.V. and F.P. Sassi, 1976: Muscovite as a petrogenetic indicator mineral in pelitic schists. Neues Jahrbuch Mineralogie 127: 97-142.

Harland, W.B., W.T. Horsfield, G.M. Manby, and A.P. Morris, 1979: An outline pre-Carboniferous stratigraphy of central western Spitsbergen. Norsk Polar- institutt Skrifter Nr. 167: 119-144.

Harte, B. and C.M. Graham, 1975: The graphical analysis of greenschist to amphi- bolite facies mineral assemblages in metabasites. Journal of Petrology 16: 347- 470.

Hjelle, A., Y. Ohta, and T.S. Winsnes, 1979: Hecla Hoek rocks of Oscar I1 Land and Prins Karls Forland, Svalbard. Norsk Polarinstitutt Skrifter Nr. 167: 145-169.

Horsfield, W.T., 1969: The geological history of Western Oscar II Land, Spitsbergen. Ph.D. thesis, University of Cambridge.

Hoschek, G., 1969: The stability of staurolite and chloritoid and their significance in metamorphism of pelitic rocks. Contributions to Mineralogy and Petrology

Hudson, N.F.C., 1980: Regional metamorphism of some Dalradian pelites in the Buchan area, NE Scotland. Contributions to Mineralogy and Petrology 73 : 39-5 1.

La Tour, T.E., R. Kerrich, R.W. Hodder, and R.L. Barnett, 1980: Chloritoid stability in very iron-rich altered pillow lavas. Contributions to Mineralogy and Petrology 74: 165-173.

9 : 309-3 18.

22: 208-232.

-5 1-

h i t c h , E.C., 1976: Zonation of low grade regional metamorphic rocks, Nambucca slate belt, Northeastern New South Wales. Journal of Geological Society of Australia 22 (4):413422.

Manby, G.M., 1978a: The geology of north-central &ins Karls Forland, Svalbard. Ph.D. thesis, University of Cambridge.

- 1978b: Aspects of Caledonian metamorplusm in central western Svalbard with particular reference to the glaucophane schists of Oscar 11 Land. Polar- forschung 48: 92-102.

Mather, J.D., 1970: The biotite isograd and the lower greenschist facies in the Dal- radian rocks of Scotland. Journal ofPetrology 11: 253-275.

Metz, P. and V. Trommsdorff, 1968: On phase equilibria in metamorphosed siliceous dolomites. Contributions to Mineralogy and Petrology 18: 305-309.

Misch, P. and J.M. Rice, 1975: Miscibility of tremolite and hornblende in the pro- gressive Skagit metamorphic suite, North Cascades, Waslungton. Journal of Petrology 16: 1-21.

Moms, A.P., 1979: The geology of south-central Prim Karls Forland, Svalbard. Ph.D. thesis, University of Cambridge. 1981 : Competing deformation mechanisms and slaty cleavage in deformed, quartzose meta-sediments. Journal of the Geological Society London 138: 455462.

Ohta, Y., 1979 : Blue schists from Motalatjella, Western Spitsbergen. Norsk Polar- institutt Skrifter Nr. 167: 171-217.

Reed, S.J.B. and N.G. Ware, 1975: Quantitative electron microprobe analysis of silicates using energy-dispersive X-ray spectrometry. Journal of Petrology 16:

Robie, R.A. and D.R. Waldbaum, 1968: Thermodynamic properties of minerals and related substances at 298.15'K(25OC) and one atmosphere (1.013 bars) pressure and at higher temperatures. United States Geological Survey Bulletin 1259: 256.

Segonzac, G.D. de, 1969: Les mineraux argileux dans la diagenese passage au meta- morphisme. These pour obtenir le grade de Docteur des Sciences Naturelles, St rasb ourg .

Seidel, E. and M. Okrusch, 1975: Chloritoid-bearing metapelites associated with glaucophane rocks in western Crete, Greece. Contributions to Mineralogy and Petrology 49: 105-1 15.

Slaughter, J . , D.M. Kerrick, and V.J. Wall, 1975: Experimental and thermodynamic study of equilibria in the system CaO-MgO-Si02-H20-C02. American Journal of Science 275: 143-162.

Statham, P.J., 1976: A comparative study of techniques for quantitative analysis of X-ray spectra obtained with a Si (Li) detector. X-ray Spectrography 5 : 16-28.

Storre, B. and K.-H. Nitsch, 1972: Die Reaktion: 2 ZOISIT 4- 1 C02 = 3 ANORTHIT 4- 1 CALCIT 4- 1 H20. Contributions to Minerdogy and Petrology 35: 1-10.

Thompson, A.B., 1976a: Mineral reactions in pelitic rocks: i. Prediction of P-T-X(Fe- Mg) phase relations. American Journal of Science 276: 401-424. 1976b: Mineral reactions in pelitic rocks: ii. Calculation of some P-T-X phase relations. American Journal o f Science 276: 425-454.

Thompson, J.B., 1957: Graphical analysis of mineral assemblages in pelitic schists. American Mineralogist 42: 842-857.

