Aspects of Pressure Solution As a Deformation

ASPECTS OF PRESSURE SOLUTION

- by -

Robert Kerrich

Imperial College, London. 1974. CONTENTS

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=AUER 1 Introduction to concepts of pressure solution . • 1- 10 CHAPTER 2 The historical development of research on pressure solution and related topics 11- 60

PREFACE TO CHAPTERS 3 AND 4 .. • • • • 61- 63 CHAPTER 3 Fluid inclusion studies of tectonic veins 64-120 CHAPTER 4 Oxygen isotope geochemistry applied to tectonic environments 121-174 `CHAPTER 5 Pressure solution and deformation

mechanism fields • • • • 175-214 * CHAPTER 6 The interrelationship of pressure solution and tectonic veining 215-220

CHAPTER 8 The distribution of chemical elements and mineral species in the deformed matrix around rigid cylindrical inclusions 231-243 CONCLUSIONS. APPENDIX I Density measurements on fluid inclusions 244

APPENDIX II H20/C02 determinations on fluid inclusions 244 APPENDIX III Measurements of salinity of fluid inclusions 244 APPENDIX IV Preparation of mineral separates for 245-250 isotopic determination • • • • APPENDIX V Fluorination and mass spectrometry 251-254 APPENDIX VI 180/160 values of mineral separates 255 analysed • • • ACKNOWLEDGMENTS

I would like to thank my supervisor, Dr N.J. Price, for help and stimulating comment throughout this research. Dr 1. 3uoh made representations of the oxygen isotope

project to the 'Universities Stable Isotope Comrnitt 0.1.1.-Lv kindly ap17)7..oved the work. Isotopic analyses were done at n.':`]. R. Aldermarston by J. Durham, to whom my thanks are ex-

tended for instruction in fluorination 1. - mass spectrometers. Dr R.D. Beckinsale has given much useful advice concerning applications of stable isotope geochemistr71 and has critically read t'ne manuscript of chapter 4. Carbonate determinations were kindly made by Dr Y. Shackleton and Er E. gall at the department of Pleistocene Research, Cambridge. Mr D. Brig s and Er D. 3ernard provided help and instruction in mineral separation procedure. Facilities at the Institute of Geological Sciences have been used, with the kind permission of Dr IT. Snelling. The 1120 /CO2 determinations on fluid inclusions were done by Dr T. Sheppard of I.G.S., who has been generous in allowing use of his microscope freezing stage apparatus. The X-RF at ire nerial College has been used for geochemical analyses by kind permission of Dr G. Borley, who has also advised on theoretical aspects of X-ray spectrometry and geochemistry. Er G. Bulland gave practical instruction in sample .preparation and running of the spectrometer. Data reduction was kindly done by Mr R. Parker. Samples for analysis by XRD were run by Mr R. Curtiss, who helped with interpretation of the output. Thanks are due to Mr J. Blount and Mr D. Bailey for rock cutting and polishing services respectively. Dr G. Borley has given freely of facilities for making photomicrographs. Mrs B. Richardson and her staff in the photographic department have provided help and instruction in producing the plates. Professor J.G. Ramsay, Drs A. Siddans, P.R. Cobbold, A. Ries, S. Schmidt, and M. Casey, A. Bell and C. Taylor have kindly suggested field areas suitable for investigating press ure solution. Much benefit has accrued from discussions with Drs E. Rutter and S. White, to whom I am most grateful for critical reading of the manuscript. In particular I would wish to extend my thanks to Martin Casey. Much of our research has run in parallel, and an in- valuable interchange of ideas has taken place during the two years in which we have shared a room. This research has been undertaken during tenure of a Natural Environment Research Council grant. ABSTRACT

Pressure solution, which may be defined as the solution, diffusion, and precipitation of rock forming minerals in re- .sponse to sl3ress fields, bas lon7 been reco— ed as an im po tant mode by :hich low grade metamorphic rocks deform. forraton on the PT conditions of crustal environments in which pressure solution is the dominant deformation mech,sni= has been obtained using oxr:Ren isotoT:e and fluid inclusion thermometry. In addition, data from the former technique has been used to define empirically the thermal stability limits of some metamorphic mineral assemblages. Field work for this research has been conducted in the Dalradian Series of Southwest Scotland, and Palaeozoic rocks of the Central Pyrenees. Estimates of the crustal conditions over which pressure solution is an important deformation mechanism in dominantly quartz bearing assemblages are 30°C to 450°C ± 50°C, ir- respective of the crustal depth ( =mean stress). Oxygen isotope geochemistry has been used to assess the equilibrium relations between tectonic veins and their host rocks. In addition, thermometric data on veins, obtained from this method, has been related to the mineral assemblages and deformation mechanisms characteristic of tectonic veins at different crustal levels. Some information is adduced from data on fluid inclusions and oxygen isotopes, in minerals from tectonic veins, concerning the nature and sources of fluids in tectonic environments. Attention has been directed towards the problem of non- equilibrium isotopic fractionation in stress fields. Quantitative data relating to the distribution of chemical and mineral species around cylindrical inclusions with pressure shadows, has been obtained from XRF and XRD. The measured distribution of chemical and mineral species is related to stress fields, and to considerations of kinetic effects. Aspects of tectonic striping are discussed with regard to the physics of diffusion. 3.

CHAPTER I

Introduction to concepts of pressure solution

Terms and definitions.

Pressure solution is a term which is widely used to describe the phenomenon of solution, diffusion, and precipitation of rock forming minerals. Together, these three elements of the process constitute an intercrystalline mechanism by which rocks may deform. Although use of 'pressure' imparts a genetic sense to the term there can be little serious doubt that 'pressure effects', or 'stress' play an important role in this process. In addition, 'solution' bears the implication that pore fluids are present in the rock system, and that the diffusive element of pressure solution takes place through such a fluid phase. There is an ill-defined state between pore fluids and water adsorbed at grain boundaries. This has led to use of the term 'water assisted diffusive processes' to describe the phenomenon, in preference to pressure solution. However, the former term is sufficiently general to include many metamorphic re-

)actions, in addition to at least one intracrystalline deformation mechanism. Pressure solution, as understood above, bears a close relation to a process known in the metallurgical literature as 'Coble Creep'. Coble creep is a mechanism whereby deformation is achieved through a transfer of atoms from points of high stress to points of low stress along grain boundaries. No fluid phase is involved in this process of 'grain boundary diffusion'. There are disadvantages to the use of either 'Coble Creep' or 'grain boundary diffusion' in reference to the geological •

phenomenon. First, there is a certain knowledge that in some instances diffusion takes place through a fluid phase, rather than exclusively along grain boundaries. Second, the latter term bears no inference concerning the role of stress or deformation in this process. These are two factors which have been in- timately associated with pressure solution at a descriptive level, since inception of the concept. Solution transfer also suffers from this inadequacy. Grain boundary diffusion could equally be used to describe the growth of authigenic minerals during diagenesis, or granitisation by grain boundary ichors. In conclusion, it would seem that pressure solution remains the most appropriate term to cover the process of solution, diffusion, and precipitation, of rock forming minerals in response to differential stress. It should be noted, however, that some workers do not consider that such a mechanism exists at all.

Boundaries of the phenomenon. There is a concensus of opinion in the` literature, arrived at by subjective means, that pressure solution is an important mode by which rocks at a low grade of metamorphism deform. Pressure solution is known also to play a significant role in the diagenesis of some sediments. At deeper crustal levels, - there is no satisfactory distinction between mineral segregation by pressure solution, and what is loosely termed metamorphic differentiation. There may in fact be no real difference, because they are probably both the result of diffusion. Such confusion reflects the lack in detailed understanding of physical processes in the earth. ,In this thesis pressure solution is taken as a deformation mechanism that generally leads to small scale mineral segregation, and which is commonly active •

in high and intermediate crustal levels. Features such as gneissic layering, ultrabasic pods, or pegmatites, which are formed in high grade metamorphic terrain, are not included within the context of pressure solution. The term pressure solution does not extend to include hydrothermal transfer of minerals in a moving fluid s7stem.

Pressure solution - the problem.

As a prelude to any research in this thesis a comprehensive literature survey on pressure solution and related topics was undertaken. This survey served two main purposes. First, to establish the state of the science with regard to pressure solution. Second, as an aid to providing a broad perspective from which worthwhile research problems could be evaluated. A detailed account of the historical development of research on pressure solution and related topics was written directly from the synthesis of literature, and is presented as a review of the subject in chapter 2. Although the survey covers a wide range of topics, some of which are discussed individually in different chapters, it was decided that a self- contained review would make for a more unified presentation of the subject. Further, some continuity is lent to a thesis in which the research theme would otherwise be disjointed by numerous introductory and qualifying passages of historical content, that would be difficult for the reader to interrelate. What are the problems posed by pressure solution? In the author's opinion these fall into four broad categories -

I. Descriptive - is it possible to describe and classify the distinctive changes in rocks that generally result from pressure solution activity w 4

2. Strain - can the shape change of a rack that has deformed by pressure solution be measured? 3. Experimental - under what controlled conditions could rock forming minerals, or analogue materials, be made to deform by pressure solution in the laboratory? Mechanisms - what are the detailed physical processes by which pressure solution takes place, and what are the factors that govern these processes?

To become engaged in descriptive geology did not seem a productive approach because major systematic works on pressure solution have already been undertaken. In addition, variations in structures are as numerous as the localities at which they may be found. The-problem of strain determination within structures was not entertained because a general requirement for strain analysis is to establish the presence or absence of a chemically closed system for each measurement. In the case of rocks deforming by pressure solution this is almat a pro- hibitively difficult task. The author decided that a rewarding line of research would be to study pressure solution at the level of detailed physical processes, with the possibility of complementary experimental work in mind. A search was then conducted in a number of fields such as solid state physics, thermodynamics, kinetics, surface chemistry (etc.), with two objectives in mind. First, to ascertain the theoretical work in these fields that had been done directly in relation to pressure solution. Second, to discover if some of the theoretical descriptions of atomic and molecular level processes, which had been developed within these fields, could usefully be brought to bear on research in pressure solution. • s

From the authors readinc and field experience several interesting lines of inquiry relating to the mechanism of treasure solution were evolved. The three main problems that were chosen for detailed investigation are as follows -

1. What are the physical conditions in the crust that determine whether or not pressure,solution is a viable deformation mechanism? 2. The nature of the interrelationship between tectonic veining and pressure solution. 3. Which parameters govern the distribution of chemical elements and mineral_species in a structure which has formed by pressure solution?

These three questions were selected because answers to the problems were considered central to a broad understanding of pressure solution as a crustal process. This type of problem has not been approached empirically before; and it was judged leasible to obtain the necessary answers within the time available. The analytical methods that vere used to obtain quantitative information are introduced in chapters 2 and 4. The author's observations relating to the application of these methods to some aspects of structural geology are included. Chapters 5, 6, and 8 are devoted to the fields ar investigation as outlined above. These chapters represent original research work done by the author, excepting those passages where reference is made specifically to other works.

Research philosophy.

In approaching this research problem the question arose as to the scale at which the process should be analysed. Classic regional studies of structural geology have been undertaken in areas where the rocks deformed largely by pressure solution. Such work yields little information on detailed mechanisms, d serves to illustrate the problems associated with interpreting strain data from rocks which have deformed by pressure solution. One of t'le ersential facets of pressure solution is the diffusion of mineral species over varying distances. As a broad- principle, therefore, this process has been analysed at a ,cale where mineral fluctuations are contained within a system of millimetre to metre scale; i.e. of a size where the limits of fluctuations may be traced with a reasonable degree of certainty. Aspects of this point are argued more closely in the paragraphs which follow. A striking feature that came to light from the literature survey was that, of the,many hundred contributions on pressure solution, few presented any quantitative data against which theoretical concepts could be tested. Hence problems were selected for this research which reauired a large amount of analytical information for their solution. With the main areas of investigation in mind,critical field tests were devised in order to obtain the necessary analytical information. So complex are the problems of deformation mechanisms in crustal rocks that, in the author's opinion, it is essential to start with the most simple and fundamental proposition in the field. Therefore, in designing this field work three important principles were applied -

1. The simplest possible case was investigated initially, in order to reduce the complexity of the problem to a minimum. 2. Only field situations with highly defined boundary conditions were analysed, so that the number of variables was reduced to a minimum. 7

3. Control samples were always talcen in order to make a comparison with the test samples.

With re-2ard to te first point, when pressure solution in folda was the object of investigation, then a single layer buckle in an infinite matrix was selected. in preference to a more 'impressive' structure that exhibited evidence of multiple refolding. The second point is perhaps best illustrated by the boundary conditions that were applied in selecting samples for analysis, as enumerated in chapter 8. In relation to the third point, control samples provided a useful basis against which the test samples could be compared; and yielded much information that was of interest in its own right. Geological processes are always viewed 'through a glass darkly'. The best opportunity for understanding the fundament- als of these processes is by recourse to analysis of their most simple manifestation.

Elementary concepts of pressure solution.

This section gives a brief introduction to some of the more important concepts relating -to pressure solution, within the field of structural geology. It is written as an aid to the reader who is not familiar with the idea of pressure solution; and intended as a basic framework to which some of the detailed analysis of structures in later chapters may be referred. A common feature of structures formed by pressure solution is the spatial relationship of domains of gain and loss - areas where minerals have been removed by pressure solution and migrated by diffusion to a locus of precipitation. Plate 1.1 Zones of pressure solution, and tectonic veins,

associated with formation of a shear zone. Lower

Devonian grits, Marloes, Pembroke

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Plate 1.2 quartz pressure shadows adjacent to a euhedral pyrite

crystal. Innellan grits. Dalradian Series, Southwest • 8

Plate 1.1 illustrates a shear zone formed in quartz grits. Arraya of quartz veins are associated with the shear zone, and normal to the veins are dark zones. One may'infer that during deformation the maximum principal stress was parallel to.the long axis of veins, and hence perpendicular to the dark zones. The dark zones are discrete loci where quartz has been removed by pressure solution activity, to give a concomitant concentration of Minerals that have a low response to deformation by pressure solution. These zones have propagated out through the rock, normal to the maximum principal stress. quartz dissolved by pressure solution diffuses back into a site of low stress formed by the vein, to give vein filling. One of the central concepts relating to pressure solution is that of the 'stress raiser', a rigid object or structure in a deforming matrix. Pressure solution is commonly observed to be associated with the stress perturbations arising from such objects. Plate 1.2 is a photomicrograph of a euhedral pyrite crystal: in a deformed greywacke. The sense of shortening, i.e. the direction of principal stress, is normal to the long axis of the plate. Quartz grains have been dissolved in regions of high stress, adjacent to faces- of the pyrite crystal which are exposed to the maximum principal stress. The dissolved quartz, in solution, has diffused into the regions of low stress, or' pressure shadow, where it is reprecipitated. Another example of a 'stress raiser' initiating pressure solution is given in plate 1.3. In this photomicrograph of an impure calcite slate those margins of a relatively rigid fossil which are exposed to the maximum stress, have initiated planes of pressure solution of a stylolitic type. Calcite has dissolved by pressure solution within the discrete planes to give a dark residual seam of insoluble material, in a manner exact- ly analogous to formation of the pressure solution zones mentioned in relation to plate. 1.1. 9

The fossil itself has been dissolved in the direction of principal stress, which illustrates the strain that results from pressure solution. The dark planes of pressure solution, which are spaced, compositionally different from the rock, and non penetrative, define the mesoscopic cleavage of the slate. Here we have two additional concepts relating to pressure

- solution; first, that of strain in a rock by dissolution under stress, and second, the association of pressure solution and cleavage structures. L characteristic feature of many cleaved rocks that have deformed by pressure solution is that pressure shadows tend to lie in the plane of cleavage, and parallel to the direction of stretching. Plate 1.4 is a photomicrograph of a fractured detrital grain of calcite in an impure calcite slate. Calcite has been dissolved in the dark cleavage zone at the bottom left hand corner, and precipitated in part as calcite pressure shadows on the detrital grain. Dark organic material, micas, and iron oxide minerals db not diffuse into the pressure shadow, which explains its clear colour, but remain in the cleavage zone. Hence, we note that some rock forming minerals such as quartz and calcite respond to deformation by pressure solution, whilst others do not. It is of interest to note that the pressure shadow has resulted in an effective shape change of the detrital grain, which is in no way due to internal deformation. The shape change of grains in this manner is a. very common feature of rocks deformed by pressure solution. Pressure solution activity is often associated with fold structures. Plate 1.5 illustrates cusps which have formed at the interface of folded quartzite and pelite layers. Planes I3.

Hate 1.5 Tectonic stvlolites initiated at the margins of fossil, in an impure carbonate slate, 1-embroke.

•

klate 1.4 Calcite pressure shadows on a fractured detrital grain of calcite. Impure carbonate slate, Ilfracombe, North Devon. Note dark zones where calcite has been removed by pressure solution, assrciated with the pressure shadows. t o of pressure solution have been initiated at the cusps, which are sites- of volume loss; and have •propagated throughout the fold. to give a. fanning cleavage structure.- Offset of the pelite layer across cleavage planes is due to removal of quartz by pressure solution. Some of the quartz would appear to have been precipitated in tectonic veins• associated with the fold structure. Lithology has a strong influence on pressure, solution activity. There is evidence of stronger deformation by pressure solution in the pelite layer than in quartzite (plate 1.5). In the deformed polygenetic conglomerate of the Swiss Helvetic molasse, illustrated in plate 1.6 the grey coloured dolomite pebbles have been dissolved at point contacts in preference to the dark granite, banded gneiss, or ,white quartzite pebbles. Thus we may conclude that pressure solution is an important mode by which rocks at a low grade of metamorphism deform. We have seen how structures formed by pressure solution generally bear a constant orientation to the directions of principal stress, and are associated with perturbations in the stress field. compositional variations arise because some minerals respond to deformation by pressure solution more readily than do others. This factor leads to a number of types of compositional cleavage, and to structures characterised by volume change. Shape change, or strain, of grains takes place by the external subtraction and addition of material at boundaries, rather than by intracrystalline deformation. Finally it has been noted that lithology has a strong influence on pressure solution. late 1.5 Folded layers of impure cuarzit,e and pelite, Ilfracombe, North Devon. ilanes of pressure solution, a and tectonic veins, associated with the fold.

• slate 1.6 iressare solu:,ic)n in deformed conlomerates. Tertiary, lielvetic rolasse, Niederurnen, Glarus, 11,

CHAPTER 2

The historic@ development 'of resarch, on pressure solution'and related- topics

The principles of stress directed melting and solution in -- crystal aggregates were brilliantly formulated and experimentally confirmed by Thompson in 1862. These principles form the essential basis of pressure solution. It is of interest to ' npte that the development of theoretical concepts relating to this deformation mechanism preceded field descriptions of structures formed by pressure solution. Because some degree of deformation is almost universally present in rocks, pressure solution has been described in many branches of the geological sciences; principally in the domain of structural geology, but also in sedimentary petrology, and the geochemistry of ore deposits. As a consequence research literature on pressure solutiom is scattered among both general and specialist journals. The theoretical treatment of pressure solution has evolved in four directions -

1. thermodynamic theories of non-hydrostatically stressed solids. 2. phenomenological theories. 3. theories founded in solid state physics, which are based on a statistical mechanical description of rate processes. 4. computer based numerical studies.

There has been limited success in the use of non- hydrostatic thermodynaMics to describe the deformation of simple monocrystalline aggregates. This approach is not applicable, howsver, to treatment of the complex polycrystalline iaggregates which constitute natural rock systems.

• 12

Phenomenological theories (Weyl, 1959) take into account some physicochemical processes in a generalised way, to give a semi-quantitative description of pressure solution. Theories founded on solid state physics deal with the atomic level at which many deformation mechanisms occur. Pre- dictions of deformation structures in crystals, based on these theories, may be directly compared with crystal defect structure as observed using transmission electron microscopy. Further, the statistical mechanical treatment in solid state physics, is based on factors which determine the kinetics of processes; and hence this approach has a special relevance to the important subject of deformation mechanism-fields. The finite element method is an important numerical technique which may be used for obtaining solutions to problems such as stress fields in complex anisotropic materials, and non-linear diffusion. This- technique enables the effect of a variety of boundary conditions and material properties to be investigated in the development of a given structure. Solu- tions. may then be compared with structures' observed in the field. Useful qualitative results have been obtained from simple experiments on pressure solution, using crystals of high solubility as analogues of rock forming minerals. Continuing attempts are being made in rock deformation laboratories throughout the world to produce pressure solution in natural mineral assemblages. To date, such attempts have met with little or no success. Because research into pressure solution and related topics is conducted over such a broad range of observational and theoretical sciences, the historical development of this subject is treated in four sections. 13 1. Elementary theoretical concepts and experimental work. 2. Early field observations and empirical rules derived from them, 3. The specific role of pressure solution in various structural features. 4. Recent advanced theoretical treatments of deformation mechanisms.

2.1 ELEMENTARY THEORETICAL CONCEPTS AND EXPERIMENTAL WORK RELATING TO PRESSURE SOLUTION.

James Thompson first noted the effect of hydrostatic pressure on the melting point of ice in 1850 and went on to distinguish the case of stressed ice in water (1.862, a, b, c). He wrote "I came also to the more general inference that stresses tending to change the form of any crystals in the saturated solutions from which they have been crystallised must give them a tendency to dissolve away, and to generate, in substitution for themselves, other crystals free from the applied stresses or any equivalent stresses". Thompson made quite clear that this effect is as true for single crystals as for crystal aggregates, and that material is recrystallised in the same state of stress as that of the crystal or crystal face upon which it is being precipitated. He attributed the plasticity of ice, as in glaciers, to such solution effects, and stated explicitly that it is quite distinct from crystal plasticity due to internal atom movement. Experiments by Thompson on the compaction of loosely pack- ed salt in a saturated solution confirmed his predictions, and he considered that diffusion through liquid might cause a "great slowness" of the process. 14

Poynting (1881) derived an expression for the change in melting point of hydrostatically stressed ice in contact with uncompressed water, using the property of vapour tension. He did a series of experiments on stressed ice which confirmed his predictions and those of Thompson (1862a, b, c). Poynting stated that the increase in solubility due to pressure on solid and liquid is small compared with the solubility change from pressure on the solid alone. In experiments which involved the compression of crystal aggregates Spring (1888) discovered that moistened powders compacted and welded better than dry powders; an observation that long predated understanding of the role of water in the kinetics of grain boundary diffusion. Spring suggested that the cementation of sandstones and conglomerates was due to solution at points of high stress giving supersaturated solutions, from which there was deposition where there was no stress. Le Chatalier (1892) derived mathematically the result that non-hydrostatic stress has a greater effect on solubility than hydrostatic stress, independently confirming Poynting's formu- lation. Hostetter (1917a, b) considered the converse to be true, despite the results of his own experimental work, which was in agreement with the evidence compiled by Thompson, Poynting, and Le Chatalier. Le Chatalier accounted for the consolidation of rock salt, gypsum, and carbonates by this process of stress directed solution-transfer, and demonstrated stress cementation of crystal aggregates by compressing sodium nitrate crystals in contact with its saturated solution at 200 bars for 8 days. With the theoretical prediction and experimental verifica- tion of solution transfer in crystals under differential stress • 15 by Thompson (1862a, b, c); and the application of this mechanism to explain several geological processes by Sorby, 1863a, b, 1865; Spring, 1888; and Le Chatalier, 1892, it is indeedotortu- nate that this phenomenon is known as the Riecke principle, an historical inaccuracy which has been perpetuated by several authors (Knopf, 1933; Durney, 1972a; 1972b). Riecke (1894, 1895, 1912) merely proposed that if two prisms of the same solid are originally in equilibrium with a common saturated solution, that application of stress to one of the prisms would cause it to go into solution, while the unstressed crystal would grow at its expense. Riecke made no mention of this tendency to go into solution being greater on certain crystal faces than on others, or of simultaneous solution and deposition on different faces of a single crystal under non-hydrostatic stress; although his principle has often been stated as implying such an effect. The expression which Riecke presented in 1912 for the depression of the melting point of ice under stress

T a 2

where ck = 0.0036, and 6 is the stress in Kg.cm-2 is not in agreement with the formula derived by Poynting, whose theoretical result was confirmed by experimental work. The advanced theoretical treatment by Gibbs (1878) of the thermodynamic relations between a strained elastic solid and its solution proved too difficult for many workers to adapt to geological situations. Gibbs envisaged boundaries separating different crystal faces in contact with a solution at different stress, because he could not conceive of a fluid supporting differential stress; and as a consequence Gibbs' theory has limited direct application to geological problems. • 16

After the turn of the century several authors proposed thermodynamic theories for crystal aggregates under non- hydrostatic stress, having a common intergranular fluid, because this is the situation most representative of natural rock systems. However, these theories made no advance on the work of Gibbs-, because of unfounded and even tacit assumptions. A series of papers concerning the effect of high pressure ,on the physical and chemical behaviour of solids, with reference to metamorphic processes. ere written by Johnson (1912), Johnson and Niggli (1913), and Johnson and Adams (1913). They showed from general thermodynamic reasoning that the effect of pressure'on a solid, in a solid-liquid system, is always to raise the vapour pressure, and hence solubility, of the solid in the liquid by an amount which is much greater than the effect produced by stressing both phases simultaneously; confirming the results of Poynting (1881) and Le Chatalier (1892). Johnson and others denied the validity of Riecke's Prin- ciple. They rederived" the Clausius-Clapeyron equation with reference to melting points and solubilities, but neglected the qualification of negative molal volume of solution for solubility increase with pressure, explicitly stated by Sorby (1863a). It was the contention of Johnson and Niggli (1913) that solution and recrystallisation may occur on crystal faces under different normal stress- for the case of a single crystal, or between crystals under different normal stress; and that crystal anisotropy will be an important factor in the process. They considered that the stress dependent increase in solubility is responsible for the shape orientation of grains in rocks having slaty cleavage. Objections to Johnson's (1912) theory of 'plasticity' by melting at stressed faces in crystalline aggregates were raised 17 by Bridgman (1916), who considered that plastic flow may occur by preferential melting on unstressed, crystal faces, but that this was a minor mechanism of deformation, as indicated by the temperature dependence of flow being exponential rather than linear, as suggested by Johnson. Bridgman derived a complex equation for the temperature change necessary to restore equilibrium to a solid-liquid system subjected to non-hydrostatic stress. He extended his treatment for isotropic homogenous crystals to take account of anisotropy involving the twenty-one elastic compliances. As for the formula of Johnson and Adams (1913), Bridgman's formula reduces to the Clausius-Glapeyron equation (which describes the effect of hydrostatic pressure on melting temperature) for the hydrostatic case. Williamson (1917) considered that Johnson's (1912)'equations were inapplicable to real physical processes, but that his ideas were qualitatively correct. He stated that Bridgman's assumptions were contentious, for it is indeed difficult to visualize the physical reality of temperature changes restoring equilibrium in a natural system under differential stress. Williamson approached the problem, in the same way as Gibbs, by considering the change in chemical potential of a given mass of solid going into solution, in relation to the normal stress across the solid-liquid interface. At the turn of the century controversy raged over the reality or otherwise of the force of crystallisation, a phenomenon closely associated. with pressure-solution (Taber, 1916, 1917). Boydell (1926) put the concept on a theoretical basis by relating the true energy change of a solid and its solution to the work done during a given stress increment. He derived I i8

a linear dependence of concentration upon pressure, although Ramberg (1947) arrived at a different result, in a derivation which is not clear to the author. A unified thermodynamic theory for pressure solution and the force of crystallisation has been developed by Khaimov Mal'kov (1959). This theory resolves earlier conflicting results- by relating the magnitude and sign of the force of crystallisation to values of supersaturation or under saturation relative to equilibrium saturation at a given stress. Experi- mental work done at low values of supersaturation for single crystals has confirmed the initial theoretical prediction of Boydell.

Experimental work relating to pressure solution.

A great deal of simple but elegant experimental work has been done on pressure-solution and the force of crystallisation. Becker and Day (1905, 1916) reported experiments showing that a crystal of alum growing in a saturated solution would lift a weight (force of crystallisation), but that as the weight was increased the loaded surface would dissolve, with growth on unstressed faces (pressure solution). By placing loaded potash alum crystals in chrome alum solution they demonstrated stress solution surfaces with dark chrome alum overgrowths on free crystal faces. Becker and Day (1916) considered the force of crystallisation to be of importance in the opening of fissures and veins. The experiments of Becker and Day were repeated by Bruhns and Mecklenburg (1913), but with an unloaded, as well as a loaded crystal, in the same saturated solution. Not unnaturally they failed to reproduce the results of Becker and Day, and 19 disclaimed the reality of 'force of crystallisation'. Bruhns and Mecklenburg's experiments were identical to the situation envisaged by Riecke, whereby an unstressed crystal would grow at the expense of a stressed crystal in the same saturated solution. Using different types- of crystal Taber (1916) and Wright and Hostetter (1917) achieved similar results to those of Becker and Day. Wright and Hostetter (1917) demonstrated optically that overgrowths on alum crystals occur at the same state of stress as that of the parent crystal. The formal statement of Riecke's Principle led both Knopf (1933) and Russell (1935) to infer that the preferred shape orientation of minerals in metamorphic rocks was not to be explained by a shape change in response to the Riecke Principle; which provides only for solution of material in a stressed system, and redeposition in the pressure-shadows of larger grains. Russell (1935) used photoelastic techniques to.demonstrate stress concentrations in parts of loaded halite crystals, and achieved shape change by solution of stressed alum. Russell considered, however, that a solution-transfer- precipitation mechanism would not work in natural systems; and pointed out that this process failed to explain the preferred crystallographic orientation which is present in many deformed metamorphic rocks. This latter observation does not take into account the contention of Johnson and Niggli (1913) that crystal anisotropy is an important factor in stress directed solution. Experimental work on non-hydrostatically stressed crystals was developed by Correns (1926, 1939, 1949) who showed that 8 growth and dissolution of alum crystals depends on stress, 20

solution concentration, temperature, and nature of the interface. Correns discovered that crystal faces would not grow if bounded by certain materials such as mica or glass. He proposed the concept of a phase boundary force (f) from which the necessary condition for growth of a crystal face may be defined -

fac > fbc + fab

are the crystal-solution, crystal- where 1'8.01 fbe and fab plate, and plate-solution phase boundary forces respectively. The phase boundary effect cannot be readily quantified. It does, however, have significance in a qualitative consideration of the importance of pressure solution in different lithologies, as related to the interface of various mineral constituents. By calculating the work necessary to transform, iso- thermally and reversible, a solution (pressure p, concentration c) into a supersaturated solution (pressure ps, concentration

es), Correns derived an expression of the form RT" (1n. c - in cs) =- Vs (p Ps)

for the change in concentration of solute at a crystal face due to a given change of hydrostatic pressure; where Vs = molar volume of solid; R =-universal gas constant; T = temperature °K). He obtained good agreement between theoretical prediction and experimental work at low values of p. Mosebach (1950, 1951) investigated the influence of crystal anisotropy on growth under conditions of differential stress, and concluded that growth or dissolution depends on the crystal face, its orientation in the stress field, and the magnitude of differential stress. In relation to the study of rock deformation Goranson (1940) revived Johnson's (1912) concept of creep in solids due 21 to solution at point contacts of high stress, which continually shift, and recrystallisation at free crystal faces. Extending this concept to a solution-diffusion-, recrystallisation mechanism for creep in rocks he derived an expression relating equilibrium activity to stress at a crystal face

a 1 - a 2 = e RT in (ai/a2)/M

stress. at 2 crystal faces; where a 1 and a 2 are values of are activity of solute at the crystal feces under stress a1, a2 C71 and CT 2 respectively; T = temperature °K, e = density, and M = mole volume of the solid., From the above equation Goranson inferred that, for the system gypsum-water at 24°C, a solubility increase of 3 times per 1000 bars would occur at stressed crystal faces. From this result it was suggested that supersaturation at stressed faces in natural rocks would cause diffusion along grain boundaries, and recrystallisation on free faces acting as nuclei. Goranson then assumed that the creep rate (a) was proportional to the change in activity of the solute with respect to stress, and derived an equation

1 - C72 K ln( /v) where K and v are physical constants. A logarithmic stress- strain rate relation of the same type was found by Griggs (1940) to be consistent with experimental results he obtained on water saturated gypsum under conditions of differential stress. Goranson considered that point contacts between grains would be 'ironed'out' by solution and that preferential solution would occur at crystal faces having the highest free • 22 energy. Goranson further suggested that permanent deformation of a solid will only take place by the change of phase mechanism outlined above, provided that the confining pressure is sufficiently high; although intracrystalline flow may also contribute to the overall deformation.

2.2 EARLY FIELD OBSERVATIONS CONCERNING THE ROLE OF PRESSURE SOLUTION IN DEFORMED ROCKS.

Henry Sorby, who made great advances within so many fields of geological research, was the first to describe, and recognise, the significance of pressure solution phenomena in rocks. No advances comparable to those achieved by Sorby have subsequently been made in understanding the mechanism of pressure solution, since his detailed field descriptions and inspired interpretation of these features. Sorby was interested in the nature and formation of cleavage structures in deformed rocks, and initially favoured a purely mechanical rotation of grains, until they were perpendicular to the maximum pressure, to explain the shape texture of Cleaved rocks. He stated that cleavage was due neither to chemical action nor to recrystallisation, and that there was no preferred crystallographic orientation (Sorby 1853, 1856a, b). However, Sorby had observed features• of mineral migration, for he wrote "in slates quartz seems to have either remained disseminated amongst the mica or to. have been removed and deposited in other situations, for quartz veins are of very general occur rence where this has occurred" (Sorby, 1858). • 23 This observation concerning volume change during deformation was to contrast with the later contention of Becker (1904) that no volume loss occurs in the formation of cleavage. The presence or absence of volume loss during deformation has recently been shown to be one of the single most important factors which determine the geometry of structures (Ramsay and Wood,.1973). One year after Thompson's discovery of stress effects on crystal solubility Sorby realised the significance of this principle in geological processes. He explained the shape change of crystals in cleavage as being due to pressure solution, and wrote poignantly "mechanical force has been resolved into chemical action" (Sorby, 1863a). Sorby described several features of deformation such as slickensides, veining, stylolites, and impressed pebbles with associated insoluble residues; attributing their formation to pressure solution of certain mineral species relative to others, with diffusional transfer in liquid. It was one of his conten- tions that the role of liquid in the deformation and metamorphism of rocks was"to greatly facilitate" the transfer of mineral constituents (Sorby 1865, 1879, 1880). One of the first estimates of strain in deformed rocks was made by Sorby, who determined the shape change of oolites deformed by pressure solution, and noted that material removed from certain faces now formed pore space filling (Sorby, 1908). Sorby demonstrated that the solubility of a salt is linearly increased by differential stress (over a limited range of values), and qualified this result for the case of negative molal volume only. a 24 By loading salt aggregates with a glass rod Sorby simu- lated indented pebbles of the Swiss Nagelfluh, and obtained impressions with insoluble residues, as in the field examples, that were not due to plastic flow (Sorby, 1865). Describing folded limestone beds at Ilfracombe Sorby (1879) attributed the missing inverted limbs and striping around hinges to pressure solution, for he wrote "solution has taken place where there was great pressure and stress, and deposition where they were relatively less". He described how crinoid stems dissolve and mutually penetrate when in the direction of pressure, whilst the dissolved material has crystallised out on each side which is more protected; or, for the case of mica chists, as quartz overgrowths on original quartz grains. Sorby (1908) considered that a similar process was operatiVe in uncleaved limestones, due to superineumbent load, which assisted in producing a consolidated rock, and stylolites with insoluble residues. In a classic paper on the geology of the South Devon Coast T.G. Bonney (1884) noted the association of folding, cleavage, and veining. He described for the first time cleavage striping parallel to, and normal to bedding, suggesting a purely mechanical origin of the lens-shaped tectonic stripes; as had Sharpe in 1849. Harker (1885) was more specific in describing the mechanism of formation of tectonic lenses, and followed Gorby's interpretation of such features as being due to pressure solution. He wrote "where a mass of rock is subjected to unequal pressure, some of the constituents may be dissolved at points where the pressure is greatest and redeposited where it is least . . . . ; another possible explanation for the obliteration of stratifi- cation by pressure, and its replacement by lenticular and 25 elongated structures". "Individual crystals of a rock may thus change their form by molecular rearrangements without possessing in themselves the property of plasticity in any degree". In his treatise on 'Metamorphism of rocks and rock flowage' Van Hise (1904) states explicitly the role of Thompson's Principle in explaining the shape texture of quartz in schists. He considere&that when minerals are in a state of strain, solution is favoured in the Subcapilliary spaces, and shape change occurs through solution and recrystallisation (also Van Hise, 1898). Van Hise distinguishes two cases, of which the second one overcomes the limitations of Riecke's Principle, which was to prove such a stumbling block for some workers.

1. Particles of quartz under great stress are taken into solution, and thematerial transported to boundaries of less stressed grains.

2. Material dissolved from the more stressed part of a grain may be deposited on another part of the same grain.

The observations made in these classic papers at the end of the nineteenth century, and the principles which were evolved from the observations have subsequently been restated and- elaborated by several authors (Wright, 1906; Adams and Coker, 1910; Backlund, 1918; Knopf, 1933; Brehier, 1951). Sorby's pressure solution mechanism for deformation was extended by Buerger and Washken (1947) to include the case of stress directed grain boundary diffusion in the solid state. There was little progress in research on the detailed mechanism of pressure solution, or its relation to structural s 26

features, during the first half of the twentieth century: a situation which parallels the lack of advance in research on slaty cleavage (Siddans, 1972). Interest in pressure solution as a deformation mechanism was revived by Voll (1960) and Ramsay (1967), who made detailed observations on the structural petrology of pressure solution features. The development of ideas concerning the role of pressure solution in certain specific structures such as crenulations and pressure shadows is considered in the following section.

2.3 THE ROLE OF PRESSURE SOLUTION IN THE FORMATION OP STRUCTURES.

In this section of the development of ideas concerning the role of pressure solution in the formation of structures is discussed. Structural features are considered under three headings -

1. Pressure shadows. 2. Pressure solution in crenulations and folds. 3. Pressure solution in diagenetic environments.

1. PRESSURE SHADOWS

Pressure shadows provide a well defined example to illustrate the effects of pressure solution as a deformation mechanism, and as the cause of mineral segregation. In this section a short description is given of pressure shadows-1 and how they form. The historical development of ideas concerning the significance of pressure shadows-is traced. This is followed by a more detailed discussion of textural • _2 7

features in shadow zones. An account is then given of the manner in which such textural features have been used to define the microscopic strain history of deformed rocks. Finally, mention is made of the interrelationship of the stress field and chemical distribution around pressure shadows.

