Fractography: fracture topography as a tool in fracture mechanics and stress analysis

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Fractography: fracture topography as a tool in fracture mechanics and stress analysis

Geological Society Special Publications Series Editor A. J. FLEET

G E O L O G I C A L SOCIETY SPECIAL PUBLICATION NO. 92

Fractography: fracture topography as a tool in fracture mechanics and stress analysis EDITED BY

M O H A M M E D S. A M E E N Independent Consultant, UK

1995 Published by The Geological Society London

THE G E O L O G I C A L SOCIETY The Society was founded in 1807 as The Geological Society of London and is the oldest geological society in the world. It received its Royal Charter in 1825 for the purpose of 'investigating the mineral structure of the Earth'. The Society is Britain's national society for geology with a membership of 7500. It has countrywide coverage and approximately 1000 members reside overseas. The Society is responsible for all aspects of the geological sciences including professional matters. The Society has its own publishing house, which produces the Society's international journals, books and maps, and which acts as the European distributor for publications of the American Association of Petroleum Geologists, SEPM and the Geological Society of America. Fellowship is open to those holding a recognized honours degree in geology or cognate subject and who have at least two years' relevant postgraduate experience, or who have not less than six years' relevant experience in geology or a cognate subject. A Fellow who has not less than five years' relevant postgraduate experience in the practice of geology may apply for validation and, subject to approval, may be able to use the designatory letters C. Geol (Chartered Geologist). Further information about the Society is available from the Membership Manager, The Geological Society, Burlington House, Piccadilly, London W1V 0JU, UK. The Society is a Registered Charity, No. 210161. Published by The Geological Society from: The Geological Society Publishing House Unit 7, Brassmill Enterprise Centre Brassmill Lane Bath BA1 3JN UK (Orders: Tel. 01225 445046 Fax 01225 442836)

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Contents AMEEN, M. S. Introduction

1

Part 1: Fractographic studies Methods of observation, data collection and processing]quantitative fractography AYDAN, O. & SHIMIZU,Y. Surface morphology characteristics of rock discontinuities with particular reference to their genesis

11

JESSELL, M. W., Cox, S. J. D., SCHWARZE,P. & POWER, W. L. The anisotropy of surface roughness measured using a digital photogrammetric technique

27

Application of fractography to fracture analysis from core BANKWITZ, P. & BANKWITZ, E. Fractographic features on joints of KTB drill cores (Bavaria, Germany)

39

Fractography applied to the analysis of experimentally produced fractures in rocks and non-rocks KULANDER, B. R. & DEAN, S. L. Observations on fractography with laboratory experiments for geologists

59

BUTENUTH, C. & DE FREITAS, M. H. The character of rock surfaces formed in Mode I

83

Fractography applied to fracture analysis in field studies aimed at understanding regional tectonics AMEEN, M. S. Fractography and fracture characterization in the Permo-Triassic sandstones and the Lower Palaeozoic Basement, West Cumbria, UK

97

AMEEN, M. S. Fracture characterization in the Chalk and the evolution of the Thanet monocline, Kent, southern England

149

ROBERTS, J. C. Fracture surface markings in Liassic limestone at Lavernock Point, South Wales

175

Part 2: Non-fractographic studies Miscellaneous studies of fractures COSGROVE, J. W. The expression of hydraulic fracturing in rocks and sediments

187

GOODWlN, A. M. Spatial change in joint geometry in the Chalk of eastern England

197

GROSS, M. R., FISCHER, M. P., ENGELDER, T. & GREENFIELD, R. J. Factors controlling joint spacing in interbedded sedimentary rocks: integrating numerical models with field observations from the Monterey Formation, USA

215

Index

235

From Ameen, M. S. (ed.), 1995, Fractography: fracture topography as a tool in fracture mechanics and stress analysis

Geological Society Special Publication No. 92, pp. 1-10

Fractography: fracture topography as a tool in fracture mechanics and stress analysis. An introduction M O H A M M E D S. AMEEN Independent Consultant, 122 Tachbrook Street, Pimlico, London S W1 V 2ND, UK

What is fraetography? Fractography can be defined as the science which deals with the description, analysis and interpretation of fracture surface morphologies (fracture topographies) and links them to the causative stresses, mechanisms and subsequent evolution of the fractures. Fractography can be classified according to the field (scale) of observation into: (a) macrofractography, dealing with features which can be studied using the naked eye or a hand lens, and (b) microfractography, dealing with features which can be studied with the aid of optical microscopes .or electron microscopes only. The term fractography was first used by the metallurgist Carl A. Zapffe (Zapffe & Clogg 1944), and was introduced to geology by Byron Kulander in the 1970s (Kulander et al. 1979). The discipline of fractography (as a science) in geology is therefore new, even though nearly a century ago Woodworth (1896) was the first to describe joint surface morphology from field observations. Fractography is hardly mentioned in text books on structural geology, consequently many geologists are unfamiliar with this term and its implications and may confuse the term with fractals. The subject is largely neglected by those who deal with fracture characterization of rocks (e.g. core laboratories) in spite of its obvious potential. In addition, the interpretation of fracture surface features in rocks has been almost exclusively qualitative, and frequently controversial. The state of infancy and neglect of fractography in geology is further reflected by the fact that in this field there has been no clear definition of fractography. In addition there is no universally accepted, unified nomenclature for fractographic features. All these factors prompted this special publication, the first thematic volume on fractography in geology.

The evolution of fraetography in geology Fractography in geology has evolved to its present state through four stages: (a) The embryonic stage started with Woodworth's studies (1896) and the classification of surface structures on joint faces. (b) Nearly half a century during which no studies on fracture surface features in rocks were recorded.

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(c) A phase of fracture surface studies started in the early 1940s (Parker, 1942) and culminating in the publication of 2 key papers by Hodgson (1961) and Roberts (1961). Hodgson reviews the definitions and implications of those features identified by Woodworth (1896). (d) The final phase was stimulated by Hodgson's and Roberts' publications (1961) and included several studies (e.g. Bankwitz 1966; Syme-Gash 1971) and reached its climax when Kulander et al. (1979) introduced the term fractography to the studies of surface features in rocks. This initiated quantitative study of the topic both by using the already established name from metallurgy (i.e. fractography of Zapffe & Clogg 1944), and by leading to a thorough description and explanation of macrofractographic features in both field and core studies. This phase continues to the present and is characterized by relatively greater interest than the previous stages, though is still limited (mostly field macrofractography). For a review of the subject of fractography and comprehensive bibliography, the reader should refer to Kulander et al. (1979) and a recent review by Bahat (1991).

Where does fractography stand within the wide field of fracture studies? An understanding of fractures and their mechanics is a matter of great interest in both material and geological sciences. Workers within these two fields are driven to determine fracture mechanics for two different reasons. A metallurgist for example is concerned with producing alloys which are less susceptible to failure. However, an earth scientist's or engineer's interest is considerably more complicated in that it embraces mechanisms for enhancing fracture development (e.g. in a hydrocarbon reservoir), methods of reducing the chances of fracture development (e.g. in a nuclear waste repository), and the quantification of fracture surface morphologies as they influence fluid flow and rock stability. In addition to the considerably complex nature and wide spectrum of a geologist's task compared to that of a metallurgist, the former is faced with dealing with a vastly heterogeneous natural material (rocks) under circumstances of both spatial and temporal scale that fall well beyond the well constrained time/space scales in the studies of metals and alloys. This difference has undoubtedly affected the pace of progress within the disciplines of fracture mechanics of rocks. Deprived of the chance of direct observation of natural fracture development in rocks, geologists followed two approaches to infer fracture mechanics: they study experimentally induced fractures in rocks, and observe fractures and their host regional/local structures in the field. Both approaches involved analysis of fractures as a 'population' and as 'individuals'. Fracture population studies generally separate a particular population of fractures in a rock mass into sets according to their 3D orientation and spatial and temporal distribution, and try to link the sets to their causative stresses. Individual fracture studies deal with the evolution of an individual fracture and its causative stresses. An overlap of the two approaches involves studying how a fracture interacts with neighbouring fractures during its history of evolution and propagation, and how that affects the statistical observation of the fractures population (e.g. their frequency, spacing, dimensions, etc.). Fractography primarily falls within the last two of these areas (Fig. 1).

INTRODUCTION

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Fig. 1.

Do all fractures develop fractographic features? The answer to this is yes. All fractures possess fractographic features at the time of their inception although they vary in the degree of development depending on the rock type and the environment in which they have been initiated and propagated. However, consequent history may obliterate or change these fractures, e.g. erosion, mineralization, stylolitization, shear movement. Although the last three processes may introduce new surface features (e.g. slickenlines) these are not usually considered as fractographic features by geologists (Kulander et al. 1979). In addition some authors include some of the features shown by the trace of fractures on a free surface, e.g. hooking and bifurcation, as part of the fractographic features. The present author is of the opinion that all aspects of surface morphology are fractographic features in that they are diagnostic characteristics of the fracture history and its evolution from the time of inception of a fracture to the time of observation. Therefore fractographers should incorporate all aspects of fracture surface morphology in their observations. This practice is essential to the widening of the fractographer's task from the limited sphere of the moment of inception of a fracture to the universal sphere of the temporal evolution of a fracture/fractures, e.g. the development of faults and fault rocks. This universal application of fractography will speed up and enhance our understanding of fracture mechanics and related phenomena in rocks (e.g. faulting and fault rocks development).