Thompson, P.H., 1973: Mineral zones and isograds in 'impure' calcareous rocks, an alternative means of evaluating metamorphic grade. Contributions to Mineralogy and Petrology 42: 63-80.

-

499-5 19.

-

-5 2-

Tyrrell, G.W., 1924: The geology of Prince Charles Foreland Spitsbergen. Trans- actions of the Royal Society of Edinburgh 53: 443478.

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Weiss, L.E., 1953: Tectonic features of the Hecla Hoek formation to the south of St. Jonsfjorden, Vestspitsbergen. Geological Magazine 90: 273-286.

Winkler, H.J.F., 1976: Petrogenesis of metamorphic rocks. Springer-Verlag. Fourth edition.

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40: 225-238.

57: 524-553.

PLATE I

1. Specimen M2045. Two textural modes of Muscovite, parallel to bedding (so) and parallel to the pressure-solution tectonic cleavage (SI). Plane polarised light.

2. Specimen M2958. Predeformational biotite grain (bi) retrogressing to and crosscut by muscovite (mv) parallel to S1. Plane polarised light.

3. Specimen M2752. Predeformational chloritoid lath (cd) with quartz pressure-solution overgrowth (qz-og). Crossed polarisers.

Specimen M2868. Chloritoid rosette growth retrogressing to chlorite. Plane polarised light.

4.

5. Specimen M2752. Pre-deformational pyrite grain surrounding chloritoid laths and with pressure-solution quartz overgrowths. Crossed polarisers.

Specimen M2901. Fine-grained, cleavage-parallel biotite development. Plane polarised light.

6.

All scale bars are 0.5 mm long.

-5 3- PLATE 1

Carbonates

T o t a l s

61.0 6.0

61.0 6.0 -

57.8

54.3 6 .O

54.2 - 6.0

50.1 6.0

96.0 14.9 96.7 15.0 97.1 15.0-_ 97.1

97.8 8.0

94.8 8.0

15.0 87.6 15.2

-6-0 .

-

15.1 -

- $1.7 -

- SpNo

2064

2864

3041

3004

2853

2853

-

2864

2864

2862

2862

3004

300 4

3041

2853 -

N.

- -

.

3.i 0.2

O X 0.2

P e l i t e s

Hs 0.9 0.1 0 .4 0.1

-

13.7 2.2

16.6 2.5

19.7 4.2

21.1 4.4

19 .O

194% I --4.o_

b.6 0.1 2.0 0.5

- nin -

cc

cc

cc -_ cc

do

do

t r

tr

t r -_ tr

CP - e p - mv

mv -

ret Fez

0.4

0.5 P. o... . .

.O.o_ .-

0.4 0 .o 0.5 0 .o 3.2

5.6

5 .O

. .- - -

.. -.

0.3- -. .

327.. Q,? 0.6 -Or6

_02-. -0 -7 6.6

- 0.8 -!LI._.P* 8.3..

.. 0.5 7.4 . . 0.5

0.5- - - - - 0.1 0.1

5.. i

Fe3 N'O'

6

6

6

6

6

6

23

23

23

23

13

13

24

24

- . ...

___ _ _

- - . . . . . ._~

. ..

-.

. . .

- .

0.5.

3 L -~

- -

IpNo

z y v i

ZRliII

28HO

2 i 0 0

2766

2"O'4

2 i 5 2

l HOO

2827

2RdO

2 7 6 1

2958

2892

2958

2HHO

2H80

2 8 R l

1146

2 7 4 ' ~

0.4-- 0 :l

1.4 _0_.2 . 1.3 0.2 26;T-

2.4- 26.9

2.5 55;5--

6.2 29.5

5.5 -

M i n i Na Hg

2 . I 0.5 2.2 0.5 2.0 0.4

w

Ill"

IllV

3.n . !?.e Ill"

2.9 0 . 6 5.1 1.1

nlv

!,,,f

0.7 0.2

w

10.5 11 i 2.6

10.9 b i 2.7

10.6 b l

10.2 b i

b i 9.6 2.3 8.1 b i

-2 -0 13.6 4,4

c h l

14.5 4: 5 ch l

15.6 4.7 c h l

21.3 6.2 ch l

21.1 C I I I

18.8 5.3 ch l

. .-

2.5

2.. 5

- -

'I . 9

A l t

27.9 4.8

211.1 4.8

29.2 4.9

29.0 4.8 30.0

4.9.

-

29.0 4.9

36.7 b.3

17.9 3.5

17.2 3.4

23.7 4,4

17.5 3 . 4

18.4 3.5

i i . 1 3.4

18.7 . 4.7. 20.7

-- 5.A 21.4 5.1

21.7 5.0

24.1 5.3

28.1 6.2

Fc2

0.4

0.5

0 . 4

9.2. 0.2

0 .4

0.2

-,

0.2

0 2 - ~

5.1

4.2 -

4.5

2.7

2.4

3.0

-5 4-

Appendix 1. Minera l a n a l y s e s

Fe3

_ _ .