Introduction

Many rocks contain rigid particles which are large in relation to the grain size of the rock matrix.' These rigid particles, or inclusions, may form a ductility contrast with the enclosing matrix due to the difference in material properties. Thus rigid inclusions are generally considered. to act as 'stress raisers' in a deforming matrix (Hossain, 1973; Stromgard, 1973). An area of high stress is generated in the matrix adjacent to boundaries of the particle which are normal to the maximum principal stress. Conversely, there are areas of low stress, or pressure shadows, adjacent to boundaries which are normal to the minimum stress. This variation in the stress field around rigid inclusions generates a stress gradient in the matrix. The stress gradient acts as a driving force, or potential, to cause migration of soluble mineral species into the shadow area; resulting in textural and mineralogical variations adjacent to the stress raiser. Rigid particles- which are commonly observed to have associated pressure shadows are pyrite and magnetite crystals, fossil fragments, large detrital grains; sedimentary nodules; isolated fold hinges and boudinaged layering; pebbles; and porphyroblasts. 28

Secondary minerals which have grown in pressure shadows may be closely. related to the mineralogy of the host rock. In the general order of abundance these sec'ondary minerals are quartz, carbonates, chlorite, muscovite, and feldspar (Pabst, 1931; figry, 1969). Harker (1889) described quartz pressure shadows around pyrite cubes from the North Wales slate belt, and carbonate pressure shadaws around bellemnites from the Ardennes (Harker, 1939). He concluded that the pressure' shadows were due to non- uniform pressure around the pyrite crystals; and that the quartz iifilling was concurrent with deformation, occupying space as the matrix pulled away from the pyrites. There are many reports in the literature of various types of pressure shadows, and the use of these structures as an indicator of stretching direction in regional tectonic studies (Knopf, 1929; Mugge, 1930; Pabst, 1931; Bain, 1933; Fairburn, 1949, 1950; Kopp, 1955; Thome, 1955, Frankel, 1957; Ramsay, 1967). Much of this work has been summarised by Spry (1969, p. 240-247). Thome (1955) has given a useful qualitative description of the stress distribution around rigid objects in a deforming matrix; and the resultant formation of pressure shadows. In addition, Thome has given a comprehensive account of the relationship between deformation, mineral migration and pressure shadows in structures of the Rheinisches Schiefegebirge.

Classification of pressure shadows

Str8mgard (1973) distinguishes two types of pressure sha- t dow on a genetic basis. 29 Secretionary type, resulting from metasomatic infilling of a progressively opening fracture or void at the inclusion- matrix interface.

2. Replacement type, caused by metasomatic replacement of the matrix in zones of low pressure adjacent to the inclusions.

Spry (1969, p. 240-246) gives a classification of pressure shadows based on the texture of secondary minerals, and their relationship to the matrix foliation. The shape of pressure shadows is extremely variable. They may develop either symmetrically or asymmetrically about the rigid inclusion. Bain (1933) illustrates the growth of shadows on four faces' of a euhedral pyrite crystal.

Textures of secondary minerals in pressure shadows.

Pressure shadows are generally composed of minerals which have grown with a fibrous crystal habit. Fibrous minerals may exhibit various degrees of curvature. These fibrous minerals nucleate at the crystal-matrix interface where they continue to grow as the fibre is progressively drawn out together with the deforming matrix.

Crystallographic and shape orientation of secondary minerals:

There has been a great deal of contention concerning 1) the significance of the host crystal structure in controlling the shape fabric of secondary minerals; 2) the mechanisms which determine the crystallographic orientation of secondary minerals. Several authors have commented on the relation between fibre axis and the crystal face of the rigid particle, on which . nucleation of the fibre took place (Pabst, 1931; Bain, 1933; 30

Frankel, 1957; Spry, 1969; Durney, 1972b). Thus, fibres of quartz generally grow normal to (111) or (100) of magnetite or pyrite; and the basal plane (001) of chlorite or muscovite is parallel to (111) or (100) of magnetite or pyrite. These geometrical relations cause the formation of compound pressure shadows, with segments of different mineral species, each having a different fibre orientation (Spry, 1969). Magge (1930) suggested that chlorite nucleated on magnetite with an epitaxial relationship, due to structural similarity of (001) chlorite and the (111)lattice of magnetite. Spry (1969, p. 240) does not consider that chlorite nucleation is controlled apitaxially, because chlorite may also form parallel to (100) of magnetite. According to Mugge (1928) quartz nucleates with a random orientation at the matrix boundary. Mugge proposed that under certain conditions a preferred orientation of crystallographic axes may arise as a result of the growth kinetics of different crystal faces. Phillips (1974) has suggested that a preferred orientation of optic axes' in fibrous minerals may be attributed to crystal growth under tensile strain; and that this mechanism is explained by Kamb's (1961) theory of crystallisation under con-

• ditions of nonhydrostatic stress. Hydrothermal quartz grows most quickly in the c-axis direction. Hence support has been lent to the theory of MUgge (1928) by the findings of Pabst (1931) and White and Wilson (1975), who report a c-axis preferred orientation of quartz fibres in some pressure shadows. Several authors have reported that no strong preferred orientation of optic axes occurs in quartz pressure shadows (Pabst, 1931; Mugge, 1928, 1930; Frankel, 1957; Spry, 1969).

Phyllosilicates are generally oriented with (001) parallel to the stretching direction in the host rock.

Syntaxial, pressure shadows on detrital grains. As was related in an earlier section, many authors invoked a pressure solution mechanism to explain the preferred shape orientation of grains in deformed rocks. Williams (1972, p. 22) has described mica beards growing in the extension direction on detrital quartz grains, in cleav- ed_greywackes. Misch (1969, 1970) postulated that the elonga- tion of certain syntectonic amphiboles may, be due to microboudinage of the grains, together with growth in shadow areas. Durney (1972a) commented that the precipitation of numerous small syntaxial pressure shadows on rock grains probably ac- • counts for the majority of removal by pressure solution on stressed: faces, and hence for the shape change of grains in deformed rocks.

Deformation of rigid inclusions.

The rigid inclusions around which pressure shadows form often show evidence of deformation. Fracturing and abrasion • of pyrite crystals during deformation has been noted by Marker (1889) and Frankel (1957). In addition, inclusions may be boudinaged, veined, recrystallised, or exhibit features- of internal distortion such as strain shadowing and deformation bands (Spry, 1969). Stromgard (1973) has- discussed the stress condition which is necessary to produce a tensile opening in the shadow area of a rigid inclusion. The ratio of maximum to minimum principal stress must be 3 : 1. • 32

Pressure shadows and deformation paths.

Rocks deform with a history of successive strain increments which may be coaxial or non-coaxial (Ramsay, 1967). Elliott (1972) has termed this strain history the deformation path. Fibrous- minerals in pressure shadows may record the orientation of successive incremental strains. Hence, in some instances these structures can be used to reconstruct the microscopic deformation paths within rocks, from which the strain history of mesoscopic structures may be deduced. Since the original observations of Sorby (1863) and Harker (1889) it has generally been accepted that pressure shadows grow in the stretching direction of the host rock (Mtigge, 1930). Fairburn (1950) related the microscopic sense of shear, as deduced from the deformation history of pyrites with associated pressure shadows, to the sense of movement in a shear zone. Badoux (1963) computed the finite strain ellipses for two stages of deformation represented by fibrous mineral growth, first of quartz and then of calcite, in the boudin necks of deformed belemnites from the Morcles Nappe. There are three main theories which account for the orientation of fibrous minerals in pressure shadows -

1. According to Fairburn (1950) and Spry (1969) fibre orientation is controlled by rotation of the rigid crystal during deformation. The direction of the first fibre growth increment remains constant with respect to the original orientation of the rigid crystal, whilst the fibre nucleates continuously perpendicular to the rigid crystal as it rotates. 33

2. Hills (1963, p. 130), Durney (1972b), Elliott (1972) and Wickham (1973) have postulated that fibrous minerals in pressure shadows have grown in the direction of relative displacement of the host rock.

3. In contrast to the above two theories, White and Wilson (1975) contend, on the evidence of dislocation structures, that some curved fibres- are due to deformation, during or after growth, In addition, they suggest that certain types of 'indistinct' fibre are not: growthfeatures, but are tectonically induced subgrain structures.

Durney (1972b) and Elliott (1972) present evidence from which it may be inferred that in some cases fibrous minerals do record the incremental strain history of rocks. For instance, there is a close correspondence between strain measurements on deformed oolites, and the strain computed from fibrous minerals in pressure shadows from the same rock (Durney, 1972b). Elliott (1972) has pointed out that the relative frequency of straight as against curved. fibrous minerals in pressure shadows may be taken as evidence that coaxial deformation paths are not uncommon, e.g. Ramsay (1967, p. 181); Spry (1969); Burger and others (1970). Examples of pressure shadows with curved fibres, which have been interpreted as recording non-• coaxial incremental strain, are illustrated by Langheinrich (1964); Zwart and Oele (1966); Ramsay (1967, p. 181); Spry (1969), and Durney (1972b). Clearly one may contend that fibre orientation in pressure shadows may be controlled by several mechanisms, which in turn depend on the teatonic and metamorphic environment. From a study of nodules in slate, Glover (1973) has recognised that pressure shadows which are associated with such 34

nodules, may play an active role in determininc the rotation of rigid particles in a deforming matrix. This feature has important implications in relation to the use of particles rotated during deformation as a measure of shear strain.

Chemical variations around pressure shadows.

Berglund and Ekstrom (1974) have analytically determined the FeS content of sphalerite in relation to the theoretical stress distribution around a boudin neck. They concluded that the effect of variations in the fugacity of S2 are almost as large as the effect of pressure on the composition of sphalerite, which corresponds to less than 1 mole "t. FeS per 0.5 Kb.

• 2. PRESSURE SOLUTION IN CRENULATIONS AND FOLDS

Many cleavage structures which form in low grade metamorphic rocks are heterogenous. In particular, structures associated with crenulation cleavage and folding are characterised by domaii which are defined by differentiated fabric and varied mineral composition (Williams, 1972). Several authors have suggested that these variations of composition and texture, in buckled • mineral fabrics, may be attributed to pressure solution activity, driven in response to stress gradients during folding (Mosebach, 1951; Thome, 1955; Voll, 1960; Rickard, 1961; Dewey, 1965; Nicholson, 1966; Roberts-, 1966; Ramsay, 1967; Talbot and Hobbs, 1968; and others). Sorby (1879) noted features of mineral migration around folds, from Ilfracombe, N. Devon. He considered that calcite had been dissolved from areas of high stress on fold limbs; and 35 precipitated in regions protected from stress, behind fold hinges. In addition, Sorby (1908) described mineral differentiation on a microscopic scale in the 'slip surfaces', or limbs, of microfolds in slate. Coarser grained, uncleaved, material remains between the slip surfaces. Thus, mineral segregation in folds has long been explained in terms of the pressure solution mechanism of deformation.

Mineral differentiation in buckled microfabrics.

Helm (1969) noted the effect of mineral fabric on differentiation in microfolds. In buckled quartz-mica fabrics, quartz migrates from fold limbs to hinge regions, giving a concentration of muscovite and chlorite on the limbs. In pure mica fabrics, chlorite is localised in hinges, with resultant enrichment of muscovite on limbs. Helm attributes this mineral differentiation to ionic diffusion over limited distances. microfold controlled mechanism of metamorphic differentiation, as a result of pressure gradients in folds, has also been proposed by Misch (1969, 1970). Misch gives petrographic descriptions of mineral segregation in folds similar to those of Helm (1969). Segregation of chlorite into hinges, with concomitant mica concentration in limb regions: and, in a folded crossite schist, residual crossite enrighment in limbs, due to movement of mica into hinge zones. A comprehensive account of the formation of metamorphic layering and cleavage structures, in relation to mineral migration under stress, has been given by Williams (1972). He describes the progressive development of crenulations, from weak buckles, into microfolds with differentiated limb and • 36

hinge zones, which he denotes 'A' and 'B' domains respectively. 'A' domains are characterised by the micragion of quartz out of fold limbs, with a resultant passive concentration of layer silicates and iron oxides. The layer silicates are rotated into a high degree of preferred orientation as quartz is progressively removed. Those quartz grains which remain in 'A' domains have a strong shape orientation due to the formation of overgrowths parallel to the domain boundaries. Quartz migrates from 'A' domains into 'B' domains. The latter are characterised by a low concentration of layer silicates, and by equant quartz grains. Williams (1972) attributes the offset of bedding across fold limbs to removal of quartz in 'A' domains. He has postulated that structures with elongated lenses- of quartz, bounded by anastomosing 'A' domains rich in layer silicates, may be a transitional stage from crenulation differentiation to slaty cleavage.

The relationship between stress distribution and mineral migration in folds.

As noted above, mineral differentiation is commonly observed to have taken place in folded layers and buckled microfabrics. It is widely accepted that this segregation has resulted from pressure solution, which is driven by stress gradients during folding. Gresens (1966) proposed that stress variations, which are set up during the folding of rocks, might act as a potential gradient, to cause the migration of minerals from areas of high stress to areas of low stress. Cosgrove (1972) has applied the same reasoning to explain the features of quartz and mica • 37

segregation in crenulated mica schists, that were described by Helm (1966) and Misch(1969, 1970). A finite element approach has been adopted by Stephansson (1973), to make numerical estimates of the stress variations around buckled multilayers, in relation to mineral migration.

Analytical chemistry of mineral segregation in folds.

There has been little quantitative work done on the distribution of mineral species and chemical elements in fold structures. The information that is available comes from studies of economic mineral deposits located in tectonic structures. McDonald (1970) has measured a 40% increase in the content of galena in the hinge zone of a fold, relative to the percentage of galena in undeformed layers. Migration of galena has taken place from fold limbs into the hinge area. McDonald considers that the mineral migration is a solid state process, and that enrichment of galena in the hinge zone has not resulted from the introduction of material in hydrothermal solutions.

3. PRESSURE SOLUTION IN DIAGENETIG ENVIRONMENTS

Introduction

Pressure solution is considered to play an important role in the diagenesis of some sediments. The mutual penetration of pebbles, in elastic rocks, and the widespread development of stylolites in carbonate and other fine-grained sediments, provide convincing evidence from which it may be inferred that the selective removal of soluble minerals under conditions of differential stress is a characteristic feature of diagenetic environments. 38

There is an extensive literature on the role of pressure solution in the diagenesis of sediments. The object of this section is to review the more significant contributions in this field of sedimentary geology, and to present the conclusions in relation to the role of- pressure solution in rock deformation. The four main questions which arise from a consideration of pressure solution activity in sediments are as follows -

1. What types of structure are present in sediments that have undergone diagenesis by pressure-solution?

2. What are the factors which control pressure solution activity in diagenetic environment?

How do these structures, and the factors which control their development, relate to structures in deformed rocks which have resulted from pressure solution?

4. Is there a zone in the crust which separates diagenesis from tectonic activity; or may these two processes, each resulting in part, from pressure solution, be considered as end points- of a continuously varying state of stress in the upper crust?

Pressure solution structures in sediments which have undergone diagenesis are considered under two main headings in this section. They are -

1. Solution surfaces in detrital rocks, and 2. Stylolites.

Pressure solution surfaces in detrital rocks.

As we have seen, Sorby (1863) first explained the origin of impressions in pebbles by the mechanism of pressure solution. He attributed the chemical removal, at point contacts in the 39

Swiss Nagelfluh conglomerates, to increased solubility of the material in the "direction of mechanical force". Sorby considered that the dark film present along many pebble contacts was an accumulation of insoluble material, due to selective removal of the more soluble components under the influence of normal stress. Reade (1891) studied impressions in Permian conglomerates of the English Midlands; and followed Sorby's interpretation of these features as being due to solution at point contacts. An alternative explanation for the impressions was given by Gresley (1895), who proposed an hypothesis of mechanical indentation. In a cogently argued reply to Gresley's paper, Reade (1895) pointed out the field evidence for solution at point contacts, with precipitation and associated veining in the surrounding matrix. Pressure solution phenomena in detrital rocks have been described and classified in detail by Trurnit (1968), with particular reference to impressed conglomerates. Trurnit gives a comprehensive review of the literature on this subject. The most common features of pressure solution are contacts between impressed pebbles, and between impressed grains. Trurnit (1968) defines two partners and a contact surface. The partners- may have a) equivalent or different relative pressure solubility in the direction of differential stress, b) equivalent or different radii of curvature at the contact. These factors, taken in combination, lead to 14 types of contact surface; which may be plain or curved, and smooth or sutured. The surface geometry is influenced by a) and b), but surface texture is influenced by b) alone. • 44

H8fer (1879) postulated that if two pebbles are in contact with one another the impression will be made by that one which has the smaller radius of curvature at the contact. Trurnit (1968) has restated this principle as follows - for partners of equivalent relative pressure solubulity, the stronger curved partner indents the less curved partner along a sutured pressure solution contact. At an advanced stage of solution, if much residue has been accumulated, the residue behaves as au insoluble partner. Morawietz (1958) recognised that only one partner is dissolved along a smooth contact, but that both partners are dissolved along a stylotitic contact. Trurnit (1968) suggests that grain size of the partners, and the solution residue which accumulates along the contact, play an important role in press-

• ure solution effects. According to Morawietz (1958), material adjacent to the contacts of impressed pebbles remains unaltered. However Trurnit (1968) contends that in pebbles of inhomogenous composition, a contact zone is present, with the intensity of pressure solution decreasing towards the interior of the pebbles. Microslickensidea on the surfaces of impressions are continuous within the zone of pressure solution activity, inside such pebbles. The columns of stylotites, at sutured contacts, are not necessarily oriented perpendicular to their respective pressure solution surfaces; but they are always parallel to the direction of maximum principal stress. Trurnit (1968) has defined a series of terms to describe pressure solution surfaces increasingly oblique to the principal stress. These are as follows - stylolites, oblique stylolites, slickolites, and slicktnaides. * 41

Asymmetrical impressions, having steep rims which grade into striated surfaces, due to the relative sliding between pebbles, have been described by Trurnit (1968). According to Nitecki (1962), such striations are often fibre-like structures which develop along discrete movement planes, due to pressure solution and recrystallisation. The association of fracturing with pressure solution activity in the deformed Spanish Buntsandstein conglomerate, has been noted by Trurnit (1968). Behrens and Wurster (1972) have made estimates of bulk strain in the deformed Bavarian molasse from measurement of the volume loss vector in pebble impressions. There are no criteria by which pressure solution contacts between pebbles, formed in diagenetic and non-metamorphic tectonic environments may be distinguished with certainty. In the latter situation there may be a greater density of pressure solution surfaces-, and a higher frequency of fracturing and tectonic veining associated with pressure solution.

The nature of stylolites as a pressure solution surface.

Stylolites have claimed: the attention of sedimentologists largely because of their role in the diagenesis of sediments. Pettijohn (1957, p. 213) has defined stylolite as "a surer face marked by interlocking or mutual interpenetration of the two sides. The teeth-like projections of one side fit into sockets of like dimensions on the other". Stylolites are one of a series of pressure solution surfaces defined by Trurnit (1968), which are generally characterised by the presence of an insoluble residue. Park and Schott (1968) have classified stylolites on the basis of a) morphology, and b) relationship to bedding. 42 Fuchs (1894) explained the origin of stylolites by pressure solution. He distinguished two types, diagenetic stylolites which formed in response to overburden pressure; and tectonic stylolites, which resulted from lateral tectonic pressure, according to whether they are parallel or oblique to bedding respectively. The pressure solution theory for the origin of stylolites- has never been seriously challenged except by Schaub (1939), who proposed a 'contraction-pressure' mechanism of plastic extrusion. The following characteristic features of stylolites are accounted for by the pressure solution theory (Heald, 1955, 1959; Park and Schott, 1968;- Trurnit, 1968).

1. The apices of projections are frequently composed of minerals having low pressure solubility relative to the rock matrix. Examples of this feature are cherts and fossils in carbonates (Bastin, 1933; Mossop, 1970).

2. The seams represent insoluble material once disseminated throughout the rock, and now concentrated by removal of pressure-soluble minerals along the stylolite surface.

3. Ooliths, fossils, mineral grains, pebbles, etc. are frequently observed to be truncated across the stylolite surface.

4. Mineral grains are generally neither fractured nor deformed at the apices of stylolites, but progressively dissolved in one direction.

5. Material which is removed along stylolite surfaces is precipitated in pore spaces, and as overgrowths on grains in adjacent rock. 43

Stylolites are especially well developed in limestone and dolomites (Bastin, 1951; Park and Schott, 1968; Mossop, 1970), but have also been observed in marbles (Schaub, 1953); sandstones (Heald, 1955, 1959); orthoquartzites (Conybeare, 1949); tectonic breccias (Herbert and Young, 1957); conglomerates (Bastin,1940; Trurnit, 1968); porphyries (Bloss, 1954); gypsum (Stockdale, 1930) and in pegmatites (Bailly, -1954).

Nature of the residual seam.

As stylolites propagate through the rock mass, and pressure soluble minerals are removed, disseminated insoluble material is concentrated along the surface of discontinuity to form a residual seam. The thickness of seams varies from a few microns to several centimetres (Mossop, 1970) and is directly proportional to the volume of rock which the stylolite has swept through. In carbonates, seams are generally composed of clay minerals, carbonaceous and bituminous material, pyrite, zircon and other insoluble minerals (Trurnit, 1968; Park and Schott, 1968): Bloss (1954) has noted an accumulation of iron oxides along stylolites in rhyolite porphyry. Platy and acicular minerals have a preferred orientation parallel to the stylolite surface (Heald, 1955; Dunnington, 1954, 1967; Trurnit, 1968). Soluble detrital grains within the seam may have a shape fabric due to pressure solution acting preferentially in the direction of maximum stress (Heald, 1955). Pressure shadows have been observed on insoluble grains (Park and Schott, 1968). Brown (1959) measured the c-axes of authigenic quartz grains in Carboniferous rocks of North Wales. c-axes had a 44 random distribution in the matrix, but lore found to have a preferred orientation parallel to the limbs of stylolites, and normal to this direction at the apices. Whole beds composed of relatively soluble grains may go into solution without adjacent beds showing solution effects. The only remnant of this process is a thin seam of insoluble residue, which requires a different interpretation from primary sedimentary features (Bathurst, 1958; Barrett, 1964; Mossop, 1970). Manten (1966) and Mossop (1970) both report that stylolites tend to deamplify as the residual seams become very thick.

Factors which govern the formation of stylolites.

Factors which are considered to be important in the formation and development of stylolites have been discussed by Heald, 1955; Trurnit, 1968; and Park and Schott, 1968. The most significant factors may be summarised as follows -

1. Depth of burial. The magnitude of stress concentration at grain contacts, is related to depth of burial. Moraweitz (1958) reports the development of impressions in limestone pebbles under 30-40 metres of overburden. The onset of pressure solution contacts in sandstones has been variously estimated at 800 m., 1000 m. and 1050 m. by Taylor (1950, p. 715), Bushinskiy (1961), and Trurnit (1968) respectively. Decrease of porosity, and intensity of stylolitisation have been related to depth of burial by Pittman (1972) and Siever (1959) respectively. Overburden would appear to be the single most important factor which dictates the extent to which stylolitisation 45 develops. However, Heald (1955) presents evidence from which it may be inferred that pressure solution does not necessarily require a high level of stress.

2, The effect of pore fluid pressure on stress at grain contacts. High pore fluid pressure lowers the effective normal stress at grain boundaries. Conversely, pressure solution is promoted by high grain boundary stress under conditions of low pore fluid pressure.

3. Shape and size distribution of grains. Areas of fine or angular grains are preferentially removed by pressure solution.

4. Relative pressure-solubility of minerals. Grains of high pressure-solubility are invariably impressed by those of low pressure-solubility; or preferentially removed along stylolites.

5. Rate of precipitation and permeability. The kinetics of precipitation may be the rate controlling step governing the development of stylolites. Alternatively, a rock of high permeability will allow the diffusion of solute, in a supersaturated solution, away from the site of pressure solution.

6. Proportion of argillaceous and insoluble material. It has been observed that stylolitisation occurs preferentially in sandstones and carbonates which have a high primary content of clay minerals, carbonaceous material, and other impurities (Heald, 1955, 1959; Thompson, 1959; Lerbekmo and Platt, 1962; Pittman, 1964; Oldershaw and Scoff in, 1967; Trurnit, 1968; Mossops 1970). 46

The explanation for this phenomenon is not fully understood, but it probably relates to interfacial energy, and to ease of diffusion, particularly within the clay lattice. Mossop (1970) has shown that argillaceous lagoonal carbonates have a higher frequency of stylolites in comparison to clean reef limestones.

7. Texture. The importance of texture to the development of stylolites in sediments and igneous rocks has been emphasised by Whisonat (1970) and Bloss (1954) respectively. Both Heald (1955) and Park and Schott (1968) distinguish between grain boundary stylolites, and mesoscopic stylolites in granular aggregates as described above. Different factors may affect the development of these two classes of stylolite. A series of minerals with declining relative pressure solubility is listed by Trurnit (1968). This series has been compiled from the work of Heald (1955), Moraweitz (1958), Bushinskiy (1961), Schidlowski and Trurnit (1966), and Trurnit (1967b). The relative pressure solubility of minerals is ascertained from a study of grain boundary contacts in pressure solution planes. The series is as follows - I) halite and potassium salts, 2) calcite, 3) dolomite, 4) anhydrite, 5) gypsum, 6) amphibole and pyroxene, 7) chert, 8) quartzite, 9) quartz, glauconite, rutile, and haematite, 10) feldspars and cassi- terite, 11) mica and clay minerals, 12) arsenopyrite, 13) tourmaline and sphene, 14) pyrite, 15) zircon, 16) chromite. 47

Experimental investigations on the diagenesis of sediments.

Fruth and others (1966) have studied the compaction of carbonate sediments experimentally. Interpenetrating grain boundary contacts developed which were not due to pressure solution. Experiments on the diagenesis of quartz sediments have been reported by Renton (1969), and Renton and others (1969). Solution pits in quartz were obtained at contacts with zircon crystals. The conditions of temperature and differential stress were in the range of 30000 and 2.500 psi respectively. Without evidence from transmission electron microscopy it is not possible to say unequivocally whether the pits are due to pressure solution or to plastic deformation of quartz.

Tectonic stylolites.

In 1894 Fuchs proposed that stylolites develop by a pressure solution mechanism, inxesponse to sedimentary or tectonic loading. Since that time there have been many reports of stylolites which are attributed to tectonic activity (Wagner, 1913; Sloss, 1948; Blake, 1949; Rigby, 1953; Heald, 1953; Wagner, 1967; Arthaud and Mattaeur, 1969; Glover, 1969). Lindstrom (1962) ascribed cleavage structures in folded carbonates of the French Jura, to the progressive development of tectonic stylolites. He described the geometrical relationship between folds, stylolites, and associated tectonic veining. Chouckrane (1969) and Jarouwaski (1969) have noted that tectonic stylolites have the same orientation characteristics as fracture cleavage. 48

Heald (1953) commented that tectonic stylolites in sediments and low grade metamorphic rock have many features in common with stylolites formed in diagenetic environments. In.a study of the Helvetic Alps Durney (1972b) also recognised this similarity between diagenetic pressure solution and certain types of tectonic cleavage. Durney noted that fossils, ooliths, and other rigid objects appeared to have initiated planes of pressure solution, which form spaced, non penetrative cleavage structures. This coarse tectonic cleavage described by Durney (1972a, b) and Williams (1972) has some of the petrographic characteristics of anastomosing diagenetic stylolite seams.

A comparison of tectonic and diagenetic pressure solution surtces

From the above descriptions it is now possible to make a comparison of structures formed by pressure solution in tectonic and diagenetic environments. The following characteristic features are common to both seta of structures -

1. Surfaces- have an orientation which is statistically sub- perpendicular to the inferred direction of maximum principal stress. 2. Residual seams have the same origin and significance, and have similar composition. 3. Mica and clay minerals have a preferred orientation parallel to seams. 4. Relict soluble minerals have a shape fabric in the plane of seams. 5. Insoluble grains may develop pressure shadows. 6. A similar order in the relative pressure solubility of minerals is observed. , 49

Material removed along pressure solution seams may form overgrowths in the adjacent matrix, or veins. 8. Surfaces of pressure solution impart to the rock a non penetrative fissility which has a preferred planar orientation. 9. Pressure solution surfaces have a higher frequence in rocks with clay- or phyllosilicate minerals. 10. A preferred crystallographic orientation of grains may result from the selective development of grains which have the axis of fastest growth in a given direction.

One may infer from evidence presented in the above discussion that there are few certain criteria by which pressure solution surfaces, formed in diagenetic and high level tectonic environment respectively, may be distinguished. Further, pressure solution activity in diagenetic and high level tectonic environments would not appear to be confined to separate levels in the crust.

Pressure solution and slaty cleavage.

The role of pressure solution in the formation of non- penetrative and compositionally heterogenous cleavage has been discussed in some detail. Slates are characterised by a penetrative cleavage and homogenous mineral composition. Because of the fine-grained nature of slates it is difficult to ascertain the deformation mechanisms which have been operative in producing tectonic cleavage. However, Von Plessman (1964, 1965) and Siddons (1971, 1972) have argued that pressure solution plays an important role in the formation of slaty cleavage. ' S0

2.4 THEORETICAL TREATMENTS OF PRESSURE SOLUTION.

Thermodynamic theories of non-hydrostatically stressed solids have been reviewed by Kamb (1961), McLellan (1970), and Patterson (1973). The general result of reversible thermodynamics applied to stressed crystal aggregates may be summarised by the normal stress theorem of McLellan, which states that 'the chemical potential of a stressed solid in solution is proportional to the normal stress across the interface'. Elliott (1973) has summarised some of the metallurgical literature relatins to grain boundary diffusion, and applied the results to geological systems. However, his assumption that the activation volume of diffusion is negligible in de- riving a stress exponent of unity may not be valid. Elliott has noted on the basis of field observations and theoretical predictions that an intercrystalline diffusional mechanism will give way to intracrystalline mechanisms in the deformation of rocks at high grades of metamorphism. 51

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• •

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TAYLOR, J.M. 1950. Pore space reduction in sandstones. Bull. Am. Assoc. Petrol. Geologists, v. 34, 701-716. THOME, von K.N. 1955. Die Tektonische Frdgung des Pennsattels und Seiner Umgebung. Geol. Rund., 266-305. THOMPSON, A.F. 1959. Pressure solution and porosity. Soc. Econ. Palaeontologists Mineralogiste, Spec. Publ., 7, 92-110. THOMPSON, J. 1862a. Theoretical consideration on the effect of pressure in lowering the freezing point of water. Trans. Roy. Soc. Edinburgh, v. 16, 575-580. THOMPSON, J. 1862b. On crystallisation and liquifaction as influenced by stresses tending to change the form in crystals. Phil. Mag. (Ser. 4), v. 24, 395-401. THOMPSON, J. 1862c. On crystallisation and liquifaction, as influenced by stresses tending to change of form in crystals. Proc. Roy. Soc. Lond., v. 11, 473-480. TRURNIT'l P. 1967. Morphologie und Entstehung diagenetisches Druck-lbsungserscheinungen. Geol. Mitt. Aachen, v. 7, 173-204. TRURNIT', P. 1968. Pressure solution phenomena in detrital rocks. Sedimentary Geology, v. 2, 89-114. VAN RISE, C.R. 1898. Metamorphism of rocks and rock flowage. Geol. Soc. America Bull., v. 9, 269-328. VAN RISE, C.R. 1904. Treatise on metamorphism. U.S. Geol. Surv. Mon., 46, 693. VOLL, G. 1960. New work on petrofabrics. L'pool Manchr. Geol. Soc. Jour., v. 2, 503-567. VON PLESSMAN, W. 1964. Gesteins18sung, ein Hauptfaktor beim Schieferungsprozess. Geol. Mitt. Aachen, v. 4, 69-82. VON FLESSMAN, W. 1965. L8sung, Verformung, Transport und GefUge (Beitrage zur Gesteinsverformung im nord8stlichen Rheinischen Schiefergebirge. Z. Dtsch. Geol. Ges., v. 115, 650-663. WAGNER, G. 1913. Stylolithen und Drucksuturen. Geol. Paldontol. Abhandel., Neue Folge, v. 11 (2), 1-30. WAGNER, G.H. 1964. Kleintektonische Untersuchungen im Gebiet des N8rdlinger Rieses. Geol. Jahrb., v. 81, 519-600. WRITE, S.H. and C.J.L. WILSON. 1975. Microstructure of some quartz pressure shadows. (In press). WICKHAM, J.S. 1975. An estimate of strain increments in a naturally deformed carbonate rock. Am. Jour. Sci., v. 273, 23-47. 60

WILLIAMS, P.F. 1972. Development of metamorphic layering and cleavage in low grade metamorphic rocks at Bermagui, Australia. Am. Jour. Sci., v. 272, 1-47. WILLIAMSON, E.D. 1917. On the effect of strain on heterogenous equilibrium. Ihysics Reviews, v. 10, 275-283. WRIGHT, F.E. 1906. Schistosity by recrystallisation. Am. Jour. Sci. (4th Ser.), v. 22, 224-230. WRIGHT, F.E. and J.C. HOSTETTER. 1917. The thermodynamic re/brsibility of the equilibrium relations between a strained solid and its liquid. Jour. Wash. Acad. Sci., v. 7, 405-417. ZWART, H.J. and J.A. OELE. 1966. Rotated magnetite crystals from the Rocroi Massif (Ardennes). Geologie en Miinbouw, v. 45e, 70-74.

Theoretical aspects

RLT,IOTTI D. 1973. Diffusion flow laws in metamorphic rocks. Geol. Soc. America Bull., v. 84, 2645-2664. KAMB, W.B. 1961. The thermodynamic theory of nonhydrostati- cally stressed solids. Jour. Geophys. Research, v. 66, 259-271. MeLELLAN, A.G. 1970. Non-hydrostatic thermodynamics of chemical systems. Proc. Roy. Soc. London, Ser. A, 314, 443-445. PATTERSON, M.S. 1973. Nonhydrostatic thermodynamics and its geologic applications. Reviews of Geophysics and Space Physics, v. 11, 355-389. WEYL, P.K. 1959. Pressure solution and the force of crystallisation - a phenomenological theory. Jour. Geophys. Research, v. 64, 2001-2025.

Addenda COSGROVE, J.C. 1972. Ph.D. Thesis, London. DEWEY, J.F. 1965. Nature and origin of kinkbonds. Tectonophysics, v. 1, 459-494. HoFER, H. 1879. Die hohlen Gerolle and Geschiebe-Eindracke des Sattnitz-Konglomerates bei Klagenfurt. Tschermaks Mineral. Fetrog. Mitt., N.F. 2, 325-349. RIGBY, J.K. 1953. Some transverse stylolites. Jour. Sedimentary Petrology, v. 23, 265-271. 61

PREFACE TO CHAPTERS_ 3 AND 4.

It has long been recognised that pressure solution, the sOlution, diffusion, and precipitation of, rock forming minerals in response to stress fields is an important mode by which low grade metamorphic rocks deform; whilst deformation of grains, by internal dislocation flow, becomes progressively more significant at higher grades of metamorphism. This relationship between crustal temperature and deformation mechanism has been subjectively assessed from observation of the geometry of, and degree of chemical redistribution within, structures developed under different metamorphic conditions. It may be deduced from theoretical considerations of deformation mechanism fields (Ashby, 1972), and experimental work on rock deformation (Price, 1964; Donath, 1971; Rutter, 1974) that temperature, in addition to other parameters such as confining stress,fluid pressure, and strain rate, plays a fundamental role in determining the behaviour of rocks under differential stress in the earth's crust. This is due to the temperature dependent kinetics of many deformation mechanisms which operate in the complex polycrystalline aggregates which constitute crustal materials, and as such forms the basis of the subjective observations noted above. The questions which arise from these considerations of deformation mechanism fields and crustal environment are -

1. What is the crustal temperature range over which pressure solution is an important deformation mechanism?

Is it possible to place limits- on the range of temperature or depth of other deformation mechanisms, such as dislocation 82 flow and grain boundary sliding: and to what extent do these mechanisms occur simultaneously with pressure solution?

If it is possible to make estimates of the temperature range of deformation mechanisms in crustal materials, how do the results relate to theoretical models of deformation mechanism fields, and to experimental tests on rock deformation?

Having defined the problem it was then necessary to in- vestigate the best techniques for solving the problem, and to devise critical field tests to which these techniques could be applied. In order to answer these and other questions it is necessary to define conditions such as temperature and pressure in crustal environments where pressure solution was observed to play an important role in deformation. The two techniques which were selected as most suitable for obtaining such data in this study are fluid inclusion and oxygen isotope thermometry. These are the most accurate and widely applicable techniques available for geothermometry. When used in combination they provide a powerful method for investigating crustal environment. These techniques have been applied to solution of the problems in three areas of prograde metamorphism, the Dalradians of Southwest Scotland, and Palaeozoic rocks of the Central Pyrenees and Southwest England. The above three' areas were selected to provide a critical test of the deformation mechanism field hypothesis, in that each represents a range of lithology and different thermal gradient, so that effects of rock type and mean stress could be separated from the effect of temperature alone. 63 The theory and application of fluid inclusion and stable isotope thermometry to the investigation of crustal conditions is dealt with in chapters 3 and 4 respectively. Since these . two aspects are closely associated, and both relate directly to the main problem of crustal temperature, much of the interpretation of data obtained by the use of these techniques is treated together in chapter 5, which deals specifically with deformation mechanism fields and crustal environment. 64

CHATTER 3

Fluid inclusion studies of tectonic veins

'Bubble, bubble, toil and trouble'.

3.1 INTRODUCTION

When crystals grow from a fluid phase growth irregulari- ties of many types, and subsequently the healing of fractures, cause small portions of the fluid phase to be trapped within the solid crystal. These portions of the trapped phase are known as fluid inclusions. The study of fluid inclusions has a history reaching back over one hundred and fifty years. Early research has been summarised by Smith (1953a) and Roedder (1972): whilst recently four comprehensive reviews of the subject have been published (Deicha, 1955; Ermakov, 1965; Roedder, 1967b; Roedder, 1972). Sorby (1858) proposed the hypothesis that the vapour bubble which is present in the fluid of most inclusions is the result of differential shrinkage of the liquid and enclosing host mineral, during cooling from the higher temperature of trapping to the temperature of observation. Hence the temperature of trapping may be estimated by heating the sample to the point at which the vapour phase disappears. This point is known as the temperature of homogenisation, or filling temperature. This property of fluid inclusions makes available a technique for estimating the temperature of formation of crystals, which have grown from a fluid phase, from a wide variety of geological environments. In addition, from further studies of 65 fluid inclusions, it is possible to obtain data on the physical and chemical properties of hydrothermal fluids in the crust. Whilst fluid inclusion research has largely been directed towards a study of the thermal and physico-chemical environment of'hydrothermal ore deposits (Roedder, 1967b), the technique has been used on tectonic vein material in order to determine the thermal history of crustal rocks which have undergone deformation.