This volume The eleven papers in this volume range from those directly linked to fractography (Part 1) to those which, although not addressing the subject explicitly, have significant implications for fracture mechanics, the main objective of fractography (Part 2).

Part 1: Fractographic studies The eight papers in this part fall mainly within the subject of macrofractography and cover the following areas:

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Methods of observation, data collection & processing/quantitative fractography This area of studies is important in two ways. Firstly it offers methods and techniques to quantify the morphological characteristics of a fracture surface and their mechanical/physical implications (e.g. modelling the strength and fluid flow characteristics of fractured rock masses). Secondly it is of vital importance to develop fast and automated ways to record and analyse fracture surface topography for a statistically representative number of fractures. The latter is essential considering the fact that earth scientists deal with a 'natural laboratory' that is extremely difficult to access on the spur of the moment due to financial, environmental and time constraints. Two papers by Aydan & Shimizu and Jessell et al. discuss two different techniques of quantitative fractography. Aydan & Shimizu present a profilometer technique for quantifying fracture surface morphology in rock masses. Their approach incorporates three main components: classification of natural discontinuities in rock masses, choice of functions or parameters which represent the quantitative aspects of the topographical features, and use of a specially built device to measure, process and digitize the data on site (in the field). They classify discontinuities according to their origin into primary (eigen), secondary and tertiary. The first includes discontinuities associated with the origin of the rock, e.g. bedding planes in sedimentary rocks. The second includes discontinuities resulting from desiccation, cooling and freezing of rocks. The third is of tectonic origin and includes shear and tension fractures. The authors choose linear profile parameters (height of asperities, inclination of asperity walls, length of asperity wall relative to base length, and periodicity of asperities) to express the geometry of surface morphology in a group of functions. These parameters are measured for two mutually perpendicular profile directions (the eigen directions) one parallel (X) and one perpendicular (Y) to the axis of elongated protrusions (ridges) on the fracture surface, which are considered as the eigen directions. A specially built, manually operated, profilometer is used in this work in conjunction with a digitizer and a microcomputer for digitizing and storing data which are analysed by a processing sub-system. The technique is applied to different rock types in central Japan. It is concluded that the fracture surface morphology can be represented using profiles along the two eigen directions (i.e. X and Y). Asperity inclinations have their minimum values along the X-axis and maximum values perpendicular to the ridge axis. The centre line average height (CLAH), which is a function of the asperity heights for a specific type of discontinuity, does not show significant variations with the variation of rock-types. However, in a given rock type CLAH is generally higher in secondary and tertiary discontinuities than in the primary discontinuities. The study shows a proportional relationship between the angle of a weighted asperity inclination (WAI) and the ratio of profile length (RPL). However, the correlation between WAI and CLAH is poor. Finally this study illustrates quantitatively the anisotropy in the surface morphology functions measured in different directions on the same fracture surface. ,Iessell et al. present a digital photogrammetric technique for the quantitative determination of fracture surface morphology. They have developed a software package that can automatically reconstruct surface morphology information from stereo pairs at any scale, using edge correlation and point matching. After initial alignment of the images, the software does not require user intervention, making it well suited to the production of large datasets. To test the accuracy and reliability of their technique the authors made a comparison between a traditional profilometer and the new digital photographs of a quarrying-induced fracture face in granodiorite (Alexander Quarry, Harcourt, Victoria,

INTRODUCTION

5

Australia). They chose a test area of 256 by 116 mm of the fracture face. The results show a strong degree of correlation between the sets of data obtained from the two techniques. The digital technique is found superior to that of the profilometer in that it is applicable to a wide range of horizontal scales from satellites to microscopes with a consistent methodology. The main drawback of the technique is that its vertical resolution is smaller than that achievable using profilometers. This limitation can be reduced by either increasing the magnification of the photographic techniques, or by scanning the resultant negative at a higher resolution; both methods are easily achievable with the right equipment. However, with the present technology the best vertical resolution achievable still falls short of the 0.1/~m achievable by the profilometers.

Application of fractography to fracture analysis from core This field of fractographic studies is of great importance to analysing and understanding fractures observed in cores (Kulander et al. 1990). Its significance lies in its application to distinguish natural fractures from those induced by drilling, coring and handling of core. As a result the analyst is able to discard induced fractures when considering the effects of the natural fractures and their surface morphologies on the mechanical and physical properties of rock masses, e.g. permeability, stability, strength. In addition, the topographical features of the drilling/ coring-induced fractures are very good indicators of the orientation of the maximum and minimum horizontal in situ stresses (all and Crh respectively). This application of fractography is largely neglected by core analysts and core laboratories who instead use quick core scanning and classify fractures population according to their orientation into 'sets'. This can lead to unrealistic assumptions as to the origin of fractures. Bankwitz & Bankwitz present an excellent illustration of the application of fractography to the determination of the maximum horizontal in situ stresses (SH). Their study is particularly interesting because it covers a great range of depths and two different core diameters. The study is of coring-induced fractures in core sections recovered from the superdeep (8000m) borehole in northern Bavaria (KTB: Continental Deep Drilling, Windischeschenbach). The borehole cuts the crystalline basement of the Bohemian Massif, which includes amphibolites, gabbro and layered paragneisses. In cores with a diameter of 10 cm the surface morphology of fractures is not easily recognizable due to the coarsegrained texture of the basement rocks. However, the overall shape of the coring-induced fractures is used to determine their origin and propagation direction. In cores of larger diameter (23.5 cm) the fracture surface morphology is easier to recognize. Two main types of drilling-induced disk fractures (according to their geometry) are recognized. The first are referred to as core disking surfaces (CDS) which are characterized by saddle and trough structures, and are restricted to fine-grained rocks. The second group lack the saddle and trough structures, show clear fractographic features, occur more frequently than the first and are referred to as the usual fractographic surface (UFS) structures. The coring-induced fractures are initiated by the coupling of the stress states in turn related to the combined effect of the in situ stresses within the rock mass and the stress release due to the removal of overburden pressure. The direction of the maximum horizontal stress axis Sn is determined from the pattern of fracture surface morphology on the UFS. These data are mostly in agreement with those obtained from core disking. The SH orientation is NW-SE (165~ This direction coincides with the results from other techniques, e.g. borehole breakout analysis and this agreement further supports the

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feasibility of fractography as a tool in stress analysis for both checking the reliability of other techniques and for use on its own with a high level of confidence.

Fractography applied to the analysis of experimentally produced fractures in rocks and nonrocks Unlike natural fractures, experimentally induced fractures in rocks and non-rocks are easier to analyse, because they are formed within a better defined/constrained environment. Under such circumstances a clear and direct link between the causative factors (e.g. type and configuration of applied stresses) and the morphology of the resulting fractures can be established. This technique is a double-edged sword in that it can be used by both teachers and students of fractography and fracture mechanics to illustrate and better understand these subjects. Furthermore, researchers can manipulate the science of fractography to solve particular problems within the field of fracture mechanics in rocks. In the present collection of papers these two applications of fractography are illustrated in two papers by Kulander & Dean and Butenuth & de Freitas respectively. The first paper is by two of the pioneers of fractography in geology, Kulander & Dean. The paper highlights the principles of fractography through fascinating and easy to follow experiments in which fractures are induced in cheap, readily available material. To achieve their objectives the authors used microscopic slides, solid glass rods and capillary rods, loaded by bending and torsion. Point loading and thermal stresses are also used to induce fractures in slides. In discussing the results, Kulander & Dean use quantitative inference based on more complicated, controlled 'quantitative' experiments of others in the field of fracture studies in glass. The authors give a detailed anatomy of the fracture surfaces produced in the laboratory and illustrate the different stages of evolution of a fracture as can be recognized from its topographical features, e.g. origin, mirror, mist, velocity, hackles and twist hackles. In doing so they draw attention to the common occurrence of fractographic features which are linked to rapid fracturing, e.g. velocity hackles, Wallner lines and forking. Their discussion of the results also focuses on the lack of these velocityrelated features in natural and drilling-induced fractures in rocks. They conclude that such a contrast in brittle fracture morphology between glass and rock is a witness to the different properties of the two materials and the manner in which they respond to stresses. The authors also stress their belief that shear joints do not exist. The second paper by Butenuth & de Freitas illustrates the uses of fractography to investigate the effect of boundary conditions during hoop tests (to determine tensile strength) on the mode of failure. The hoop test used in this study is relatively new, and its operation is yet to be fully analysed and assessed. The tests involve the insertion of two hemi-cylindrical loading platens into the central hole of a hoop of rock. Then the platens are expanded until the hoop fails. The tensile strength is obtained 'indirectly' from plotting the ultimate load line force and the surface area of the fracture at the time of failure. Their plots for two rock types used in their experiments show deviation of some hoop tests from the otherwise linear force v. area relationship. In addition there exists a constant intercept on the area axis for each of the two rocks. The authors address these two features using fractography. The deviation of some of the results is found to be related to the fact that the samples were not fully separated by fractures at failure. Subtracting the area that has not been broken in tension brings the deviated results on line with the majority of samples. Fractographic studies show that the induced fracture terminal end (in the part of the hoop that is not fully broken in halves by the fracture) is characterized by rougher and more branching surface. This evidence was used to re-examine samples which are already broken in halves (after the test) to recognize the boundaries of the fracture surface at

INTRODUCTION

7

failure and correct the area data on the force/area plots. The constant intercept is linked to imperfections in the configuration of the hoop samples and the platens and/or the dynamic branching of the fractures. The former is proved to exist and induces compressive stresses in certain parts of the hoop. In spite of that, fractographic inspection on microscopic and macroscopic scale indicates that the experimentally induced fractures lack evidence of shear movement (e.g. slickenlines or debris). Therefore they are of Mode-I. However, using fractographic evidence the authors could not decide on whether dynamic branching and the bending due to imperfections in the samples and the testing equipment are acting separately or in combination.