2.6

2.6

2.6

2 . 7

- 2 . 5

2 2

-

4.2

4.2

4.3

4.2-

4.3

4.0

5.0

1.6

1.5

2.2

1.6

1.7-

__ 1.6 -

-_ 2.4

2.4

2.7

2.7

2.9

3.4

57.4 0.5 1 . 5

51.1

52.5 0.6 7.4

5 ) . 1 - o-.a -2.4

53.4 0.6 7.4

' 49.1 0.9 1.1

46.3 1.3 6.7

36.0 2.0 -6.0

36.2 -119 6.1

36.1 2.2 5.8

37 .o 1.8- 6L2

38.0 , 126 6.2

37.2 1.8 6.2

26.4

27.2 -2.L 5.6

27.3 2 . 5 5.6

29.2 2.3 5.1.

29.6 1.5 4.3

27.7 2.8 5.2

-0.k- 7.4

-2.3 5.7

0.2 0.0 2.7

10.4 1.9

10.2 1.9

10.2 1.8

10.1

10.3 1.8 8.9 1.6 8.4 1.6 9.3 2.0 9.3

1.n.

0.1 0.4 3.2 0.0 0.0 0.4

0 .5 4.1 0.1 0.5 0.3 3.8 0.0 0 .4

0.3 0.5 2.0

0.3 0.6 1.4 0.0 0.1 0.2 0.2 0.5 3.7 0.0 0.1 0.4 0.2 1.1 0.0 0.2

1.4 18.5 0 .2 2.6 1.5 18.1

o.o..oo,i. 0 . 2

1 I 3.0

N' 0'

24

24

24

24

24

24

24

24

24

24

24

24

24

28

28

28

28

28

2 8

-

. _ _

. .

~

Totals

95.5 15.1 Yb.4 15.1 97.9 15.0 98.7 15.1 - 98.7 15.0 Yb.4 15.2 93.8 15.0 9 3 . b 17.0 93.5-

- 17 LO- 95.4

- 16.2 94.0 16.9 97.2 16.8 94.3 16.7

20 .o 89.8 19 .L 90.8 19.9 89.1 19 .a 94 .5 19.8 93.8 19. I

- -_ 87.6

-5 5 -

- SpNo

2859

__

2 n ~

2 1 5 2

2898

2970

3000

2800

2760

2892

2092

2761

79 70

2970

2161

~

- bI i n

c t d

c t d

c t d

___.

c c

c c

c c -

c c

ank

ank/ do

ankl d o

dnk I do

C P

ab

ab

P s a m m i t es

- Hs

0.9 0.1 0.8 0.1 2.2 0.3 1.0 0.1 0.4 0.1 0.5 0.1 1.4 0.2 3.1 0.4

17.0 2.4

17.7 2.5

13.4 1.9

-

--

-

A l t

40.9 4.2

41.0 4.5

40.1 4.5

-

0 . 3 0 .o

25.0 2.5

19.5 4. I

181.2

4.0 -

si 34.9

3.1 27.2

2 .6 26.8

2.5 6.3- 0 .o 0.4

0 . 4

0.5 0 .o 2.5 0.2 0.2 0 .o I .o 0.1 0.5 0.0

3R.4 3.2

67.8

6H.'I 12.1

0 .O

0.0 -

12.0

-

K

3.2 0.4

0.3- 0.1

-.

:a T i Fet

15.6 1.2

1.9 l.-3 21.6

2 K 2

0.1 .1.7 5? ;a 6.8 5.1 .- 0.1

54.9 0.6 5.8 0.1

54.5 1.1

55.7 2.2

40.9 14.7

29.7 4.7 3.1 0.4

28.5 4.1 2.9 0.3

29 .O 9 .O 3.0 0.7

21.R 9.5 2.2 0.7 0.2 0.) 0.0 0.0

5.1.- ...- 0.1

5.5.- . _ 0.2

4.9 .- 1.1

N'O

14

14

14

6

6

6

6

6

6

6

6

13

32

32

-

-.

-

Tota l

96.2 9.0

94.0 9.2

93.0 9.2

5b.5 6.0

56.4 6.0

56.9 6._0

50.7

61.9 5.8

52.3 6.0

5.Y 55.1

6.0 96.Y

8.5 Y V . I

' I Y . > 19.9

50

57.8

20 .IJ

-56-

Mafics

Abbreviations1

SpNo - specimen number LcNo - l o c a t i o n number ana lyses l i a t s are t h e c o r r e c t e d oxide percents . The Hin - minera l lower numbers are t h e foruula ion va lues c a l c u l a t e d on - ans lysed us ing the microprobe t h e banis of t h e number of oxyBen u n i t s i n the mineral x - maJor minera l conmtituent fornula (as denoted i n column “ 0 ’ ) . o - minor minera l c o n s t i t u e n t Discrepancies i n t h e t o t a l s ire due t o rounding e r r o r s

i n t h e c a l c u l a t i o n a .

The upper numbers i n t h e elemental columns of t h e mineral

low grade (greenschist facies) metamorphism in southern prins karls forland, svalbard

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