Use of fluid inclusions in tectonic studies.

Temperature determinations have been made on tectonid veins in low grade rocks from the Helvetic Alps and Prealps of Western Switzerland (Martini, 1968; Stalder and Tournay, 1970). Fluid inclusion studies undertaken at Imperial College have been used to determine regional temperature variations in tectonic provinces, and to,measure thermal fluctuations in tectonic veins which exhibit incremental growth features (Durney, 1972b; Clark, 1972). In addition, this technique has been used in an attempt to assess the relationship between crustal temperature and deformation mechanism fields (Durney, 1972b; Cosgrove,1972).

Purpose of chapter.

The general objectives of the research which is presented in this chapter are --1.bareViewthe extensive literature on fluid inclusion research which is relevant to problems in structural geology; 2. to apply the method of inclusion thermometry to study the influence of temperature on the formation of structures in the earth's crust; 3. to use further analytical techniques of fluid inclusion study, which have b een developed 66 to gain an understandina7 of the physical and chemical environment of hydrothermal ore deposits in order to obtain a similar understanding of the properties of hydrothermal fluids in tectonic provinces. The approach which has been adopted for this research is to bring to bear on a few well defined problems in structural geology the most powerful available analytical techniques of fluid inclusion study, including combined stable isotope - fluid inclusion methods, in order to answer those specific problems. This approach of combined analytical techniques has met with considerable success in studies of hydrothermal ore deposits (Ohmoto and Rye, 1970; Rye and Sawkins, 1974). Because of the complexity of crustal systems, data from several techniques= serves to eliminate some of the uncertainties that are inherent in a more limited approach. With the type of information in mind which may be gained from the combined analytical techniques, a field sampling programme was designed in order to test those hypotheses relating to the specific problems of crustal environment and deformation mechanism fields in tectonic provinces. The general objectives have. been outlined above. More spe- eifically, the problems which were chosen for analysis are am follows -

1. To define the crustal conditions of temperature and pressure obtaining in environments where pressure solution and associated tectonic veining are observed to play an important role in deformation. 2. To assess the effect of deformation of tectonic veins on the use of fluid inclusions as a thermometer. 67 To carry out.a- comparison of inclusion and oxygen isotope temperature determinations on quartz-carbonate tectonic Veins, in order to establish the accuracy of inclusion geothermometry applied to structural problems. 4. The determination of physical and chemical properties of hydrothermal fluids from which vein material was precipitated: and a comparison of these parameters within and between. tectonic provinces. 5. To obtain further information on the physical environment of actively deforming rocks from combined use of data obtained from the analytical techniques mentioned above.

A brief outline and description is given of tectonic veins, in which the inclusions are found. This is followed by a detailed description of the types of inclusions which occur in tectonic veins, and of the relationship of these inclusion types to features of crystal growth and deformation. The effect of deformation on leakage from inclusions is considered in some detail, with special reference to recent ideas concerning deformation mechanisms in quartz. Descriptions are given of sample preparation and analytical method for each of the techniques which have been used for fluid inclusion study of tectonic veins. An analysis is presented of errors which may arise in homogenisation determina-

. tions, and of errors in temperature corrections which are applied for factors which influence the true temperature. Some elementary statistical operations have been performed on the data in order to place objective limits of confidence on the results. To date there has not been a sufficiently critical ap- praisal of the ways in which the physical conditions obtaining 68 in tectonic environments may cause both real and apparent temperature variations in homogenisation data. In the light of more complete information, which has been obtained in this study, relating to these physical conditions, a critique is presented of previous fluid inclusion work in structural geology, where homogenisation determinations alone have been made. Whilst this chapter deals with the theory, application, and problems of fluid inclusion work in tectonic provinces, and with the results that have been obtained from a specific study of tectonic veins, much of the interpretation closely relates to the associated stable isotope work on veins and their host rocks(which is discussed in Chapter 4), and to considerations of crustal environment. As a consequence, these several aspects are treated together- within the general context of deformation mechanism fields and crustal conditions in Chapter 5, and where relevant in Chapter 6 on tectonic veins. In summary, conclusions are drawn regarding the advantages, the limitations, and problems of data interpretation that arise from application of the techniques to tectonic veins. With these points in mind, it is possible to make an objective assessment concerning the value of fluid inclusion studies in structural geology.

3.2 TECTONIC VEINS

The field of investigation of this research concerns fluid inclusions in minerals from veins which have developed in a tectonic environment. Plate 3.1 Tectonic veins of calcite, and stylolites associated with deformation of a coral limestone.

•

Plate 3.2 Arrays of quartz-carbonate veins, Pic de Sarradets, 6 9 The majority of research to date, has been carried out on minerals from economic deposits. These deposits are predominantly veins with an epigenetic, cross-cutting relationship to the country rock. In addition, certain stratiform deposits hexe been studied which are considered by, some authors to be syngenetic. some examples of these deposits have complex mineral paragenesis, and are of coarse grain size. Tectonic veins always have an epigenetic relationship to the original country rock. They are generally of medium to fine grain size and simple mineral paragenesis. Tectonic veins usually bear a well defined geometrical relationship to structural features such as folds, shear zones, and planes of pressure solution. Tectonic veins are widespread throughout areas of deformed low grade metamorphic rocks, but, as we shall see, are formed less commonly within 2-3 Km of the earth's surface, and in high grade metamorphic terrain. The processes of deformation which result in tectonic veining may be responsible for the mobilisation of certain mineral species into solution, and the formation of reservoirs of hydrothermal fluids enriched in these mineral species. As a consequence, there may be a relationship between deformation, tectonic veining, and the formation of veins in regions external to the tectonic activity. There is a wide range in scale of tectonic veins, from microscopic features, to structures tens of Metres in thickness, and up to 100 Km in lateral extent. Plate 3.1 is a photomicrograph of calcite veins oriented perpendicular to planes of pressure solution, in a deformed coral limestone. Regular arrays of veins three orders of magnitude larger than •

Plate 3.3 Tectonic veins normal to cleavage planes in folded greywacke units, together with a sigmoidal vein array associated with a hinge shear zone.

•

late 3.4 Tectonic veins associated with deformation of a conglomerate ty pressure splutlnr. 70 those illustrated occur over hundreds of metres (Plate 3.2). The latter vein arrays are related to major fold structures. On an intermediate scale, Plate 3.3 shows ellipsoidal veins with a regular geometrical relationship to cleavage planes in a fold; and a sigmoidal vein array in a hinge shear zone. Tectonic veins generally exhibit some degree of deformation because they must accommodate the same strain that develops in the host rock as a result of tectonic processes. The minerals most commonly present in tectonic veins, in the general order of abundance are quartz, carbonates, phyllosilicates, and feldspar. There are three reasons why quartz has been used almost exclusively for fluid inclusion work in this research -

1. High frequency of inclusions. 2. Facility of observation due to the efficiency of light transmission. 3. Quartz has a high resistance to leakage from inclusions during deformation; in contrast to minerals which exhibit well developed cleavage.

The data from fluid, inclusions may be directly related to the physico-chemical environment obtaining during specific tectonic episodes, because of the well defined geometrical relationship between tectonic veins and structural features. a 71_

3.3 DESCRIPTION OF FLUID INCLUSIONS IN TECTONIC VEINS.

An inclusion is any space isolated within the host mineral during crystallisation (primary inclusions), or during, the healing of microfrcturen (pseudosecondary and secepdry inclusions), and which formsa phase boundary with the host mineral. elusiona of t mineral-forming solution are by far the most commonplace. Fluid inclusions from tectonic veins which have been observed o-oticolly are generally between 2 and 20 microns in size: the majority fall within the range of 2 to 10 microns. A Size continuum down to C-2 microns has been revealed by electron microscopy (Sella and. Deicha, 1962a; 1962b; 1963; White, 1973). The frequence of primary inclusions in crystals probably reflects growth conditions. Clear crystals with a few large primary inclusions have grown slowly under uniform conditions. The cloudy white appearance of some minerals such as quartz and calcite is due to the presence of numerous small inclusions, up to 109 inclusions per cm3, which is indicative of relatively fast growth under fluctuating conditions (Roeder, 1967b). Ermakov- (1950) has classified primary inclusions into four categories on the basis of the trapped phase, These are - Type 1, solid. Type 2, fluid plus daughter mineral which has crystallised from the trapped fluid. Type 3, gaSeous. Type 4, liquid. Type 3 gaseous, and type 4 liquid, are by far the most common. True hydrothermal minerals always have inclusions originally of type 4, in which the Vapour contraction bubble occupies less than 50% volume. Inclusions in tectonic veins belong to this category. 1:11 4 , S 5,t, -s ' hs of pl. fluid inclusions in 72

PRIMARY- INCLUSIONS,

Any imperfections in the growth of a crystal which is forming from a solution may cause the trapping of fluid as primary inclusions. Rapid dendritic growth, covered subsequently by complete layers, may trap inclusions. Changes in the concentration of minor constituents in the solution affecting crystal perfection; crystals which grow hollow and later become complete; fluctuation in conditions of growth by bulk flow or diffusion; temporary starvation of centres of crystal faces relative to edges; the presence of solid particles or im- miscible fluids, all may cause the trapping of primary fluid inclusions (Roedder, 1967b). Hence primary inclusions represent the physical conditions which obtained during growth of the host crystal. Most crystals have 'lineage structure', due to growth as a series of parallel blocks. If there is a difference in rate of growth of these blocks, the crystal face may become ir- tregular and form negative crystal cavities (plates 3.5 A and 3.6 R). Individual blocks form by their own growth spiral, and where three such spirals meet, cylindrical inclusionsl. elongated in the direction of growth may form (plate 3.5c) (Roedder, 1967b). Spherical, tabular, and irregular primary inclusions are also common (plates 3.6,C, 3.5p, and 3.6E1 respectively). A series of inclusions which can be directly related to a former crystal shape, or growth feature, may safely be considered as of primary origin. Vein minerals may grow in a variety of crystal morphologies'. Commonly observed are minerals of well developed crystal ilate5.bg, 5.7 ihotomicrograpb of euhedral quartz and dolomite crystals in a tectonic vein. :Plain light above, • crossed iolars below. Note primary inclusion arrays parallel to F7rowth zones in quartz. x 50.

• 7 3 habit, resulting in euhedral crystals (plate 3.68,7) Such crystals may develop as a series: of parallel growth layers and L3, 4 ) which form a continuous lattice and re- 11 L2 sult in optical continuity. Alternatively they may develop as overgrowths (G2) that are not in optical continuity with the host crystal (G1). Of other crystal morphologies which are present in vein minerals those of irregular shape, anhedral crystals; and crystals which grow as needles, or 'fibres' (plate 3.8 ) are commonly observed. Inclusion arrays A, B, and C may be directly related to the growth outline or growth layers of the host crystals, and are thus judged to be of primary origin. However, inclusions which lie on the actual boundary of growth layers, or which are trapped due to mismatch along adjacent crystal faces (arrays D, E, F) have not been used for homogenisation determinations, since such boundaries are high diffusivity pathways, and consequently the inclusions are liable to leakage. One qualification that relates to the use of inclusion arrays parallel to growth features as a criterion for primary origin, when inclusion thermometry is contemplated, concerns the nucleation of vapour bubbles on crystal faces. Such bubbles form in a boiling solution, due to pressure drop or change in fluid chemistry, and may become incorporated as inclusions. It is not advisable to use inclusions formed from a boiling fluid for meaningful homogenisation temperature determinations unless it is known that boiling has occurred. Large isolated inclusions, which have no obvious planar relationship to other inclusions, are probably of primary origin, since fractures within a crystal are unlikely to heal 74 into a single large inclusion. Primary inclusions are most readily identified in_crystals free from microfractures. Primary inclusions represent the physical conditions of hydrothermal fluids in the crust during crustal growth. Secondary inclusions which are formed at a later date subsequent to growth are almost certain to represent different environmental, conditions. As a consequence, it is of criti- . cal importance to identify correctly primary inclusions. The most useful subjective guide lines for identifying primary inclusions may be listed as follows -

1. Large 2. Isolated. 3. Of spherical, cylindrical or negative crystal shape 4. Related to features of crystal growth 5. Occur in euhedral crystals which project into vugs

Unfortunately none of these criteria for distinguishing primary and secondary inclusions are absolute, and on many inclusions it is not possible to use any of the criteria.

PSEUDO SECONDARY INCLUSION$

If a crystal fractures during growth, due to internal or external stresses, fluid from wnich the crystal is growing will enter the fracture and become trapped. Ermakov (1950) has termed' these pseudosecondary inclusions; they have been observed in both natural and synthetic crystals. In addition, during the process of hydrothermal leaching, fluid may become trapped in deep etch pits to yield pseudosecondary inclusions (Barton and others, 1963). •

Plate 3.8 Arrays of pseudosecondary inclusions which terminate at the margins of quartz fibres. Tectonic vein, Castel Marim, Southern Portugal.

Plate 3.9 Arrays of pseudosecondary inclusions transverse to quartz fibres. Quartz-calcite tectonic vein. 7 5 The characteristic feature of pseudosecondary inclusions is that the inclusion array terminates against crystal growth layers. Plate 3.S is a photomicrograph of a quartz vein which has elongate zones defined by the presence or absence of parallel arrays of pseudosecondary inclusions. In crossed polars each zone is seen to be a single crystal, but adjacent zones are not in crystallographic continuity. These zones are considered to be growth 'fibres'; pseudosecondary arrays do not cross from one crystal fibre to another. The presence or absence of inclusion arrays transverse to the fibres is presumed to be a function of local variation in growth conditions. Plate 3.9 is a photomicrograph of a quartz-calcite vein with transverse and longitudinal arrays of inclusions in quartz 'fibres'. The inclusion arrays do not traverse from quartz to calcite crystals. This may be due to healing of fractures in the calcite. However, the preferred hypothesis is that they are true pseudosecondary inclusions formed by fracturing within the quartz crystals during growth, as a brittle response to deforma- _ tion, and that the calcite has'responded to stress in a different manner. Pseudosecondary inclusionsa are not generally used for re- liablehomogenisation temperature measurements. This is for reason of necking and leakage, which are discussed in later sections. However, it has been found that pseudosecondary inclusions may be used to throw some light on the response of quartz crystals in tectonic veins to deformation, and to distinguish between 'crystal growth fibres' and 'deformation induced fibres'. This aspect is treated in the section on deformation of tectonic veins in Chapter 5. •

a * MP 104 Ibt

Plate 3.10 Secondary fluid inclusions in quartz, lying in a healed fracture. Note differences in the degree of filling of inclusions, and the corresponding variation in homogenisation temperatures.

•

Plate 3.11 Sub-planar arrays of secondary inclusions in quartz. 76

SECONDARY ircLusIors'

Secondary inclusions result from the healing of microfractures which postdate crystallisation of the host mineral. If a, crystal fractures in the presence of a fluid in which it has finite solubility the fluid will enter the fracture allowing the mineral to dissolve and recrystallise, thus trapping secondary inclusions in the process (Roedder, 1967b). The original fracture generally heals into a configuration of minimum free surface energy, resulting in a planar array of equant inclusions (plate 3.10). Secondary inclusions are present in most crystals of all tectonic veins, as a result of deformation, and several generations of secondary arrays are generally present. Unlike nseudosecondary inclusions, secondary arrays frequently cut across grain boundaries, and bear no relationship to growth layers (plate 3.11). Secondary inclusions haye been used in tectonic studies as an indicator of stress orientation in structures (Tuttle, 1949), and for temperature determinations of late brittle fracture of pressure shadows and veins (Durney, 1972b). The former is unlikely to yield meaningful results due to the incremental nature of most deformation, and to the complex causes of microfracture in crystals (Roedder, 1967b, p. 523). There is no basis for the use of secondary inclusions as a reliable geothermometer due to the fundamental nature of their origin whereby variable phase ratios may be trapped during necking, and to the possibility of leakage (Ermakov 1950 ; Roedder 1967b, p. 534:: and section 3.4). '77

Plates 3.1c and 3.11 are photomicrographs of a strongly deformed quartz aggregate, from a tectonic vein, with planar arrays of sub-spherical secondary inclusions which, unlike pseudosecondary arrays, traverse grain boundaries. The inclusions have trapped different liquid-vapour phase ratios during neckinc_71 and as a result give widely different hono- genisation temperatures.

3.4 NECKING OF INCLUSIONS

Fluid inclusions may be originally trapped in cavities with a large surface area-volume ratio: for instance-irregular, cylindrical, and tabular inclusions. Such cavities possess a high free surface energy, and if the host mineral is soluble in the fluid, the cavity will either 'neck' down by diffusion and recrystallisation into a more stable configuration of smaller equant inclusions (Roedder 1967b, p. 534), or simply adopt such a configuration without necking. Evidence of the early stages of necking in a cylindrical and irregular primary inclusion may be seen in plates 3.s c and 3.6 B respectively. Inclusions may even change their position within a host mineral due to recrystallisation or migration. These effects concerning the change in shape and position of inclusions are minimised for large inclusions, and low temperature (Roedder, 1968). This process has serious implications for inclusion thermometry, since an inclusion which has already nucleated a vapour bubble may neck down, and give rise to two or more cavities with different liquid-vapour phase ratios. Hence inclusions will give anomalous homogenisation temperatures (Th). Necking may cause both anomalously high and low values of Th. 78 From this it follows that the highest measured Th does not necessarily represent the closest approximation to trapping temperature, if a temperature correction for pressure is applied, as has been stated by Clark (1972). For instance, the difference between the upper limit and mean Th in fig. 3.44s 30t. If these values of Th arecor- rected for pressure on the basis pf a 30°C/Km thermal gradient, then the difference becomes 7 O. • Because secondary and pseudosecondary inclusions have almost invariably formed by necking, it is not advisable to use them for obtaining thermometric data. In addition, such inclusions are prone to lekkage as a consequence- of the healed fractures, along which they are located, being high diffusivity pathways. The variation in density and homogenisation temperature from an array of secondary inclusions along a healed microfracture is indicated in plate 3. lo. It is not possible to make any meaningful interpretation of such data. Necking is the cause of at least part of the scatter in data that is present in most homogenisation studies. 79

3.5 'LMTAGE INTO OR OUT OF INCLUSIONS

Leakage of fluids into or out of inclusions is considered to be one of the major problems in the interpretation of inclusion data, and reaches the most serious proportions in deformed crystals. The experimental evidence concerning leakage from inclusions is conflicting. Kennedy (1950) subjected quartz samples to 1.5 Kb hydrostatic stress, for 1.2 x 106 seconds, at 10°C below Th" He inferred that fluids had entered inclusions along lineage structures (Carstens, 1968) after pressure reduction in the inclusion consequent upon cooling. Skinner (1952) demonstrated that external pressure caused a consistent decrease of -,,10°C in Th from leakage into inclusions, and that excess internal pressure caused an increase of o C in Th from leakage out of inclusions. However, criticism has been levelled against the experimental method of Kennedy and Skinner, which gave high pressure gradients across inclusion walls in thin specimens. Richter and Ingerson (1954) were unable to reproduce Skinner's results, and in later tests run at 4Kb hydrostatic stress, 400°C for 6 1.5 x 10 seconds Roedder and Skinner (1968) observed no evidence for leakage from inclusions. Whilst the concensus of- opinion on the experimental evidence has come.out against leakage from inclusions, it should be pointed out that none of the tests were run under conditions of differential stress, or for a significant duration. Khetchikov and others (1968) have loaded quartz samples at 0-4Kb differential stress without causing leakage, but the conditions of temperature and strain rate are not documented. 80 One of the strongest lines of evidence against leakage of fluid into or out of inclusions is found in the numerous reports of a systematic relationship between Th of primary inclusions and zoning of the host mineral (Ermakov, 1950; Kalyuzhnyi, 1958c; Roedder, 1967b; and plate 3.11.). A similar systematic relationship is found between freezing temperatures and zoning (Reedder, 1953). Other reports indicate that these systematic relationships do not always hold, because of leakage (Ermakov, 1950; Bailey, 1951; Roedder, 1967b). Brine inclusions with a vapour contraction bubble which is almost a vacuum may exist for millions of years without dilution by surface waters, and adjacent inclusions of two different generations, separated by a few microns of host mineral, may maintain extreme differences of internal pressure and composition. The inference from such evidence is that the quartz lattice does not have significant permeability to protons or hydroxyl ions, even along dislocations, under conditions of moderate temperature 350°C) and hydrostatic stress. It is evident from the P.T. curves of systems representative of hydrothermal fluids in inclusions (Kennedy, 1950b.; Lemmlein and Klevstov, 1961) that inclusion pressures rise rapidly above confining pressure, for small temperature increments above Th. As a consequence, heating of veins above the temperature of formation by metamorphism may cause rupture of inclusions (Deicha, 1961; Khetchikov and others, 1968). This decrepitation of inclusions may be observed using the microscope heating stage, and generally occurs, 30°C above Th. Birkeland and Carstens (1969b) and Carstens (1969b) have noted that dislocations, twin boundaries, and microfractures in quartz are generally decorated by bubbles; and White (1973) has suggested that small fluid inclusions are pushed along by dislocation arrays during deformation. 81

The migration of fluid inclusions in halite and quartz, which are maintained under conditions of high thermal gradient, has been observed by Bradshaw and Sanchez (1969). However, for the case of quartz, migration of inclusions in the thermal gradients which obtain in geological environments is unlikely to be a significant process. Over lonc, periods of time it is suggested that fluids may be driven irreversibly into the host crystal lattice by 7:ode- rate overpressure, and not return to the inclusion on release of the pressure.

Intracrystalline mechanisms and leakage.

It is well established that alkali ions, hydroxyl ions, and protons can readily diffuse through the quartz lattice, especially under conditions of differential stress. Hydrolitic weakening of quartz has been attributed to the diffusion of hydroxyl ions, which form SiOH bonding in the quartz lattice (Griggs and Blacic 1965; Griggs, 1967; White, 1971); and small fluid inclusions are dragged along by dislocation arrays during deformation (White, 1973). The mobility of such components in a host mineral may change the phase ratio and composition of fluid inclusions. With the use of infra red spectrometry it has been revealed that natural and synthetic quartz may contain large amounts of 'water' held in the lattice as el CH-, and H2O (Kats, 1962a, b). Experiments on the annealing of such wet quartzes have shown that fluid inclusions may form from the breakdown of silanol bonds (McLaren and Phakey, 1966; Griggs, 1967). The reaction is written as follows (White, 1973) -

2( Si - OH) +H20 82

McLaren and Fbakey observed inclusions of 1000 A0 in cuartz annealed at 600oC for a few hours. Griggs noted that inclusions of 0•2 microns formed in quartz annealed at 300°C for 24 hours. Whilst many natural quartzes have not been annealed at temperatures as high as 300°C or under realistic conditions of hydrostatic pressure, it is probable that this process would occur at lower temperatures over longer periods of time. ',Laite (1973) has suggested that the release of hydrostatic pressure may cause the formation of such inclusions, which characteristically decorate dislocation arrays in pressure solution veins. It is possible that large numbers of inclusions formed by annealing or pressure release of wet quartz may coalesce to form an optically visible inclusion that could be misidentified as being of primary origin;' or migrate into primary inclusions and change the phase ratio. Since the volume of inclusions is proportional to the cube of the radius, "leakage-in" of 105 inclusions of C.1 microns would be necessary to give a change in the phase ratio of a 10 micron inclusion. Hence one may conclude that "leakage-in" of small inclusions is unlikely to be a source of significant error. The migration of fluid inclusions in quartz is a slow process (Petrovic, 1969) and, as for the diffusion of water through the lattice, is- a thermally activated process with the rate an exponential function of temperature. As a consequence, no corrected Th data greater than - 350°C has been used in this study, and the most reliable results are considered to come from lower temperature veins. Leakage may occur along fractures induced by sawing during sample preparation. The thin sections used in this and other fluid inclusion studies On tectonic veins of Imperial College are more liable to this source of error than are standard 83 double polished thick sections. In order to overcome this problem the sections were checked optically for fractures and hermetic so7lins of inclusions. In addition inclusions near the surface were not used: for determinations. Whilst an impressive array of evidence for the validity of fluid inclusion techniques has been gathered in studies of ideal, undeformed crystals, from hydrothermal ore deposits, there is a concensus of opinion among all the leading workers in fluid inclusion research that deformed material is unsuit- determinations (Ermakov, 1950; Ypma, 1963; Roedder, able for Th 1967b). Indeed, the wide spread of data that has been obtained from deformed veins (Durney, 1972; Cosgrove, 1972; Clark, 1972) provides equally convincing evidence to suggest that fluid inclusions have leaked under conditions of differential stress; and that no meaningful interpretation may be placed on data obtained from such material. The implications of leakage have been completely ignored in the above works. To illustrate the effect of deformation on leakage an analysis was made of inclusions in quartz from different parts of a vein that had experienced varied intensity of deformation. A sequence was observed from single Th data maxima with low variance, in undeformed grains, to several data maxima above

Th' with high variance, in deformed grains. This result is taken as unequivocal evidence that the inclusions in strained grains have leaked as a result of deformation, and as a consequence give anomalous homogenisation temperatures. Indeed,

Th data from deformed veins may be in direct conflict with other geological evidence. For instance, Durney (1972b) obtained temperatures of 510°C mean, 575°C maximum (corrected from 340°C mean, 400°G maximum, for 1.5Kb pressure, 10% NaC1 84 solution; Obmoto and Rye, 1970, p. 528) from veins in host rocks of lower greenschist facies mineralogy, where one would anticipate temperatures of 300°C - 350°C. The effects of deformation on tectonic veins is discussed more fully in Chapter 5. In addition to leakage, variable phase ratios in inclusions ) may be due to the following factors: (end hence variable Th 1. Trapping of a heterogeneous fluid. 2. Yecking. 3. Misidentification of primary and secondary inclusions. 4. Variation in temperature, pressure, density, and chemistry of the hydrothermal fluids.

Factors 1, 2 and 3 may be recognised or eliminated by examining large numbers of inclusions in successive crystal growth zones. It is required to know all the properties listed under 4 in order to compute true temperatures by correcting homogenisation data. These properties are treated individually in the following sections.

Criteria used to overcome problems of leakage.

With these limitations in mind regarding the interpretation which may be placed on fluid inclusion data from tectonic veins, due to deformation and leakage, as discussed in this section, it was considered essential for this research that the following requirements should be satisfied for homogenisation work -

1. Samples must be seledted from undeformed euhedral crystals projecting into vugs because such crystals cannot have been subjected to differential stress. 85

2. Consistent measurements. should he made in successive growth zones of crystals free from microfractures. 3. All measured values of T, should, after application of the ,pressure correction, yield temperatures below ,-35000.

3.6 FLUID Ii,CLUSICIT- THERECEETRY

This section deals first with the principles, assumptions, and limitations of fluid inclusion thermometry, with special reference to tectonic environments; and second, with the application of this technique to some specific problems in structural geology. Sorby (1858) in a classic paper on fluid inclusions suggested that the vapour bubble present in most inclusions was the result of differential volumetric contraction of the liquid and enclosing host mineral, during cooling from the higher temperature of trapping to the temperature of observation. Sorby demonstrated experimentally that the coefficients of volumetric contraction for a variety of liquid solutions, representative of the fluids in inclusions, was one to three orders of magnitude greater than the coefficients of contraction for the host minerals. Hence the temperature of trapping may be approximated by heating the sample to the point at which the vapour phase disappears, and the fluid becomes homogeneous. This technique is known as the homogenisation method, and is now accepted as a reliable method for estimating the minimum formation temperature of minerals which have grown from solution. Sorby worked out the effects of pressure, temperature, and chemistry on compressibility of fluid phases. Together with thi 86 data, and information on the volumetric ratio of the two phases he derived a theoretical model for calculating original trapping temperatures. Because of the difficulty in accurately measuring phase ratios, and hence in estimating the density of the fluid (see section 7, density of fluid inclusions), Sorby's method has not been widely used. The decrepitation method of thermometry (Ermakav, 1950) is also beset by serious errors, which are listed by Roedder, 1967b. Almost all accurate inclusion thermometry is- now done by the homogenisation method. The limitations and assumptions of the homogenisation method are relatively simple in nature; difficulties arise in attempts to estimate the sources and magnitudes of errors for a given sample and geological setting. The assumptions are listed and discussed in detail by Roedder (1967b), from which a brief summary is given below, with special reference to factors important in tectonic veins.

I. The origin of the inclusion is known. Only primary inclusions, i.e. those which formed during crystal growth, are representative of the conditions under which the host mineral formed. Conclusive evidence for primary origin is rare.

II. The fluid trapped upon sealing of the inclusions was a single homogeneous phase.

If a large number of inclusions in a sample all have the same composition and phase ratio (and hence Th) it may be assumed that a homogeneous fluid was trapped. Individual inclusions formed from a heterogeneous system (e.g. a boiling solution) can be expected to trap different ratios of the phases present. 87

Trapping of vapour from a boiling solution, or of a separate gaseous phase such as CO2, may cause a large positive error in Th, but this effect should be evident in many inclusions. In addition to the trapping of a heterogeneous system, variation of phase ratio between primary inclusions may be due to fluctuations in temperature, pLessure, and chemistry of the hydrothermal fluids; to leakage or tecking.

III. The cavity in which the fluid is trapped does not change in volume after sealing.

Thermal contraction of the host mineral is' reversed on heating to Th, excepting the additional temperature correction required for pressure effects. Crystals may be formed under considerable hydrostatic stress, and the inclusion will experience elastic strain as the host mineral is uplifted to the surface,-of the earth. Further expansion of the inclusion on heating is to be expected from excess pressure inside the inclusion. Because the compressibility, and coefficient of volumetric expansion of most minerals is one to three orders of magnitude lower ban that of aqueous brines, the total effect of thermal and stress induced volumetric changes of inclusion cavities is insignificant. The author suggests that deformation of the host mineral may cause a shape, and hence volume change, of inclusion cavities strained at constant surface area. If this occurs before nucleation of the vapour phase leakage will result: if a bubble has formed, then the phase ratio will be altered. In the analysis of deformation of quartz, in a tectonic vein, a sequence was observed from large primary inclusions of 88 spherical shape and negative crystal habit with uniform phase ratios in undeformed grains, to ellipsoidal and irregular inclusions with variable phase ratios in strained grains. It may be in erre from this evidence that the shape Ch an-e and associated leakage of inclusions in strained r;rains is a conequence of deformation of the host mineral. Permanent non-elastic deformation of inclusion cavities in

uorite, (7wri- ti cave cumulative e-n-r‘,rs ,-40°C ° to +60 C in Th from three successive homo7enisation runs has been reported by Larson. and other (lW3). This process is un-

likely to have a significant effect in ouartz below 35C0C because the value of Younf:-ys Lodulus of quartz is greater than for fluorite, and because intracrystalline 1,rocesses in quartz are generally insignificantbelow this temperature over short periods of time. Small changes in volume may occur due to crystallisation of daughter minerals in inclusions (Ermakov, 1950). Only a few daughter crystals which were believed to be NaC1 were found during the observation of several thousand primary inclusions in this study (see also Durney, 1972b, p. 250). Daughter minerals generally redissolve on heating, but if this does not occur due to metastbility then inclusions containing such crystals are not used for Th determinations, because anomalous results may be obtained.

IV. Nothing is added to or lost from inclusions after saa]ling.

The addition to or loss from inclusions has been discussed in the sections on leakage and necking. Leakage from inclusions is almost universal in deformed crystals,but is rare in minerals from hydrothermal ore deposits. 89 0.2 0.4 0.6 0.8 1.0 Density

Fig . 3.1. Temperature-density-pressure diagram for the system H2O , adapted from Roedder (1967 b, p. 531 L is the boiling curve for pure water; L, is the boiling curve for a

20% NaCl solution. 90

Necking is common, especially at higher temperatures, and probably accounts for a large part of the scatter of data that is apparent in most thermometric studies (Roedder, 1967b, p. 558).

V. The effects of pressure are insignificant or are known.

Pressure controls the density of fluids above the boiling curve (Fig. 3.1). If a fluid inclusinn is trapped along the boiling curvel Th equals the true temperature of trapping, and no pressure correction is required. Should an inclusion be trapped at P.T. conditions above the boiling curve, a vapour bubble does not fOrm on cooling until the pressure and temperature have dropped to the boiling curve. Four inclusions which formed at 360°C but at different pressures, and therefore densities, are depicted in Fig. 3.1. Each will homogenise at different temperatures along the boiling curve: the vertical lines represent the pressure corrections (Roedder, 1967b, p. 559). The magnitude of discrepancies between measured temperature and the true temperature of formation, with respect to pressure and fluid chemistry, has been estimated by Safronov and Khetchikov (1950); Roedder (1967b), and Shaposhikov and Ermakov (1968). The author has computed the magnitude of errors which may arise in estimating true temperature of formation from Th data alone. This has been done for a range in the values of those variableswhich determine corrections of Th (namely P, X, H20/CO2, density), table 3.1. Even if the crustal level at which a vein formed may be accurately inferred from stratigraphic data, there can be no knowledge of the depth of water which was present above the sedimentary column; and it remains purely subjective as whether ` to assume that the hydrothermal fluids were at lithostatic, hydrostatic, or some intermediate value of pressure for appli-

cation of corrections to Th values. 91

Clark (1972), using the method of Ingerson (1947), assumed a 30°C/Em thermal gradient in the crust in order to calculate the depth of vein formation for pressure corrections. This method is the best available for situations where there is no independent evidence on the depth of cover, but it may involve very lar^ e errorsat moderate pressures (see Fi7. 3.1), and does not resolve toe question of lith static versus hydrostatic fluid pressure. Whilst the assumption of uniform thermal gradient and fluid pressure may be valid for hydrothermal ore deposits in stable crustal environments, these assumptions are not tenable in tectonic provinces for the following reasons:

1. Pressure fluctuations may arise from tectonic overloading or underloading as a consequence of nappe movement. 2. Stress release or reversals from hydraulic fracture, during the formation of veins and folds (Price and Hancock, 1972). 3. Pressure release from hydraulic fracture of impervious layers in the crust which have -ponded fluids at anomalous fluid pressure (Price, 1975). 4. Normal dissipation of fluid pressure from lithostatic to hydrostatic in a magmatic hydrothermal system (Toulmin, 1967). 5. Dynamic pressure changes due to fluid flow through channels of varying cross-sectional area (Toulmin, 1967).

Evidence that hydraulic fracture and filling of many tectonic veins occurs slowly is adduced in chapter 6 on veining. From this it follows that dynamic pressure changes are negligible due to slow rates of fluid movement. However, any process which results in pressure release may cause boiling in hydrothermal fluids, and as a consequence give rise to anomalous values of Th, as discussed under trapping of a homogeneous phase (II). 92

Variation in thermal gradient will arise primarily from tectonic overloading, and this effect has been measured within a nappe unit of the. Central Pyrenees (see chapter 5). Thermal fluctuations due to endothermic or exothermic reactions in the hydrothermal fluids will be negligible at slow rates of fluid movement. Boundary layer effects (Bailey, 1951) give rise to errors which are insignificant in relation to those listed above. All the processes which it is suggested may give rise to variation of thermal and pressure•gradients in tectonic environments probably operate within, or over, the same time scale as the formation of veins. Pressure correction involves the further assumption that the chemistry of hydrothermal fluids is knowns becausAethe vapour pressure, thermal expansion, and compressibility of water is markedly affected by the presence of alkali halides and CO2. Pressure corrections to Th, for inclusions the composition of which approximates to a complex H2O - CO2 - alkali halide system, may be either greater or less than for inclusions of pure water which homogenise at the same temperature, depending on the exact fluid composition. The most up to date sources of data on the thermodynamic properties (pressure, temperature and density) of fluid systems representative of hydrothermal fluids in the crust, which have been used in this research, are as follows:

1. Pure water Kennedy, 1950b; Burnham and other s,1969. 2. H2O - NaC1 Lemmlein and Klevstov, 1961. Souriajan and Kennedy, 1962. 3. H 2O - NaC1 KC1 Samoylovich and Khetchikov, 1968b.

4. H2O - CO2 Takenouchi and Kennedy, 1964. H2O - NaC1 - CO2 Takenouchi and Kennedy, 1965. 93

In addition, data is available for the thermodynamic properties of potassium chloride solutions (Samoylovich and Khetchikov, 1968a, b). Hence, in order to apply meaningful pressure corrections to homogenisation determinations it is necessary to make prior measurements of the physical and chemical properties of fluid inclusions. This is achieved first .by the freezinc; technicue, to determine fluid salinity; second by H20/002 determinations from crushed samples; and third by measurement of phase ratios to give density. These techniques, and the results obtained from them, are described in subsequent sections. Khetchikov and others (1970) have grown synthetic crystals under known conditions of temperature, pressure, and fluid chemistry; and subsequently' compared the measured Th to true trapping temperature, for a known pressure. Using data from this technique exact corrections may be applied to T1,, values from natural crystals if the pressure is accurately known. How- ever the data is available for a limited range of T.P. and fluid chemistry (X). Errors may arise in the application of this data (Durney, 1972b) to inclusions in geological environments outside the available T.I.X. range.