Fractography applied to fracture analysis in field studies aimed at understanding regional tectonics Field studies utilizing fracture analysis can benefit greatly from systematic observations of fractographic features. Field analysis of fractures (apart from rare studies mostly in the USA) has been based largely upon fracture orientations and subsequent grouping into sets of parallel fractures, interpreted as 'conjugate shear' or 'tension/ extension' fractures on geometrical assumptions. The studies presented in the current volume (two by Ameen and one by Roberts) present contributions to the newly emerging school incorporating fractographic observations and interpretations as an essential part of fracture/joint observations and analysis. These studies also offer an important contribution to the universal aspects of fractography by adding relatively large amounts of data to those collected in other provinces. Such a database is a vital requirement to test the persistence of fractographic features and their implications for fracture mechanics and regional tectonics. Furthermore these fractographic studies permit observation on the evolutionary and tectonic history of the studied areas. The three studies discuss different regions in the United Kingdom. Ameen's papers cover West Cumbria in the English Lake District, and the Isle of Thanet, Kent, southern England. Roberts' paper covers part of South Wales. In the first study Ameen presents the results of a comprehensive fracture characterization of both the sedimentary cover rocks (Permo-Triassic) and the Lower Palaeozoic basement rocks in West Cumbria. This study is mainly based on work carried out on behalf of UK Nirex Ltd as a part of their ongoing geological investigation into the suitability of the Sellafield area for an underground nuclear waste repository. The author combines fractographic analysis with the conventional analysis of fracture orientations to determine the fracture characterization and related stresses in the area. Three main fracturing phases are recognized in the Permo-Triassic part of the cover rocks and their causative stresses are discussed. In contrast, the Lower Palaeozoic basement rocks reflect a complex fracturing history. The results suggest that the geometry and degree of development of fractographic features in the sedimentary cover are controlled by sedimentary layering, lithological heterogeneity, the presence of open fractures, pore fluid pressure, orientation and magnitudes of the principal stresses, and their temporal and spatial variations across the fractured horizon. The work also investigates exposed macrofaults in both basement and cover rocks and their relationship with the mesofractures. It is suggested that 3D configuration of the well exposed mesofractures zones should be considered when modelling the scarcely exposed macrofaults and their interpretation from subsurface data. In his second paper Ameen uses systematic studies of fractures (including fractography) in the Chalk and seismic data to establish the evolution of the regional structure and related fractures on the Isle of Thanet, Kent, southern England. The area has received little attention with regard to its fracture pattern and tectonic evolution. Burial, tectonic

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and uplifting stresses, mostly related to basement block faulting, including extensional and transtensional/transpressional movements, caused the evolution of the fractures and the host regional forced fold. The Thanet monocline is divided into a minimum of four zones (domains), according to fracture characterization. The study illustrates the effectiveness of combining seismic data with field observations to determine the tectonic evolution of a region. In addition the results are significant in that they touch on the fracturing of the Chalk, an important subject to a number of engineering and hydrocarbon projects (e.g. North Sea reservoirs, the Channel Tunnel, nuclear power stations). Roberts describes the joint system in a 2.5 m thick slice of interbedded limestones and mudstones located in the upper part of the Saint Mary's Well Bay Formation (Lower Liassic) exposed locally between Lavernock Point and Saint Mary's Well Bay, South Wales. Three systematic joint sets are recognized which strike 190~, 280 ~ and 335 ~ respectively. The three sets show fracture surface markings including hackle plumes, point sources (origins), ribmarks and twist hackle plumes. The origins are consistently located in the limestone layers from which the hackle plumes spread with bedding-parallel axes. The latter indicate that the leading edge of the joints propagated parallel to the bedding. The joints terminate in the overlying and underlying mud rocks with deep borders of twist hackles. The persistent location of joint origins in the limestone layers and their termination in the mudrocks are attributed to the different elastic moduli of these two rock types. The limestone acts as a stress concentrator and provides sites for fracture initiation. Although limestone lithifies earlier than the mud sediments, there is no evidence that joints in the studied limestone developed early and later propagated to the surrounding mud rocks. The three sets of joints show the same value of twist angles (10~176 The sense of twist in each set is consistent, counter-clockwise for the 190~ joint set and clockwise for each of the other sets. The relatively small twist angle is attributed to jointing under relatively low confining pressure. The sense of twist is used to deduce the orientation of the maximum and minimum principal stresses during the propagation of these joints. Joint architecture suggests that the 190~ joint set is younger than the other two sets. No firm conclusion is reached concerning the relationship of the joints to the local structures.

Part 2: Non-fractographic studies The second part of the volume includes three studies. The first by Cosgrove addresses tensile fractures in unconsolidated, unlithified sediments which tend to have their delicate surface patterns (e.g. hackle plumes) obliterated by weathering, filling, compaction and buckling of the fractures. In such circumstances the dynamic evolution of the fractures cannot be documented by fractographic features. Therefore some explanatory mechanism must be found. To illustrate this problem Cosgrove briefly considers sedimentary dykes. In some regional field studies fractographic features are not clear due to the quality of exposure, direction and degree of sunlight, etc. Such studies rely mainly on statistical fracture population analysis. As an example, Goodwin presents the results of a study of joints in the flat lying layers of the Upper Cretaceous Chalk of eastern England. Orientation data were collected from 20 localities (no morphological features on joint surfaces were identified by the author) and analysed using stereographic projection and cluster analysis. The studies reveal the occurrence of two regional dominant sets (according to their orientation) striking N W - S E and NE-SW respectively. The author suggests that these sets are the same as those that occur in southern England, and recognizes a change in the joint plane geometry of each set from steep conjugate 'hybrid'

INTRODUCTION

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joints in Norfolk to sub-vertical extension joints in Humberside. Such a change is mainly referred to northward increase in rock strength and hardness and a possible decrease in differential stress towards the north. Stratigraphic and regional structural control on jointing are considered to play a minor role in controlling joint geometry and their variations. Spacing is one result of the interaction of neighbouring and parallel fractures that do not interact by cross-cutting/abutting. Instead they affect the spatial and temporal distribution of their neighbouring penecontemporaneous fractures. In cases where the fracture traces are the only source of information available, and there is a lack of exposed fracture surfaces for fractographic observations, a theoretical explanation is needed for the factors controlling the spacing between the fractures of one set. This is thoroughly discussed by Gross et al. who give a detailed explanation of Hobbs' (1967) analysis of joint spacing and attempt to clarify the assumption and implications of the one-dimensional model offered by Hobbs. Gross et al. use a two-dimensional finite element numerical simulation of a crack confined within a bed, to describe the stress and displacement fields around the profile section of a crack in an interbedded sedimentary rock mass. They use field studies in the Monterey Formation of California to illustrate their findings.

Suggested target areas for future research in the field of fraetography in geology Fractography in geology is a 'young science' and can therefore benefit from more studies into any of its aspects. However, to achieve balance, key areas must be recognized and given priority in future studies. In need of most urgent attention are: 1. A comprehensive review of the relatively vast nomenclature used in fractography to establish a universally accergted 'fractographic code' for fractographers. 2. Microscopic studies of fracture surfaces (microfractography). This has been the most neglected aspect of fractog~: aphy in spite of its great potential in advancing our knowledge and understanding of fracture mechanics. 3. Systematic and detailed 3D mapping of fractographic features of fractures that are induced experimentally under well constrained conditions, to determine reliable criteria for quantitative interpretation of surface features. 4. Analysis of surface features of fractures as a part of core logging. The amount of work and data collected on this subject is insignificant compared to that carried out on field exposures.

This volume stems from a joint Petroleum Group-Tectonic Studies Group meeting held at the Geological Society in September 1993. I would like to thank these Groups for sponsoring the meeting. The efforts of the Conference Officer at the Geological Society during and before the meeting are greatly appreciated. Thanks are due to the authors who contributed to the meeting and this volume, and to the reviewers of the papers.

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References BAHAT, D. 1991. Tectonic-fractography. Springer-Verlag, Berlin. BANKWXTZ,P. 1966. Uber Klufle, II. Geologie, 15, 896-941 HOBBS, D. W. 1967. The formation of tension joints in sedimentary rocks: an explanation. Geological Magazine, 104, 550-556. HODGSON, R. A. 1961. Classification of structures on joint surfaces. American Journal of Science, 259, 493-502. KULANDER, B. R., BARTON, C. C. & DEAN, S. L. 1979. The application offractography to core and outcrop fracture investigations. Report No. METC/SP-79/3, United States Department of Energy, Morgantown, West Virginia, USA. , DEAN, S. L. & WARD, B. J. 1990. Fractured core analysis: interpretation, logging, and use of natural and induced fractures in core. American Association of Petroleum Geologists, Methods in Exploration Series, 8, Tulsa, USA. PARKER, J. M. 1942. Regional systematic jointing in slightly deformed sedimentary rocks. Bulletin of the Geological Society of America, 53, 381-408. ROBERTS, J. C. 1961. Featherstone-fracture and the mechanics of rock jointing. American Journal of Science, 259, 481-492. SYME-GASH P. J. 1971. A study of surface features relating to brittle and semibrittle fracture. Tectonophysics, 12, 349-391. WOODWORTH, J. B. 1896. On the fracture system of joints, with remarks on certain great fractures. Proceedings of the Boston Society of Natural History, 27, 163-184. ZAPFFE, C. A. & CLOGG,M. Jr. 1944. Fractography: a new tool for metallurgical research. Preprint 36, American Society of Metals.