Pressure Estimates

A lower limit to the pressure of formation of inclusions may be established from the vapour pressure of the solution at 1 the temperature of homogenisation. This pressure may be estimated from thermodynamic data on the relevant system, if in dependent evidence on the fluid salinity and CO2 content is available (Dolgov and others, 1968). , 94

200 300 400 500 600

Fig. 3.2. Temperature- pressure-density relationships of water and 10 wt. % NaCt solution, adapted from

Ohmoto and Rye (1970, p. 42 8 ). 95

Two methods have been suggested for using L20 - 00,, in- elusions in order to determine both the pressure and temperature of'inclusion formation (Nacken, 1921; Smith and Little, 1959; Roedder, 1967b, p. 565). However, such inclusions are rareland. large errors are inherent in both methods. Sharkov (l965) has estimated pressure and temperature from the difference between observed Th' and the true temperat ure trapping-calculated from phase ratios using Sorby's method. This technique requires inclusions in which the phase ratios may be accurately measured (see section 3.7). It is apparent from the graphs of the thermodynamic properties of a fluid system of specified composition (Fig. 3.2) that the homogenisation temperature cannot give both the temperature and pressure of inclusion formation. If one is known the other may be derived. Errors which arise in pressure corrections become very large at moderate temperatures and pressures ( 25000 and Kb, Table 3.1) and place severe limitations on the interpretation which may be placed on Th data. As a consequence a new method was devised by the author for making pressure estimates, in order to apply valid corrections to Th data, and to obtain further information on crustal conditions. This method depends on an independent temperature measurement, Tt, established from oxygen isotope mineral pair determinations on a vein (quartz-carbonate). If the homogenisation temperature is known, and the physical and chemical properties of the hydrothermal system is defined (density, salinity, H20/002) then the pressure may be derived by reference to data on the relevant fluid system, by assuming that the correction due to pressure is Tt - Th (Fig. 3.2). 96

Summary of requirements for thermometry,

Fluid inclusion thermometry provides one of the most wide- . ly'applicable and accurate methods of determininL; crustal temperatures. In order to make use of this technique on tectonic vein material the main conditions that must be satisfied, which relate to t7riose discussed abo7,,- may be suratari follows:

1. Material must not be deformed. This required euhedral crystals projecting into vugs. 2. Temperatures must be lower' thane," 350°C due to the diffusion of 'water' through the lattice, and formation of inclusions during annealing. 3. Consistent data within growth zones of crystals must be obtained, in order to eliminate the possibility of leakage. 4. Independent evidence on the physical and chemical properties of the hydrothermal fluids is required. 5. Necking down processes must have been negligible.

These conditions are difficult to satisfy simultaneously in tectonic environments, and require careful field sampling and experimental technique. No Th measurements have been made in this study unless all the conditions are satisfied, or unless it could be demonstrated that negligible errors are introduced by not having information on condition 4. 1 97 CV 415

NV 347 TV 178

IV* 240 '7°

LOT 294

78 too

Fig. 3.4. Histogram of homogenisation temperatures C 0C 1, of inclusions -identified as being of primary origin, from SOME tectonic veins.

the number of observations is indicated. 98

Statistical methods.

The filling temperature data are presented in the form of histograms, Fig. 3.4. Statistical tests have been used in order to establish the internal consistency of data, and the significance of measured variations of Th. Each group of data was tested first for normality usin7 the Kolmogorov-Smirnov method. Homogeneity of variance was then established using Bartlett's test. Both of these conditions were satisfied at a 95% level_of confidence for all data from undeformed material. There is a marked departure from normality of data obtained from deformed veins. The F test was used to make a comparison of Th measurements 1. from different area of a given section; 2. from different sections from the same vein; 3. from veins in different areas. The first two points were satisfied at a 95% level of confidence. The third point could be established within 99% confidence limits. NaCl and density determinations have been reduced to arithmetical means and standard deviations for each group data. These results are listed in Table 3.2. 99

PREFACE TO SECTIONS ON DENSITY, SALINITY, AND MEASUREMENTS ON FLUID INCLUSIONS. H20/C02

As was stated in the introduction to this chapter the approach which was adopted for -rluid inclusion-research on tectonic veins was to use a combination of analytical tecl-lnioues which, served to eliminate some of the uncertainties in interpreting homogenisation data alone; and to gain an understanding of the physico-chemical environment in tectonic provinces. prom these considerations folic* two major reasons for making estimates of density, salinity, and H20/CO2 ratios on fluid inclusions.

1. As a necessary requirement for applying pressure corrections to homogenisation data (section 3.6). 2. To make a comparison of the physical and chemical properties of hydrothermal fluids within, and between, tectonic provinces; and to enable a further comparison of these fluids ' in tectonic provinces with established data on the properties of fluids in hydrothermal ore deposits.

Inclusion data connected with the second point are closely related to associated studies on the oxygen isotope characteristics of hydrothermal fluids in tectonic provinces, and these aspects are treated collectively in chapter 6, which is concerned with 'hydrothermal fluids in tectonic provinces'. 100

3.7 DENSITY IdEli.SU are ryTS. O FLUID INGIUSIC1:S.

Fluid inclusions may be used to make reasonably accurcte estimates of the density of hydrothermal fluids in the crust, and are the only directly available source of this information. - if the relative volumes of the liquid phase and inclusion cavity are measured at room temperature, in inclusions of regular geometrical shape, then the density of the originally homo- -genemm3fluid may be calculated (Roedder, 1967b, p. 565). The reasons for making density estimates on fluid inclusions in this study, in addition to those listed in the preface to this section, are:

1. To demonstrate the trapping of variable phase ratios in secondary inclusions (see plate 3. ). 2. To demonstrate leakage of fluid from primary inclusions in deformed host minerals (section 3.5 and chapter 5).

Both of these points have been conclusively established by the author.

Method

The technique adopted for density determinations by the author was to use photomicrographs of primary inclusions, because the facility for direct measurement under the microscope was not available. The inclusions were projected on to graph paper from negatives placed in an enlarger, and the image of 'inclusion and vapour bubble traced out. Area ratios of liquid and inclusion cavity may then be computed, translated into volume ratios, and thus into values of density. The shape of primary inclusions was approximatect•to spherical, 'cylindrical, or rectangular, for purposes of calculating volume ratios. 101

Photosraphs were tal,en from three slides of each vein, from several grains within each slide, and of several inclusions within each grain; the same sampling, method as was used for homogenisation determinations. This served to establish the consistency or otherwise of measurements, and the sampling level at which inconsistencies mi7ht arise. In practice it was found that the data from each undeformed vein was in-5ern- ally consistent. This is because the same rigorous conditions in choice of sample and measurement technique which were met for homogenisation determinations were also met for density measurements. In meeting these conditions the requirement of random sampling for statistical comparisons is not fulfilled; however, as for homogenisation determinations, unrepresentative inclusions are being subjectively rejected. The data does meet the test of normality (at a 955 level of confidence), so that one may conclude the sampling is largely of a normally distributed population of primary inclusions, with variation due to fluctuation in growth conditions.

Errors

Volume changes from external or internal pressure, and, from thermal contraction of the host mineral, are negligible compared to the errors involved in volume measurements. In addition to sources of error such as leakage and necking which should be largely eliminated by choice of sample and measurement technique, errors inherent in density measurements are as follows:

1. Underestimation of vapour volume may occur due to the curved upper surface of inclusions acting asa strong negative lens (Reedder, 1972, p. 13). , 102

2. The maximum diameter of bubble and inclusion cavity may not occur at the same focuSsing point. 3. Uncertainty of the shape of inclusions in the third dimen- sion, particularly for the case of flat inclusions. 4. Small vapour bubbles may become non spherical due to in- ertial forces during Brownian movement.

5. Operator variance in tracing the images due . to poor defini-. tion at high magnification. 6. Subjective assessment as whether an inclusion approximates to cylindrical, spherical or tabular shape for calculation of volume ratios.

All these sources of error are minimised for measurements on large inclusions, and standard deviations on the data are within reasonable levels. The data is presented in Table 3. . 103

10 solution NaCI. solution

Na CI •2H,20 + solution ice 4- solution

-21.1 ice + NaCl•2Hp

-30 I . 0 5 10 15 20 25 30 23.3 Weight percent NaC1

Fig 3.3 The system H2O — NaCt . Adopted from

Roedder 1962 p. 1059 104

3.8 SALINITY MEASUREMENETS ON FLUID INCLUSIONS.

The salinity of hydrothermal fluids present in the crust during vein formation may be estimated by recording the depression of the freezing point of fluid inclusions. A microscope freezing stage, cooled by boiling nitrogen gas, is used to make the determinations (Roedder, 1962, 1963, 1967). The measured freezing point of an inclusion may be directly converted to salinities from data on the system NaCl - E20 (Fig. 3.3). This method assumes that the NaCl - H20 system is representative of hydrothermal fluids in the crust. Data obtained. on the salinity of primary fluid inclusions in quartz from five tectonic veins is listed in Table 3.2. All inclusions studied exhibited metastable behaviour. during cooling, which results in freezing points of -6000 to -90°C. This metastable phenomena is due to the high purity of some alkali halide hydrothermal fluids in the crust, and absence of microscopic particles that are necessary for nucleation of an ice phase (Roedder, 1962, 1963). Hence one may conclude that the hydrothermal fluids present during formation of the tectonic veins studied were moving too sloIrly to contain microscopic particles in suspension. It follows from tlis result that no large volumes of hydrothermal fluids rsre mi';ratin7 rapid 7- t::--0111,. the crust dur- inr: deformation, and as a consequence 1=terial for vain filling must have been locally derived. •

ilate 3.11 temi:erature measurements within succes ive growth layers cf a zoned quartz crystal. 1 0 5

3.9 H20/002 DETERMINATIONS ON FLUID INCLUSIONS.

The gases and fluids from inclusions are released by crushing the host mineral in a vacuum extraction line (static O. and CO are frozen down in a liquid. pressure 10-5 torr). H2 2 nitrogen trap (- 196°C); 002 is then allowed to "boil off" by replacing the liquid nitrogen with a mixture of dry ice and acetone (-4/7°C). The volume of has evolved_ is measured in a calibrated dibutyl phthalate oil manometer. Water is converted to hydrogen by reaction with hot uranium metal at 700°C, and volumetric estimates made as above. For, amounts of CO2 below the sensitivity of the manometer (-•,714,kg 002) the gas is transferred to the inlet system of a gas source mass spectrometer. The CO2 is equilibriated with a known volume of neon gas and then analysed accordingly. Data obtained for IT20/002 ratios on fluid inclusions are listed in Table 3.2.

CONCLUSIONS RELATING TO TEE IEYSICO-CEEEICAL rRCFERTIES CF HYDROTNERNAL FLUIDS IN TECTONIC EROVINCEa

Estimates have been mule of temperatures obtaining during vein formation, using the fluid inclusion homogenisation method. Temperatures of the hydrothermal fluids vary from <1000, to

> 350°C de7enditc- on- the crustal environmen from which samples were collected. 72m2,-I'atutres ta-d to deoe during the history of vein growth (plate 3.12). This result contrasts with the hypothesis of Cosgrave (1972), who contends that fluids present during later stages of vein growth may have come from greater depth, and as a consequence be hotter. Hence one may conclude that energy consumed by endothermic reactions is greater than the heat evolved from -exothermic reactions in Density measurements on fluid inclusions

Sample Density Standard deviation Number of observations

PV 0.90 0.01 47

TV 0.92 0.01 31

NV 0.87 0.008 54

CV 0.89 0.02 20

LV 0.94 0.01 42

Analytical results for H2O and CO2

Sample Water CO Wt% CO 2 2 H 0 / CO 2 2 ratio content content

PV 2.19mg 0.001mg - 1500;1

TV 0.82 ' 0.07 8.5% 12:1

NV 0.30 0.02 6.7% 15:1

CV 1.44 0.26 18% 6;1

LV 0.44 0.003 - 1500:1

Determinations of fluid salinity

Sample Weight % NaC1 Freezing point

PV 5.4 -85 0C

TV 9.7 -54

NV 10.3 -78

CV 6.7 -80

LV 3.8 -95

Table 3.2 Density, H20/CO2 , and salinity data on fluid inclusions

from tectonic veins. 106 hydrothermal fluids in many tectonic provinces. The measured decrease of temperature during vein formation is consistent with available information on the thermal history of fluids in hydrothermal ore deposits (Ohmoto and Rye, 1970; Roedder, 1968a, 1971a, 1971b). The measured densities of fluids in tectonic provinces is consistent with a subcritical H20-002-Nan system, that one would anticipate in high and intermediate crustal levels at temperatures below fv360°C. Low ratios of H20/CO21 and hence relatively high CO2 content, wore Irycent in inclusions from veins formed in predominantly carbonate environments. The converse is true of CO2 content in inclusions from veins formed in dominantly quartz bearing rocks. Thus lithology controls the chemistry of hydrothermal fluids in tectonic provinces; a result that is consistent with data on hydrothermal ore deposits (Roedder, 1972).

Low values of salinity, < 10 wt/o YaC1 may he inferred for primary fluid inclusions. Temperatures and salinities fall into the range ouoted by Roedder (1967) for'veins of hydrothermal affinities'. Mississippi Valley-Alpine type deposits, and mineralisation associated with magmatic bodies, are associated with hydrothermal fluids having inferred temperatures and salinities significantly outside the range of values obtained for fluids associated with tectonic veins. Hence one may conclude from quantitative data that hydrothermal fluids in many tectonic provinces originate from dewatering associated with compaction (diagenetic and tectonic) and metamorphic dehydration reactions. Further evidence to support this contention is presented in chapter 4. 107

3.10 FIELD SAMPLING

The field sampling programme for obtaining specimens was designed specifically to test those hypotheses arising from the problems of crustal environment as outlined in the introduction to this chapter. This proRramme covered two major points:

1. 'To sample throughout the range of crustal levels at which pressure solution was obberved to play an important role in deformation, in order to define the crustal temperature interval of this deformation mechanism. 2. To make comparisons of the, relative importance of pressure solution (in the generalised temperature interval) within a number of tectonic provinces, each characterised by different litholo7ies and tl-ermal gradients. In this manner the effects of rock type and mean stress on the deformation mechanism field of pressure solution may be separated from the influence of temperature alone.

With respect to the first point there are many tectonic provinces in which the rocks have been deformed at low to intermediate metamorphic grade, such as the Dalradians of Southwest Scotland, and Palaeozoic rocks of the Pyrenees, Rheinisches Schiefegebirge, and Southern England. _Hence there was no shortage of field areas suitable for assessing the higher temperature zone in which pressure solution becomes of decreasing importance as a deformation mechanism. There are, however, few localities in young mountain belts in which rocks deformed within 1-2 Km of the surface have sur- vived erosion. Considerable time was spent in searching out such localities in order to evaluate the role of pressure 108 solution in high level, low-temperature, tectonic regimes. Those field areas which were selected as most suitable for study of this objective are as follows - Palaeozoic cover west-Central Pyrenees; Dinant anticlinorium, Southern Belgium; Tertiary Helvetic Eolasse, Northern Switzerland; and Alpine structures in Mesozoic rocks of Southern England. The second t rresented few difficulties with resrect to sampling localities. Those field areas which were selected on the basis of varied lithology and thermal gradient from literature surveys and personal experience are the Mona Series of Anglesey, which was involved in a late Precambrian Orogeny; the Dalradians of Southwest Scotland, Palaeozoic rocks of Northwest Scotland, and the Lake District, which were involved in the Caledonian Orogeny; Palaeozoic rocks of Southwest England, the Central, Pyrenees, Rheinisches Schiefegebirge, Southern Portugal and Southern Belgium, which were involved in the Hercynian Orogeny; riesooic rocks of the Swiss Alps, French Jura, and Southern England in which Alpine structures formed. Areas with a complex history of repeated deformation were not considered, so that ambiguities in the relationship of veining to a given structural episode could be avoided; and the possibility of veins having been deformed ai a later date, post formation,was largely eliminated. For instance, tectonic veins developed in Cambro- Ordovician sediments of the Heilam Nappe, Sutherland, may be related to one major structural episode. This Nappe has not been involved in repeated deformation during subsequent tectonic episodes, as is the case for higher Nappe Units (Kerrich, 1971). An additional example is- given by the chevron folds at Hartland, N. Devon, which do not exhibit features of refolding such as are evident in folds of the same generation at Bude. 109 A further major requirement, concerning the thermal history of each field area considered was first, that the structures and associated veining should have formed synchronously with the attainment of peak temperatures; and second, that the rocks should not have been overprinted by a later metamorphic event. These points were determined using the s dard techniques of structural petroloy. The former condition allows inclusion temperatures to be directly related to metamorphic mineral assemblages of the host rocks. The latter condition ensures that inclusions have not leaked as a conseauence of heating beyond the temperature of formation; and lowers the possibility that annealing effects have altered the distribution of 'water' within the vein auartz lattice (see section 3.5). Having selected field areas which met the general conditions outlined above, it was then necessary to exercise critical standards of sampling at individual localities, with respect to the requirements concerning quality of vein material necessary for inclusion thermometry, as listed in section 3.5. Tectonic veins with a definite geometrical, and where possible chemical, relationship to structures of a given tectonic episode were selected. In order to satisfy all these conditions simultaneously only I. in 1000 of veins examined in the field were finally used for inclusion thermometry.

Sample Preparation.

Standard double polished thick sections were used for those samples on which both heating and freezing determinations on the same inclusionewere made. 110 50 micron, unpolished sections were utilised for the re- f mainder of the work. These sections give excellent transmission of li7ht when used with silicone oil (MS. 710). They are sufficiently thick not to be delicate; and enable measurement of inclusions which are not within one diameter of the surface, in case cf leakage induced by 7rindin7. Details of the method for preparing these thick sections are given by DUrney (1972b) and Clark (1972). The slide is immersed in methylated spirits for 12 hours, during which time the Lakeside mounting medium is dissolved, and the section becomes detached from the mount. This procedure prevents Lakeside vapour obscuring vision during heating runs.

Instrumental Calibration

The heating stage is calibrated using a fusion point kit (Clark, 1972). Calibration was done over the relevant temperature range before and after each series of measurements, and temperatures recorded on a mercury thermometer accurate to 4- - 0.5°C. A thin film of silicone oil was introduced between the heating stage and glass mount, to increase the efficiency of thermal transfer. It was considered essential, however, to determine the thermal gradients between heating stage and sample, at a given heating rate of 2°C/minute. This was achieved by noting the apparent difference in fusion temperature, with respect to the mercury thermometer, of a substance placed above the cover slip, on the glass mount, and on the heating stage. Estimates made of thermal gradients obtained using this method were - 111 0'5°C at 150°C 1.0°C at 200°C 2.0°C at 300°C

All filling temperatures were corrected using these data. In order to establish the effect of heating; rate on measured T7a1 a series of tests was run at heating rates of 1°C/minute - 5°C/minute (in 1°C/minute increments) on some large inclusions with precisely determined filling temperatures in the range 210°C - 250°C. It was found that errors increased rapidly above 4°C/minute, and a heating rate of 2°C/ minute was used throughout the thermometric studies.

All measurements of Th on individual inclusions were made in duplicate, and in triplicate when there was cause for uncertainty. This procedure increases the accuracy of results, and provides a check against leakage of inclusions. Success- ively higher temperattresare recorded in each heating run if leakage takes place. Most small inclusions exhibit a varied degree of metastability in nucleation of the vapour phase below Th (as a consequence, no measurements were made on reappearance of the vapour phase, in contrast to the method of Durney, 1972b, p. 257). All work was done in a closed room, using the heat shield to reduce thermal fluctuations caused by air currents. The combination of instrumental errors coupled with uncertainty in the ohservation of Th-gives an accuracy in measured filling temperatures of ± 5°C up to 300°C. This. figure is comparable with errors for data obtained by inclusion thermometry quoted in the literature (e.g. Ohmoto and Rye11970). 112

Some caution should be attached to the trial and error method of Durney (1972b, p. 257) for which an accuracy of ± 0.25°G in T1, is given. This level of accuracy is more than an order of magnitude better than has been obtained by leading workers in the field of fluid inclusion research. Durney ignored the problem of metastable nucleation, did not make heatinr7 rat calibrations, and based his measurements on the tacit assumption that thermal gradients in the heating stage were negligible.

Measurement Technique

The 50 micron sections were separated into four subsections and heating runs done on each piece individually. This procedure was adopted so that internal consistency of the data could be checked by comparison of four sets of results (see section on statistics), and so that if one subsection was overheated, causing leakage from inclusions, the whole slide was not lost. In addition it is easier to identify and traverse inclusion arrays within individual grains viewed at high magnification by manoeuvring small slides. If a vein is formed over a temperature interval of more than ti 30°C then heating to homogenise an inclusion in the upper part of this range may cause leakage from inclusions formed at lower temperature (for reasons discussed in section / 3.5). Leakage will generate inclusions of anomalous density, which will give rise to spurious values of measured As a result of this factor it is necessary always to make measure-/ ments from the lowest filling temperature upwards. In order to achieve this a visual estimate was made of / fluid-vapour ratios in inclusions at the start of heating rush 113

Initial measurements were done on those inclusions with the highest density, and therefore lowest filling temperature; subsequently Th determinations were made on inclusions in a 30°C interval, before moving on to measure inclusions which homogenised in consecutive 30°0 increments. As acheck against leakage, inclusions with low Th, on which the first filling temperature determinations had been made, were remeasared at the end of each series of heating runs. This condition was satisfied for all sections on which the author made heating determinations. In practice no large temperature fluctuations were observed in those veins which had been selected for inclusion thermometry, and the only systematic variations are such as those measured in successive crystal growth zones, Plate 3. . Attempts have been made to reconstruct the thermal history of veins which exhibit features of incremental growth, by making measurements of Th along growth fibres (Durney, 1972b; Clark, 1972). It is important to realise, however, that any

measured fluctuations in Th may be due not only to real temperature variation, but also to changes in pressure, fluid chemistry, fluid salinity, CO2 content, leakage, or to inter- mittent boiling. It is not possible to determine all the variables, but fluctuations in these parameters are to be expected in tectonic environments (as discussed in Section 3.6), and systematic variations of density, salinity, and Th in hydrothermal systems have been discovered from detailed fluid inclusion studies (Ohmoto and Rye, 1970). One may conclude that carefully designed field sampling, which is related to problems that are capable of solution by inclusion techniques, coupled with rigorous experimental methoa are essential prerequisites for obtaining useful information from fluid inclusion studies. 114

REFERENCES - CHAPTER 3

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ROEMER, r. 1967c. Metastable superheated ice in liquid-water inclusions under high negative pressure. Science, v. 155, 1413-1417. ROEDDER, E. 1968a. Temperature, salinity, and origin of the ore-forming fluids at Pine- Point, Northwest Territories, Canada, from fluid inclusion studies. Econ. Geol., v. 63, 439-450. ROEDDER, E. 1971a. Fluid inclusion studies on thejporphyry- type ore deposits at Bingham, Utah, Butte, Montana, and Climax Colorado. Econ. Geol., v. 66, 98-120. ROEDDER, E. 1971b. Fluid inclusion evidence on the environment of formation of mineral deposits of the Southern Appalachian Valley. Econ. Geol., v. 66, 777-791. ROEDDER, E. 1972. Composition of fluid inclusions in Data of Geochemistry. U.S. Geol. Surv. Prof. Papers 440JJ. ROEDDER, E. 1968c. Fluid inclusion shape, a transient feature of little diagnostic value (abst.) in Fluid Inclusion Research. Proceedings of COFFI 1968, 4. ROEDDER, E., A.V. 11-,YL and T.P. CREEL. 1968. Environment of ore deposition at the Mex-Tex deposits, Hansonburg district, New Mexico, from studies of fluid inclusions. Econ. Geol., v. 63, 336-348. ROEDDER, E. and B.J. SKINNER. 1968. Experimental evidence that fluid inclusions do not leak. Econ. Geol., v. 63, 715-730. RUTHERFORD, M.J. 1964. Geothermometry of liquid inclusions in quartz. Coronation Mine, Flin Flon area, Saskatchewan (abst.). Can. Min. Jour.,v. 85, 1220. RUTTER, E.M. 1974. The influence of temperature, strain rate and interstitial water on the experimental deformation of calcite rocks. Tectonophysics, v. 22, 311-334. RYE, R.O. and F.J. SAWKINS. 1974. Fluid inclusion and stable isotope studies of the Casapalca Ag-Fb-Zn-Cu deposit, Central Andes, Peru. Econ. Geol., v. 69, 181-205. $APRONOV, G.M. and V.Ye. KHADZI. 1950. Thermometric investigations of liquid inclusions in quartz, in Ermakov, N.P. and others, 1965. Research on the nature of mineral-forming solutions, with special reference to data from fluid inclusions. Ed. E. Roedder, Pergamon Press, New York. 743 p. SAMOYIOVICH, LA. and KHETCHIKOV, L.N. 1968a. Correction factors for homogenisation pressure and temperature of aqueous solutions of sodium and potassium chlorides from experimental data, (abst.) in Fluid Inclusion Research - Proceedings of COFFI. 1968, 65. SAMOYIOVICH, L.A. and L.N. KHETCHIKOV. 1968b. On the corrections between pressure, temperature, and density in aqueous solutions of sodium chloride and potassium chloride. (in Russian). Translated in Doklady Acad. Sci. U.S.S.R., v. 180, 155-157. 119

SAY, P.J. 1964. Lead-zinc ore deposition in the li7ht of fluid inclusion studies, Providencia, Zacatecas, Mexico. Econ.'Geol., v. 59, 881-919. SCI EDT, R.A. 1962. Temperatures of mineral formation in the Miami-Eicher ditrict as indicated by liquid inclusions. Econ. Geol., v. 57, 1-20. SELLA, ;C. and G. DEICHA. 1962a. Etude au microscope electro- ' nique,-, des pores intergranulaires des gangues-et des roches. Acad. Sci. (Faris) Comptes rendus, v. 254, 2796- 2793. SETiTiA, C. and G. DEICHA. 1962b. Lacunes de cristallisation et pores intergranulaires du quartz. In Breese, Sydney, Eds. Electron Microscopy, Reports of the Fifth International , Congress for Electron Microscopy, no. 664. Academic Press, New York. SELLA, C. and G. DEICHA. 1963. Importance des cavites intra et intercristallines dans 1'architecture des mineraux et des roches. Jour. Microscopie, v. 2, 283-296. SHATOSHIKOV, A.L. and N.P. ERMAKOV. 1968. Magnitude of the discrepancy between homogenisation tamperatures of gas- liquid inclusions and the true temperatures of crystallisation of the synthetic quartz in which they are sealed off. (abst.) in Fluid Inclusion Research - Proceedings of COFFI 1968, 17. SHARKOV, A.A. 1965. Thermometry and composition of liquid inclusions in crystal quartz from deposits in Dagestan High- lands, in Ermakov,- N.F. and others, 1965. Research on the nature of mineral forming solutions with special reference to data from fluid inclusions. Ed. E. Roedder. Pergamon Press, New York. 743 p. SKINNER, B.J. 1953. Some considerations regarding liquid inclusions as geologic thermometers. Econ. Geol., v. 48, 541-550. SMITH, F.G. 1953a. Historical development of inclusion thermometry, Univ. of Toronto Press, Toronto. 149 p. SMITH, F.G. and W.M. LITTLE. 1959. Filling temperatures of H2O-0O2 fluid inclusions and their significance in geothermometry. Can. Mineral., v. 6, Part III, 380-388. SORBY, H.C. 1858. On the microscopic structure of crystals, indicating the origin of minerals and rocks. Quart. J. geol. Soc. Lond., v. 14, Part I, 453-500. SOURIAJAN, S. and G.C. KENNEDY. 1962. The system H2O-NaC1 at elevated temperatures and pressures. Am. Jour. Sol., v. 260, 115-141. STADLER, H.A. and J.C. TOURNAY. 1970. Fensterquartze mit Methen- Einschlttssen aus dem westlichen Til der Schweizerischen Kalkalpen. Schweiz. Min. Petr. Mitt., v. 50, 109-135. 120

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Introduction The purpose of this chapter is to discuss some applications of oxygen isotope geochemistry to an investigation of the' crustal environment of pressure solution. A theoretical background to oxygen isotope thermometry is presented. The advantages and limitations of this method are discussed in order to give a basis on which results may be judged. The results of this investigation are presented in three sections. The first part is concerned with the thermal environment of pressure solution as a deformation mechanism. This aspect of the research was conducted in the Dalradian series of Scotland. Specific reasons for choosing this area, with respect to the conditions of structure, metamorphism, and lithology, are argued in detail. The work has been extended in connection with the thermal stability limits of some metamorphic mineral assemblages. The second part of this research is concerned with the equilibrium -relations between tectonic veins and their host rocks. Some conclusions are drawn concerning the nature and sources of fluids in tectonic environments. Finally, some attention has been directed towards the problem of non-equilibrium isotope frationation in stress fields. 122

INTRODUCTION TO OXYGEN ISOTOIE GEOCHEMISTRY.

The nucleus of an atom, which contains a fixed number of protons for individual elements, may have a small range in the number of neutrons, which results in isotopes of the element each having a different atomic mass. For instance, the three isotopes of oxygen have approximate atomic mass, as identified 16 17 18 by the superior prefix, 0, o and 0. In contrast to the isotopes of elements which may undergo radioactive decay, to form radiogenic daughter isotopes, and which thus change their abundance; the stable isotopes of light elements such as 02, S, N2, H2 and Si have approximately fixed relative proportions in natural substances. For example, the relative abundance of 1601 170 and, 180 in atmospheric oxygen is 99'76%, 0°00 and 0.2% respectively (Nier, 1950). Oxygen is the most common element of the earth's crust, and is present in the structure of most rock forming minerals. Because all the isotopes of a given element have identical electronic structures, the bonds which they form in any given compound are also nearly identical. Hence the various isotopic species of an element exhibit broadly similar chemical behaviour. However, the energy levels for atomic vibrational, translational, and rotational modes depends on mass as well as on bond strength, and although the potential energy surface for two isotopes of the same element are identical, the zero point 1 energy ( 2 hv) is greater for a molecule composed of light atoms 18 (Fig. 4.1). As a consequence a molecule containing 0 differs in its-thermodynamic properties such as internal energy and entropy from a molecule containing 160 . For instance, the 18 free energies, and hence the solubilities of Si 0 and 2 16 Si 02 are slightly different. 123

These- differences in the thermodynamic properties of isotopic species lead to small systematic variations in the relative abundancies of light stable isotopes in natural materials. The variations occur as a result of two effects, equilibrium and kinetic isotope fractionation. Equilibrium fractionation of stable isotopes is a temperature dependent effect, and as such forms the basis for isotope thermometry. The theory of temperature dependent equilibrium isotope exchange reactions is firmly established, and an outline of this theory is given in the following section. Little is known about the frequency or magnitude of kinetic isotope fractionation for solid state processes in geological environments. A direct test, having well defined boundary conditions, has been designed in order to establish the degree of isotopic disequilibrium resulting from kinetic fractionation in a diffusional field. Kinetic isotope effects are discussed in Section 4.5

Isotopic standards and notation.

The abundance of oxygen isotopes in natural materials is analytically determined and reported as 6: values, where

( 0 18 / 160 ) sample 1 . 103 4_.1 (180 / 160 ) standard

This value expresses the enrichment of the sample in 18 0 relative to the standard in permil (°/oo ). The international oxygen standard which has been adopted is Standard Mean Ocean Water (abbreviated SMOW I S = 0 permil. As is the normal procedure, the analytical results in this 124 a) >,

Distance between atomic centers

Fig. 4.1. Schematic diagram showing the relation-

ship between zero-point energy and atomic mass

for hydrogen (Hi) , deuterium( Di), and HD frequencies

4405 cm , 381 7 cm , 311 9 cm , respectivly. . Adapted

from Broecker and Oversby (19 7 0 ) p.151 125 study on silicate and iron oxide minerals have been converted from the laboratory standard to the snow scale. Most carbonate 180 analyses in the literature are based on the PDB standard (Craig, 1961). The carbonate results obtaine in this study were determined relative to the PDB standard, and 18 16 in order to use the results in combination with 0/ 0 analyses on silicate and iron oxide minerals, for mineral-pair thermometry, the carbonate data were normalised to the SNOW scale (see Appendix V ). Isotopic fractionations between phases are quoted as values: thus, for the mineral pair quartz-magnetite

4.2 ,A (quartz-magnetite) = 8 quartz s magnetite

Isotope exchange reactions

The oxygen isotope equilibrium fractionation factor, co( between two phases such as quartz. (Q) and magnetite (M) is given by ( 0 18 / 160 ) quartz C( QM 4.3 ( 16 180 / 0 ) magnetite which expresses the degree to which the isotopes of a given element are partitioned among the different phases at equili7 brium in an isotope exchange reaction. In order to express the equilibrium constant of an isotope exchange reaction in terms of the 8 values which are analytically determined the fractionation factor,44 may be related as a close approximation to the isotopic equilibrium constant, KQM, by the expression QM' 1/n K 4.4 °C QM QM

120 where n is the number of atoms involved in "the exchange. For example, the isotope exchange reaction between quartz ) may be written such that only one (SiO2) and magnetite (Fe304 atom is exchanged

1 . 1 18 1 .18 16n Fe3 04 7 Si 02 3 `14

.16 16 16 0 etc. (O'Neil and Clayton, where SI 02 indicates Si 0 1964). It then follows that + 8 si02/1000 4..6 1 + ( Fe304/1000) which, using certain mathematical approximations (e.g. Clayton and others, 1972) reduces to

0 1000 lno(= 8SiO2 - 6 Fe3 4

The theoretical basis of oxygen isotope geothermometry.

The general theory of isotope exchange reactions was developed by Urey (1933, 1935a, b) on the basis of statistical mechanics, for the case of ideal gases. Urey calculated the fractionation factors for some reactions involving hydrogen and duterium molecules. It followed from this theory that the fractionation factors of isotope exchange reactions are temperature dependent (Urey, 1947; Bigeleisen and Mayer, 1947). Urey proposed the use of this property to define a palaeotemperature scale, from measurement of oxygen isotope exchange between calcite and water. The science of stable isotope geochemistry has subsequently expanded to cover a wide range of applications. 127 In order to understand the cause and mechanism of temperature dependent isotope fractionation, and to interpret experimentally determined data on natural systems, it is useful to consider the theoretical basis of isotopic distribution between phases. The partition function is an expression, derived on the basis of statistical mechanics, which defines how atoms will . distribute themselves among the available energy levels. It may be demonstrated that for a diatomic gas the translational, rotational, and vibrational partition functions each contribute an isotopic fractionation, but that only the vibrational compo- nent gives rise to a temperature dependent effect (Rushbrook, 1949; Broecker and Oversby, 1971). The distribution of isotopes between coexisting solid phases is controlled entirely by vibrational frequencies. Hence, when computing equilibrium fractionation between minerals for oxygen isotope thermometry it is necessary only to consider the temperature dependent vibrational partition function. The equilibrium constant, Km for the isotope exchange reaction between quartz and magnetite (eqn. 3.6) may be written in terms of partition function ratios of the two isotopic forms of each reactant in the system:

( Q2/Q1) Si02 4.8 (Q2/Q1) re304 where Q2and Q1 are the partition functions for the totally substituted heavy and light isotopic species respectively. The ratio Q2/Q1 is given by

3/2 Q2 S exp(-E2/kT) 1 . 2 4.9 Q1 [m1 Rexp(-El/kT)

• where S 1 and S2 are the symmetry numbers of the two molecules; 128

MI I m2 their molecular weights; E is an energy term; k, Boltzmann's constant; v, a vibrational frequency term: and T is absolute temperature. The temperature dependence of fractionation factors for isotope exchange reactions is evident from this expression. In order to make a quantitative evaluation of the partition function ratio, it must be expressed in terms of atomic vibrational frequencies, w, which can be directly measured using spectroscopic techniques, equation (8), Bottinga and Javoy, 1973, p. 253. In the high temperature limit, their equation simplifies to-

In C7t = AT-2 4..10 and in the low temperature limit to

In ck = A 4 BT-1 4..11

,The general expression for isotopic exchange between silicates (s) and an aqueous phase (w) in the temperature range 500°C - 800°C is

103lnc4(sw) A + BT-2 4-.12

Equation 3.10, however, is strictly valid only for ideal gases. The calculation of fractionation factors for isotopic distribution among solid and liquid phases requires certain assumptions and simplifying approximations, which introduce errors of an unknown magnitude. In particular, the atoms in solids vibrate at many different frequencies: for most solids the exact spectrum of these frequencies is unknown. Further, the effect of anharmonicity in atomic vibration, and of the interaction between internal atomic vibration and crystal - lattice vibrations is neglected in the calculations. 129

Despite these limitations of the theory, calculation of fractionation factors between divalent metal carbonates (O'Neil and others, 1969) are in reasonably good agreement with some experimentally determined mineral-water fractionation curves. The isotopic partition function due to internal atomic vibrations is treated using frequencies obtained from infra- red and Ramon spectra. The contribution of lattice motions to the partition function are evaluated using a combination of Debeye and iinstein functions, with parameters taken from heat capacity and Raman spectra. The heat capacity of a solid provides a measure of its average vibrational frequency. Quartz, with a heat capacity per atom (Cp/n) of 3.54 cal/mole-atom is characterised by a higher vibrational frequency (wq = 1235cm 1) than is haematite (wh = 1000cm 1) with a heat capacity Cp/n .4°96 cal/mole-atom. In addition to the vibrational effect, as inferred from heat capacity data, the Fe-0 bonding in haematite is weaker than the Si-0 bonding in quartz, and Fe has a higher atomic weight than does Si. These factors combine to account for the stronger bonding of oxygen atoms in quartz than in haematite. The heavy isotope is always concentrated in that mineral in which the oxygen is most strongly bonded, which explains the 18 enrichment in 0 of quartz relative to haematite. The calculations of Broecker and Oversby (1971) on the 18 relative tendency of minerals to concentrate 0, which follows the approach outlined above, correlates with the sequence obtained from experimental measurements on natural mineral assemblages (Taylor, 1967, p. 128). Bottinga and Javoy (1973) have used a combined theoretical - empirical approach to derive the constants A and B in 130 equation -13, for several mineral-water exchange reactions. (Their equations are given in table 1. 4. .) The equations of Bottinga and Javoy give better agreement between temperatures calculated from measured isotope ratios in minerals, and temperatures inferred from other geological criteria, than do many of the equations obtained by experimental calibration. Further, their equations suggest a greater degree of concordant isotopic equilibrium between coexisting phases in natural assemblages than is indicated by use of the experimentally derived equations. In all of the calculations discussed above it is necessary to make assumptions which introduce errors of an unknown magnitude (Taylor, 1967; Bottinga and Javoy, 1973; O'Neil and others, 1969). Therefore it is necessary to measure experimentally mineral-water fractionation factors as a function of temperas ture in order to callibrate mineral pairs for oxygen isotope thermometry. This experimental approach, together with its problems and limitations, is discussed in the following section. The value of statistical mechanical calculations of stable isotope fractionation among phases, in addition to those aspects which have already been discussed, is twofold:

1. To derive equations such as 4.12, which give a theoretical basis for extrapolating calibrated mineral-water curves beyond the temperature range over which experiments were done. 2. To provide independent estimates of fractionations for comparison with the experimental data. 131

Experimental calibration of mineral pairs for oxygen isotope geothermometry.