From Arneen, M. S. (ed.), 1995, Fractography: fracture topography as a tool in fracture mechanics and stress analysis

Geological Society Special Publication No. 92, pp. 11-26

Surface morphology characteristics of rock discontinuities with particular reference to their genesis OMER

AYDAN

l & YASUHIRO

SHIMIZU 2

1Department of Mar&e Civil Engineering, Tokai University, Orido 3-20-1, Shimizu 424, Japan 2Department of Civil Engineering, Meijo University, Shiogamaguchi, Tenpaku-ku, Nagoya 468, Japan

Abstract: A research program was undertaken to characterize the surface morphology of natural discontinuities with particular reference to their genesis. For this purpose a special profilometer, which could easily be transported and operated in the field, and was capable of measuring 480 mm long profiles, was designed. The recorded profiles were digitized and analysed to obtain surface morphology characteristics of the discontinuities in different rock types found in Central Japan. Natural discontinuities have an anisotropic surface topography which may be characterized by measuring profiles along their eigen directions. Consequently, surface morphology parameters are directional and differ depending upon the discontinuity type.

Rock discontinuities in the form of cracks, joints, bedding planes, schistosity planes, etc., are commonly found in rock masses in the upper part of the Earth's crust (Pollard & Aydin 1988). Discontinuities in rock masses influence the stability of rock engineering structures as well as thermo-hydrological characteristics of rock masses, which have become a field of interest in relation to radioactive waste disposal projects (Aydan & Kawamoto 1990). There have been numerous experimental and numerical studies for the characterization of the surface morphology of discontinuities and for relating the surface morphology parameters to their thermo-hydromechanical properties in recent years (Sezaki et al. 1993). When the authors undertook a project to investigate the hydromechanical characteristics of rock discontinuities, it was realized that there were two areas of study missing in most of the previous work: firstly, hardly any attention had been paid to the genesis of discontinuities; and secondly most experiments were carried out in laboratories by splitting rocks or rock-like m a t e r i a l s - i.e. experiments on natural discontinuities were limited to those which could be transported and tested in the laboratory. The purpose of this study was to present a procedure for quantifying surface morphological characteristics of discontinuities in rock masses with particular reference to their genesis. Firstly, a classification for discontinuities in rock masses is proposed, then a system developed for measuring and quantifying the surface morphology characteristics of the discontinuities is described. Finally, in situ measurements carried out on various kinds of discontinuities in different rocks, found in Central Japan are presented and discussed.

12

6.

A Y D A N & Y. S H I M I Z U

Classification of discontinuities in rock masses Discontinuities in rock masses can be classified in various ways (Pollard & Aydin 1988; Aydan & Kawamoto 1990). Proposed here is a classification based on the order of occurrence as follows: (1) Primary (eigen) discontinuities which are associated with the genesis of rocks. Such discontinuities are termed flow planes in igneous rocks, bedding planes in sedimentary rocks and schistosity planes in metamorphic rocks. These planes are easily distinguished in sedimentary and metamorphic rocks while they may be difficult to recognize in igneous rocks. Nevertheless, observations on platy minerals (e.g. micas) and elongated minerals (e.g. turmaline and hornblend) generally help in the recognition of such planes. (2) Secondary discontinuities which develop in rocks as a result of desiccation, cooling and freezing of rocks. Desiccation discontinuities are generally observed in sedimentary rocks while cooling discontinuities occur in initially hot igneous or metamorphic rocks. Freezing discontinuities can be observed in any kind of rock. These planes are usually perpendicular to primary (eigen) discontinuities. These planes can generally be classified as grain planes, hardway planes and sheeting joints. The cause of these discontinuities is mainly tensile stresses. (3) Tertiary discontinuities which develop in rocks due to tectonic forces. They occur during folding and faulting processes, and are classified as shear and tension planes.

Surface morphology characterization The characterization of the surface morphology of rock discontinuities is merely a geometrical procedure. In other words, it is an identification procedure of the topography of the discontinuity surfaces, which specifically involves height, shape and periodicity of protrusions, and the ratio of the surface area to the base area. It will be appropriate if a function which can represent the topography of discontinuity surfaces can be found. This is extremely hard not only because of the difficulty in finding an appropriate function but also because of the enormous efforts required in measuring and processing the data. As a result of this, most of available techniques are based on linear profiles and various characterization parameters are proposed (Myers 1962; Sayles & Thomas 1977; Tse & Cruden 1979; Thomas 1982; Tfirk et al. 1987). The parameters associated with linear profiles are height of asperities, inclination of asperity walls, length of asperity wall relative to base length, and periodicity of asperities. These parameters are briefly explained in the following: Centre-line average height (CLAH) is defined as;

x=L Ir CLAH = Z1 fx =o

(1)

where L is the measurement length, x the distance from the origin and r the height of the profile from the reference base line.

SURFACEMORPHOLOGYCHARACTERISTICSOF ROCKDISCONTINUITIES

13

Mean standard variation of height (MSVH) is defined as: x=L r MSVH -- ~1 fx =o

(2)

Root mean-square of height (RMSH) is defined RMSH =

r

=0

(3)

Ratio of profile length (RPL) is defined as:

1 fxz=L

RPL = ~

1 fxx=L (

=o ds = ~

=o

de) 1/2

1 + dxx

dx.

(4)

Weighted asperity inclination (WAI*) is defined as: 1 fz x=L de WAI* = ~ =o I ~ xx I dx, WAI = tan-l(WAI*).

(5)

Mean standard variation of inclination (MSVI) is defined as: MSVI = ~

o

dxx

dx.

Root mean-square of inclination (RMSI) is defined as: RMSI=

=0

dxx

(6)

]12 dx

(7)

Auto correlation function (ACF) is defined as:

1 ACF = ~

fx =o x=L r162

+ ~)dx.

(s)

+ ~)]2d~,

(9)

Structure function (SF) is defined as: x=L [r SF = ~1 fx --o

- r

where T is a measure of the periodicity of asperities.

Fractal dimension is defined as: N = CU D,

(10)

where N is the number of steps, C is a constant, l is step length and D is a fractal dimension. Since the following relation holds between the total length of the profile and the step length:

14

O. AYDAN & Y. SHIMIZU

I. 81

I

530

L

_,

Rec~ Pen-Recorder

~1

l~

'l

,I1

~

Supports

Drum-likeHandle

= ~ tR~

SIDEVIEW

FRONT VIEW

i'01111

0

I.

olo

,W

PLANVIEW Fig. 1. Views of the profilometer.

L = N1,

(11)

then the above equation can be rewritten as: L=

Ct ~-~

(12)

The most crucial aspect of characterization using linear profiles is how to select the appropriate direction of measurement. Although there are no guidelines for this, the usual practice seems to be a line which is perpendicular to the axis of elongated protrusions by taking into account the maximum shear resistance due to geometrical dilatancy of the surface. Nevertheless, it is difficult to say that such profiles are representative of the discontinuity surfaces. It is well known in fracture mechanics of rocks and metals that the ridges of discontinuity surfaces develop almost perpendicular to the crack propagation direction.

SURFACE MORPHOLOGYCHARACTERISTICSOF ROCK DISCONTINUITIES

136E |

137E

138E

|

t

/

139"~E 'l

37N" JAPANS E A j ~ HONSHUISLAND

|

36N'

15

~Tsuge (granite) (~)Otake (andesite, basalt) | (granite) (~Kitamatado(shale) @Okumino(rhyolite, sandstone, chert, limestone, basalt) (~)Chita(mudstone) (~Echizen(tuff) |

35N

34N

PACIFICOCEAN l

I

l

Fig. 2. Locations of the field measurements.

Furthermore, the ridges in sedimentary rocks are also perpendicular to the flow direction of fluids, as is well known in sedimentology. By taking these facts into account it is possible to propose that there are generally two mutually perpendicular eigen directions; one perpendicular to the ridge axis and the other parallel to it. Measuring profiles along these directions may be sufficient to characterize the surface morphology of discontinuities. Consequently, one may introduce an elliptical coordinate system so that the principal axes of such a coordinate system coincide with those eigen directions. Such a coordinate system would be appropriate for many discontinuity types found in rock masses.

Surface morphology characterization system To measure a linear profile there are two principal measuring techniques. (1) Mechanical devices of the contact type, which can be further subdivided into two groups - stylus of needle type and stylus with roller tip. (2) Optical devices of the non-contact type which utilize reflected light beams, i.e. laser beams. Measurements and data acquisition are generally carried out manually, semiautomatically or automatically. The automatic systems are desirable in that measuring is a laborious procedure. However, such systems are unsuitable for measurements under various in situ conditions as they require heavy equipment for their power supply, driving

16

O. AYDAN

& Y. SHIMIZU

BEDDING

PLAN - -

ES ~

~

~ P 1

BP3

B P 6

BP7 BP8 TENSION

D ISCON

TIN

U ITIES

TD1 TO2

TD3 TD4 TD5 TD6 TD7

SHEAR

PLANES SD1

--

SD2

=

S.D4

SO5

"""'---'~"-~'~_...