As was axplained in the previous section it is not possible to make an accurate quantitative evaluation of the partition functions, which define equilibrium constants, for most isotopic exchange reactions involving solid and liquid phases. As a consequence it is necessary to make direct measurements of mineral-water fractionations over a range of temperatures in order to calibrate mineral pairs for oxygen isotope geothermometry. For experimental purposes it is convenient to determine equilibrium constants between each mineral and an aqueous solution, or 002, rather than to measure directly the fractionation between solid phases. In most studies of oxygen isotope fractionation between minerals and an aqueous solution or 002, over a wide temperature range ( 000 - 800°C) the experimental data closely approximates to an equation of the form:

103 ln ,J A -F. BT-2 4.13 where is the mineral-water equilibrium fractionation factor; A and B are constants; T is temperature, °K. The expressions fittedito measured isotopic fractionations between mineral phases and aqueous solutions over a range of temperatures may thus be combined to give an equation for the isotopic fractionation between mineral pairs as a function of temperature. For example

1031n of (quartz-water) = -2.73 + 3.57 . 106T 2

1031nd (muscovite-water) = -3.10 + 1•90 . 106T-2 132

These equations may be subtracted according to the relation:

1031n (X (quartz-muscovite) = 1031nc4(quartz-water)-1031n (m-w) which gives the quartz-muscovite fractionation:

1031n ok(quartz-muscovite) . 1.16 1.19.. 106T-2

Isotopic equilibrium between amineral and the aqueous solution at a given temperature is tested by approaching the final 180/160 ratio in the mineral from both directions. However, it has not yet been possible to obtain complete isotopic equilibrium with calcite, dolomite, or quartz below 200°C, 400°C and 500°C respectively, for kinetic reasons (O'Neil and others, 1969; Clayton and others, 1972). The experimental problems of calibrating mineral pairs for oxygen isotope thermometry are discussed by Taylor (1967, p. 120-122); O'Neil and others (1969); and Clayton and others (1972). It is not practically possible to attain equilibrium fractionation below a certain temperature (which differs for each, mineral) due to kinetic reasons. A technique involving partial exchange and extrapolation to an equilibrium value has been used to overcome this problem. However, systematic errors in the data which have been obtained from use of this technique have not yet been resolved. There have been successive attempts to calibrate mineral- water fractionations as a function of temperature (see, for instance Taylor, 1967; O'Neil and others, 1969; Clayton and others, 1972); but variation among the equations which have been obtained for quartz-water fractionation are so large that o significant errors (",100 C) may result from combining the quartz-water equations with another calibrated system such as muscovite-water. 133

The most,recent quartz-water, and muscovite-water fractionation curves (Clayton and others, 1972; O'Neil and Taylor, 1969) have been obtained from several synthesis paths, using different starting materials, which give broadly consistent equilibrium exchange data. In addition, theoretical calculations of fractionation factors from partition function ratios for the system quartz-calcite are in good agreement with the latest experimental results for this system at all temperatures (Clayton and others, 1972). It is possible in some cases to derive an expression for the isotopic fractionation of a particular mineral-water pair as a function of temperature, when there is no experimental data on the system. This may be achieved by empirical measurement of three cogenetic minerals, two of which have been experimentally calibrated. For instance, Clayton and Epstein (1961) obtained the equation

1031n 0( (quartz-haematite) = 1.388 1031n o( (calcite-haematite) from isotopic determinations of cogenetic quartz, calcite and haematite. From this relation the haematite-water equation may be computed. Relationships between other mineral pairs have been established using this technique:

1031n0( (quartz-magnetite) = 0-49 1031n o( (quartz-biotite) Sheih and Taylor (1969)

1031n of (quartz-chlorite) = 0.59 1031n of (quartz-ilmenite) Wenner and Taylor (1971)

This technique extends the range of mineral pairs which are available for oxygen isotope thermometry, and it has been 134 used in this study to derive the equations for haematite-water, chlorite-water, and dolomite-water fractionations (substituting the most recent equations for calibrated mineral-water fractionations). The empirical measurements have generally been made over a limited range of temperature, and this method has the implicit assumption that the expression derived for mineral- water fractionation follows equation 3.13.

Advantages and limitations.

One of the major advantages of stable isotope thermometry is that fractionations are independent of pressure, because the volume changes associated with isotopic substitutions are negligible. Other methods auch as muscovite-paragonite or pyrrhotite-sphalerite salvus type thermometers, and the fluid inclusion technique are -pressure dependent. Some modification to the calibrated mineral-water fractionations may be necessary because of the non-ideality of oyy:f,en isotope fractionations in hot saline brines, which constitute fluids in the crust (Truesdell l974). •

25 Km

Palaeozoics granites igneous limestone schist phy II i te Dalradians 9 ri t chlorite schist M Moines

Fig. 4·3 Lithologies of the Iltay Dalradians, Southwest Highlands of Scotland. Sampling localities - s pec i min s for isotopic determin-

ations. Place names listed in table 4.3 135

4.2 THE DIRADI.17 97=1 OF THE SOUTHWEST EIGTILArDS OF SCOTLAND.

Introduction.

The D1rEldians a-e a sl7elf seouence of sediments and volcanics t7t were de-posited daring late I-recan-brian and early Combrian times. Accordin7 to Bailey (1922, 1934) and subsequent wori,ers l'i:37; Johnson, 1965) the Dal-r-ans can he divided into two major structural and stratiF:rarhic units: the Ilitay 1:appe Conplex a d the Ballapel Foundation. Each of these units is characterised by a distinct stratiEraphical succession and structural history; and the contact between the two is invariably a powerful tectonic discontinuity, the Illtay Boundary Slide. This work is connected with the Illtay Dalradians, a succession of rocks which is effectively continuous for 350 Km along strike, from Kintyre, through the West-Central Highlands, to Banffshire (Fig. 4.3). An approximate average age for sedi- mentation is given by an Rb-Sr isochron of 690'2:30 my (Dewey and Pankhurst, 1970). The Illtay succession is about 10 Em - 15 Km thick in the south-west. A wide range of lithologies are present, from coarse grits, through arkoses and Quartzites, to fine grained pelites and carbonates. A correlation of the stratisraphy between various areas is given by Roberts (1974, p.-_96, 97). 0

1 Primary structures

3 Secondary structures

Fig 4.5 Major structural elements of the ,Iltay Da(radians, Southwest Highlands of Scotland.

* sampling localities for isotopic determinations • Sampling localities for structural petrology 136 Structural history of the Illtay Nappe Complex,

The overall disposition of Illtay lithologies is controlled by major naple structures, with axes oriented in a 173-3W direction. These nappe structures have evolved in three tectonic episodes of widespread extent;. although many fold phases, which are of purely local significance, have been recognised in certain areas (Clough, 1897; Rast, 1963; Treagus, 1964; Roberts, 1966c, Roberts, 1974). A large tract of the Illtay Dalradians lies on the flat, upper inverted limb of the early Loch Tay recumbent synform. This synform closes to the SE in a complementary recumbent antiform, the Aberfoyle anticline, which is preserved only in the Highland Border Region. To the NW of the Loch Tay 'flat belt' lies a 'steep belt' of the Loch Awe - Sron Mhor syn- clines, which passes into another flat belt of recumbent, NW facing anticlines. Minor folds and cleavage structures associated with the early recumbent nappes are widely developed, excepting those areas which have suffered penetrative deformation at a later date. Major second folds are developed only in the West-Central Highlands, where their NW-SE_ attitude has caused a marked swing in the regional strike, for example at Schiehallion. However, folds, with associated axial planar schistosity, are minor F2 common throughout the Illtay Dalradians. Major third folds (F )are approximately cylindrical with 3 the early nappe structures. These F3 folds, such as the Ben Lawyers syncline and Tarbert-Ben Ledi monoforms are responsible for deforming metamorphic zones established before F3, and for the downward facing of the Aberfoyle and Cowal anticlines. 137

Structural and metamorphic history of the Illtay Dalradians

Structural Episodes NetamorlYc Grad.,

Formation of major 17E-SW cuartz-chlorite-muscovite-albite name structures. Einar fine rained th7llites rold4,nG and cleave. D,) Ninor to interrete 11,,,rtz-1-)iotite-muccovite-alhite scale foldinc on axes: formation of garnet-staurolite-yonite scistos'ty Lare scale and ascocia- ted maTcor folds, 7-3'::- Lnnec.1.:-; and matth co,aronL[1. Defortion of D2 metamorpilic zones Minor structures, locally Local retro7ra'le meta7oriarlt develoLed. c1711oritisation of bictite an:. garnet .

Table 4..1. Generalised sec.uence of structural e;isodes and metamoric events for the Illta7 Dalradians. Cam iled fro17, Rast, 1950b;_Zturt and Harris, 1961; Johnson, 1962, 1963, 1965;

arte and Johnson , Dewey -07,-u-, t lo7r, garnet • kyanite

garnet

+ Glasgow

Metamorphic Zones

garnet • kyanite

garnet Arran biotite

chlorite

Kintyre

Fig. 4.4 Metamorphic index mineral zones in the Iltay Dalradians, with sampling localities. 138

Einor structures of the third generation are characterised by chevron folds in well -bedded lithologies, and in pelites by stron,ly developed crenulations. The traces of major fold structures in the Illtay Dalradians is illustrated in Pi7.

r an - jsn of the IlItay Dalradians.

The Deiradians provide a clasic and well documented example of reional r:letamorphisn. The distribution of metamorphic index mineral zones is now well established, together with a reliable correlation of structural episodes and metamorphic grade between different parts of the Ilitay outcrop (Bast, 1958b; Sturt and Harris, 1961; Johnson, 1962, 1963; Harte and Johnson, 1969). A wide range of physical conditions (PT) is represented in the gradation from low temperature mineral assemblages of chlorite grade, to high temperature sillimenite bearing rocks, within both the low pressure and intermediate pressure metamorphic facies series (Den Tex, 1971). Johnson (1963) has -1 estimated that thermal gradients of r.. 16000 rim and 100°0 Km-1 obtained during metamorphism in the Highland Border Region, and West-Central ighland respectively. Nuch lower thermal o -1 gradients of 30-40 C Km may be inferred from stratigraphic data available for the Southwest Highlands (Borrodaile, 1973). Metamorphic zones cut across stratigraphical boundaries and across the primary recumbent nappe structures. The metamorphic zones or- isograds, have subsequently been deformed by major third folds, as described in the previous section. A highly generalised correlation of structural episodes and metamorphic grade for the W-Central Highlands is given in 139 Table 4.1. This sequence represents the progressive increase and decrease of metamorphic grade as related to the tectonic history, determined using the criteria of structural petrology. The maximum metamorphic grade attained varies throe the Tlitay Do1r 143r9; for instance, from chlorite Trcle in. Bute, throuh bfotfte schists of 7.7. Coal, to ;arnet bearinr; as- semb7a7es at Tumrel (see Pig. ). ac=s for n.otaorhism of r3n-e frcm 480 my in the south to 410 my in the north-west. This 1 a,narept1 are pattern orlIses from differential coelin7, rates (Broun c2nd others, 1965). 140

4.3 TEMPERATURES AND DEFORMATION MECHANISMS

Boundary conditions for lithology, structure, and metamorphism.

Large areas of the illtay Dalradians, such as knapdale and Southern Gowall have remained at chlorite grade througbout the hisGory of Oeformation. Ctler areas have boon deformd at a low grade of metamorphism, but exhibit evidence of progressive deformation under conditions of increasing metamorphic grade throughout the structural-metamorphic history (Ben Lawyers, Schiehallion, Ben Lui - Fig. 4.4.). Metamorphic zones are oblique to the strike of stratigraphic units (Fig.4j.I-). Therefore the various lithologies may be examined for given features of deformation at low metamorphic grade, and a comparison made of these features at increased temperatures as the lithologies are traced through areas where the rocks contain successively higher grade metamorphic assemblages. Further, because the axis of deformation is broadly parallel to the strike of lithologies (Figs. 4.3 and 4.4), a given episode of deformation may be examined in a number of rock types under a variety of metamorphic conditions. This reliable correlation of structural and metamorphic episodes between different parts of the Dairadian outcrop (Table 4.1) has enabled a rapid study to be made by the author of the relative importance of deformation mechanisms at different temperatures, in a variety of rock types.

Field of investigation.

As was stated in the preface to chapters 3 and 4 the major objective of this study is to establish the temperature range 141 over which pressure solution is an important deformation mechanism; and if possible to place limits on the range of temperature or depth of other deformation mechanisms, such as dislocation flow and grain boundary sliding. With the boundary conditions relatinr to the Illtay Dalradians outlined above in mind, the foliowin specific features concerning deformation mechanisms were investigated

1. The importance of pressure solution in a given rock type (e.g. quartzitic) deformed under different conditions of metamorphic grade. 2. The importance of pressure solution in other lithologies (e.g. carbonate, pelite) deformed under different conditions of metamorphic grade; and hence to establish the temperature range over which a number of different minerals respond to deformation by pressure solution. 3. The importance of pressure solution in a number of different lithologies (e.g. carbonate, quartzite, pelite) at a given metamorphic grade. 4. The effect of grain size on pressure solution in quartzites at a given metamorphic grade (i.e. in fine and medium grained quartzites, quartz grits, and conglomerates). 5. The effect of mean stress (.-crustal depth) on the role of pressure solution as a deformation mechanism, in a given lithology, at a specified metamorphic grade.

It should be implicitly understood that an intercomparison was made concerning the relative importance of pressure solution and other deformation mechanisms, with respect to conditions of temperature, mean stress, and lithology in each of the five aspects of deformation enumerated above. 142 Method

Oxygen isotope thermometry was used to estimate the temperatures obtaining during syn-metamorphic deformation, over a range of metamorphic grades, from lower greenschist facies to garnet-kyanite grade of the amphibolite facies. Two main criteria governed' the choice of locality for collecting specimens for isotopic determinations. They were -

The samples were collected at field sites where features of deformation that developed contemporaneously with the peak of metamorphism were investigated. This enabled a direct relation to be made between temperature and deformation mechanism. The time relation of deformation to metamorphism was established from use of the standard techniques of structural petrology. B. Localities were selected where the lithologies had the re- quisite grain size and mineralogy, in order to obtain sepa rates for isotopic determinations (see Appendix IV).

In areas which were deformed at the lowest grades of regional metamorphism some temperature estimates were made using the fluid inclusion technique on material from syntectonic veins. Two fursther aspects concerning the thermal environment of deformation mechanises and metamorphism were investigated, in addition to the five points listed above. These aspects relate to the data which was obtained from the analytical methods of thermbmetry used in this study. The points are as follows -

6. To make a comparison of temperatures obtained for a given metamorphic assemblage using oxygen isotope thermometry, with temperatures inferred for that assemblage on the basis of thermodynamics and experimental equilibrium relations. 143

7. To combine temperatures obtained from oxygen isotope and fluid inclusion thermometry, in order to make estimates of the temperature and fluid pressure obtaining during syn- metamorphic deformation; using the method devised by the author, as discussed in chapter 3.

Field and petrographic studies.

A number of critical localities were chosen on the basis of the boundary conditions and specific objectives listed above. Structures at these localities were examined in the field, and subsequently in thin section, in order to assess two factors -

1. The extent to which pressure solution or other deformation mechanisms had played a role in the formation of structures. 2. The time relation of deformation to metamorphism, as determined using the standard techniques of structural petrology.

Difficulty was experienced in assessing the mineral assemblage and deformation mechanism associated with areas which were deformed initially at a low metamorphic grade, but which subsequently experienced tectonic episodes under conditions of amphibolite facies metamorphism. Some tentative information concerning the early deformation and metamorphism was obtained from a study of petrographic textures within garnets, which were pretectonic in relation to the higher temperature deformation (for instance, at Tummel, and Schiehallion). The early fine-grained textures within such garnets, survive the matrix coarsening which is characteristic of rocks that subsequently experience higher grades of metamorphism (Harte and JohnSon 1969). 144 Metamorphic index minerals.

Ectomorphic mineral assemblages were recorded for each locality from studies of thin sections, and mineral separates.

Criteria used to assess the presence or absence of pressure solution in structures.

The -c,resence or absence of pressure solution in structures was assessed in the field, and from a study of thin sections, using the criteria listed below -

1. Pressure shadows. 2. Coarse anastomosing cleavage zones.

In coarse grained lithologies such as grits and gray- wackes, pressure shadows together with associated zones of depletion, characteristically develop around large elastic grains and iron oxide minerals. There is generally a well defined boundary between pressure shadows which form as overgrowths on large detrital grains of quartz, feldspar, or calcite, and the detrital host grain (Plates 5.%,5..10). In thin section such pressure shadows may be distinguished from the host grain by the absence of inclusions (Plates (:15c7 ). Detrital grains are bounded by zones of pressure solution which are defined by the depletion of pressure-soluble minerals such as quartz or calcite, and the concomitant enrichment of phyllosilicate and iron oxide minerals, which have a relatively lOw mobility (Plates 5.1, and 14). Within the zones, phyllosilicate minerals exhibit a high degree of preferred shape orientation, as do residual grains of quartz or calcite. In addition, the residual grains of quartz or calcite are of significantly smaller size than such grains in the matrix, which have not undergone pressure solution (Plate 5.2). i45

The large detrital grains act as "stress-raisers" initiating the process of pressure solution in discrete zones which propagate through the rock. These zones are aperiodic, typically anastomosing, and form the mesoscopic parting, or non- penetratiVe compositional cleavage which is characteristic of grits that have been deformed, at least in part, by pressure solution (Plate.S S. 1),, S.10). In the matrix, detrital grains tendto be eauant, and phyllosilicates do not exhibit the high degree of parallelism which is apparent within discrete zones of pressure solution (Plates S Some of the large detrital grains are fractured in a direction which is oriented normal to the cleavage planes. The fractures are generally infilled with quartz or calcite, in a manner which is analogous to the formation of pressure shadows.

Petrographic features of grits deformed at higher grades of metamorphism.

coarse-grained lithologies were traced progressively from areas which had experienced deformation at low temperatures, to areas where structubes developed synchronously with the formation of biotite and then garnet bearing assemblages. A. change in the nature of the deformation mechanism, from that described above, became apparent in rocks which were deformed at higher grades of metamorphism.

1. Detrital grains deform by an intracrystalline mechanism. The presence of subgrains in quartz, together with deformation bands and strain shadowing may be taken as evidence to infer deformation by a dislocation process (Plate 5.7). Hence, the shape orientation of grains is due to internal deformation, not to pressure solution with associated overgrowths. •

Plate 4.1 Buckled slate with pressure solution cleavage. Kilmory, Knapdale, Southwest Scotland.

Plate 4.2 Buckled bictite schist, Tichrabruaich, Southwest Scotland. 146

2. The deformation is generally homogeneous, which results in a smaller range of size and shape orientation of grains. 3. Phyllosilicates are of larger grain size, and more uniformally distributed throughout the rock. As a consequence the schistosity is not defined by tectonically indaced cori:osi- tional variations. 4. Detrital .;rains do not have pressure shadows together with ctures.

At the highest grades of metamorphism, the rocks that were originally grits, are characterised by equant polygonal grains, which do not generally exhibit evidence of intracrystalline deformation, such as strain shadowing, subgrains, and deformation bands. Presumably, at higher temperatures, either the defect density has been lowered to a point where their effects are no longer optically apparent, or deformation has taken place by Nabarro-Herring creegpiatz c.ci)

Further criteria used to assess the presence or absence of pressure solution in structures.

1. Mineral segregation in buckled sedimentary or metamorphic mineral fabrics. 3. The detachment of small-scale fold hinges along planes of pressure solution, with the associated formation of compositional tectonic striping around the isolated hinges by pressure solution. These features are illustrated in Plate 4.1 . In contrast to the above structure, Plate 4./ shows an example of a buckled rock in which the layers have deformed internally, remaining continuous. Further, the axial plane schistosity is homogeneously developed and is not associated with mineral differentiation. Flate 4.3 Tectonic striping formed by pressure solution, Ardrishaig, Southwest Scotland. 147 3. 'Displacement' of bedding or schistosity across zones of pressure solution. This feature is illustrated in' Plate 4.1 for a buckled layered carbonate in a pelite matrix. 4. Volume loss in cusp structures. Plate 4.1 shows an example of volume loss, by pressure solution of calcite, in cusps and associated planes of pressure solution. 5. Tectonic stylolites in carbonates. 0. Fitting of deformed conglomerates.

Petrographic studies of pressure solution.

The criteria used to assess the presence or absence of pressure solution in structures from a study of thin sections are precisely the same as numbers 1 to 8 enumerated above, but on a microscopic scale. There are two further criteria used from observations made on thin sections, which related to pressure solution in structures. These are -

1. Zones of pressure solution are characterised by a smaller grain-size and strong preferred shape orientation of the pressure soluble mineral, in addition to the condentration of layer silicates and iron oxides (Plates 5.7 0—nci 5 .9 ) 2. A coarse, non-penetrative, cleavage may be present in mono- mineralic rocks, which is defined by discrete zones in which the grains have a preferred shape orientation that develops in response to pressure solution. This structure is discussed and illustrated in Chapter 7.

The information concerning lithology, metamorphic assemblage, deformation mechanism, and tectonic episode, for each locality is compiled in Table 4.3,

LOC,,L1TY REGION 3TOCTUUE-LiTfluLOGY

Skipness SK Cleaved (-nit Q-Alb-Chl-Ms-Ca-Hm-fy Carradale Cleaved grit Q-Alb-Chl-Ms-Ca-Hm-FY Kilchnhtie KINTYRE Crenulated mica schist Ms-Chl-Q-Bi-Hm-ly Clachan CLA Crenulated mica schist Ms-Chl-Q-Bi-Ca-Hm-ly Bilberry KB Folded mica schist Q-Ca-Ms-Bi-Ga-Mg-lyrr Tarbert Crenulated mica schist Q-Ms-Bi-Ga Clachbrech LX I:BAIT/ALE Folded mica schist Achahoish LC Deformed quartzite and pelitos ,-Ms-Chl-Alb-Pm-1;yrr Kilmory ICI Folded carbonates, grits, quartzites Ca-c-Ns-Chl-Alb-Rm-Fy Tayvallich Cleaved grit, carbonate Q-Ms-Alb-Chl -Ca Craignish Taint CK Crenulated quartz - carbonate-petite Ca -Chl -Q-Ms-Hm-Yy Ardfern CRAIGNISH Crenulated quartz-carbonate-pelite Ca -Q-Chl -Ms

Loch Nant LOCH AWE Crenulated quartz and carbonate slates Ca -Q -Ms-Chl Loch Gair Cleaved grit Q -Alb -Ca -Ms-Chl Inverary LF LOCH FYNE Folded quartzite-pelites Q-Ca-Ms-Chl-Bi-Hm-Pyrr Cairndow CA Folded mica schists, grits Q-Ca-Ms-Cld-Bi-Ga-Ng-Hm N. Tichnabruaich TI Crenulated schists Q-Ca-MsrBi-Mg Tichnabruaich Cleaved grit Q-Ca -Alb -Chl-Ms Innellan IN COWAL Cleaved grits and slates Q -Alb-Ca-Ms-Chl -Hym-Hyrr Dunoon Deformed phyllites Chl-Ms-Q Cairletter Point CL Folded quartzite-pelites Q-Ca-Ms-Chl-Fyrr Glen Finart GF Crenulated mica schists Q-Ca-Fld -Ms -Bi -Mg-Hm Ardlochy LL LOCH LOMOND Crenulated mica schists, grits Ms-Chl-Q-Fy-Hm N. Loch Lomond Crenulated mica schists Q-Ms-Bi-Fld Ben Lui Folded mica schists Ms -Bi -¢Ga Aberfoyle Cleaved grits; slates Q-rs -Chl -Alb

Inverlochlairg CRIANLARICH Crenulated mica schists Ms -Bi-Q-Ep Ben Lawyers Folded mica schists Q -Ca-Ms -Bi -Ep-Pl -To Ben Lawyers BL Folded mica schists Q-Ca-Ms-Bi-Ga-Ep-Fl-Mg

Schiehallion LOCH RANNOCH Folded quartz and carbonate mica schists Q-Ca-Ms-Bi-Ep-Ga-St

Tummel TU Folded mica schists Q-Ms -Bi -Ga -1(y -Mg

Ben Vrackie KILLIECRANKIE Folded quarter.lend carbonate mica schists Q-Ms-Bi-Ga-Hb

Wale 4.3 Structure, lithology, mineral assemblage, and peak metamorphic grade for localities studied. The relationship between metamorphic grade and the presence or absence of pressure solution during deformation, through successive tectonic episodes is listed. There is not necessarily a correlation of primary, secondary and tertiary struc-, tures between the localities compiled above.

' indicates presence of pressure solution during deformation. and C- indicate high and low chlorite grade metamorphism respectively. ▪•

PEAK METAMORPHIC RELATIONSHIP BETWEEN METAMORPHIC GRADE AND

GRADE - PRESSURE SOLUTION DURING SUCCESSIVE TECTONIC EPISODES

Primary Secondary Tertiary Chlorite-(C) C PS C PS Chlorite C PS Biotite (B) 0 PS B weak PS Biotite C PS B weak PS Almandine (A) C PS A - C weak PS Almandine C PS A - Biotite C PS B weak PS Chlorite C PS Chlorite C- strong PS C strong PS Chlorite C PS Chlorite C strong PS C strong PS C weak PS Chlorite C PS Chlorite C. PS C PS Chlorite C PS Biotite B weak PS Almandine A - Biotite C PS B Chlorite C PS Chlorite G- strong PS C strong PS C weak PS Chlorite C PS C PS Chlorite C PS C+ weak PS Biotite Be - Chlorite a :Xs C- weak PS Biotite a PS Be - Almandine C PS A - C weak PS Chlorite C PS C PS Biotite C IS Be C weak FS Biotite C PS B weak PS Almandine C PS A - C weak PS

Almandine-Staurolite (AS) C PS AS

Almandine-Kyanite (AK) AK weak PS Almandine-Hornblende (AH) C PS - AH Q = quartz Alb albite Chi = chlorite Ca calcite Hm = haematite Py = pyrite Pyrr pyrrhotite Mag = magnetite Ms = muscovite Bi = biotite Ga = garnet Pi& feldspar PI = plagioclase Ep = epidote To = tourmaline St = staurolite Ky = kyanite 148 Qu rtz-water 6 - 1. 1031n o4 . -3-40 + 3'38 . 10,T (20*-50*) 2. 103lnc4 = -1-96 2.51 . 106T_ (500'0-70Q-C) 3. 1031nnc4 = -2.73 + 3'5? . 10 T 4. 10.11nci. = -3.70 + 4.10 .. 106 T-2 (500°C-800°C) Dolomite-water 6 -2 5. 101n 0( = -2.94 + 3.34 104 T_2 6. 101no( = -3.79 + 3.23 . 10'T6 -2 7. 1031u d. = -3-39 + 3.28 . 10 T Calcite-water o 8. 1031n of = -3.39 2.78 106T2 (0 C-500°C) Muscovite-water 9.10 31n of .-3.89 + 2.38 . 106T 2 (400!C-650°C) 10. 10-11n0( = -3.10 + 1,90 . 106T-2 (500°C-800°C) Chlorite-water 11. 1031n = -4.69 + 1.56 . 106T-2 Haematite-water 12. 1031n0( = -5.08 + 0.75 . 106T-2 Magnetite-water 13. 1031n of . -3.41 - 1.18 . 106T-2 14. 1031n of--/ = -3.70 - 1-47 . 106T-2 (5000C-800°C) 15'.. 10-lino( = -6.38 - 0.13 . 106T2

1. Clayton and others, 1972. 2. Clayton and others, 1972. 3. O'Neil and others, 1973. 4. Bottinga and Javoy, 1973. 5. O'Neil and Epstein, 1966. 6. Sheppard and Schwarcz, 1970. 7. Northrop and Clayton, 1966. 8. O'Neil and others, 1969. 9. O'Neil and Taylor, 1969. 10. Bottinga and Javoy, 1973. 11. Wenner and Taylor, 1971; O'Neil and others, 1973.* 12. Clayton and Epstein, 1961; O'Neil and others, 1969; O'Neil and others, 1973.* 13. Anderson and others, 1971. 14. Bottinga and Javoy, 1973. 15. Wenner and Taylor, 1971. *Empirically derived equations. Table 4.4 Equations for the isotopic fractionations (1031n olk) between some minerals and water, as a fUnction of temperature (106T-2); together with sources of data. 149

RESULTS

Isotope data.

Temperatures obtaining during syn-metamorphic deformation 18 were interpreted from the analysed 0/160 values of mineral separates using the equations listed on Tables 4.4 and 4.5. These equations are presented in graphical form on Pig. 4.5. Temperatures quoted are related to the expression used by a numerical suffix (see Table 4.5). Some additional mineral-pair equations were derived from empirical relations given in the literature; for instance, for A (quartz-dolomite). One of the major advantages claimed for oxygen isotope thermometry is the ability to demonstrate the presence or absence of isotopic equilibrium in a given mineral assemblage, from three or more analysed co-genetic minerals (Taylor, 1967). However, this method requires an accurate calibration of mineral- water fractionations as a function of temperature. A high degree of isotopic disequilibrium may be inferred from the data for cogenetic quartz-muscovite-iron oxides and quartz-chlorite-iron oxides, at the majority of localities sampled in the Dalradians (using equations given in Table 4.5). For each of the samples analysed there was petrographic eviden&e for equilibrium relations in the mineral assemblage. Hence, one interpretation which may follow from these isotopic data is that the mineral pair fractionation equations are incorrecti R.D. Beckinsale has come to a similar conclusion in connection with isotopically analysed muscovite and chlorite minerals from the Eastern Alps, that were interpreted with respect to equations 1 and 3 respectively, Table 4.5 (personal communication). 150 6 2 1. A ,(quartz—muscovite) + 1.16 + 1.19 . 10 T 6 -2 2. A (quartz-muscovite) = - 0-60 + 2.20 . 10 T 6 -2 3. A (quartz-chlorite) = + 1.96 + 2.01 . 10 T 6 -2 4. A (quartz-calcite) + 0'66 + 0'79 . 10 T 6 -2 5. A (quartz-magnetite) = + 3.66 + 3.70 . 10 T -2 6. Q (quartz-magnetite) . 0 + 5.57 . 10T 7. A (quartz-haematite) .+ 2'35 + 2.82 . 106T-2

1. O'Neil and others, 1973; O'Neil and Taylor, 1969. 2. Bottinga and Javoy, 1973 3. O'Neil and others, 1973; O'Neil and others, 1969. 4. O'Neil and others, 1973; Wenner and Taylor, 1971. 5. O'Neil and others, 1973; Wenner and Taylor, 1971. 6. Bottinga and Javoy, 1973 7. O'Neil andthers, 1973; O'Neil and others, 1969; Clayton and Epstein, 1961.

Table 4.5 Equations for the isotopic fractionations (A) between some mineral pairs as a function of 106T=2; together with sources of data. 151

With the above considerations in mind new equations for muscovite-water and chlorite-water fractionations were derived as follows -

1. Temperatures were computed from quartz-iron-oxide fraction- 18 16 ations (for which there was also a 0/ 0 value for muscovite), and these temperatures assumed to be correct. 2. TheA(muscovite-iron oxide), and Z,(quartz-muscovite) fractionations were plotted as a function of 106T-21 using the temperatures as computed above. 3. A best fit line was constructed through the data points, and an equation calculated for the mineral-water fractionation.

The two new equations presented by the author for muscovite-water and chlorite-water fractionations are -

1. 1031nck(muscovite-water) 1.3%. 106T-2 3 2. 10 1nck(chlorite-water) =--1•10 -1- 140 106T-2

Use of these equations gives concordant isotopic equilibrium for the cogenetic minerals analysed at most of the localities sampled, because the equations have been derived with that assumption.- However- one may infer isotopic disequilibrium between co-existing minerals in some instances, for example at Tummel (TU). The isotopic composition of water in equilibrium with the mineral assemblages during metamorphism varies between 18 16 o/ 0 + 11 and + 660 (Fig. 4.6). These data are consistent with the isotopic composition of water expelled during prograde metamorphism of a sedimentary pile (O'Neil and Clayton, 1964; Epstein and Taylor, 1967; Garlick. and Epstein, 1967; Devereux, 1968; Schwartz and others, 1970; Eslinger and Savin, 1973). 152 12

0 4 6 -2 10 T H-W.1 2 -2

-4

-6 M-W 15

-8 M-W 14

Fig. 4.6 Mineral-water fractionations plotted as 103 Ln ck against 106 T4 °K . The numerical suffix relates to equations listed in

table 4.4 0-0 Neil and others,1964 ; Ma-Matthews and others,1973 F—Freiderichson , t173 153 18 There is a slight relative depletion in 0 of water in equilibrium with the higher grade assemblages, as compared to the chlorite grade rocks. A similar trend in the 180/160 values of auartz is present (Fig. 4.6). It is suggested that t 1Dcl. of intern,?1 consistency for sore of the data may be re ult the lower Trne assembla es equilibratin7 'tip 130 depleted water. Tbe depletion may be due to excIlanr,e of c=en bete7:n _t`c,-er and rocs of - metamorphic :7rade, at 180 These trends in the depletion of water and quartz have been obb-:ino'l by othe- author studies of metamorphic environments (Ofeil an-2 Clayton, 19..54; Epstein and Taylor, 1907; Garlic"; and Epstein, 2957; Devereux, 1968; Schwarcz and otbers, 1970; Eslinger and ';'ovin,. 1973).

The thermal environment of pressure solution.

With recard to the specific features investigated concerning the thermal environment of pressure solution as a deformation mechanism, the interpretations placed by the author on the data, as listed in Tables 4.3 and }ELI , and Fig. 4.7, are as follows -

1, 2. Quartz and calcite may deform by pressure solution at temperatures up to 450 ± 50°C and 350 = 50°C respectively.

Below $'1, 300°C pressure solution is commonly the dominant mechanism by which quartz and calcite respond to deformation. One may infer that for the rare examples of muscovite and chlorite which have deformed by pressure solution that temperatures were less than 450 - 50°C. These statements are qualified for effects of strain rate and grain size in chapter 5. 3. At a given metamorphic grade the presence of some phyllosilicates increases the extent to which auartz and calcite 154 deform by pressure solution. In addition, there is an increase in the size of grains which can deform by pressure solution, relative to rocks with no phyllosilicates. Further, the presence of - micas raises the temperature limit at which , quartz and calcite deform by pressure solution by 50°C. This point is dfiscuseed in chapter 5. 4. At 300°C grains of less than 100 to 1000 microns in

quartzite deform, at least in part, by fressuro- so L ion. At 500°C this size limit is reduced to > 10 microns. 5. Mean stress does not appear to exert a strong influence on the transition temperature above which quartz or calcite respond to deformation largely by pressure solution. The transition temperature for quartz was estimated to be 450 - 50° 0 under conditions of both 5 Kb and 2 Kb,(i.e. lower biotite to lower almandine, and lower biotite to lower andalusite isograds of the low and intermediate pressure facies series of regional metamorphism).

Metamorphic P.T. conditions.

6. Estimates for the temperature limits of metamorphic isograds, within the intermediate pressure facies series of regional metamorphism, made by the author from isotopic data are as follows -

Chlorite 300°C (± 10) - 400°C Biotite 400°C (t 15) - 5000C Almandine >500°C (± 22) Staurolite > 560°C Kyanite >580°C (± 30) Sillimanite >625°C .155 16

IN CK KI LL SK TI LX GF BL CA TU

14 -

12 -

10-

0 - 305 338 358 373 384 422 424 -474 484 532 588 low high garnet kyanite - chlorite chloteri biotite garnet

o quartz • muscovite A chlorite • haematite • magnetite ✓water

Fig. 4.7 180/0 16 values of coexisting minerals from

individual localities sampled in the Dal radian Series.

Temperatures in °C, and metamorphic grade are indicated, 156

The author's data (see Figs. 4.7 and 4. 8) are broadly consistent with the PT field for the intermediate pressure facies series suggested by Turner (1967) and Miashiro (1970). The data are not in agreement with PT conditions for that facies series as contended by Winkler (1967) and Den Tex (1971), who quote temperatures -,.100°C greater than those determined above (Fig. 4. 8), for the thermal stability limits of some index minerals. Devereux (1968), using oxygen isotope thermometry, obtained results similar to those of the author for the temperature range of the chlorite and biotite zones.

7. The prevailing conditions of temperature and pressure during metamorphism at Innellan and Kilmory are estimated to be ^a300°G and 4 Kb. One may infer from data obtained on fluid inclusions in tectonic veins at those localities, that the value of X (fluid pressure / overburden pressure) obtaining during deformation was between 0.7 and 1'0 (i.e. the fluid pressure was between 2.8 Kb and 4.0 Kb).

General observations.

No evidence was found from which it could be inferred that pressure solution is an important deformation mechanism at almandine, and higher grades, of metamorphism. However, two rare examples of small quartz pressure shadows on syntectonic garnets in a fine-grained matrix were collected from within the lower almandine isograd. In some rock systems diffusion may be the rate limiting step for pressure solution. Under these conditions the path length over which deformation by diffusion is faster than deformation by intracrystalline flow decreases from several cm at ,v 30000 to a few microns at A■550°C. This point is discussed further in chapter 5. 157 1000

2 - Pyroxene facies group

BOO 2

Amphi balite facies group -1 600 till; man; he

almonclic.4 Greenschist facies gro

loietits Glctucophane schist 400 facies group C for at

L-P-P -L facies group 200

5 10 15 Depth in Kb

1 Haptogranitic water-saturated liquidus -solidus

2 7 Opx + 1 Ab + 2 An +1W = 1 Ho 4. 1 Alm + 5)0 3 2 Zo + 1 Py + 6W = 4 La + 2Q 4 1Py + 1W = 1 Kcto +2Q

Fig. 4 .8 P T graph of the facies groups, with equilibrium curve s for boundary reactions. Adapted from Den Tex 1971 . The authors data on the thermal stability limits of some index minerals is superimposed in heavy type. 158

During deformation of high grade metamorphic rocks at low temperature the transition-temperature for pressure solution would appear to be the same as for the prograde case. Towever, in the former instance a much lower intensity of pressure solution is observed. This plisnoncnon E.nly be due to the cu_sos

1. Lar7,er grain size (e:,:cert for the case of m7lonites, etc.). 2. A lower water content he assist d-iffusior. 3. The reel: has lower bull: anisotropy, and hence is loss susceptible to bucklinL instabilities which generally ini tiate pressure solution. 159

4.5 127=TIC =ECM TIT STRESS-IEDUCED DIF7TSTC7AL =DS.