0

Fig. 3. Measured profiles.

I

SD6

50 mm

I

systems for the measuring stylus and data acquisition. As a result of this, their use is generally restricted to the laboratory. Since our main purpose was to make measurements on actual discontinuities, lightweight equipment was needed. Therefore, a measuring system consisting of a manually operated profilometer for site measurements, a digitizing system for recording profiles, a microcomputer for data storage and a data processing subsystem for the digitized data has been developed. Figure 1 shows several views and

SURFACE MORPHOLOGY CHARACTERISTICS OF ROCK DISCONTINUITIES

17

dimensions of the profilometer; it is ma n u a l l y operated, weighs c. 5 kg and is made mainly of aluminium. The diameter of the roller-type stylus is 3 mm, which is the smallest one available on the market. Profiles are recorded on a sheet using a pen-recorder. Profiles up to 480 mm, with a m a x i m u m asperity height of 90 mm, can be measured. The apparatus can be used on surfaces having an inclination varying between 0 and 90 ~ and it can be carried to any site to which h u m a n s have access. All recorded profiles are digitized using a digitizer and stored in files using a m i c r o c o m p u t e r at the authors' universities. The digitized data are then used to obtain surface m o r p h o l o g y parameters as defined in the previous section.

Table 1. Rock and discontinuity types at measurement sites No.

Location

Rock type

Discontinuity type

1 2

Tsuge Otake

3 4 5

Nakatsukawa Kitamatodo Okumino

6 7 8

Chita Echizen Chausuyama

Granite Andesite Basalt Granite Shale Rhyolite Sandstone Chert Limestone Basalt Mudstone Tuff Hornfels

Flow plane, grain plane, hardway plane Flow plane, grain plane, hardway plane Flow plane, grain plane Flow plane, grain plane, hardway plane, sheeting joint Bedding plane, tension joint Flow plane, grain plane, hardway plane, shear discontinuity Bedding plane, tension discontinuity Bedding plane Bedding plane Cooling joint Bedding plane, shrinkage joint Bedding plane, tension discontinuity and shear discontinuity Schistosity plane, grain plane, hardway plane

Measurements and discussion Measurements have been ma d e on discontinuities in different rocks found in Central Japan. Figure 2 shows the locations of measurements and Table 1 gives rock and discontinuity types in the respective localities. Figure 3 shows several of the profiles

l Fig. 4. Notations for measuring profiles of discontinuities.

18

O. A Y D A N

& Y. S H I M I Z U

T BEDDING PLANE (BP6)

TENSION DISCONTINUITY

! 0

5 mm 0

1 20 mm

(TD4)

SHEAR DISCONTINUITY (SDE)

Fig. 5. Profiles of selected discontinuity types shown in Fig. 3.

measured. Measurements on discontinuity surfaces were carried out along the eigen directions of surfaces; the x-axis was selected to be parallel to the ridge axis and the y-axis perpendicular to the ridge axis (Fig. 4).

Table 2.

Surface morphologyparameters

Discontinuity plane

BP6 TD4 SD5

CLAH (mm)

Height parameters MSVH RMSH WAI (mm 2) (ram) (o)

2.301 2.174 2.212

6.312 6.696 6.097

2.512 2.586 2.469

Inclination parameters Periodicity parameters RPL WAID MSVI RMSI T for ACF r for SF ( o) (mm) (mm)

1 0 . 7 2 3 1 . 5 6 4 0.052 11.45 1 0 . 0 2 2 0.063 8.92 0.547 0.036

0.229 0.251 0.189

11.50 14.54 23.62

12.13 13.85 19.65

1.025 1.030 1.017

Fractal dimension

1.009 1.010 1.006

Figure 5 shows digitized profiles of a bedding plane (BP6), tension discontinuity (TD4) and shear discontinuity (SD5) whose original profiles are shown in Fig. 3. The sampling interval for digitizing profiles was 2 mm in these particular examples. Table 2 gives the values of surface morphology parameters defined in the preceding section. However, it should be noted that computed results are generally influenced by the sampling interval and profile length. Figures 6 & 7 show the effect of sampling interval and profile length, respectively, on parameters CLAH, WAI and RPL for each discontinuity. The effect of the sampling interval on parameter CLAH is almost negligible; on the other hand, parameters WAI and RPL decrease in magnitude as the sampling interval increases (Fig. 6). Therefore, it is desirable to make the sampling interval as small as possible. However, it should be noted that the sampling interval in digitizing profiles manually is generally restricted to >I 1 mm.

SURFACE MORPHOLOGY CHARACTERISTICS OF ROCK DISCONTINUITIES

4

19

.1 04~- BEDDINGPLANE 1 0 ~ , 9

[

~: 1.02

5

o

(BP6)

F a

o

~

,.o, F ~ 0

10 20

30

40

S A M P L I N G I N T E R V A L (ram)

11

'

"0

I

8~~

I

]-~i

10 20

L

30

, J

/

,

40 0

S A M P L I N G I N T E R V A L (mm)

~-~ I

10

~

I

20

,

I

30

,

I

40

S A M P L I N G I N T E R V A L (mm)

TENSION DISCONTINUITY

4-

1.04 r

_,3

(TD4)

10

1.03 ~-

E E v

~-2 _ 5 o 1

~

0

-

~

10 20

30

6- 8

~: 1.02

-~64 ~:

1.01

2

1

, 40 4 0

40

S A M P L I N G I N T E R V A L (mm)

S A M P L I N G I N T E R V A L (mm)

i

I

t

I

10

i

20

I

30

,

I

40

S A M P L I N G I N T E R V A L (mm)

SHEAR DISCONTINUITY

1.04 -

IB

3

1.03

z2 5 O 1

~ 1.o2

E E v

I

0

I

i

I

10 20

I

I

30

,

I

40

S A M P L I N G I N T E R V A L (mm)

10

(SD5)

10 F o-~8 ~

10 20

- I

30

,

I

40 0

S A M P L I N G I N T E R V A L (mm)

,

I

10

,

I

20

30

40

S A M P L I N G I N T E R V A L (mm)

Fig. 6. The effect of the sampling interval on surface morphology parameters CLAH, WAI and RPL.

20

0.

AYDAN & Y. SHIMIZU

1.04-

BEDDINGPLANE

12-

(BP6)

1.03

9

Q.

,.,- 1.02 o 1

1.01 I

40

80

I

,

I

120 160

3

10,

I

4O

PROFILE LENGTH (mm)

,

I

,

80

I

,

I

o

2

1.02

I

,

,

I

120 160

!

16

,o4

..3

I

80

PROFILE LENGTH (mm)

TENSION DISCONTINUITY

(mD4)

i

4O

PROFILE LENGTH (mm)

1.06 - ,

4

I

,

120 1600

4 !

0

40

80

120 160

10

PROFILE LENGTH (mm)

A

E

4!i

40

80

120'1600

40

PROFILE LENGTH (ram)

80

120 160

PROFILE LENGTH (mm)

SHEAR DISCONTINUITY

1.04r-

(SD5)

10

-r-

5 o

1.01 ~ ,

0

I

40

,

I

80

I

I

I

I

120 160

PROFILE LENGTH (mm)

4/

"0

~

!

40

a.

I

80

,

1

9 9

120 1600

PROFILE LENGTH ( m m )

I

I

I

40

t

I

80

,

I

.I

I

120 160

PROFILELENGTH (mm)

Fig. 7. The effect of the profile length on surface m o r p h o l o g y parameters C L A H , W A I and RPL.

SURFACE MORPHOLOGY

CHARACTERISTICS

OF ROCK DISCONTINUITIES

21

30*

6o"

120 ~

180~._._

_

240 e

o

50 mm

Fig. 8. Profiles of a sheeting joint. The length of profiles may also influence computed results. For this reason computations were carried out varying the length of digitized profile for a sampling interval of 2 mm. As seen from Fig. 7, if the profile length is greater than the wavelength of the main asperity of a given discontinuity, then the influence of the profile length becomes less pronounced. As pointed out in the preceding sections, we discuss whether eigen directions for discontinuity surfaces exist or not. For this purpose, the direction for measuring profiles was varied by 30 ~ on several discontinuity types. As a specific example, the profiles of a sheeting joint in Nakatsugawa granite for respective directions is shown in Fig. 8. The profile length was 100 mm for all directions. The sampling interval was taken as 1 mm during digitization and computed results for surface parameters CLAH, WAI and RPL are shown as a function of measuring direction in Fig. 9. As seen from Fig. 9, peaks and troughs were observed in computed results and they were mutually orthogonal to each other. The peaks were observed when a profile was perpendicular to the ridge axis and troughs were observed when a profile was perpendicular to the ridge axis and troughs were observed when a profile was parallel to the ridge axis. This tendency was also observed in

22

6. m

E E

A Y D A N & Y. S H I M I Z U

SHEETING JOINT

m

4

"1"- 3 (3

21

0

I

I

,,I

I

I

I

60

120

180

240

300

360

I

I

15-6"-

~

MEASURING DIRECTION (o)

10.

I 0 1.2+

I

I

I

60

120 180 240 300 MEASURING DIRECTION (o)

360

60

120

360

rrl.1

1.0 0

180

240

300

MEASURING DIRECTION (o) Fig. 9. Variation of surface morphology parameters CLAH, WAI and RPL as a function of direction of measurement.

other types of discontinuities measured. The directions of peaks and troughs may be taken as spectra of eigen directions. Furthermore, these measurements implied that surface morphology parameters were strongly anisotropic. As a result of this, it is most likely that the associated hydromechanical characteristics of discontinuities will be also anisotropic. Next, interrelations between parameters CLAH, WAI and RPL for discontinuities found in Nakatsukawa granite, Okumino rhyolite, Otake andesite and basalt were investigated. Discontinuities measured were classified as either primary or secondary.

23

SURFACE MORPHOLOGY CHARACTERISTICS OF ROCK DISCONTINUITIES

25

I

i

I

i

I

X Y

9 o 9 [] 20

111 W

15

eo

-

PRIMARY DISCONTINUITY SECONDARY DISCONTINUITY SECONDARY DISCONTINUITY

9

0 111 0 O

<>

10

i 00~0 ~eO 9

I

0

1

2

o

(~0 9

!

I

4 3 CLAH (turn)

5

6

Fig. 10. Relation between parameters CLAH and WAI.

Figure 10 shows the relation between parameters WAI and CLAH for all discontinuities found in these igneous rocks. The values of WAI for primary discontinuities seem to be smaller than those for secondary discontinuities. Nevertheless, the data plotted do not support any correlation between WAI and CLAH. This may be a natural consequence, since WAI depends on the ratio of the height to the wavelength of asperities. Figure 11 shows the relation between RPL and CLAH. The data plotted are highly scattered, which may be a result of the difference in grain sizes of the rocks and the direction of measurement. The correlation between these two parameters is not so well defined, but it seems that RPL increases as CLAH increases. Finally, we checked the relation between WAI and RPL (Fig. 12). A well-established relation between these two parameters was observed. As RPL increased so WAI

24

,3. AYDAN & Y. S H I M I Z U

1.10

I

I

I

I

Y 0

PRIMARY DISCONTINUITY

[]

SECONDARY DISCONTINUITY

,O, 0

SECONDARY DISCONTINUITY

X

I

.O.

_1

n

rr

1.05

0

oO

P oop eO

[] O

~ 9

0O

o

0

q) o

e

I

I

I

3

4

5

6

CLAH (ram)

Fig. 11. Relation between parameters CLAH and RPL.

proportionally increased. Furthermore, the width of scattering was small as compared with the data plotted in Figs 10 & 11.

Conclusions In this paper, the quantification of surface morphology of discontinuities in rock masses with particular reference to their genesis has been considered and the following conclusions are drawn. (1) The discontinuity surfaces can be characterized using the profiles along the eigen directions of discontinuity surfaces. These eigen directions are mutually perpendi-

25

SURFACE MORPHOLOGY CHARACTERISTICS OF ROCK DISCONTINUITIES

25

'

x

20

UJ W

-

'

'

'

"'

I

'

'

'

'

Y

9 O

PRIMARYDISCONTINUITY

[]

[]

SECONDARYDISCONTINUITY

~

SECONDARYDISCONTINUITY

15

4~

.

(.9

O

00

~1 . 0 5

1.10

RPL Fig. 12. Relation between parameters RPL and WAI.

(2) (3)

(4) (5)

cular to each other. One eigen direction is parallel to the ridge axis and the other perpendicular to it. Asperity inclinations have a minimum along the ridge axis and a maximum perpendicular to the ridge axis. There is no great difference between rock types with regard to the function CLAH, which is a measure of asperity height. However, C L A H is generally higher in the case of secondary and tertiary discontinuities as compared with that in the case of primary discontinuities. The relation between W A I and C L A H for all measured data is poor and does not support any correlation between these two parameters. The relation between WAI and RPL is well established - as RPL increased so WAI increased proportionally.

26

<3. AYDAN & Y. SHIMIZU

The authors acknowledge partial funding for this research by the Research Foundation for Electrotechnology of Chubu (REFEC), Nagoya, Japan.

References AYDAN,O. • KAWAMOTO,T. 1990. Discontinuities and their effect on rock mass. Proceedings of the International Conference on Rock Joints, Loen, 149-156. MYERS, M. O. 1962. Characterization of surface roughness. Wear, 5, 182-189. POLLARD, D. D. & AYDIN, A. 1988. Progress in understanding jointing over the past century. Geological Society of America Bulletin, 100, 1181-1204. SAYLES,R. S. ,~ THOMAS,T. R. 1977. The spatial representation of surface roughness by means of the structure function: a practical alternative to correlation. Wear, 42, 263-276. SEZAKI, M., AYDAN,6. & SHIMIZU,Y. 1993. Development of a data-base system for the surface morphology characteristics and mechanical properties of rock discontinuities. Proceedings of the 25th Japan Rock Mechanics Symposium, Tokyo, 11-15. SHIMIZU, Y., AYDAN, (~. & KAWAMOTO, Z. 1993. Surface morphology characteristics of rock discontinuities. Proceedings of the 14th Western Japan Rock Engineering Symposium, Fukuoka, 65-70. THOMAS, T. R. 1982. Rough Surfaces, Longman, London. TSE, R. & CRUDEN, D. M. 1979. Estimating joint roughness coefficients. International Journal of Rock Mechanics and Mining Science, 16, 303-307. TORK, N., GERIG, M. J., DEARMAN, W. R. & AMIN, F. F. 1987. Characterization of rock joint surfaces by fractal dimension. Proceedings of the 28th US Symposium on Rock Mechanics, 12231236.

From Ameen, M. S. (ed.), 1995, Fractography: fracture topography as a tool in fracture mechanics and stress analysis

Geological Society Special Publication No. 92, pp. 27-37

The anisotropy of surface roughness measured using a digital photogrammetric technique MARK

W. JESSELL, 1 SIMON

J. D . C O X , 1'2 P H I L L I P

& WILLIAM

SCHWARZE

1

L. P O W E R 2

1Victorian Institute of Earth and Planetary Sciences, Department of Earth Sciences, Monash University, Clayton 3168, Victoria, Australia 2CSIRO Division of Exploration and Mining, 39 Fairway, Nedlands, WA 6009, Australia Abstract: The characterization of fracture surface morphology is important for the understanding of shear and normal deformation behaviour of rock masses, and for their transport properties. It has been suggested that natural fracture surfaces might be fractal. If true, this would have dramatic implications for these rock properties and would provide a sound basis for future modelling. Evaluating this hypothesis requires measurement of fracture surface morphology over a large range of scales. A software package has been developed that can automatically reconstruct surface morphology information from stereo pairs at any scale, using edge correlation and point matching. The data derived using this technique compare favourably with results generated from a traditional surface profilometer survey, and results are presented from the analysis of the anisotropy of blasting-induced tensile fractures in granite from Mt Alexander Quarry, Harcourt, Victoria, Australia.

The mechanical and physical properties of low porosity rocks at low pressures are strongly influenced by the presence of fractures. Examination of normal and shear deformation (e.g. Bandis et al. 1983), and fluid and electrical transport (e.g. Brown 1989) have highlighted the effects of the shape or roughness of the fracture surface. This is most important at shallow levels in the crust, which is within the realm of most engineering studies, or in zones of low effective confining pressures. Traditional engineering analyses have considered surface roughness within the scales of 0.01-1 m as affecting frictional strength (Thomas 1982). However, recent studies of rock surfaces suggest that roughness of the surfaces occurs at a wide range of scales, and the surfaces may be approximately self-affine or fractal (Power & Tullis 1991). In this context, fractal means that surface irregularities are present at all scales, with longer wavelength irregularities having a larger amplitude and that some form of power law describes the change in roughness with scale. There are some reasons why rock fractures might be fractal or chaotic, namely that they develop under high energy conditions where the small-scale local fracture orientations might depend on a wide variety of influences, including material heterogeneity, preexisting fractures and damage features, or other flaws. Alternatively, rock fractures might be essentially Euclidean at large scale because they develop under the influence of a continuous, well-behaved stress field. In the latter case, it might be expected that the large scale (1-10 m) fractures will be essentially flat, while at scales approaching the rock grain size the surface will be relatively rougher. Furthermore, once fracture surfaces are

28

M . W . JESSELL E T AL.

Fig. 1. Locality of the quarry where samples were collected for this study. Harcourt is located at 144~ 36~ activated as the locus for sliding, there is ample evidence to suggest that their roughness characteristics will evolve and that the anisotropy of roughness will then become a major consideration in the analyses of fracture properties. Evaluating these relationships requires measurement of fracture surface morphology over a large range of scales. This is very difficult to do using a consistent method, particularly in 2D, rather than profile, form, since most current techniques have an inherently narrow range of the power spectrum over which reliable data can be collected. We have developed a software package capable of automatically reconstructing surface morphology information from stereo pairs at any scale (Schwarze et al. 1992). After initial alignment of the images the software does not require user intervention, and thus is well suited to the production of large datasets.