Introduction

In a7n7 non-eQuilibrium process isotopically lis-:hter atoms react faster than the heavy atoms. Such processes display phenomena are collectively known as kinetic isotope effects. The separation of isotopic species is a highly developed field of modern technology. However, in the earth sciences, kinetic isotope effects in solid state processes are not well understood, and rate constants have been established for only a few reactions. During the bacterial reduction of sulphate to sulphide in sedimentary diagenesis, the ratio of rate constants, K, for 32 production of SO4- and 34504 has been measured as 34 32 32 K2 W/K1 S =- 1.020. The rates of oxidation of 5 and Fe32S are 0.1% faster than for 34S and Fe34S2 respectively (Naki, 1964). Kinetic fractionation of oxygen and carbon isotopes occurs in some oxidation reactions and the growth of algal mats, respectively (Nizutani, 1969; Weber, 1970; Seckbach and Kaplan, 1973). At deeper crustal levels, Jensen (1967) has invoked kinetic effects to account for gradients in the concentration of 32 H2 S in country rock, away from hydrothermal systems in which H2S is produced. Sheih and Taylor (1969) have contended that kinetic fractionation of 160 and 12C has taken place in decarbonation reactions of calc-silicate minerals in contact metamorphic skarns. Kinetic isotope effects will not be preserved if minerals re-equilibriate with a fluid phase, or if the reactants of a 160 given process are completely converted into products. In order that kinetic fractionations are retained in a given system it is a prerequisite condition that the products be continuously removed from the reactants.

Pine tic ir,otol-e effects in tectonic environments.

The central assumption of many studies within the field of oxygen isotope geochemistry is that equilibrium isotooe fractionation has been achieved between minerals and fluid ox en reservoir. One may infer from examples cited above, that this assumption does not always hold. Whilst studying the diffusion of minerals in response to stress, the author predicted on theoretical grounds, from consideration of the pnysics of diffusion, that kinetic isotope fractionation should take place in differential stress fields. This process may be understood in simple qualitative terms as follows: when minerals diffuse down a chemical potential gradient, which has been established in response to'stress differ- entials, ions with isotopically light atoms have greater mobility than those with heavy atoms, and as a result will dif7• fuse faster.-- Fractionation in a diffusional process may be considered on a theoretical basis with reference to the Nernst-Einstein relation, for the mobility (p) of a diffusing particle

va2exp (-Em/kT)(q/kT)

where v is the vibrational frequency of the diffusing particle; a is the path length of an activated jump; Em is the activation enthalpy for diffusioh; k Boltzmann's constant; T absolute temperature; and q is the charge on the particle. 161 This formula has universal validity for a particle diffusing in an external force field Nett and Thompson, 1970). As we have seen, isotopic species of a given element have different vibrational frequencies (Fig. 4.1). It is this difference in the frequencies (as included in the Nernst-Einstein relation) which results in a separation of the diffusion rates of isotopically light and heavy atoms. In order to test the author's hypothesis of a stress-induced kinetic fractionation, and to establish the magnitude of the effect, samples were selected for isotopic analysis from the deformed matrix and pressure shadows around rigid undeformed particles. The reasons for sampling these structures are as follows -

1. Strong gradients of differential and mean stress are present in the matrix around rigid inclusions during deformation. The form of these fields has been established analytically, and from numerical techniques. 2. The relatively short diffusion path between defined. zones of associated mineral gain and loss, is a situation which offers the closest available approach known to a chemically closed system. 3. Kinetic fractionations are most likely to be retained in coarse-grained minerals from pressure shadows.

Two sets of samples were collected from the zone of depletion, and the associated area of gain in pressure shadows,

- at the following localities -

1. Dolomite around clay pellets in Trias Marls of the Aar parautochthon. 162

1. Dolomite pressure sUadow, Aar (PDB)

°"1 -4.50

-2.45 -2.09 C33

-2.93

2. Quartz pressure shadow, Trebar with Strand (SHOW)

23-14

23.14 03

22.25

3. Quartz pressure ,shadow., N. Wales (SNOW) a, 14'99

14.61 15.93 0/

Fig. 4.10 180/ 16 0 values of analysed quartz and dolomite from areas of high stress and loss stress in the deformed matrix around rigid inclusions. 163

2. Quartz from around euhedmi pyrites in deformed Carboniferous vclaanics, Trebarwith Strand, North Cornwall. 3. Quartz adjacent to euhedral pyrites in Lower Cambrian green slates, Penrhyn, North Wales.

The third case was taken as a control sample because there did not appear to be any evidence of quartz depletion in regions of high stress, from which the pressure shadow had been formed.

RESULTS_

In low stress areas dolomite is isotopically lighter than in regions of high stress; thus the shadows have a high content of light 160 atoms. However the magnitude of differences are not consistent (Fig. 4.10). For the second case, of quartz adjacent to euhedral pyrites, the shadows are 3 to 4°/0O lighter than in zones of quartz depletion (Fig. 4.(0). As for the dolomites above, the magnitude of differences are not consistent. This may 4e a function of pressure shadows developing asymmetrically about the rigid inclusions. No petrographic evidence was found that pressure shadows had formed by diffusive transfer of quartz out of high stress areas, in.' the samples analysed from North Wales. Quartz in 16 pressure shadows is depleted in 0 relative to the host rock; hence the results for this control sample are in contrast with data obtained for the two examples described above. The implications of this result are discussed in Chapter 5. One may infer from data obtained from the first two samples that kinetic fractionations of oxygen isotopes may take place in 164 stress induced diffusive processes. However, it has not been possible to obtain sufficient data from which an answer to this problem may be given with a statistically defined level of confidence. It is clear that either the effect is small, or that minerals re-equilibriste after the diffusionel rrocess has tor,red. Carbonate minerals are more susceptible to re- eouilibriation than is quartz. This factor may explain the greater jouoric differences of sa le 2 in comparison 11,t,11 sample 1. With these considerations of kinetic effects in mind, spe-

cial care was observed during the field work not to collect samples for isotopic analysis from within structures with which there is associated a powerful stress gradient, or mineral segregation. For non-equilibrium reactions in anhydrous conditions it is suggested that kinetic effects may be frozen in more strongly than is the case for minerals which can exchange with a fluid, oxygen reservoir. However, at the higher temperatures associated with many anhydrous mineral assemblages oxygen may exhibit a significant degree of self diffusion in oxygen bearing minerals. Plate 4.E Fibrous quartz and calcite in a tectonic vein. Ouartz- calcite host rock, Marloes, Pembroke.

Plate 4.7 A tectonic vein composed of quartz, calcite, and chlorite, in a quartz-calcite- chlorite host rock, Marloes, Pembroke. 165

4.4 OXYGEN AND CARBON ISOTOPE GEOCHEMISTRY OF TECTCNIC VEINS.

Method.

Samples were selected for which there was field evidence of deformation by pressure solution in the host rock and associated vein filling in a fracture. It was necessary that both vein and host rock contained a given mineral pair, e.g. quartz-calcite, ro provide an isotopic thermometer (plates 4.6 and 4.7). In addition, the criteria that governed the selection of material for fluid inclusion work were applied to this aspect of the research in order that a comparison could be made between temperatures derived from inclusion homor7enisation and oxygen isotope methods.

RESULTS

Temperatures obtained from ouartz-iron oxide pairs in the host roc:s are consistent with temperatures obtained from fluid inclusions in ouartz. The temperature differences yield fluid pressures obtaining during deformation of 1 to 3Kb, depending on the locality (rig. 4.9). Cuartz carbonate pairs in the veins and host rocks do not yield meaningful temperatures. Carbonate minerals are significantly and systematically enriched in 180 compared with quartz. However, equilibrium of oxygen and carbon isotopes has taken place between carbonate minerals in the vein and host rock. Two suggestions are made by the author to explain the isotopic disequilibrium between host and vein carbonates - 166

PV CM ND TV LOT h v h v h v h v h v

219°C 317°C 340°C' 276°C 97C

quartz •C) calcite 411' calcite (carbon isotopes-PDB) A dolotite" h - host V dolomite(carbon isotopes-PDS) v - vei n m chlorite muscovite x haematite

18 16 Fig 4.9 0/ 0 values of mineral separates analysed from veins and host

rocks in tectonic environments. The second syMbols for calcite and

dolomite are carbon isotope ratios referred to PDS standard. 167

1. Pressure solution of carbonates in the host rock and associated vein filling took place subsequent to, and at lower temperature than, precipitation of the vein quartz. 2. The carbonates were precipitated by low temperature hydrothermal fluids, subsequent to the quartz filling. Carbonates in the host rock achieved isotopic equilibrium with the hydrothermal fluids.

The first hypothesis is preferred because of the field evidence for pressure solution of carbonates, and because of the high degree of isotopic equilibrium achieved between host and vein carbonates (Fig. 4.9).

There is a systematic enrichment of ,-.J.1°/o°- of vein quartz compared to quartz in the host rock. This result may be explained in two ways -

1. Vein quartz was precipitated from hydrothermal fluids that were hotter than the host rock, throughout the period of vein formation (by ", 25°C). 2. Vein quartz equilibriated with fluids that were 180 enriched in comparison to fluids with which quartz in the host rock equilibriated.

One may infer from the field evidence that the source of vein quartz lies in the adjacent host rock, and is not of hydrothermal origin. In addition, it is unlikely that a 25°C thermal gradient would obtain over a few, ‘centimeters between hydrothermal fluids and the host rock during long periods of time_1(see-_discussidmin-Chapter 7). Hence I contend that a steady expulsion of water from sediments in tectonic environments takes place during deformation. As a consequence, minerals in host rocks will equilibriate with fluids initially in the pore spaces. Subsequently, fluids expelled from greater 168 depths will pass through fracture systems, and vein quartz will equilibriate with such deeper seated hydrothermal fluids. 18 The inferred 0 enrichment of hydrothermal fluids with which the 'vein quartz equilibriated is consistent with the isotopic character of fluids that 1:',:),ve ori,7,inatod from deeper o-,Id ' V e _tore f^ :ironments.

7inolly, ono 2 note tht bec'ause ouartz in the

1?C017; 71'S not e,7ujli'-)riated -it'zz the inferre& e71-r,ichd hydrothermal fluids I.:list the ratio of fluids to rock was low in all tl2e tectonl envi:non :rents studied. TZence, one may concluae that the 1=2e volumes of vein quartz in these t2c- tonic environments was not precipitated from still larr:er volumes of hydrothermal fluids flowins throuyh the system. 18 The calculated 0/160 ratios of hydrothermal fluids in equi- 30/ to 100/ . librium with the vein quartz varies from 4. 00 0o These data are consistent with the isotopic character of fluids expelled from low to intermediate grade metamorphic 16 rocks; but are definitely not consistent with 180/ 0 ratios typical of meteoric or magmatic waters. 369

REFERENCES - CHAPTER 4

ANDERSON, A.T., R.N. CLAYTON.. and T.K. MAYEDA. 1971. Oxygen isotope thermometry of mafic igneous rocks. Jour. Geology, v. 79, 715. BIGELEISEN, J. and. M.G. MAYER. 1947. Calculation of equilibrium constants for isotopic exchange reactions. J. Chem. Phys., v. 15, 261. BOTTINGA, Y. and N. JAVOY. 1973. Comments on oxygen isotope thermometry. Earth and Planetary Science Letters, v. 20, 250-265. BROECKER, R. and V. OVERSBY. 1971. Chemical equilibria in the earth. McGraw Hill, New York, 318 p. CLAYTON, R.N. and S. EPSTEIN. 1961. The use of oxygen isotopes in high-temperature geological thermometry. Jour. Geology, v. 69, 447-452. CLAYTON, R.N., J.R. O'Neill and T.K. MAYEDA. 1972. Oxygen isotope exchange between quartz and water. Jour. geophys. Research, v. 77, 3057-3067. CRAIG, H. 1961a. Isotopic variations in meteoric waters. Science, v. 133, 1702-1703. CRAIG, H. 1961b. Standard for reporting concentrations of deuterium and oxygen-18 in natural waters. Science, v. 133, 1833-1834. DEVEREUX, I. 1968. Oxygen isotope ratios of minerals from the regionally metamorphosed schists of Otago, New Zealand. New Zealand Journal of Science, v. 11, 526-547. EPSTEIN, S. and H.P. TAYLOR. 1967. Variations of 180/160 in minerals and rocks. In J. Wiley. New York. Re- searches in geochemistry, v. 2, 29-62. ESLINGER, E.V. and S.M. sAvir. 1973. Oxygen isotope geothermometry of the buerial metamorphic rocks of the Pre- cambrian Belt Supergroup. Geol. Soc. America Bull., v. 84, 2549-2560. GARLICK, G.D. and S. EPSTEIN. 1967. Oxygen isotope ratios in coexisting minerals of regionally metamorphosed rocks. Geochim. et Cosmochim. Acta, v. 31, 181-214. JENSEN, 1967. Sulphur isotopes and mineral genesis. In 7.T, Barnes (ed.), The geochemistry of hydrothermal ore deposits. Holt, Rinehart and Winston. New York. Nail N. and M.L. JENSEN. 1954. The kinetic isotope effect in the bacterial reduction and oxidation of sulphur. Geochim. et Cosmochim. Acta, v. 28, 1893-1912.

170 NORTHROP, D.A. and R.N. CLAYTON. 1966. Oxygen isotope frac- tioation in s7stems containing dolomite. Jour. Geology, v. 74, 147-196. O'NEIL, J.R. and R.:7. CLAYTON. 1934. Oxygen isotope geo- thc=emtry. Tn F. Filler, S.L., 7Tas,:erl)ur7, G.J. (e.(7.), Cosmic and isotopic chemistry. Forth Company, Amsterdam. 553 p. C'E:IL, J.R. and S. E]'072=7. 1966. A method for ox-yen iso- to e rrnzT117,,si.s of milli ram coantities- of water and some of its applications. Jour. Geophys. Research, v. 71, 4955-4961. • O'NEIL, J.R. and H.P. TAYLOR. 1969. Oxygen isotope equilibrium between muscovite and water. Jour. Geophys. Re- search, v. 74, 6012-6022. O'NEIL, J.R., R.N. CLAYTON and T.K. MAYEDA. 1969. Oxygen isotope, fractionation in divalent metal carbonates. Jour. Chem. Phys., v. 51, 5547-5558. O'NEIL, J.R., M.L. SILBERMAN, B.P. FARBI and C.W. OHESTERMAN. 1973. Stable isotope and chemical relations during mineralisation in the Bodie Mining District, Mono Country, California. Econ. Geol., v. 68, 765-784. RUSHBROOKE, T.S. 1949. Introduction to statistical mechanics. Oxford University Press. London. 334 p. SCHWARCZ, H.P., R.N. CLAYTON and T.K. MAYEDA. 1970. Oxygen isotope studies of calcareous and pelitic metamorphic rocks, New England. Geol. Soc. America Bull., v. 81, 2299-2316. SHEIH, Y.N. and H.F. TAYLOR. 1969. Oxygen and carbon isotope studies of contact metamorphism of carbonate rocks. Jour. Petrology, v. 10, 307-331. SHEPPARD, S.N.F. and H.P. SCHWARCZ. 1970. Fractionation of carbon and oxygen isotopes and magnesium between coexisting metamorphic calcite and dolomite. Contr. Mineral. and Petrol., v. 26, 161-198. TAYLOR, H.E. 1967. Oxygen isotope studies of hydrothermal mineral deposits. InBarnes, H.L. (ed.), The geochemistry of hydrothermal ore deposits. Holt, Rinehart and Winston, New York, 109-142. TRUESDELL, A.H. 1974. Oxygen isotope activities and concentrations in aqueous salt solutions at dlevated temperatures: consequences for isotope geochemistry. Earth and Planetary Science Letters, v. 23, 387-396. TURNER, F.J. 1968. Metamorphic Petrology. IcGraw-Hi11. New York. 403 p. UREY, H.C. 1933. The separation and properties of the isotopes of hydrogen. Science, v. 78, 566-571. 171

UREY, F.C. and L.J. GRIP. o935. Isotopic exchange equilibria. Am. Chem. Coe. Jour., v. 57, 321-327. F.C. and D. RIT.T747)U1r., 1935. Some thermodynamic pro- porti of t7.-(A 717(-=, YY moleculen ar:1 comroun,ls cor- tainins7 -ne T.7_ atom. Jour. Cheri. Mys., v. 1, 137-23. wErrm, D.7. ard F.D. TOTCR. 1-71. Terpel'at=esof s4,-nrcAn- tinistion of ultramafic rocks bcsed on J-c)0/1-=-0 fracon- ation betlecn corentine and 7c-Tletite. Contr. Linerca. and letrol., v. 32, 165-105. =1, C.A. and 1WO. nysics 01 solids. Ttew York. 522 p. TI.-.F. 1267. Petro7enesis of metamorphic rocks. SprinTer-Verla;. 1.:e York. 257 y. 172

Dalradians of the Central and Southwest Highlands• of Scotland

ALLISON, A. 1940. Loch Awe succession and tectonics: 1iilmartin-Tayvallich-Danna. Quart. J. geol. Soc. Lond. v. 96, 423- BAILEY, E.B. 1922. The structure of the Southwest Highlands of Scotland. Quart. J. geol. Soc. Lond. v. 72, 132- BORRADAILE, G.J. 1970. The west limb of the Loch Awe Syncline and the associated cleavage fan. Geol. Mag. v. 107, 459- BORRADAILE, G.J. 1972. Variably oriented co-planar primary golds. Geol. Mag. v. 109, 89-98. BORRADAILE, G.J. 1973. Dalradian structure and stratigraphy of the northern Loch Awe district, Argyllshire. Trans. Roy. Soc. Edinb. v. 69, 1-21. G.J. , J.I. RCBERTS and J.E. TREAGUS. 1971. The Loch Nant area, Argyllshire, and its significance in the structural history of the Loch Ae Syncline. Scott. J. Geol. v. 79, 79-85. BROWN, E.E. and others 1965. Potassium-arson ages of some Dalgadian, Maine and related Scottish rocks. Scott. J. Geol. v. 1, 1/0-151. CLOUGH, C.T. 1897. In Gunn. W. and others. The geology of Cowal. Mer. Ge 1. Surv. Scotland.

CUMMINS, W.A. and R.F. SHACULETON. 1955. The Ben Lui recumbent synOline (S.W. Highlands). Geol. Mag. v. 92, 355-363. DEWEY, J.F. and R.J. FANKHURST. 1970. The evolution of the Scottish Caledonides in relation to their isotopic age pattern. Trans. Roy. Soc. Edinb. v. 68, 361-389. HARRIS, A.L. 1963. Structural investigations in the Dalradian rocks between Pitloehry and Blair Atholl. Trans. Edinb. geol. Soc. v. 19, 256-278. 1 7 3

HARRIS, A.L. AND N. RASP. 1960. The evolution of quartz fabrics in the metamorphic rocks of Central Perthshire. Trans. Edinb. geol.Soc. v. 18, 51-78. HARTS, B. and N.L.W. JOHNSON. 1969. Metamorphic history of Dalradian rocks in Glens Cleva,-Esk and Lethnot, Angus, -Scotland. Scott. J. Geol. v. 5, 54-80. JOHNSON, Y.P.W.- 1962. Relations of movement and metamorphism in the Dalradians of Banffshire. Trans. Edinb. geol. Soc. v. 19, 26-64. JOHNSON, m.a.w. ,1963. Some time relations of movement and metamorphism in the Scottish Highlands. Geol. Eiinbouw, v. 42, 121-142. JOHNSON, ri.a.w. 1965. Dalradian. In Craig, G.Y. (Ed).. The GeoloFy of Scotland. Oliver and Boyd, Edinburgh. KING, B.C.- and N. 1955. Tectonic styles in the Dalradians and Moines of part of the Central Highlands of Scotland. }roc. Geol. Ass. Loud., v. 66, 243-269. KING, B.C. and N. RAST. 1956. The small-scale structures of south-eastern Cowal, Argyllshire. Geol. Nag., v. 93, 185-195. KNILL, J.L. 1960. The tectonic pattern in the Dalradian of the Craignish-Kimelfort District-Argyllshire. Quart. J. geol. Soc. Lond, v. 115, 339-364. RAST, N. 1958a. Tectonics of the Schichallion complex. Quart. J. geol_. Soc.Lond., v. 114, 25-46. RASP, N. 1958b. The metamorphic history of the Schichallion complex, Perthshire. Trans. Roy. Soc. Edinb., v. 63, 414-431. RASP, N. 1963. Structure and metamorphism of the Dalradian rocks of Scotland in The British Caledonides (Eds) Johnson, M.R.W. and Stewart- Oliver and Boyd, Edinburgh ROBERTS, J.L. 1959. Fold structures in the Dalradian rocks of Knapdale, Argyllshire. Geol. Mag. v. 96, 221-229. 174

ROBERTS, J.L. 1966c. The formation of similar folds by inhomogenous plastic strain, with reference to the fourth phase of deformation affecting Dalradian rocks in the south-west Highlands of Scotland. Jour. Geol. v. 74, 831-855. ROBERTS, J.L. 1974. The structure of the Dalradian rocks in the S.W. Highlands of Scotland. Quart. J. geol. Soc. Lond. v. 130, 93-124. ROBERTS, J.L. and J.E. TREAGUS. 1964. A reinterpretation of the Ben Lui fold. Geol. Na. v. 101, 512-519. SHACKLETON, R.M. 1958. Downward facing structures of the Highland Border. Quart. J. geol. Soc. Lond. v. 113, 361-392. STURT, B.A. 1961. The geological structure of the area south of Loch Tummel. cuart. J. zeol. soc. Lond. v. 117, 11- 156. STUaT, and xl.L. Ela2TS. 1961. The metamorphic history of the Loch Tummol area, Central Perthshire, Scotland. Jour. Geology v. 2, 689-711. T=GUS, 1964. 1:otes on the structure of the Ben Lawyers synform. Geol. r . v. 101, 260-270. VOLL, G. 1960. New work in petrofabrics. Jour. Geology v. 2, 503-567

4 175

Chapter 5

PRESSURE SOLUTION AND DEFORMATION MECHANISM FIELDS

This chapter is concerned with the nature of deformation processes in the crust. A brief introduction is given to the different mechanisms by which natural rock systems may deform. This is followed by a discussion of specific aspects of tress-. ure solution. A detailed account is presented of the deformation mechanisms in tectonic veins, in relation to temperature. This information is combined with data obtained for the limits of deformation modes in Dalradian rocks (see chapter 4), together with additional field observations from other localities, to give a quantitative description of deformation mechanism fields. Finally, conclusions are drawn regarding some general aspects of physical conditions obtaining in tectonic environments.

Deformation mechanisms in rocks.

The polycrystalline aggregates that constitute natural rock systems may deform by a number of different mechanisms. These mechanisms may be divided into intracrystalline and intercrystalline processes. The five main modes by which rocks are considered to deform are illustrated in Fig. 5.2. The first intracrystalline mode is rabarro-Herring creep, a rechanism wI-.ich accounts for the shape change of crystals by flow of point defects into regions of high stress. This is essentially a high temperature mechanism (0.8 T melting), but there is no unequivocal evidence that rocks deform by this mechanism, even within the mantle, where one may anticipate temperatures in excess of 1000°C (Raleigh and Kirby, 1973; Stocker and Ashby, 1973). 176 INTRACRYSTALLINE

Herring- Nabarro creep Dislocation creep

IN TERC RYSTALLINE

Grain boundary sliding Grain fracture

Pressure solu t ion

-.constant mass mass toss

Fi g 5 .2. Schema tic diagram of major deformation mechani sms in rocks 176 Dislocation processes include a number of intracrystalline processes. These may be conservative and athermal, such as slip; or non-conservative and thermal, such as climb. Strain energy in crystals that results from dislocation flow may he released by intracrystalline subgrain formation, which is a coherent process, or by recovery which is an incoherent 'orodess. 1s aadit-Thn to dilocation mchanisms there are other modes of intrac7- ele'orma';ion GUCh as ts:inin and kinking, by which a limited amount of strain may tape place. IntercrystFaline mechanisms fall into three groups. Grain boundary sliding describes the change in mutual spatial relations of Trains by frictional sliding and rotation. Cataclastic deformation involves the brittle fracture and rupturing of individual grains into smaller units. There is some confusion of terminology in relation to these mechanisms. In the metallurgical literature grain boundary sliding is understood as a high temperature mechanism, which is not necessarily pressure dependent, whilst in the geological literature it is often implicitly taken as part of cataclasis, which is dependent on effective confining stress. confusion is further extended because it would appear that what are known by the field term 'cataclasites' are not necessarily rocks formed by grain fracture, but by intracrystalline dislocation flow leading to tectonic grain refinement. Hence, grain boundary sliding is used here to describe frictional sliding between grains, whilst cataclasis is taken to imply grain fracturing. Under certain conditions the two processes are linked. Pressure solution, or grain boundary diffusion is - the third intercrystalline mechanism. This is the one mechanism which may lead to systems that are either chemically open or 177 closed during deformation, because of intercrystalline diffusion. In the latter case diffusion is limited to individual grains, and hence grains change shape at approximately constant mass. In the former case shape change takes place with loss of mass from the system or gain into the system. In this chanter attention is directed towards a Quantitative evaluation of the physical conditions in the crust which determine the nature of the -oredont deformation mechanist-- 1 7 8

5.11 PRESSURE SOLUTION AS A DEFORMATION MECHANISM.

Introduction to aspects of pressure solution.

A number of questions arise in relation to rocks which defcrn beT.,resrture soltion -

1. °!h_ minerals deform by thts mole? 2. Over u' at -'a-ces does dif-. ten tal:e . fac:-)ors are responsible for initiatinz fresq,Lre- solution? 4. -That are the physical conditions necessary for pressure solution?

The nature of pressure solution in natural rock systems is com;:lex. f-uar.tz and calcite are the two minerals which re- sIDond most readily to deformation by pressure solution. In addition, they are two of the most abundant rock forming minerals. Pressure shadows of other minerals such as chlorite, muscovite, and feldspar are observed less frecuently. No evidence has been obtained by the author that magnetite or pyrite respond to deformation by pressure solution. In general there is a well defined sequence in the relative pressure solubility of minerals (see chapter 2); however, anomalous cases have been observed. Plate 5.1 shows the sym- metrical development of quartz, chlorite, and finally muscovite in the pressure shadow of a detrital calcite grain. The rock is a chlorite-muscovite-calcite-quartz slate. During deformation the mutual pressure solubility relations of these minerals must have changed. No convincing explanation for this phenomena can be advanced by the author; although a significant change may have taken place in the physical conditions obtaining during an incremental strain history. elate 5.1 Successive p;rowths of quartz, chlorite, and calcite, in the pressure shadow of a detrital quartz grain. Quartz-calcite-muscovite-chlorite slate, qouth Devon.

Plate 5. ;)etrl:ai- rains of calcite wltr• pressure shadows, in a fine-cfrained calcite mylonite. Lochseitenkalk mylonite, Glarus, Switzerland. 179 The diffusion of minerals by pressure solution may take place over distances from a few microns up to several. meters. The deformation of individual grains in a slate is representative of the lower end of the scale, whilst migration into veins and the formation of large scale tectonic striping is typical of diffusion over larger distances. It has been established empirically (chapter 4) that the distance over which diffu6ion by pressure solution is viable (i.c. favourable in relation to other deformation mechanisms) is temperature dependent. Pressure solution is frequently associated with perturbations in the stress field. There may be several other factors which are responsible for initiating pressure solution activity. One of these factors is fluctuations in grain size. Plate 5.1 is a photomicrograph of a fine grained pure calcite mylonite. Two lare grains which have escaped tectonic grain refinement characteristic of the matrix have well developed calcite pressure shadows. Although the relative pressure solubility of calcite in the matrix and the detrital calcite grains must be equal because they have the same composition, matrix grains only have deformed by pressure solution and diffused into the pressure shadows. Variations in grain'size effectively cause spatial fluctuations in the rate of pressure solution. Variations in the contents of mica, for instance, will cause fluctuations in the rate of pressure solution. These aspects of grain size and mineralogy are discussed more fully in the sections on lithology and grain size effects. Perturbations in the stress field, grain size, or mineral content tend to occur randomly, at loci, or in discrete zones. This factor is fundamental to initiation of structures by pressure solution in localised domains of loss or in cleavage :t late 5. elite defnrming internally; and externally, by pressure solution along stylolites, together with forma- nf rressure shadows. Ilfracombe, North Devon.

Plate 5.4 Photomicrograph of deformed grit, Cairletter Point, Cowal, Southwest, Scotland. The large detrital grain of quartz exhibits evidence of recrystallisation at the margins. Internally, rectilinear arrays of inclusions are undeformed. Small grains of quartz in the matrix have deformed by pressure solution. 180 lamellae. This leads in turn to the development of heterogeneous strain and compositional variations that characterise structures formed by pressure solution. The sical conditions which are necessary for rocks to -'7form by pressure solution are discussed in the section on Cerorratf, '__ gilds.

Dtforn:ti= b7 t-zo or more r,echanisrs.

In sone_ caez; it 4s tos7ible to adduce evidence a rock has de-r'ermod by two or more uechanisms simultaneously. This is probably a general rule for rock deformation, aithou'jih support - evidence is difficult to obtain. For instance, one may er'7ue that grain boundary sliding invariably accompanies deformation by cataclasis or pressure solution, but the presence of grain boundary sliding is almost impossible to detect. The reasons for deformation taking place by two mechanisms simultaneously are complex. Deformation mechanism fields over- lap to a considerable extent in natural rocks. In addition, variations in lithology and grain size will affect the kinetics of processes, increasing the range of physical conditions over which two mechanisms are viable together. Significant effects of deformation by two or more mechanisms can appear over approximately two orders of magnitude in strain rate. Some examples of rocks which have deformed by two or more mechanisms simultaneously are illustrated opposite. Plate 5.3 is a photomicrograph of an oolitic limestone. The oolites have deformed by an intracrystalline mechanism from an originally spherical shape into irregular ellipsoids. In addition, the margins of the oolites have been deformed along pressure solution stylolites, with associated precipitation of clear calcite as pressure shadows. Hence the oolites have changed 181 - log

Fig. 5.3 Schematic diagram of isothermal and isostructural plots for an intracrystalline deformation mechanism (n assumed 4) and pressure solution (n assumed 2).

The differential stress-strain rate space in which deformation—loy

take place by both mechanisms simultaneously is indicated by shading. 182 shape by two different mechanisms, internal dislocation flow and pressure solution of the margins. The effect of grain size on deformation mechanism is illustrated in plate 5.4, a photomicrograph of deformed grit. Small matrix cuartz gains have a stron: preferred shaDe, ali7ned in r;lane of cl- . and hove clearly deformed pre: sure solution. The 1cl-se =ain of vein cu-Dtz exhibits evideuce -'- -17-; - ,,t4 or_ at the mar:ins, thus is dec'ormin7 in port a dislocation -orocess. Fowever, tie centre of the grain retE.ins rectilinear arrays of inclusions which have not 'ent tosignificant de:ree, and hence this part of the grain is not reDondin,--7 to the applied stress. When a rock deforms by two or more mechanisms simultaneously the absolute rates of the processes must be ap,roximately equal. The absolute rate (E ) of a thermally activated deformation mechanism may be described by an equation of state which takes the form

E= A exp (E/RT) 6 n where E is the activation enthalpy of the process, R the universal gas constant, T absolute temperature, 6 n differential stress raised to a given power, f is a function of tile material structure, and A is a term which includes the frequency of attempts, the entropy of activation and other parameters. From experimental work on rock deformation it has been found that isothermal and isostructural plots of differential stress against - log fall on lines as shown by Fig. 5.3. A. hypothetical case has been illustrated for a dislocation process (n = 4) and pressure solution (n = 2). The intersection of isotherms then defines the differential stress-strain rate space in which both deformation mechanisms may take place simultaneously. 183 The effect of lithology on pressure solution.

Mention has already been made concerning the effect of lithology on pressure sc1ition, both in chapter 1, and in the section on deformation mechanisms in Dalradian rocks. Varia- tions in lithology make the task of defining the temperature range of a given deformation mechanism difficult, and in theory limits may be quoted only for a single rock type and grain size. (For intracrystalline deformation of quartz and calcite n ranges typically from 3 to 11; for Coble creep n has a value of 1. In the case of pressure solution n probably varies between 1 and 2.) What are the different properties of lithologies that determine the nature of deformation processes? A folded sequence of graded beds is illustrated in plates 5.5. Despite the fact that deformation is most intense on the inner arcs, it may be seen that a pressure solution cleavage is strongly developed on the outer arcs, within the fine grained pelitic unit of the graded bed. It is suggested that pressure solution has been favoured by 1. fine grained quartz, and 2. the presence of micas within this unit. Inner arcs which are composed largely of coarser grained quartz and feldspar would appear to have deformed by grain boundary sliding and intracrystalline processes. Phyllosilicates have a low degree of preferred orientation in the plane of cleavage. Hence one may infer that, under given conditions, pressure solution is faster at a quartz-mica, or calcite-mica interface than at a ouartz-quartz, or calcite- calcite interface. In addition,the presence of phyllosilicates is almost certain to increase the rate of diffusion in relation to pure rocks, for the case when diffusion is the rate limiting step. graded quartzitic unit is illustrated in plate 5.52 As for the case discussed above, the rate of deformation by ilate 5. 5 , 5 , 5 B Effect of litholoy on deformation mechanism. Iressure solution in polite of graded beds,

North ',Devon. (above); and y.ressure solution in fine grained quartz-pelite of graded bed, Rloscolyn, Anj:lesey (below . Plates 5.6 and 5.7 Photomicrographs of a grit deformed under conditions cf greenschist metamorphism. Innellan,South- west Scotland.Large detrital grains exhibit evidence of intracrystalline deformation- strain shadowing and subgrains. Small matrix grains of quartz have deformed by pressure solution. 184

pressure solution has been faster in the upper fine grained element of the unit. It is suggested that phyllosilicate content and grain size are the two dominant factors in lithology which influence the nature of the deformation mechanism. The effect of ,,rain size is discussed more fully in a later section.

The effect of 'in size on deformation mechanism fields.

It has been demonstrated in an earlier section that, at grains above a certain size will deform by intracrystalline mechanisms and below this size by pressure solution. In addition, evidence was adduced that the transition grain size:, was temperature dependent. In theoretical models of grain boundary diffusion (Coble Creep) the rate is proportional to 1/d31 where d is the grain diameter of the aggregate. Diffusion is taken to be the rate limiting step, and in turn grain size dictates the diffusional path length. Natural rock systems are exceedingly complex because there is generally a number of different crystal types, each with a range of grain size. Further, evidence has been presented that the rate limiting step may in some cases be diffusion, but in others precipitation. One may infer, as mentioned above, that in general, with increasiirisize, grains of quartz and calcite respond to deformation less readily by pressure solution, and increasingly by intracrystalline mechanisms. This fact is illustrated by , plates 545, 5-11 and 5.8. In the former, a large detrital grain in a fine grained matrix has nucleated small strain free subgrains at the margins. Areas that have not recrystallised retain evidence of strain shadowing. These features are indicative of intracrystalline deformation by dislocation flow. Plate 5.8 Photomicrograph of slate deformed by pressure solution, Cairletter Point, Cowal, Southwest Scotland.

Plate 5.9 Photomicrograph of deformed garnet-mica schist, Tummel, Central Highlands of Scotland. 185 The second micrograph shows a cleavage zone in which fine grained quartz has deformed by pressure solution to leave residual grains with a strong shape fabric. Diffusion of quartz out of the zone results in a concomitant increase in the concentration of micas, and a preferred orientation of these minerals rotation as the quartz dissolves. In re deformed 'grits, where the relations outlined abovegenerallyhold true, it i s common that large det-,4 tr1.1 grains have pressure shadows, plate 5.10. This is in spite of the fact that their size may be up to an order of magnitude larger than grains in the matrix which deform by intracrystalline mechanisms. It is suggested by the author that the pr,,ss- ure shadows are the result of two effects -

1. A region of effective tensile stress, favourable to pre- ciritation, may be present around objects which are l r in comparison to the grain size of the matrix, i.e. the stress gradient is increased. This condition would not obtain in an aggregate of uniform grain size.

2. Grain boundary diffusion, which determines the rate of deformation of grains in relation to grain size, may give way in this case to diffusion through a pore solution. Rate of diffusion is governed by diffusivity of the migrating ions and concentration gradient. A high concentration gradient probably obtains across individual stressed grain boundaries, whilst a low gradient is present in a fluid filled pore. If diffusion takes place over a distance which is large in relation to the grain size then a smooth concentration gradient obtains for the case of inter- connected pores. However, if diffusion takes place strictly 188

along grain boundaries then an irregular potential-gradient obtains between two points. Thus grain boundary diffusion is probably • tly slower than diffusion through pore fluids. These two processes have fundamentally different characteric.tics. As a consequence of this, there is some debate whether they should be considered as

tI,Jd diffcrent m,i,cbanis7s, with pressure solution limited to diffusion through pore fluids; and grain boundary diffusion retained for the former process. There is probably as yet insufficient information relating to intercrystalline diffusion during deformation to justify this distinction.

Hence one may infer that in some rocks which have deformed at least in part by pressure solution diffusion may occur either along grain boundaries or in a pore solution. This will depend on grain sorting which governs pore space geometry. The rate of the former process is h ighly sensitive to grain size, that of the latter would not appear to be so.

A general observation that has been made in relation to pressure shadows is that if the matrix grains are greater than about half the size of the rigid inclusion then pressure shadows do not form. This is a result of the scale of the stress perturbation being small in comparison to the grains.

Anomalous results were obtained for the third pressure shadow that was analysed isotopically (chapter 4). In this 187

sample there was no evidence for zones of quartz depletion from which the shadow had formed. Further, the shadow was enriched in heavier 180 atoms in contrast to the other two samples. If grains in the slate, which are a few microns in size, deformed by pressure solution, then it is possible that diffusion into the shadow area, over a distance of three to four orders of magnitude larger than the grain size, was not kinetically viable. It is the author's opinion that the ouartz pressure shadow is of hydrothermal origin, and this would explain the 180 enrichment relative to quartz in the host rock. Supporting evidence for this hypothesis comes from the fact that a carbonate mineral was detected in the shadow, which is not present in the host rock. Hence, certain structures which have the appearance of pressure shadows may not be formed by a local diffusion system from points of high stress into the shadow area. They may be formed by hydrothermal infilling as a gap in the matrix around rigid inclusions progressively opens during deformation. Finally, in relation to the effect of grain size, it is suggested that one reason for pressure solution being less common in rocks deformed at high temperatures is that matrix 1: + 4 coarsening, which is a consequence of annealing, yields grains which are too large to deform by a mechanism of grain boundary diffusion (Plate 5.9).