Method The aim of all stereo photogrammetric techniques is to solve the stereo correspondence problem, which involves the matching of the same location in both images of a stereo pair. Heights are calculated by comparing apparent displacements (known as disparities) between matched points in the two images. These disparities are proportional to the heights of individual points (relative to some datum). In a manual technique matching of points is performed by a human operator, but in the current technique it is solved by the computer using a combination of algorithms based on vision theories. Current theories of the human visual system suggest that the first visual cues used in depth perception are edges and areas of marked contrast changes (Marr & Poggio 1979), and that other parts of the image are then used to fill in the finer detail of the depth map. We have implemented a technique based on this hierarchy (Schwarze et al. 1992).

DIGITAL PHOTOGRAMMETRY

OF ROUGH

29

SURFACES

et0

,'-', "~0

~ ,,...~

e~o

2~r~

e q ~

30

M . W . JESSELL E T .4L.

Fig. 3. Scattergram comparison of surface topographic information collected by profilometer and by digital photogrammetry; > 90~ of the data fell within 1.1 pixels of an exact match.

It has been suggested that a 'short cut' method for the analysis of surface roughness may be achieved by directly analyzing the spectral content of single grey scale images of surfaces (Russ 1993). Unfortunately, this technique is not suited to the study of coarse grained rocks of the type measured in this paper, as the speckled appearance of the rocks resulting from dark and light minerals would overwhelm any shading variations due to surface topography. Conceivably, this could be avoided by coating the surface with a neutral density paint; however, this would not really be practical for the analysis of large rock surfaces.

DIGITAL P H O T O G R A M M E T R Y OF R O U G H SURFACES

31

Quantitative assessment of technique As a means of assessing the accuracy and precision of the results produced by our algorithm a direct comparison with a different technique was necessary. To achieve this a comparison was made between a profilometer survey over a quarrying-induced fracture and a digital photogrammetric analysis of the same sample. The sample was a flake of granodiorite from Mt Alexander Quarry, Harcourt, Victoria, Australia (Fig. 1), and a test area 256 • l l 6 m m was defined for comparison (Fig. 2a). The profilometer works by tracking a stylus across the sample, with the surface elevations measured on a rectangular grid at 1 mm spacing. The heights were measured with a precision of 4-0.1 #m using a linear variable differential transformer displacement sensor (Fig. 2b); the total height range measured was 13.7 ram. In order to produce a comparable grid using digital photogrammetric techniques two medium format (4 • 5 cm negative) photographs were taken at a distance c. 30 cm from the surface of the flake; 8 • 10 cm prints were then made of the negatives and each print was then scanned at a resolution of 75 pixels per inch (c. 30 pixels per cm) on a standard flatbed scanner, creating two 8 bit (256 grey level) digital images of the area of interest. These images provided the raw input to the photogrammetry program, with additional alignment information provided at this stage. The output from the program was a data file of heights for individual pixel locations, with null values for those pixels for which the program could not make a successful match. These null values were filled in using a thinplate spline interpolation scheme, producing a complete coverage of the test area (Fig. 2c). High frequency noise spikes remain where incorrect matches were made. The output from each of the data sets covered the same area of the sample and were of a similar horizontal scale. In order to allow direct comparison of the two data sets we first normalized the horizontal scales; however, there are some relative distortions in position (resulting from parallax) which have not been corrected for (as this post-processing option has not yet been developed). Since the horizontal scales were now the same, the images could be overlaid and compared pixel for pixel to assess the amount of correlation between the height values. The results of this comparison are displayed as a contoured scattergram (Fig. 3). A scattergram is compiled by taking the value of each pixel in one image and plotting it against the value at the same pixel position in the second image, i.e. the pixel in row one column one in the profilometer image v. the pixel in row one column one in the photogrammetric image, and so on for each pair of locations. In a pair of identical images normalized to a range of 0-255 pixels the plot would be a straight line. The contour level of each of the plotted points in the scattergram corresponds to the per cent of pixels out of the total which fall on that position. The scattergram of the data from the profilometer and the stereo correspondence shows a spread away from the ideal line. The points which deviate from the straight line have different height values for the two techniques, the greater the difference the larger the deviation. Nevertheless, there is a strong degree of correlation between the two data sets, with 90% of the data falling within 1.1 pixels (0.059mm) of an ideal correlation, and 80% of the data corresponding to 0.88 pixels (0.047mm). This result may actually underestimate the accuracy of the technique because errors introduced by parallax and the interpolation scheme were ignored; however, until further work is carried out it will be assumed that a vertical resolution equivalent to the distance represented by 1 pixel in the raw image can be achieved.

32

M . W . JESSELL E T AL.

Fig. 4. Quarrying-induced fracture in granodiorite plug from Mt Alexander Quarry, Harcourt, Victoria. The rock face immediately above the stepladder was the test area shown in Fig. 5. The radiating ridges and grooves on the fracture surface emanate from the rupture nucleation site at the base of the drill hole, where the explosives were sited. Scale bar below and to right of step ladder represents 2 m.

To improve the resolution, the horizontal scale that each pixel represents needs to be decreased. This can be done either by increasing the magnification of the photographic technique or by scanning the resultant negative at a higher resolution, both of which are easily achievable with the correct equipment. With improved scanning equipment a photographic negative can be scanned to 10 #m per pixel so that a 10 • 10 cm photograph would produce an image with the dimensions of 10 000 • 10 000 pixels. On a 1 : 1 scale photograph, for example, the pixel size would represent a true length of 10 #m on the ground. The vertical precision of any elevation data can therefore be determined through any combination of scale and scanning resolution for each application. It should be noted that this vertical resolution still falls short of the 0.1 #m achievable using the profilometer; however, the horizontal range of the profilometer is quite limited. The usefulness of the digital photogrammetric technique lies in its application to a range of horizontal scales (e.g. from satellites to microscopes) with a consistent methodology, rather than in its ability to measure the finest vertical resolution at a given scale.

Application to Mt Alexander Quarry fracture Photographs and calibration data from a quarrying-induced fracture in a granodiorite plug from Mt Alexander Quarry, Harcourt, Victoria, Australia have been collected (Fig. 1). This single fracture covers an area over 10 • 10m, and the radiating ridges and grooves on the fracture surface emanate from the rupture nucleation site at the base of the drill

33

DIGITAL P H O T O G R A M M E T R Y OF R O U G H SURFACES

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E T AL.

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DIGITAL P H O T O G R A M M E T R Y OF R O U G H SURFACES

35

a

b

Fig. 7. 2D fast Fourier transform images of the topographic data at two scales of measurement, based on 256 • 256 pixel subsets of the images in Fig. 6. Brightness of individual pixels is proportional to the log power spectral density for each pixel. The position of each pixel in this image is plotted in polar coordinates, with the radial distance from the centre a function of frequency and angular position a function of orientation. Lowest frequency pixels are in the centre of the image.

hole where the explosives were sited (Fig. 4). This fracture therefore represents a good test of the measurement technique as it allows analysis of the observed anisotropy of roughness. For comparison purposes stereo photographs were collected at a number of scales; however, in this paper the results from two scales only will be presented (with a scale change of 9.4: 1) which will be referred to as 'near' and 'far' photos, covering a horizontal field of view of 32 and 281 mm, respectively (Fig. 5). These two pairs of photographs were scanned using a fiat-bed scanner at 75 dots per inch and then processed using the digital photogrammetry program. Once the null matches had been filled by interpolation (which

36

M . W . JESSELL E T AL.

P

(a)

(b)

Fig. 8. Contours of log power spectral density, based on spatially averaged versions of the 2D fast Fourier transform images shown in Fig. 7. These contours clearly show the change from anisotropic roughness at lower frequencies to isotropic roughness at higher frequencies. This is apparent both within the larger-scale image and in comparison with the smaller-scale image, and can be correlated with the local orientation of radially dispersed ridges in the quarry face. Contours are drawn at levels of 94, 91, 84 and 52% of peak power (at the central pixel).

represented < 7% of the total image in the near image and < 0.3% of the total image in the far image) we had complete data for the two areas, which can be represented as two grey scale images of the topography of the surfaces (Fig. 6), with higher areas shown as lighter shades of grey. Although these spatial representations of the data provide a useful dataset, characterization of the scale-dependent and anisotropy nature of the roughness is better made in the spectral domain, and so we have selected representative 256 x 256 pixel sub-areas of each image for spectral analysis. First, a cosine mask was applied to the raw height image, which minimizes the boundary effects of the image. These areas are transformed using a 2D fast Fourier transform (FFT) algorithm and are plotted as log power spectral density (PSD) v. log spatial frequency (SF). Values are proportional to PSD for an orientation and frequency defined by position in the 2D frequency domain (Fig. 7). Of particular interest in this study is the information provided on the anisotropy of roughness. By using a 2D transform a complete picture of the azimuthal dependence of the roughness can be obtained, which is very laborious to produce using 1D sampling techniques such as profiling. By applying a spatial averaging filter to the images of the FFT spectra (which smooths out local fluctuations) contours of PSD close to the centre of each F F T image, where the majority of the power is located, can be produced (Fig. 8). Comparison of PSD contours for the two images reveals a striking contrast in the anisotropy of roughness between the small- and large-scale images. In the small-scale image the contours are close to circular for all the high power values, whereas the largescale image shows a marked axis of increased power at an angle of c. 48 ~ anti-clockwise from the horizontal, indicating an important frequency component oriented at 90 ~ to this axis, which reflects the local orientation of the surface ridges and grooves seen in Fig. 4. The anisotropy gradually gets swamped by isotopic spectral behaviour at lower power levels, which correlates with the lack of significant anisotropy in the small-scale image.