Diffusion of phyllosilicates.

Some examples have been given in earlier sections of the diffusion of phyllosilicates in response to deformation by pressure solution. The most striking example of this process

188 is the segregation that takes place in buckled muscovite- chlorite mineral fabrics. In most other cases it is difficult to ascertain the mobility of phyllosilicate minerals. Beards of sericite andchlorite may grow in the shadow area of large detrital grains. This feature is common in deformed grits, and is well illustrated from the Eriboll sandstone (plate 5.11). It is difficult to envisage a purely mechanical origin for the local concentration of these micas. In the Eriboll sandstone small grains of quartz change shape by pressure solution and associated quartz overgrowths. Grains of intermediate size deform in part by pressure solution, and also by intracrystalline mechanisms; they have small mica beards set in a quartz matrix. The largest grains deform almost exclusively by internal flow and exhibit well developed mica beards. The effect of grain size on the rate of diffusion of Quartz is as outlined above. It is sur-gested by the author that mica beards are the result of the rate of diffusion of the components of phyllosilicate minerals being faster than quartz at the larger grain size. For intermediate size grains both mineral species are diffusing at approximately the same rate.

Rate limiting steps

Pressure solution is compounded of three elements, namely solution, diffusion, and precipitation. Either one of these processes may be the rate limiting step during deformation. From experimental work on crystal kinetics it has been demonstrated for certain minerals that, at a given ratio of supersaturation or undersaturation to equilibrium saturation, the rate of dissolution is faster than that of growth (Hanevald, 1971). This result follows closely predictions made from 189 the theoretical model of crystal growth by spiral dislocations, formulated by Carrara, Banks and Frank. In an open system dissolution may be the rote limitin7L step in the formtien of structures, because diffusion often tekes -:lace aloe domains of phyllosilicoteS that form high. diffusivity 7p7tLw---- and preci-pitation is removed from the locus of solution. An ersnple ef ch an open system is Fj_ven by discrete zones: of prescare solution that e:thibit nca-linear strain in comparison to the undeformed rock (plate 1.1). In a. closed s7stem precipitation is the rate limiting step in some cases, as deduced from considerations of crystal kinetics outlined above. The size of the diffusional system is small, and individual grains change shape, at approximately constant volume, by pressure solution and the formation of syntaxial overgrowths. In other cases rate of diffusion controls the absolute rate of deformation, as indicated by the grain size control of deformation mechanism, which follows the Coble Creep model. The effect of stress is always to increase the rate of dissolution, but the stress dependence. of diffusion rate is unknown. One may infer from the heterogeneous nature, and evidence of strain softening characteristic of many structures formed by pressure solution, that the rate of dissolution is faster in domains already deformed in part by pressure solution, i.e. it is a strain softening process. If the influence of stress on rate of diffusion was greater than on the rate of "dissolution then the converse would be true. Hence one may conclude on qualitative grounds that the effect of stress on the absolute rate of pressure solution is never such as to cause diffusion to be the rate limiting step in preference to dissolution for an open system; i.e. diffusion is always faster than dissolution. (The concept expressed in this paragraph has been jointly formulated tngetheT, with M. Casey.) 12 grainod quartz cataclaste, Eriboll

ilL.te 5.13 Buckled quartzite, in quartz-i,elite matrix. Note. ho*oeneous folding of sedimentary laminae. Eriboll ,,juartzite. 190

This discussion presents an oversimplified picture of rate processes in relation to pressure solution because the concept of 'rate limiting steps' may not be valid for mechanisms which are a corhination of several different processes.

Aspects of grain boundary sliding.

lot roc'=s at high crustal levels almost certainly deform in part by grain boundary sliding it is a most elusive mechanism to detect petrographically. The necessary criterion is to establish that two grains originally in juxtaposition have changed their spatial relationship. One may infer from experimental work and theoretical considerations that frictional sliding between grains is favoured under conditions of'low effective confining stress. In the field there are certain structures for which it is difficult to envisage deformation by a mechanism other than grain boundary sliding. Plate 5.12 shows a tight small scale fold from the Eriboll sandstone, in which the sedimentary laminae have been bent in A continuous manner around the hinge. Despite high strains the deformation is relatively homogeneous, as indicated-by continuity of the laminae. In addition, there is no axial cleavage, although the rock contains detrital micas. In thin section there is no evidence for intracrystalline deformation of quartz. The margins of some grains have deformed by pressure solution. Micas are wrapped around the grain boundaries, and exhibit a low degree of preferred orientation in the axial plane. This lack of fabric in zones of high strain has also been observed in cusp structures developed in greywackes near Bastogne, Belgium. 191 Grain boundary sliding would appear to be favoured in rocks which have a high degree of textural maturity, i.e. equant grains of uniform size. In order to accommodate rotation and sliding grains must in general either change shape or the rock dilate. It is suggested that grain boundary sliding is accommodated by dilatency in high level, poorly consolidated sediments, for example Tertiary sandstones in the Isle of Wight'llonocline. At deeper levels grains may adjust shape by pressure solution. In the Eriboll sandstone, where grain boundary sliding is indicated pressure solution overgrowths on detrital grains have a near random orientation, i.e. grains have rotated after formation of the overgrowth. This is in contra- distinction to the more general case of a preferred direction within cleavage. In addition, some overgrowths have been part- ly redissolved. Such evidence is taken to support the contention that shape change of grains to accommodate grain boundary sliding, has taken place by pressure solution. Grain boundary sliding is almost certainly an important mechanism by which high level crustal rocks deform. Unlike rocks which deform largely by pressure solution in a highly non-linear manner, grain boundary sliding would appear to operate homogeneously. In addition, grain boundary sliding does not result in a strong fabric of grains, in contrast to rocks which have deformed largely by pressure solution. Effective confining stress has an important influence on the transition from grain boundary sliding to dominantly pressure solution and intracrystalline mechanisms of deformation. However, as was deduced from empirical measurements, the pressure solution intracrystalline transition is independent of confining stress. Plate 5.14 Tectonic veining, zones of pressure solution, and boudinaged layering. Hartland Quay, North Devon. 192

5.11. DE7ORMATION O TECTONIC VEINS.

In the field, tectonic veins exhibit clear evidence of deformation, which is to be expected: in view of the fact that veins subsequent to their initiation, must accommodate tl2e, same overall shorcfer- as is taLen up by tbe host =el: during rmtion- BuchliI veins under conditions of shortening is com::on, due to duct.1 7ity contrast '.:,stween is zone- rally a corse rained ouartz or calcite vein ag-7,reE;ate, and for instance, a reintivel- grrdred, imfure- host roc. In some exaLlples veins would appear to shorten across their own, 1 r-th, contractin in the direction in which t'ney have -rown. Planes of pressure solution in the host rock, which are associated with formation of the vein, may be observed in some cases to cut across the vein. Under these conditions it is su:7ested that veins may shorten across the plane by pressure solution, at the same rate as the host rock is de- formin7 by pressure solution within the same plane. This feature is illustrated in elate 5.14 for a quartz vein, and for calcite veins. Pressure solution of the vein along, discrete planes results in a necked apmearance which could be interpreted in terms of later houdinage (plate 5.14). However, in addition to the field evidence for a relationship between the planes of pressure solution and veining, the layering is boudinaged (lower left hand corner) in a sense which is consistent with the inferred stress orientation for vein formation and propagation of pressure solution planes.

Microscopic evidence of deformation mechanisms in veins.

In connection with the fluid inclusion research attention has been directed towards the textures of deformed crystal aggregates in veins. Plate 5.15 Photomicrograph illustrating intracrystalline deformation of calcite. Note bending and kinking of calcite twins.

Plate 5.16 Bent arrays of secondary inclusions in deformed quartz. Note primary fluid inclusion in negative crystal cavity( top centre ) , and area of re crystall- iqAtinn f tnn left ) 193

Bending and kinking of calcite twins is illustrated by plate 5.15. In those grains where two sets of twins are visible they form a heterogeneous strain grid. Highly localised deformation within individual grains would also appear to have taken place in quartz. The gentle bending of arrays of secondary inclusions is apparent in plate 5.16. A primary fluid inclusion within a negative crystal cavity is preserved at the upper margin. However, at the top left hand corner of this micrograph there is evidence of strong recrystallisation. The recrystallisation is a response to the build-up of strain energy which results from heavy cold working by dislocation processes. Several features that are characteristic of domains which have experienced such recrystallisation have been noted by the author. Plates 5.17 and 5.18 are photomicrographs of part of an original vein quartz crystal viewed in plain light and crossed polars respectively. Linear inclusion arrays are present only on the left-hand side of the upper photograph. Under crossed polars one may observe that the left-hand side is relatively undeformed whilst the right- hand side has suffered strong deformation and recrystallisation. Recrystallisation has 'swept out' many of the inclusions, destroying the linear arrays. In some cases inclusions have been driven into the boundaries of new subgrains (an observation that has also been made by Dr S. White, personal communication). Plates 5.19 and 5.20 are photomicrographs of quartz, taken in plain light and crossed polars respectively. The boundaries of new tectonically induced subgrains cut across arrays of pseudo-secondary inclusions that are associated with the growth of a large euhedral, and originally strain free Plates 5.17 and 5.18 Photomicrograph of deformed quartz in a tectonic vein. Plain light above, crossed polars below. Tectonically induced subgrains cross-cut arrays of pseudosecondary inclusions that are associated with growth of the original vein quartz.

( (4 Plates 5.19 and 5.20 Photomicrograph of deformed quartz in a tectonic vein. Plain light above, crossed polars below. The formation of subgrains and removal of fluid inclusions is a result of local intracrystalline deform- ation1 lower right hand corner ). 194 crystal. Hence, a general sequence may be observed in the deformation of vein quartz from original coarse grained, strain free grains having growth features; through grains which exhibit evidence of weak strain shadowing and partial obliteration of 'rowth features; to a strong ly recrystallised ag;3re;s'ate with Strain shadowinc, deformation bands, sutured subgrain boundaries and a total removal of original growth features.

Water weahening of quartz

During the measurement of filling temperatures in vein quarts it was noted that certain subareas of individual grains yielded large numbers of primary inclusions with a low variance in the homogenisation temperatures. In different subareas, primary inclusions were either absent or had leaked. Data from inclusions in these areas gave systematically high temperatures with a large variance that was generally inconsistent with data from the subareas first mentioned. The former subareas are strain free and exhibit no evidence of deformation. The latter invariably coincided with domains of recrystallisation. This observation of heterogeneous deformation at the level of a single crystal may be extended to individual grains in a crystal ag7regate. It was observed that single crystals or groups exhibited evidence of strong deformation-strain shadowing, recrystallisation, deformation bands and absence of primary inclusions. Other grains retain their euhedral habit and strain free appearance. Whilst grains with an intermediate level of deformation occur, they are not common. In the author's opinion the explanation for these features of heterogeneous deformation in quartz lies in a strain softening effect due to water assisted intracrystalline deformation . 195 It is well established that there is a significant reduction in the strength of quartz when deformed under 'wet conditions' (Grigg73s and Blacic, 1965); although there is some contention to the exact rocllanism. 7urther, there ds new informEtien, obtained by infra roc', spectrometry, tht some vein quartz may 11,-,ve a vny 7 :t.= content of u2 to several thousand e'en atoms per 10 atoms (h. Jones, personal _wniicaticn 5 .

cut: or .- ro:pc.; tiyat defcr7 7 -1 C of vein quartz fluid inclusions in some grains, or bare-s of water is driven into the cry stet. This increase in the 'mobile' water content causes a local increace in the rate of deformation. Other area31 devoid of fluid inclusions, or domains where the inclusions have not leaked into the crystal, will deform at a relatively slow rate; and thus do not exhibit evidence of recrystallisation, or other features of intracrystalline deformation that characterise ti-:e domains of hi ;h strain (see plates 5.17, 5.18, 5.19 and 5.20).

Growth fibres and tectonic fibres.

As was explained in an earlier section fibrous minerals in tectonic veins are very common. A great deal of contention has been generated over the significance of fibrous minerals, and the various theories which have been proposed to account for their formation are discussed more fully in chapter 6. In general terms these minerals are considered either to grow as fibres or to be structures induced by deformation. Certain features, noted by the author which related to fibrous structures, may be used in some cases to distinguish between these two types. The importance of this distinction lies in the L.2; Ftctomicrcgraph of a tectonic vein, which illustrates hetergeneous deformation of quartz. Grains on the left hand side retain their euhedral form. Grains on the right hand side have lost their original euhedral form, and exhibit evidence of intracrystalline deformation.

Plate 5.24 PhotomicrograH eli vein. Fracturing of quartz and intracrystallire flow of calcite. 196 .widely different interpretations that are attached to growth fibres and tectonic fibres respectively. It was explained in chapter 3, on fluid inclusions, that arrays of pseudosecondary inclusions lie along fractures that formed during growth of the host crystal. Hence, the ends of such arrays delineate growth boundaries. Plates 3.8 and 3.9 are photomicrographs of a tectonic quartz vein with fibrous quartz structures that are characterised by the presence or absence of transverse arrays of pseudosecondary inclusions. Hence these structures are considered to be true growth fibres. Quartz fibres in another photomicrograph (plate 5.13) cross cut arrays of primary and pseudosecondary inclusions, and are thus not features associated with growth of the original grain. These fibres are elongate subgrains of quartz which lie within the original euhedral crystal. Original crystals exhibit all the features of strong intracrystalline deformation-strain shadowing, deformation bands, absence of primary fluid inclusions. The elongate subgrains, have a preferred orientation to give parallel bundles of fibres. Tectonic fibres may be indistinct at an early stage of deformation, but with increasing strain the subgrain walls become more defined, of higher aspect ratio, and achieve a greater degree of parallelism (plate 5.27).

Deformation mechanisms in relation to temperature.

A considerable body of information has been compiled from oxygen isotope and fluid inclusion thermometry concerning the temperature at which some tectonic veins have formed and subsequently deformed. These data have been combined with textural studies to build up a general picture of the effect of temperature on deformation mechanisms in vein quartz, calcite, Plate 5.25 Photomicrograph of vein quartz and calcite fibres which have fractured in response to deformation.

Plate 5.26 Photomicrograph of deformed quartz-calcite vein. Calcite exhibits evidence of deformation by intracrystalline mechanisms- bending of twins and loss of euhedral form. Quartz grains exhibit no evidence of intracrystalline deformation. 197 and phyllosilicates. The sequence of deformation mechanisms, as a function of inoreasinc temr,erature for certain Minerals in tectonic veins, which has been empirically established by the author, ds i_ntonled as a broad ieneralisation. Uo atteml;t

been e ±o i t4.-7;e rl.r0771 -tion 7J-'oces'7es in veins aT2J-o,re Tr7orTfloticl_has been co71i1c,d Cron a stu-fl-L- of sevc;?n'7y-fc::? vezis t!,11.1 -en from a uido vcrietLT of cranial lovcis, w: thin (7fff'c;r:-,, t t.?3to-Jiic ,Tovinces. In the cliTleuzsdcn that follows it should be inlcil strain Tate effects caua sdTnificant shift a in the tha=...1 linus which have been empirically evaluated, and are almost certainly responsible for many of the anomalies that have been observed. Slower strain rate has the effect of lowering the interval of temperature over which a specified deformation mechanism will operate, with given boundary conditions. The converse is also true. At temperatures below/N/100°C both quartz and calcite deform in a brittle mode. Fracturing of varying intensity is common, but grains tend to retain their euhedral form (plates 5.24 and 5.25). Evidence of intracrystalline deformation mechanisms are rare. Evidence has been adduced that calcite begins to deform internally from ^/150°C. Fracturing is less common than at lower temperatures3and is replaced by bending and kinking of twins (plates 5.15 and 5.26). Quartz continues to deform in a brittle manner. Phyllosilicates, where present, tendi to be bent and fractured. For similar amounts of strain features of intracrystalline deformation in quartz - recrystallisation, deformation bands, strain shadowing, absence of fluid inclusions - are more in evidence in veins deformed above 200°C; are typic&bt250oC, and Plate 5.27 Photomicrograph of a tectonic vein. Intracrystalline deformation of quartz and calcite. Note tectonically induced fibrous subgrains of quartz.

Plate 5.28 Photomicrograph illustrating the intracrystalline deformation of quartz and calcite. 198 almost ubiquitous over 30000. Original grains lose their euhedral growth habit (such as is illustrated in plate 5.26), and increasingly intense recrystallisation takes place at the higher temperatures (plates 5.27 and 5.28). Features of brittle deformation decrease in a corresponding manner. It is in the temperature range of approximately 200°C - 300°C that examples have been found of quartz veins which have shortened across their own length by pressure solution. Calcite veins deform by pressure solution in the temperature interval of approximately 100°C - 250°C. In Quartz-calcite veins deformed above ,,, 300°C quartz generally exhibits evidence of strong ,intracrystalline flow, but calcite appears to be less deformed except at grain boundaries-. Hence it is sug7,ested that under a given set of conditions and specifically at slow strain rates, quartz is in general stronger than calcite below #, 200°C, but that above 300°0 the converse is true. The response of quartz and calcite would apl:ear to be somewhat modified in veins which have a significant content of phyllosilicetes. Grains retain their euhedral outline,T, and exhibit less evidence of intracrystalline deformation. Micas tend to be wrapped around the grain boundaries. Some tentative evidence has been obtained, from matching arrays of secondary inclusions which cross grain bbundaries, that the deformation has taken place by grain boundary sliding. The predominant deformation mechanism in veins is' rarely the same as that by which the host rock has deformed. Fea- tures of int'racrystalline deformation of quartz and calcite in a tectonic vein are illustrated in plate 5.29 (500°C). The host rock, a polite, shows all the evidence of deformation by pressure solution - dark planes of quartz dissolution and concomitant enrichment in micas and iron oxides; a prof erred shape Plate 5.29 Photomicrograph of a deformed vein and host rock.

The coarse grained vein quartz and calcite have deformed by

intracrystalline mechanisms. Quartz and calcite in the host

rock have deformed by pressure solution, which results in a

strong shape fabric of the grains, and a preferred orientat-

ion of micas. 199 of.quartz grains within the cleavage planes, and a preferred shape orientation of micas.

5.3 DETORVATION 1-7,CHAITIST FIELDS.

Geologists have long recognised that the mechanism by which structures develop is dependent upon the prevailing conditions of metamorphism. The specific physical parameters which determine the nature of the dominant deformation mechanism are -

1. Strain rate. 2. Oonfinin7 stress. 3. Fluid pressure. 4. Differential stress. 5. Temperature.

Tlie manner in which these factors govern some deformation processes has been outlined by _Trice, 1964; Price, 1970; Donath and others, 1971; Rutter, 1974, and other authors. The conce.:,t of t:e deformation-mechanism map was introduced by Ashby (1972). By utilising standard constitutive relations between tersile stress and tensile strain rate for several deformation mechanisms, Ashby (1972) and Ashby and Verrall (1773) defined the temperature-stress fields in which the individual mechanisms may take place (Fig. 5.1). Each map is for a given material of specified grain-size. Weertrnan (1968) had previously used a similar technique to construct the deformation fields for some intracrystalline mechanisms. Deformation maps are valuable because one may infer from them the stress-strain rate-temperature conditions under which a given mechanism may be expected to take place. The concept of 200 theoretic al strength

dislocation glide -101 dislocation creep

pressure Coble creep solut ion Vi abarro creep - elastic regime -10'

- 10

0 01 02 03 04 05 06 07 08 0.9 10 homologous temperature TAm

Fig.5.1 A deformation-mechanism map for silver, grain size 32".1; adapted from Ashby 1972. Strain rate cont-

ours are indicated schematically, i i> it . A possible

field for pressure solution of quartz (Tm-1500°C ) is

illustrated. 201 the deformation map may be usefully extended to geological materials and conditions obtaining in the crust. However, because of the complexity of geological materials no simple map may be constructed for deformation mechanisms in crustal environments. The constitutive equation used by Ashby (1972) for Coble creep, which bears some relation to pressure solution is -

14 crS1 1[ —2. 1 kT B where 31 is the atomic volume, d the grain size, DB the boundary self-diffusion coefficient, &the effective cross section of a boundary for diffusion, E strain rate, 6 tensile stress. Some observations relating to deformation mechanisms in geological materials have been made by Elliott (1973), Kirby and Raleigh (1973), and Stocker and Ashby (1973). 202

5. IV FIELD AND ANALYTICAL WORK RELATING" TO DEFORMATION MECHANISM FIELDS.

The methods that were evolved for estimating the limits of physical conditions in the crust over which pressure solution is an important deformation mechanism have been argued . in detail in chapter 4. These methods have been applied to further investigations of deformation mechanism fields in two other areas, namely the East Central Pyrenees and Northwest Scotland. Boundary conditions are not so well defined in these areas as is the case for the Dalradian series, but the same general approach was used. Quantitative data obtained from fluid inclusion and isotopic analyses on Dalradian rocks has been directly applied to the other areas, for instance in the case of thermal stability limits'of a given metamorphic mineral assemblage. In the East Central Pyrenees, near Andorra, Silurian, Devonian, and Carboniferous sediments have been deformed under metamorphic conditions which range from chlorite to cordierite- sillimanite grade.. Mineral assemblages are typical of the low pressure regional facies series. Zwart (1963) has estimated that the geothermal gradient obtaining during metamorphism was 1500C/Km. The upper temperature limit at which quartz was observed to have responded to deformation by pressure solution was estimated as 45000 - 500C. This temperature corresponds to a depth of less than 4 Km of overburden, according to Zwart, i.e. at less than 1 Kb confining stress. These data provide additional evidence that the pressure solution-dislocation 203 flow transition is approximately isothermal, and is not dependent upon confining stress within the limits of measurement. Unfortunately it has not been possible to examine rocks deformed within the high pressure facies series of regional metamorphism. At Eriboll in Yorthwest Scotland three nappe units are present which contain Cambro-Ordovician sediments; gneiss and mylonites deformed at greenschist facies; and rioine pelites and psammites metamorphosed under conditions of greenschist to almandine grade respectively. The nature of the dominant deformation mechanisms within these three nappe units is different. However, it is emphasised that strain rate effects may cause large shifts in the thermal limits of a given deformation mechanism. For instance, within the single uniform litholoLy of the Eriboll orthoquartzite, evidence has been adduced that in a single tectonic episode deformation has taken place by fracture, grain boundary sliding (plate 5.12), pressure solution (plates7.3,6), and dislocation processes leading to the tectonic grain refinement of a 'cataclasite' (plate 5.13). In order to compile information relating to the conditions of temperature at which deformation by pressure solution has taken place in non-metamorphic environments, investigations have been conducted using techniques described previously, at a number of localities, selected to represent a variety of crustal levels. These data are listed in Tdtgle3.1. 204

Locality Temperature- Structure depth

Tertiary cen7lomer- 60°C t 10°C . diageneLic pressure

ates, Taviva, 3. solution of Yortlu:n1 conglomerates

Cretaceous chalk, 00°C t 10°C pressure solution S. Dorset cleav,e Triassic Bunter con- 90°C ±10°C * diagenetic pressure glomerate, English solution of Midlands 4 Km + conglomerates Tertiary Helvetic 110°0 t 20°C * tectonic pressure Molasse, Niederurnen, solution of Switzerland conglomerates Carbonates, Chambery, 140°0 ± 30°C * tectonic stylolites French Jura and veining Palaeozoic greywackes, 180°C t 30°C * folding and tectonic Waves, Castel Marim, veins S. Portugal Palaeozoic greywackem, 200°C ± 30°C * folding and tectonic Tebay, Lake District Veins • '; Tries dolomite, 97°C = 10°C pressure shadows latschental, around nodules Switzerland Lower Devonian grits, Z94°C ± l0°C zones of pressure Pembroke solution -1- 1000 Middle Devonian carbon- 317°0 pressure solution ate, Coombe Martin, striping N. Devon. o Westphalian greywackes, 260 C - 10°C Hastiere anticline, Givet, Belgium 4 Km pressure solution crenulations + Palaeozoic slates, 240°C -40°C * Pressure solution Lerch Rheinische striping Schiefegebirge + Palaeozoic slates, Zell, 240°0 - 400 * Pressure shadows on Moselle nodules Dalradians, Southwest 500°C ± 50 C upper limit for press- Scotland (biotite isograd) ure solution of 15 Km + quartz Palaeozoics, Andorra, 500°C t 50 C upper limit for press- Central Pyrenees (biotite isograd) ure solution of 4 Km quartz Gambro -Ordovician 200°C t 50°C * pressure solution sediments Eriboll, zones Northwest Scotland 6 Km +

Moine pelites, 500°C upper limit for press- Eriboll, North- (biotite isograd) ure solution of west Scotland quartz

Table 5.1 Localities at which quantitative data has been obtained on the relationship of crustal conditions and deformation mechanisms.

* Data obtained by fluid inclusion thermometry for which it is 'is negligible. assumed that the effect of CO2 ,61 ' ' • 4. Ow,

;•■■,, k

**414:406.4

■ •

- . date 5.T 3 Folded layers of gneiss. Grotli, South-Central Norway.

_late 5.T1 Tectonic striping formed oy pressure solution,cross cutting original sub-horizontal bedding. Sampson's Covq, Ilfracombe, North Devon. Plate 5.T2 Folded layering in amphibolite grade gneiss, Pennine nappes, Southern Swiss Alps. 205 5- V. QUALITATIVE EVIDENCE RFIATING TO DEFORMATION MECHANISM FIELDS.

Introduction.

The relationship between deformation mechanisms and physical conditions in the crust has been subjectively assessed over many years from observations of the geometry of, and degree of chemical redistribution within, structures formed at different metamorphic grade. For instance, layering is.in general 'active' during folding of low grade rocks, but behaves in a passive and ductile manner in the development of folds at amphibolite and higher grades of metamorphism. Another example of this relationship is the invariant association of syntectonic veining with deformation of low grade rocks, whilst in deformed sediments and high grade rocks syntectonic veining is rare or absent. From such basic field observations of structures made in a variety of tectonic environments one may begin to build up a picture of which deformation mechanisms predomi- nate at a given crustal level, and how each mechanism gives rise to characteristic structures and petrofabrics. This section sets out to give a descriptive account of the features that result from deformation of certain types of structure under different conditions of metamorphism. For many of the examples cited information on conditions of temperature and pressure•has been determined analytically. These data are presented in a later section.

Folded layering..

Two types of deformed layering are illustrated in plates 5.TI and 5.T2; formed under conditions of lower greenschist and amphibolite grade metamorphism respectively. In the former case layering is weakly buckled, with a low degree of Plate 5.T5 Shear zone in Lower Devonian grits, Marloes, Pembroke.Shape fabric of quartz grains due to deformation by pressure solution.

Plate 5.T6 Eh ar zone in amphibolite grade gneiss, North Uist, Outer H :'ides. Shape fabric of grains due to intra- crysta' deformation. 206 internal deformation. It is transected by coarse anastomosing pressure solution cleavage zones, and the layering has shortened across the zones of discrete points by pressure solution. The layering has behaved in a highly ductile manner in those folds formed at amphibolite grade. Layering remains continuous and has deformed internally by intracrystalline mechanisms. These same features are expressed by the folded gneiss plate 5T3 Layering has remained continuous, despite very strains, and an axial schistosity is uniformally developed. In contra- distinction to these folds formed in high grade rock, the deformed slate illustrated in plate 11.21• had an originally sub- vertical bedding, which has been completely transposed by pressure solution into a horizontal tectonic striping with marked compositional variations.

Shear zones.

Shear zone structures formed in coarse grained sediment and gneiss with amphibolite facies mineralogy are shown for comparative purposes in plates 5.15 and 5.T6 respectively. In the sediments deformation of grains has resulted in a strong shape fabric. This is largely an intercrystalline effect resulting from pressure solution and associated syntaxial overgrowths. The same type of sigmoidal structure in the amphibolite grade rock is defined by grains which have a shape fabric inherited from intracrystalline deformation.

Deformed conglomerates.

Pitting of deformed conglomerates is almost ubiquitous in high level crustal environments. Imprints of dolomite pebbles elate 5.T9 Deformed conglomerate, Cristallina, Southern Swiss Alps.

Plate 5.T 10 Internally flattened and buckled conglomerate. lloine Series, Ben Hutig, Sutherland. liate 5:17,8 Conglomerate deformed by pressure solution. Tertiary Helvetic molasse, Niederurnen, Eastern Switzerland-. .late 5.111 Deformed matrix around quartzo-feldspathic inclusion; granulite facies gneiss, Notbodden, South- Central Morwa7.

date 5.T12 Deformed matrix around inclusions (0.4m. maximum diameter). Amphibelite facies marble, Lower Pennine Nappes, Southern Switzerland. ' 207 dissolved against a granite boulder by pressure solution are illustrated in plate 5.T7. On curved surfaces strain due to pressure solution activity may be taken up along movement planes which are decorated by slickolites (plate 5.T8). At higher temperatures pitting at pebble contacts in deformed conglomerates has- not been observed by the author. Powerful intracrystalline flow in several different lithologies of a deformed polygenetic conglomerate from the Swiss Alps is illustrated in plate 5.T9. In a like manner pebbles of the basal Moine conglomerate, Ben Hutig, have been strongly flattened and subsequently buckled by internal deformation mechanisms. Both these examples of deformed conglomerates come from metamorphic environments at higher greenschist to lower amphibolite grade.(5.T10).

Rigid inclusions.

Several examples of pressure shadows formed around rigid inclusions in slates have been illustrated in previous see- tions. At higher temperature environments there is no evidence for the selective transfer of minerals by diffusion. In both the case of the rigid euartzo-fjidspatic nodule. in a granulite facies gneiss (plate 50-11), and calcareous nodule in deformed marble with amphibolite facies mineralogy (plate 5.1- 11.) the layering wraps around the nodule in a ductile manner.

Crenulations

Crenulations are a useful structure for assessing the mechanisms by which a rock is deforming because if pressure solution has taken place it results in the characteristic .113 Buci:led chlorite-calcite schist, with mineral ation by pressure solution. Craignish, Southwest '3cotland.

1l to 14 Buckled quartz-biotit;e schist with no mineral senrEation. Harris, Cuter ITehrides. 208 segregation between hinge and limb regions. A crenulated biotite-hornblende-quartz schist and chlorite-quartz schist are illustrated in plates 5.T(L.. and 5.113 respectively. Quartz has segregated from limbs into hinges in the latter case, but the mineral fabric remains homogeneous- in the amphibolite grade schist, which has clearly responded to deformation at least in part by a different mechanism.

Pressure solution in the retrogressive case.

It was noted in chapter 4 that when high grade metamorphic rocks are involved in a tectonic episode which takes place at a lower temperature then deformation by pressure solution takes place over the same temperature interval as for the pro ride case. Plate 5.7r6is a photomicrograph of a biotite-hornblende gneiss buckled under conditions of greenschist metamorphism. Pressure solution of quartz has taken place in the fold limbs. A nodule with 2-pyroxene granulite facies mineralogy in. a matrix deformed under chlorite grade conditions has developed a pressure shadow (plate 5.-TI). A certain requirement for extensive deformation of rocks with high grade anhydrous assemblages is the presence of some water,,which increases the rate of grain boundary diffusion for deformation by pressure solution. Plate 5.T15 Crenulated hornblende-biotite gneiss. Lewissian of the Sleat of Skye. Pressure solution of. quartz in fold limbs, together with offset of gneiss layers.

1-late 5.T 16 Basic nodule with granulite facies mineralogy, Harris, Outer Hebrides. A pressure shadow has formed during deformation at chlorite grade metamorphism. 209 CONCLUSIO7S

fluantitative date on deformation mechanism fields is presented in Fig. 5.4, Table 5.1 and Fig. 4.7. The main con- clusiors follow from the aut7-ors s inte-vret:tior of this cThtf,-_ are .2c, fellos

1. Calcite deforms b pressure solution from surface .-tho ( 100 m, to .., o, Cu.:„rtrz f.efcc -r7nectsure solution from >1 1 4,0 C to 450°C - 50°C. Nuscovite and chlorite- deformby pressure solution up to 450 C - 50°C, althoull these minerals do not generJ,lly deform by pressure solution. Defo-f-mation by pressure solution is not common below 2006C although exceptional cases of chlorite pressure shodous in dia:;eretic environments have been recorded. 4. The upper temperature limits at which these minerals deform by pressure solution is indeendent of total confining stress within the limits of measurement in the range I Kb to 5 Lb. If A (the ratio of fluid pressure to overburden pressure)is R.11. during deformation in all of the field areas studied, then the effective confining stress would have been constant, and low, in all these situations. However, if the value of Xobtaining during deformation was significantly less than 1 in some of the field areas, then the upper temperature limits at which the minerals listed above deform by pressure solution may be independent of effective confining stress. One may infer from the predominance of intracrystalline mechanisms in some rocks deformed at higher temperatures (>400°G) that A 1, because rocks tend to deform by fracturing and grain boundary sliding when then effective stress is low. 210

5. In quartz-feldspar assemblages feldspar does not deform by pressure solution. The lower limit of intracrystalline deformation is estimated as 300°C - 50°C, depending on grain size. Intracrystalline deformation is ubiquitous above 500°C. 6. The transitional grain size between deformation of ouartz by pressure solution and intracrystalline mechanisms decreases from several millimetres below 300°C to a few microns at . 500°C. 7. Grain boundary sliding is an important mode by which high level crustal rocks deform. The lower limit of this deformation mechanism is dependent on confining stress. 8. Phyllosilicates do not deform to any great extent by pressure solution. Grain boundary sliding is probably the dominant deformation mechanism.

Points 1. and 2. require the qualifying statement that the limits quoted represent the upper temperatures at which pressure solution is a significant deformation mechanism. Pressure solution may be the dominant deformation mechanism below the thermal limits listed. 211 quartz chlorite albite 8000C calcite

in tracryst all ine 600 field

400

pressure solution field 200

Fig. 5.4 Suggested temperature interval over which the indicated

minerals may deform predominantly by pressure solution in

some lithologies. Grain size >101*. 212

5,VI GENERAL ASPECTS OF CONDITIONS IN THE CRUST.

The long term fracture strength of rock.

One of the problems associated with interpreting data from experimental rock deformation lies in estimating the long term strength of rock by extrapolation of data obtained at fast strain rc3tes ( E> 10') to conditions which obtain in the crust 10-14). The condition for hydraulic fracture is

0 - p > - T 3 where 0 3 is the minimum principal stress, p is the pore fluid pressure, and T the tensile strength of rock. For a horizontal vein in a tension gash formed by hydraulic fracture the overburden stress = 03. If the depth of overburden is known, and the fluid pressure may be estimated, then it is possible to compute the long term fracture strength of a rock during deformation. Under most circumstances the depth of overburden during deformation is not accurately known. However, near Givet in Belgium the depth at which the Hastiere anticline formed has been estimated as 5.7 Km from stratigrapnic data (Waterlot and heugnier, 1973). Fluid pressure is calculated using the method derived by the author as outlined in chapter 3. The following steps are necessary -

1. Isotopic data from the vein and host rock yield tempera-

tures of 340°G.(Fig. 4.9). 2. The temperature derived! from fluid inclusion thermometry is R59°C. Ria

3. Fluid salinity is estimated as 2 weight 0/0 NaCl; fluid density = 0.79 (Table 3.1).

Using the PT diagram of water (Fig. 3.2) these data give a fluid pressure of 1500b. The tensile strength of micaceous sanctoe -.,,r7,11101 to beadins t'cen follows 755. (The error. in these calculations amounts to = leob.) This is believed to be ti- e 1'. it s t attempt mane to measure enpi7-i..call7:

fract2.ra ,r thc,t Costaine con']itio. The cumulstivo errors in this LI,Dthed give a 7.7-7e uncert:aintz in the final result. In adaitioi-L, numbor'of situc,tions in which t e method mey be apnlied are of very limited , i J V t 1 J•

Thermal' fluctuations in tectonic environments.

In the Cirque d'7]staubel Central pyrenees, a 11:m vertical section through a nappe pile is exposef. A traverse was made up this section with the intention of estimating the thermal gradient obtaining during deformation, using data from fluid inclusions. F;omewhat surprisingly the measured temoeratures increased up the section. One of the interpretations which may be placed on this data, taLing into account the magnitude of pressure corrections, is that there was no temperature difference between the upper and lower part of the nappe pile during deformation. Hence one may conclude that stable thermal gradients do not necessarily obtain in tectonic environments. 214

REFE EN ES - CHAPTER 5

ALZEY, E.2. 1972. A first report on deformation-mechanism maps. Acta Met., v. 20, 887-897. ASHBY, M.F.,,and R.A. VERRALL. 1973. Diffusion-accommodated flow and superplasticity. Acta Net., v. 211 .149-163. DC7:,TH 7.A., R.T. FILL and D.G. T=N. 1971. •-Deformation pa e fields in experimentally deformed rock:. ELLIOTT, D. 1973. Diffusion flow laws in metamorphic rocks. Geol. Soc. America Bull., v. 84, 2645-2664. GRIGGS, D.T. and J.D. BLACIG. 1965. Quartz: anomalous weak- ness of synthetic crystals. Science, v. 147, 292-295. KIRBY, S.F. and C.B. RALEIGH. 1973. 'Mechanisms of high- temperature, solid-state flow in minerals and ceramics and, their bearing on the creep behaviour of the mantle. Tectonophysics, v. 19, 165-194. POW.vaill, R. and J7.. SMITH. 1973. Pressure-temperature estimates for a late metamorphic event in the Dalradian in the Scottish Highlands. Nature Physical Science, v. 244, 70-71. TICE, N.J. 1964. A study of the time-strain behaviour of coal-measures rocks.. Int. J. Rock Mech. Ein. Sci., 1 PRICE., N.J. 1970. Laws of rock behaviour in the earth's crust. In: Somerton (Ed.), Symp. Rock Mech., 11th Berkeley, Calif. Am. Inst. Min. Metall. Pet. Eng., 3-23.- RUTTER, E.H. 1974. Influence of temperature, strain rate and interstitial water in the experimental deformation of calcite rocks. Tectonophysics, v. 22, 311-334. STOKER, R.L. and M.F. ASHBY. 1973. On the rheology of the upper mantle. Reviews of Geophysics and Space Physics, v. 11, 391-42b. WATERLOT, G. and A. BEUGNIER. 1973. Guides Geologiques Regionaux, Ardenne, Luxembourg. Masson et Cie. Malesherbes, 206 p. WEERTMAN, J. 1968. Dislocation climb theory of steady-state' creep. Trans. Am. Soc. Met., v. 61, 681-701. ZWART, H.J. 1963. Some examples of the relations between deformation and metamorphism from the-Central iyrenees. Geol. Mijnbouw, v. 42, 143-154., 215

The interrelationship of pressure solution and tectonic veining

As we have seen in chapters 1 and 5, pressure solution is the one

deformation mechanism which may result in a significant migration and

segregation of mineral species. A system deforming by pressure solution

may be open or closed on a scale of millimeters to tens of meters. In

the former case, domains of quartz or calcite depletion in cleavage planes

or zones of pressure solution, are commonly associated with tectonic veins

and fractures which act as sinks for precipitation of these minerals. This

relationship between pressure solution and tectonic veining has been noted by

several authors (Sorby, 1858; Ramberg, 1961; Mackintosh, 1967;

Bacblmrev, 1969; Beach, 1974). As is the case for pressure solution, widespread veining is most commonly developed in low grade metamorphic rocks. The frequency of tectonic veining bears a direct relationship to the relative importance of pressure

solution as a deformation mechanism (chapters 3, 4 and 5). Calcite veins

are more common in carbonate environments at temperatures below ^d 150°C

than are quartz veins in quartz bearing rocks below this temperature. This feature may be explained by the temperature dependent diffusion rates of quartz and calcite. Supporting evidence for this hypothesis may be argued from the isotopic data relating to quartz and carbonate disequilibrium,

as discussed in chapter 4.

Mineralogy of tectonic veins

The mineralogy of tectonic veins is generally linked to the mineral ,assemblage of the host rock. Calcite and dolomite veins are common in

carbonate rocks; veins of quartz are common in dominantly quartz bearing rooks. Phyllosilicates and albite may be present at a low level of

concentration, particularly in eTamples where the host rock is rich in 216

these minerals (plate 4.6). Hence one may conclude that the rates of diffusion of quartz and calcite are significantly faster than the rate of diffusion of the components of phyllosilicates and albite, inaow\grade metamorphic environments. A relationship between the size of veins and vein mineralogy has been observed by the author. Veins of millimeter size generally contain the same major mineral components as the host rock. For instance quartz-calcite-muscovite-chlorite rocks may contain veins of the same mineralogy. Larger veins, of meter length or more, are commonly composed of several mineral species, one or more of which are not present in the adjacent rock. One may infer that larger veins communicate with a fracture system that extends beyond the immediate lithology. As a consequence,mineral species from other lithologies may enter into solution, to be precipitated in a fracture distant from their source. The size of a hydrothermal system over which mineral species are migrating and equilibriating with one another may be tested isotopically. This has been done at Coombe Martin, N. Devon, where the size of the equilibrium system has been determined as greater than 100 m, parallel to bedding. Quartz in tectonic veins and host rocks separated by > 100 m are in isotopic equilibrium. The mineralogy of veins is dependent on the metamorphic conditions obtaining during deformation. Chlorite is common in veins formed at 150°C - 300°C; muscovite has been observed in veins formed above 200°C; albite has been recorded in tectonic veins in greenschist facies environment (300°C - 400°C), and biotite may be present in veins formed within rocks of the biotite zone. ]1Lte .1 Veins ir ;7:Aed Freywac)7es, Hartland ioint, Irth Devcn.

lnte 6.j Vein arrays in ,craded beds. Pic de 3arradets, _(;pritral_l*rerp_n_ - 217

Rate of vein formation -

It is posible to adduce evidence concerning the rate at which tectonic veins grow. One may infer from the following points that some small tectonic veins ( m length) may form at a relatively slow rate.

1.Fluid inclusions exhibit evidence of metastability, from which one may conclude that rates of fluid' movement were slow.

2.The large primary inclusions present in some quartz crystals only form under conditions of slow and stable growth of the host crystal.

3. Isotopic equilibriation of quartz between the host and veins requires long periods of time, especially at low temperatures. •

4.If a vein has formed by diffusion of components from the surrounding rock, then long periods of time will be required because the rate of diffusion is probably slow.

5.Veins which form in graded beds do not generally extend into the fine grained pelitic element of the unit (Plates 1.1, 1.2). As a consequence,ons may infer that the coarse grained rock has deformed in a brittle mode, and the pelitic element in a ductile mode. Ductile deformation generally requires slow strain rates. Plate 6.A Shape fabric of grains in vein and host rock,

Marloes, Pembroke. 218

Unusual veins

Certain types of'Veins'have been observed by the author which do not appear to have formed as infilling of a fracture.Avein'of this type is illustrated in plate 6.A. The structure is defined by an elongate zone in which the grains have adopted a preferred shape orientation as a result of deformation by pressure solution. No distinct boundary exists between grains in the 'vein' and grains in the 'host rock'; although one- may infer that the host rock also has deformed by pressure solution, at least in part.

If these strOctures'have formed by pressure solution, then it is unusual that they should have propagated parallel to the maximum principal stress.

CONCLUSIONS

Some inferences have been drawn from the fluid inclusion and stable isotope data in chdpters 3 and 4 respectivly, concerning the physico- chemical properties of hydrothermal fluids in tectonic provinces. The results of this work, together with the observations on tectonic veins, may be summarised as follows.

Fluids in nonmetamorphic and low grade metamorphic environments are subcritical H 0- NaC1 - CO systems,of density 0.81 to 0.96. Thermal 2 2 bquilibrium has obtained brtween the fluids and host rock. The fluids have

probably originated from dewatering of sediments at depths of 2 to 8 Km

below the veins. Magmatic or meteoric sources of the fluids is considered

unlikely on the basis of the available data.

Tremendous quantities of quartz and calcite have been precipitated from

hydrothermal fluids in fractures, yet the host rocks deforming by pressure

solution, and- not the fluids themselves,are the source of this material.

Indeed, we may be more precise in stating that of four veins studied ,

less than 10% of the quartz was introduced from external sources by hydro- 219

thermal fluids. Quartz in the host rocks and veins have attained a close approach to isotopic' equilibrium. Hence it*is not necessary to invoke large volumes of fluids from which the quartz or calcite has been precipitated. On the contrary, there is direct evidence that the water rock ratio dies low, and that the hydrothermal fluids were slow moving. The size of a pressure solution source- vein sink system,in equilibrium isotopically, may be greater than 100m. An open system of this type constitutes an important mechanism for the migration of large volumes of crustal materials over relativly short distances. Pressure solution activity which takes this form represents the opposite end of the scale to a system that is closed

and in which the individual grains deform at constant mass. 220

REFERENCES - CHAPTER 6

EEACH,A.1974. A geochemical investigation of pressure solution and the formation of veins in a deformed greywacke. Contr. Mineral. and Petrol.,v.46, 61-68.

BOCHAREV,A.I. 1969.0n the source of silica in the formation of cavities with crystals in carbonate rocks. ( in Russian ) Translated in Int. Geol. Review., v.12, 1018-1021.

MACKINTOSH,D.M.1967.Veining in Namurian turbidite sandstones of the Crackington formation, North Cornwall.Geol. Mag. v.104, 7E-85.

RAMBERG,H.1961. A study of veins in Caledonian rocks around Trondheim Fjord, Norway. Norsk Geol. Tidsskrift,v.41, 1-43. 231 Chapter 8

The distribution of chemical elements and mineral species in the deformed matrix around rigid cylindrical inclusions.

Introduction

One of the central problems of pressure solution is what ' factor, or factors, control the manner in which minerals are dissolved, transferred, and precipitated. The factors are numerous and complex; and vary in their relative importance with respect to the prevailing boundary conditions. Of the few studies that have been made of the distribution of chemical elements, or mineral species within structures, only one attempt has been made directly to relate the stress field to chemical variations. Berglund and Ekstrom (1974) have measured the Fe content of sphalerite in and around a boudinaged layer. However, they concluded that variations in the fugacity of sulphur (fs2) exert almost as large an influence on Fe concentration as does the pressure effect. Durney (1972b) and StrOmard (1973) have drawn some inferences concerning the relation between formation of pressure shadows and the distribution of mean stress (a ) in the matrix around rigid inclusions.

Boundary conditions.

Stress clearly plays an important role in pressure solutions; and yet evidence has been presented in earlier chapters from which one may infer that kinetic effects have a significant influence in the formation of pressure solution structures. 232 These, and other factors are evaluated in this study, in relation to pressure solution activity. In approaching the problem as outlined above the author deofl that two main requirements were necessary. First, to mea7.':?e analytically small scale spatial variations in the relative concentration of chemical and mineral species. Second, to work with known and well-defined boundary conditions, in order to limit the number of factors which could exert an influence on pressure solution activity. With these considerations in mind the author decided that the most suitable structure to analyse was the deformed matrix around rigid cylindrical inclusions. The reasons for this choice are as follows -

1. The stress distribution in the deformed matrix around rigid cylindrical inclusions has been established analytically (Savin, 1930), and by numerical techniques (Casey, personal communication). 2. The deformed matrix around rigid objects has well defined zones of mineral depletion, and associated mineral gain in shadow areas; and hence comprises the closest available approximation to a chemically closed system.

In the field samples were selected to satisfy the following requirements, as relating to the boundary conditions listed above.

1. An effectively perfectly rigid cylindrical inclusion. 2. An initially homogeneous matrix. 3. The matrix to be effectively semi-infinite and unweather- ed, as exposed at the surface. 233 4. The matrix to be initially isotropic with respect to diffusion and material anisotropy, This condition requires a low preferred orientation of layer silicate minerals. 5. The matrix to contain a mineral or minerals of high °relative pressure solubility, and minerals of low relative pressure solubilty. 6. rrossLlre slados should be the replacement type of St-r6m -E,rd C9717)), so th ,t chemical gradients are continueus- ly variable. 7. The process of pressure solution should not have gone too far, nor pro,;ressed too little. The former situation was excluded because a high degree of diffusional and bulk anisotropy would result from the preferred orientation of micas in a matrix which had experienced a near total loss of pressure soluble minerals. The latter situation was excluded because chemical gradients would not be large enough to measure analytically. For the intermediate case there is some possibility that one is sampling a system at equilibrium. 8. The overall strain history of the matrix should be ir- rotational. The closest approximation to this condition was considered to be the hinge region of a fold, for which there was no field evidence of migration. 9. The specimens should be large enough so that a sufficiently sensitive sampling grid could be used to measure small scale variations in rock chemistry. 10. Grain size of the matrix should be more than three orders of magnitude less than the diameter of the rigid inclusion.

At the localities where samples were collected estimates have been made of the temperature obtaining during deformation. Fig 8.1 XRD Chart. Iron Co/ K4 , 27min. Smear mount

Sample 1 , Lot schental - mineralogy

Peaks i n angles of 2 8 234

Only two samples were finally selected as closely ap- proximatin the conditions outlined above. The first was a clay pellet (7 cm diameter) in a dolomite-phyllosilicate matrix, from Trias marls of the Aar Tarautochthon, Switzerland. The second was a coarse-grained volcanic bomb in the deformed matrix of Ca boriferous volcanics at Trebarwith Strand, North Cornwall. Specimens of about 0-5 x 0.5 x 0.2 m ( e•-, 120 IL-5) were collected in order to be able to establish the limits of diffusive and chemical fluctuations in the system.

Methods of analysis.

Optical methods were used to study the textures, and subjectively to assess the distribution of mineral species, around cylindrical nodules. Because the matrix was fine-grained mineral identification was assisted by XRD. In addition, XRD was used to make a semiquantitative estimate of the relative concentration of mineral species in zones of depletion and shadow areas respectively, as a preliminary to full scale chemical analysis (Fig. 8.1). The matrix was sawn up on a 1.0 cm square grid, and analysed for 8major and three trace elements by X-ray fluores- cence (calcium, potassium, magnesium, aluminium, silicon, manganese, iron, titanium, and strontium, rubidium, barium, respectively). Major elements were analysed from discs prepared at 1 'I. 7 dilution using the fusion technique. Trace elements were analysed from undiluted discs in order to obtain sufficiently high ratios of peak to background counting rate. Trace elements are quoted relative to international standard GI (Barium 1,200 ppm; Strontium 250 ppm; Rubidium 220 ppm), Flanagan (1973). 235

120 discs were analysed for the 8 major and 3 trace elements. Because Rubidium, Strontium and Barium are present at a low level of concentration the analyses were run in triplicate, in order to obtain more accurate results.

T-T

The results e s presented in contoured di es of the Con-

'") at Etion of cl,emical elements (.0 . 4. U0 If it is' as, sufmed tLt, for instance, all of the calcium and lootassium ore

.ns dolomite and muc.covite minerals resectivel7 (. 0 .1 then those diazrarls also represent the distribution of mineral species. Silicon is held in the phyllosilicates and in quartz. It is of interest to note that the diagram for this element (rig. 8.4) does not reflect the same trends as are apparent for other elements. The mineralogy of sample LOT was determined optically, and by MD, as dolomite-muscovite-chlorite--quartz-iron oxides. Optical studies showed a depletion in dolomite, and concomitant concentration of layer silicates and iron oxides in areas of pressure solution; together with the preferred orientation of layer silicates and residual dolomite grains. Pressure solution is concentrated in discrete, spaced, sub-planar zones with 'hyperbolic type' boundary conditions; and has not developed in a uniform and diffuse manner within the matrix. Shadow areas are characterised by a high relative concentration of dolomite and low relative concentration of layer silicates. Dolomite grains are equant; phyllosilicatem exhibit a low degree of preferred orientation.

CONCLUSIONS The distribution of pressure soluble minerals is broadly consistent with a mean stress grid that has been deformed in 236 the direction of principal stress. This feature results from zones of loss moving progressively toward the inclusion, and zones of gain moving progressively away from the inclusion, as deformation proceeds. Hence, the predominant loss of pressure soluble minerals is from zones adjacent to the inclusion, and precipitation of the minerals is predominantly in shadow areas also adjacent to the particle. This inference concerning the locus of precipitation is consistent with the growth of fibrous secondary minerals at the inclusion-shadow interface. Chemical gradients in the zones of depletion are steeper than predicted from solution which is crontrolled entirely by meanestress. It is suggested that a strain softening takes place at loci of depletion, and this results in discrete, sub- planar, zones with 'hyperbolic type' boundary conditions as described in the section on optical textures. It is contended that the strain softening is due to the following kinetic effects -

1. At loci of depletion there is a higher relative number of dolomite-phyllosilicate interfaces than in matrix which has experienced no loss. Solution takes place preferentially at such interfaces. 2. The higher relative concentration of layer silicates at loci of depletion gives a greater diffusivity of migrating ions. In addition, the preferred orientation of layer.:.silicates at such loci introduces a diffusional anisotropy in the matrix, which is normal to the principal stress. 3. The preferred shape orientation of pressure soluble minerals increases as solution proceeds. This shape factor results in a larger surface area normal to the principal stress. 237 Ca

3.1 Fe 34

3.0

Z-t•

Fig. 8.2. Contours of percentage concentration of oxides of major elements

in the deformed matrix around a rigid cylindrical inclusion. • Al

tif 238

barium

3i.0 340 32o 0

300

o 1Vo

rubidium

strontium

boo

Fig. 8.3 Contours of trace element concentration , ppm , in the

deformed matrix around a rigid cylindrical inclusion. 239 4. The surface energy-volume ratio of residual pressure soluble minerals increases as the grain size decreases.

Conversely, it is suggested that a strain hardening effect takes place in the shadow areas. Evidence has been adduced he rate limitin step of pressure solution in a chemically closed srvestem is the rate of precipitation.iien_ ce, precipitation will be favoured only in those areas where the minerals are at a minimum state of mean stress, i.e. to the p,erticle. One may .4, plain from this inference the rre- dominance of rrecipitation in areas directly adjacent to the inclusion, as mentioned above. Calculations of volume chance for the two systems analysed are given in Fi:. 8.7. Fluctuations in the concentration of elements about the matrix norm extend for approximately 5 particle diameters-parallel to the direction of minimum principal stress, and 1 particle diameter parallel to the direction of maximum principal stress. The data are consistent with a passive increase in the proportion of insoluble minerals within domains of depletion, i.e. there is no evidence for coupled diffusion having taken place. Estimates of the temperature obtaining during deformation for the samples from Lotschental and Trebarwith are 185 ± 10°C and 276 ± o C respectively. (These estimates were made usin both fluid inclusion and oxygen isotoye thermometry - see Fig. 4.9.) In the former case, at relatively low temperature, only dolomite has deformed by pressure solution. It would appear from petrographic studies that migration of quartz has not been significant. Hence one may infer that for this specified dolomitic lithology, at the temperature quoted, dolomite deforms by pressure solution faster than quartz. This result 240 Si

___ 10 \ ,

....

Fig. 8.1,. Contours of percentage concentration of oxides of major

elements in the deformed matrix around a rigid cylindrical inclusion.

N.Devon. 242 is inconsistent with the conclusions drawn in chapter 5, concerning the temperature dependent kinetics of pressure solution. Both quartz and calcite have deformed by pressure solution in the sample from Trebarwith. One may infer from the contours on that a ecom;uound' y.:_essure shadow has formed, with a relatively hD proportion of quartz adjacent to the particle, en relati.ve] hi7h procrtion of calcite further away from t're cle. Peace, at273 in this.volcanic litholo7,71 either the rate of diffusion of calcite is faster than that of quartz, or 'precipitation of quartz is kinetically favoured in areas of lower stress compared to calcite. Finally it is of interest to note that trace element distribution is not so well defined as is the distribution of major elements. I suggest that trace element variations follow in part the distribution of the host minerals in which they are held. However, there may be a superimposed effect due to the pressure dependent solubility of trace elements in the host mineral. 243

REFERENCES - CHAPTER 8

BERGLUND, S. and T.K. EKSTR6M. 1974. Sphalerite composition in relation to the stress distribution of a boudinage. Lithos, v. 7, 1-6. DURNEY, D.W. 1972b. Ph.D. Thesis. London. FLANAGAN, F.T. 1973. 1972 values for international geochemical reference samples. Geochim. et Cosmochim. Acta, v. 37, 1189-1200. SAVIN, F.G. 1930. Stress concentrations around holes. STR8MGRRD, K.E. 1973. Stress distribution during formation of boudinage and pressure shadows. Tectonophysical v. 16, 215-248. CONCLUSIONS

In chapters 1 and 2 it was demonstrated that research on pressure solution has largely been directed towards a description of the phenomenon. An attempt has been made in this thesis to bridge the gap between

pure field studies of pressure solution and specialised theoretical treat-

ments of deformation mechanisms, in order to draw some general conclusions concerning the role of pressure solution in crustal deformation.

Oxygen isotope geochemistry has been shown to be a powerful method for determining physical conditions obtaining in the crust during deformation. Limiting conditions for the use of fluid inclusion techniques on

material from tectonic veins have been evaluated in relation to deformation mechanisms in vein minerals.

Quantitative information has been obtained on the prevailing conditions of temperature and pressure in crustal environments where pressure solution was observed to be a dominant deformation mechanism. Estimates of the temperature and pressure intervals over which quartz bearing assem-

blages respond to deformation largely by pressure solution are>30°C to

450 ± 50°C, and >25b to,,6000b.

New data is presented on the thermal stability limits of index min-

erals within the intermediate pressure facies series of- regional metamor-

phism. Temperatures of 300, 400, 500, and 580°C are inferred for the

lower end of the chlorite, biotite, almandine, and kyanite isograds res-

pectivly. It is believed that this is one of the first attempts to define the conditions of temperature and fluid pressure obtaining during the str-

uctural - metamorphic history of a tectonic belt, together with information

on mineral assemblages and deformation mechanisms.

Tectonic veins generally form in isotopic and mineralogical equilibrium

with the host rock when pressure solution is the dominant deformation mechanism. The size of an isotopic eqilibrium system for oxygen may be greater than 100m in a tectonic environment where pressure solution is the dominant deformation mechanism. Hydrothermal fluids in low grade tectonic environments are subcritical H20-NaCl-0O2 solutions of density 0.8+ . It is sugg-

,ested, from detailed information on the physico-chemical properties of these fluids,that the hydro--thermal fluids associated with many ore deposits have

not been derived directly from tectonic environments. The water- rock ratio

in deformed rock5is low, and hydrothermal fluids maintain thermal equilib-

rium with the surroundings.

The theoretical prediction of nonequilibrium isotopic fractionation in

stress fields has been confirmed by empirical measurement.

The orientation of the principal stress tensors governsthe overall

development of structures which have formed by pressure solution. However,

it has been demonstrated that kinetic factors control the detailed distrib-

ution of mineral species and chemical elements in rocks deformed by pressure

solution.

Because rocks- generally deform by two or more mechanisms simultaneously,

it is difficult to separate the effects of pressure solution from effects

which may be attributed to other deformation mechanisms. In this context,

future studies using transmission electron microscopy to define the micro-

structural characteristics of individual deformation mechanisms should be

of particular value. Microchemical analysis of grain boundary processes

and of minerals that have deformed by pressure solution is an important

avenue of research.Oetailed information concerning the role of fluids in

deformation processes may be obtained from further use of isotope geochem-

istry.Future studies on the kinetics of deformation mechanisms are central

to an understanding of pressure solution. With this type of information,

and refined data on crustal conditions it,should be possible to construct

deformation maps for the major rock forming minerals. 244 APPENDIX I

Density measurements on fluid inclusions

Sample Density Standard deviation Number of observations PV 0.90 0.01 47 TV 0.92 0.01 31 NV 0.87 0.008 54 CV 0.89 0.02 20 LV 0.94 0.01 42

APPENDIX II

H20/002 determinations on fluid inclusions ratio Sample Water CO2 WT% CO2 H20/C02 content content

PV 2.19mg 0.001mg 1500:1 TV 0.82 0.07 8.5% 12;1 NV 0.30 0.02 6.7% 15:1 CV 1.44 0.26 18% 6:1 LV 0.44 ' 0.003 - 1500:1

APPENDIX III

Measurements of salinity on fluid inclusions

Sample Wt% NaC1 Freezing point PV 5.4 -85 °C TV 9.7 -54 NV 10.3 -78 CV 6.7 -80 LV 3.8 -95 245 Appendix IV

Preparation of mineral separates for isotopic determination

1. INTRODUCTION Both mineral separation and the determination of stable isotope abundance in natural materials by mass spectrometry are diffidult, expensive, and time consuming operations. As a consequence it is important to have a clear idea of the hypothesis to be tested, or information gained from isotopic data, before embarking on a project. For efficient sampling, and subsequently for accurate interpretation of the isotopic data, it is necessary to have detailed knowledge of the petrology, metamorphism, and structure of the area concerned. The aim, of mineral separation procedure is to produce approximately 200 milligram quantities of ultra pure quartz, calcite, muscovite or chlorite, and magnetite or haematite from the host rock. In the field, samples should be selected which have the relevant mineralogy, which are of maximum grain size, and which are of several kilogrammes mass. All samples which have been collected for isotopic:analysis are first examined in thin section for grain size, texture, and mineralogy. Specimens in which there is petrographic evidence for retrograde mineral assemblages are eliminated. Only those specimens which have equilibrium assemblages are selected for obtaining mineral separates. 246 2. CRUSHING AND SIEVING The rock is reduced by sawing to pieces of a size suitable for jaw crushing; weathered margins and small veins are removed at this stage. Part of the specimen should be conserved against accident, or for purposes of comparison at a later date. After several passes through a jaw crusher the material is assembled into thee* following four size factions by sieving -

1. 250+ microns 2. 250- to 150+ 3. 150- to 105+ 4. 105-

Coarse factions are further reduced by cone grinding to make up the necessary quantity of approximately 500 gm. in each size range. All factions are washed in water to remove fines-, which inhibit both magnetic and heavy liquid separations, then washed in acetone and dried. Examination under binocular microscope reveals whether the different mineral grains have been liberated, or whether they remain intergrown. The minerals in medium grained rocks may conveniently be separated within size ranges 2. or 3. Fine grained rocks require the material to be reground, and sieved into 105- to 63+, and 63-- to 45+ micron factions. Separations are always done at the maximum size compatible with liberation of the mineral grains, because the heavy liquid technique is faster and more efficient, and because hand picking is more rapid with large grains. The size faction which contains free mineral grains is passed through a 'Davies' large scale non-entraining magnetic separator, which divides the sample into non-magnetics, weakly: magnetics and strongly magnetics. 247

3. QUARTZ AND CALCITE Bromoform or tetrabromoethane, with dimethylformamide as diluant are used to make up heavy liquids of densities 2'68 and 2.80, measured by a Westphal balance. In:order to obtain quartz and calcite, separates the non- magnetic faction,, which contains those minerals together with feldspar, some phyllosilicates and some ore minerals, is placed in heavy liquid density 2-68. Quartz and feldspars form a float: the sink is collected, and in liquid of density 2.80 calcite floats, the phyllosilicates and ore minerals forming a sink. Each separate is washed in acetone and dried. k 'Frantz' isodynamic small scale magnetic separator is used to clean up calcite from the 2.80 float; and operated in a diamagnetic configuration to obtain pure quartz from the 2.68 float. Should the quartz be iron stained it is treated in boiling hydrochloric acid for 10 minutes. Other separates are stored for reference or for future use.

4. IRON OXIDES Magnetite is concentrated in the Davies strongly magnetic faction which is combined with material drawn off the weakly magnetic faction using a hand magnet. In methylene iodide, density 3.3, magnetite forms a sink, with impurities previously trapped in the magnetic chains floating. Final purification is done by hand picking with a,steel needle. Haematite, together with some ore minerals and garnet if present, form a sink when the Davies weakly magnetic faction is placed in methylene iodide. A variable hand magnet is used to clean up the haematite. 348 Iron oxides were present at a level of only 1 ppm. in some samples, and as3a consequence it was necessary to treat a large quantity of material in order to obtain the minimum mass of 5 mg which is acceptable for isotopic determinations. Some difficulty was experienced in separating magnetite and pyrrhotite because the magnetic properties are very similar, and the material was generally too fine grained tor effective hand picking. The presence of pyrrhotite was first suspected from low oxygen yields: in the fluorination line. Although pyrrhotite is not an oxygen bearing mineral, reaction with bromine pentafluoride gives FS6, which cannot be frozen down completely as waste, and which interferes with gas flow through the spectrometer viscous leaks. Pyrrhotite was removed' by reaction with concentrated HCI, which attacks the magnetite relatively slowly. In order to 8 ensure that this treatment did not cause any KIE ; 0/160 determinations were made on two samples of the same pure magnetite separate, one of which had been reacted with HC1. 16 The 180/ 0 values obtained for these two samples, X0 and X01101, were 2.16 and 2.20 permil. respectively. This difference of 0.04 permil. is well within the analytical error of 0.1 permil.

5. PHYLLOSILICATES Muscovite and chlorite are contained in the magnetic separate which formed a 3.3 float. In heavy liquid density 2'80 phyllosilicates form a sink, from which biotite, chlorite and muscovite may be progressively drawn off using the 'Frantz'. :ig AIV 1 X‘FR D Chart Iron Co K,„, Smear mount

0 Host and vein phyllosilicate separates Peaks in angles of 2 0 249 At each stage the separates are examined in reflected light under binocular microscope, and the mineralogy tabulated to compare with that observed in thin section.

6. "MB OF XRD In oases where the microscopic identification of minerals is-in doubt, or when the absolute purity of separates is optically indeterminate, X-ray diffraction (XRD) analysis is performed on smear slide samples of the separate in question. This procedure was necessary in three cases:

1. Some minerals could not be, identified opitcally. The phyllosilicate separates from vein samples MARV 31 and MARV 32 were confirmed to be chlorite, but with trace quantities of fine grained quartz-muscovite-calcite composites present. A mineral from sample TINTH was determined as composite chlorite- muscovite-quartz. XRD charts for these separates are illustrated in Fig. AIV 1.

2. It is common for carbonate minerals partially to replace one another or to form a solid solution series. Hence, it was important to ensure that calcite or dolomite separates from vein and host rock samples were not contaminated by ankerite or other carbonate minerals. XRD charts for some of these separates are illustrated in Figs. AIV 2 and AIV 3.

3. There was some doubt concerning the efficiency of heavy liquid separation for grains in the size range 63- to 45+ microns. Quartz peaks (2e° = 24.3, 31.0, 42.7, 46.2, 49.81 53..4) on the XRD charts illustrated in Fig. AIV 4 are stronger for the 24.68 floats than for the 2.68 sink. However, quartz is present at a level of greater than 10% in the sink, and phyllosilicates are present in the floats.

H (NI CV. lD In -4 1 C1

o wkwwoiol" P.4ftmy#0 4t.44600iiorvfloo&veoi sg Lo?, Loi.3 g .4 (N CO LO Cs O cO (0 -4 NI 0 a) 40 .4 7.4rf (.•=4) m CO 01 OCM N N

O co...t Co N N co CO r-. u5t.il a to c,.1 u5 L.() Lotn Lr) -4 at co

CO LID -4 NI Co co0 to 4 N 0 00 (0 N 0 OD CO -4 NI 0 Lo N O c0 dD tr) ti) LC) -4 4 4 04 -4 C) C) N IN `i6.1 N C..1 ,t2 IC) 5

OLCI

TI NT V o

Fig A I V. 2 X R D Chart. Iron Co Kat , 27min . Smear mount.

Host and vein calcite separates. Peaks in angles of 20 250

Such samples which did not separate efficiently using the normal heavy liquid method,were purified by centrifuging.

7. USE OF CENTRIFUGE AND HAND PICKING At each stage the separates are examined in reflected light under binocular microscope, and the mineralogy tabulated to compare with that observed in thin section. All samples are finally checked- by hand picking to ensure absolute purity, because small amounts of contaminant may give rise to large errors in isotopic determinations. 36 quartz; 12 carbonate; 10 muscovite; 12 chlorite; and 20 iron oxide separates were finally obtained from 29 whole rock samples. This required 35 'Davies' operations, 255 heavy liquid separations, 800 'Frantz' operations and 120 hours of hand picking. Average processing time per 200 mg separate was 10 hours. Vi r5TrY' Alilituivi. I H

Fig AIV.3 XRD Chart. Iron Co K, 2°/min. Smear mount Vein dolomite separates. Peaks in angles of 2 e

▪

— ' •

• ■ &e. D2 FLOAT (N, • c..7,

(N CO Cr) C (r) G) ± LT", (.° co (Q. (N LC) LC) -4.4 IN CV

o oo -4 co (0 0 ..' c0 (0 (N co LID -4 (N 0 CO LO N (.0 in Lo LI) 1.0 'AD%) -4 -4 .4 () Cr) Cr) iS) (N (N ▪ c0 co -4

Fig A IV. 4 X R D Chart. Iron Co K.( , 27min . Smear mount.

Quartz - dolomite - phyllosilicate separates- Lotschental Peaks in angles of 2 0° 251

Appendix V

Fluorination and mass spectrometry

1.4, INTRODUCTION Oxygen isotope abundance is normally derived from high precision mass spectrometric determination of the isotopic composition of carbon dioxide (002), whibh is a convenient gas to handle. Silicate and Metal oxide minerals are reacted with a fluorinating agent to extract oxygen, which is then converted. to CO 2' It is important that the extraction and conversion are reproducible and quantitative in order that kinetic isotope effects (KIE) do not swamp the small equilibrium fractionations which are to be measured.

2. EXTRACTION AND CONVERSION A mass of the mineral sample which is sufficient to yield about 100 micromolea of oxygen (02) is weighed into a steel container, and, together with a nickel reaction bomb, placed in a dry box for 24- hours. The precise mass used depends on the mineral type, but is generally about 10 mg. The sample is loaded into the bomb in a dry box, because fluorinating agents react with atmospheric water to yield contaminant oxygen. The bombs are then sealed pn to the vacuum extraction line, which is of conventional design, essentially as described by Taylor and Epstein (1962), but of all stainless. steel construction, and holds a vacuum of 105 torr. Aliquots of purified bromine pentafluoride (BrFs) are frozen down in vacuo into each of the four bombs, and reacted 252

with the mineral at 550°C. for 14 hours. Oxygen is separated and other reaction products such as Br, BrF from excess BrF5 3 etc. in a series of liquid nitrogen cold traps. Oxygen is converted to CO2 by reaction over a hot carbon - rod wiAl platinum catalyst. This procedure is done in a series "of- aliquots in order to avoid 'CIE, because, as for the oxygen attraction by fluorination, conversion of 02 to Ce2 follows the Rayleigh distillation equation; and large kinetic fractionations may occur unless the reaction goes- to completion.

The CO2 yield is measured in a mercury manometer and compared with the theoretical yield calculated from the mineral chemical formula. Isotopic results-on any sample from which the yield was not 100% ± 5%, are treated with caution, and re- peat runs done if possible. The CO2 is dried in a 2- butoxyethanol solid carbon dioxide cold trap and passed directly into the mass spectrometer.

3. MASS- SPEC TRONOVIETRY The instrument at AWRE Aldermaston which is used for high precision isotopic analysis of CO2 is a 30 cm radius, 90° sector, double collecting mass spectrometer, described by Beckinsale and others (1973a). The abundance of carbon and oxygen isotopes are such that

in CO2 prepared from natural materials the relative proportions of the molecular species m/e = 44, 45, and 46 are approximately 98-42%, 1-17%, and 0.40% respectively. The requirement of the mass spectrometer, therefore, is to measure ratios of about 1 : 100 (45/44) and 1 : 250 (46/44 + 45) with errors of less than 0°01%. (Beckinsale and others 1973b). CO2 pressure in the sample and standard reservoirs are equalised using mercury column compressors, and the CO2 253 introduced into the spectrometer through two separate viscous leaks. The sample and standard are alternately admitted to the ion source using magnetic changeover valves. m/e ratios are calculated from 100 second counting-periods, using a capacitor integrating system. (BeckL:sale and others 1973a). The, mass spectrometer is tuned firstly for the 46 ratio, from which the 0IS/016- result is calculated. The 45 ratio is also measured, primarily as a sensitive check on the presence of water, and also for computing a small correction to the 46 ratio (Craig 1957). To convert results to the Standard Mean Ocean Water (SNOW) scale a working standard (Loch Aline Glass Sand) is used, which has been calibrated at + 10°87% SNOW. This working standard is run approximately every four samples. (Beckinsale and others 1973b). Isotope ratios expressed relative to one standard (X) may be expressed relative to another standard (Y) using the expression S(Sa-Y) = d(Sa-X) + 8(X-Y) + 103 d(Sa-X)(X-Y) where 6(Sa-Y) is the isotopic composition of the sample (Se.), relative to the standard (Y) etc., and all the values are in permil. (Craig 1961). Errors which are attributable to fluorination and conversion amount to ± o.09% : the mass spectrometer is accurate to ± 0.01% , giving a total error in the results of t 0.1 % .

4. ISOTOPIC ANALYSIS OF CARBONATES The carbonate determinations were done in the department of Pleistocene Research, Cambridge. Calcite and dolomite are reacted with 100% phosphoric acid (H3PO4) at 50°C to release CO2. Only 2/3 of the carbonate 254 oxygen is converted to CO,, in this reaction. In addition, the reaction for dolomite never runs to completion. As a consequence of these factors the analytically determined 6 values are corrected for a kinetic fractionation which is 1.01008 for calcite, and 101090 for dolomite. (Perry, 1972). Apart from kinetic effects, the principles of extraction and mass-spectrometry are the same for carbonates as has been. described for silicate and iron oxide minerals.

REFERENCES: - APPENDIX V

BECKINSALE, R.D., N.J. FREEMAN, M.C. JACKSON, R.E. POWELL, and W.A.P. YOUNG. 1973a. A 30 cm radius 900 sector double collecting mass spectrometer with a capacitor integrating detector for high precision isotopic analysis of carbon dioxide. Int. Jour. Mass Spectrometry and Ion Physics, v. 12, 299-308. BECKINSALE, R.D., and others. 1973b. NERC Stable Isotope Project progress report, 1970-1973. CRAIG, H. 1957. Isotopic standards for carbon and oxygen and the correction factors for mass spectrometric analysis of carbon dioxides. Geochim. Cosmochim. Acta, v. 12, 133-149. CRAIG, H. 1961. Standard for reporting concentrations of deuterium and oxygen -L°0 in natural water. Science, v. 133, 1833-1834. PERRY, E.C. and F.C. TAN. 1973. Significance of Oxygen and Carbon isotope variations in early Precambrian cherts and carbonate rocks of Southern Africa. Geol. Soc. America Bull., v. 83, 647-664. TAYLOR, H.P. and S. EPSTEIN. 1962. Relationship between 180/1 ratios in coexisting minerals of igneous and metamorphic rocks. Pt. I. Principles and experimental results. Geol. Soc. America Bull., v. 73, 461-480. 255

Appendix VI

180/160 values of mineral separates analysed with respect to SMOW

Quartz Muscovite Magnetite LF 15.10 LL 10.89 TI 0.87 GF- 11.96 CLA 9.63 GP 1.98 LOTD 19'23 BL 8-24 LX 3.05* CA 12'21 CA 9.13 LL 0.49* XI 14.49 GP 9.55 CA 8.59 BL 114.87 9.45 BL 9.71 IN 12*47 TI 9.79 TU 0.81 LOTV 26.02 SK 9.35 0.81 TU 5..29 TU 5.55 SK 12.26 KB 9.42 CL 12-82 9-56 Haematite KB 12.50 NDH 10.61 TI 12.40 10-71 LF 7.64* PMV 15.87 KI 4.55* CMV 18.54 IN 5.12 LO 13.47 Chlorite LX 8.43* EX 14.49 MAR 0.93 13.22 LX 9.67 ZCM 6.94* LL 13.03 9-44 TH 5.70 CLA. 13.02 LC 8.40 NDH 5.25* CMI 17.30 CL3 8.03 LOT 1.85 17.10 LF 10.26 2CM 17-39 CK 10.41 ?INV 13.20 MARV 8.99 CL3 11.87 LL 7.47 Dolomite LOTA 23.13 7.59 LOTB 22.28 BL 5.80, PMV -3.99 LOTH 24.77 TH 15.31 PMH -4.33 NDV 15.88 LMH 7.22 NDV -6.15 TH 17.40 LOTH 6.44 CMV -12:!..72 PESH 15-93 TV 11.49 CMH -11.83 PEHI 14.99 LOTA -2.45 PEHO 14.61 LOTB -4.50 2NH 14.95 LOTC -2.99 15.09 LOTD -2.09 CK3 14.25 21IH 14..92 Calcite (FDB) LOT V -7.98 LOT H -6-77 TV -13.62

* Low oxygen yield in fluorination line