DIGITAL PHOTOGRAMMETRYOF ROUGH SURFACES

37

Conclusion We have used a new solution for the stereo correspondence problem as applied to studies of surface roughness. The method takes a digital stereo pair and automatically matches corresponding points in the images in order to extract height information. The package generates a dense dataset of heights, and comparison with profilometry data suggests that the matching accuracy was e. 1 pixel for > 90% of the data. Application of this technique to a quarrying-induced fracture provides a first insight into the mapping of roughness anisotropies which may be achieved using this technique. Using current image acquisition techniques, the strength of the digital photogrammetric technique lies in its applicability to a broad range of horizontal scales with a consistent methodology, rather than in its ability to measure the finest vertical resolution at a given scale.

References BANDIS, C. C., LUMSDEN, A. C. & BARTON, N. R. 1983. Fundamentals of rock joint deformation. International Journal of Rock Mechanics, Mineral Science and Geomechanics Abstracts, 20, 249268. BROWN, S. R. 1989. Transport of fluid and electric current through a single fracture. Journal of Geophysical Research, 94, 9429-9438. MARR, D. & POGGIO, T. 1979. A computational theory of human stereo vision. Proceedings of the Royal Society, London, B204, 301-328. POWER, W. L. & TULLIS, T. E. 1991. Euclidian and fractal models for the description of rock surface roughness. Journal of Geophysical Research, 96, 415424. Russ, J. 1993. Surface Fractals. Plenum Press, New York. SCHWARZE, P., JESSELL, M. & Cox, S. J. D. 1992. Surface reconstruction using digital photogrammetric techniques. 6th Australasian Remote Sensing Conference, New Zealand, Vol. 1, 298-307. THOMAS, T. R. 1982. Rough Surfaces. Longman, New York.

From Ameen, M. S. (ed.), 1995, Fractography: fracture topography as a tool in fracture mechanics and stress analysis

Geological Society Special Publication No. 92, pp. 39-58

Fractographic features on joints of KTB drill cores (Bavaria, Germany) P. B A N K W I T Z & E. B A N K W I T Z

GeoForschungsZentrum, Telegrafenberg A3, 14473 Potsdam, Germany

Abstract: The cores of the superdeep drilling in Bavaria (KTB: Continental Deep Drilling, Windischeschenbach) were used to estimate the orientation of the contemporary maximum horizontal stress by several methods in the KTB-field laboratory, including a study of core disking, borehole break-outs, cuttings and relaxation of the cores. Core disking surfaces are one type of coring-induced joint occurring in quasiisotropic parts of the rock. Fractographic features on core disking surfaces aid in understanding the core disking process. We determined the in situ stress orientation from fractographic structures on common coring-induced fractures, which are present in all sections of the cores. This analysis is based on the symmetry of joint characteristics (3D shape and surface features of joint planes). The symmetry axes of these joint features are related to the principal stress axes, which permits a determination of the orientation of the recent maximum and minimum horizontal stress axis (SHr Sh).

The World Stress Map documents a contemporary tectonic stress field (Zoback 1992). Near the southwestern border of the Bohemian Massif in the area of northern Bavaria, Germany (Fig. l a), the continental superdeep borehole KTB is located in the crystalline basement and penetrates the upper crust down to a depth of 9100 m. A pilot borehole was drilled down to a depth of 4000 m, at a distance of 300 m from the main borehole. The cores recovered from the boreholes consist of amphibolites, gabbros and layered paragneisses with more or less intense, steeply dipping foliation. Between depths of 6900 and 7200m the borehole cuts the northeastward dipping fault system of the Franconian Line between the upthrusted Bohemian Massif and the South German Block, the latter covered in its uppermost part by Mesozoic and late Palaeozoic rocks (Fig. lb). Coring-induced fractures in the KTB cores reflect the contemporary tectonic stress field. The results of the joint analysis can be checked by different methods which were used at the KTB-field laboratory by a special team to determine the maximum (Srt) and minimum (Sh) horizontal stress axis (SH ~ Sh) by means of core disking (restricted to homogeneous rocks of the drilling), borehole break-outs, cuttings and relaxation of the cores (Vernik et al. 1992). Our studies deal only with joints in cores of the main borehole. Using fractographic analysis additional data from the entire recovered core were collected. The collected observations provide new results on the fracturing process of core disking surfaces.

Method During the past 30 years surface morphology of joint planes has received significant attention. After initial studies of Woodworth (1896), Roberts (1961), Hodgson (1961), Bankwitz (1966) and others, numerous contributions from geology and physics continue

40

P. BANKWITZ & E. BANKWITZ

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41

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Co) Fig. 1. (a) Geological map of the western part of the Bohemian Massif (after Weber 1985). Broken line: boundary between Saxothuringian and Moldanubian Zones. (b) Interpretation of the drilled profile in a SW-NE section near the southwest boundary of the Bohemian Massif (after Hirschmann 1993). The drilling was finished at a depth of 9100m. SE1 and SE2 are strong reflectors in the reflection seismic profile crossing the KTB. SE1 correlates with the Franconian Line which is the main neotectonic boundary between the Bohemian Massif and the South German Block. O, location of the core with joints shown in the figures listed.

to improve our understanding of brittle failure of rocks. The surface features on the coring-induced joints are the traces of crack initiation, propagation and termination (arrest lines of the fracture). In addition, the 3D shape of fractures is useful for the determination of stress states during fracturing (Bankwitz & Bankwitz 1984). Information on coring-induced fractures can also be gained by comparing them with natural joints because the surface morphology is, in principle, the same on both types of brittle-fracture surfaces. The most important components of surface morphology on coring-induced fractures are diverging plumes with a centre of origin, hackle plumes, twist-

42

P. B A N K W I T Z & E. B A N K W I T Z

hackle fringes and arrest lines. Analysis of fracture surface morphology enables us to determine: the area of origin; the direction and relative velocity of propagation of the induced fractures; the relative timing of the development of induced fractures; and the orientation of the principal stresses during fracture initiation and growth. Surface morphology and the shape of coring-induced fractures in KTB cores formed during Mode I fracture propagation, where the surfaces of the growing fracture were opened in the direction of 0.3 (least compressional or greatest tensile stress). The direction of 0.1 and 0"2 are apparent from the distribution of fracture surface structures and their symmetry (Bankwitz & Bankwitz 1984, 1994). The correlation between stress field and surface morphology arises from the symmetry between stress field and deformation field during crack propagation, and the consequent symmetry of fractographic structures on fracture surfaces. This relationship holds where cored intervals have abundant coringinduced fractures.

Symmetry of fracture surface morphology If the overall morphology of brittle fracture is observed then some regularities concerning the shape of the fractures are found (Bankwitz & Bankwitz 1993). For example, the growing fracture planes tilt or deviate from the initial fracture orientation, depending on the local stress state. Deviation from the plane or origin generally culminates near to the end of fracture propagation into ring-like zones, seen as either a tilt or as twist-hackle fringes. Figure 2 shows schematically three different types of deviation from the initial fracture plane (Fig. 2B-D) and the accompanying symmetry (zero, without deviation; plus, upward deviation; minus, downward deviation), as it is found on both natural joints and coring-induced fractures. It should be noted that the ring-like marginal zone can deviate in the same sense from the initial plane of the fracture (Fig. 2B), or in an opposite sense at two points (Fig. 2C), or the deviation sense can change at four points (Fig. 2D). Depending on these deviations of the fringe itself, and depending on changes of hackle steps of the twist-hackle fringe (sinistral or dextral steps) a higher or lower symmetry exists and the symmetry axes are identified relative to 0.1 and 0.2. The latter corresponds to the horizontal stresses S H r Sh in a core. Basic types of symmetry (Fig. 2A-D) of the 3D shape of joints (first-order characteristics) and of fractographic structures on joint surfaces (second-order characteristics) occur as the response to development conditions. Besides the shape of joints, the arrangement of twist hackles, including changes of their overlap, indicate the symmetry of fractographic surface features. In many cases twist hackles are not restricted to the fringe zone, or they can occur without a defined fringe zone but they always appear more dense near the margin. In the types shown in Fig. 2A-D, fine lines indicate the position of a ringlike marginal zone, sometimes without a well-defined boundary to the main plane. Twist hackles would follow the direction of the radial fine lines. The scheme is derived from natural joints, which always have circular or curved natural margins, independent of their size, or parts of a curved margin. If not, then only a part of the whole joint is exposed. In cases of plume axis on planar joints at least one symmetry axis forms. If a coarsegrained core shows only roughly structured fracture surfaces, then the overall shape of the ring-like fracture zones, described above, can be used to determine the symmetry axes coinciding with the principal stress axes during origin and propagation of the joints.

FRACTOGRAPHY

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P. BANKWlTZ & E. BANKWlTZ

(a) Fig. 3. KTB cores (23.5 cm in diameter, depth 6435 m; amphibolite) with core disking spaced 2-5 cm and forking structures, which are opening towards the high points (left side). (a) The wavy-like trace of the core disking surfaces are visible. Low points near the right side; I, point of origin. Right lower comer: two relatively younger core disking surfaces | terminated at an older fracture ~ t h a t is oblique to the centre line. (b) A core adjacent to (a) which demonstrates the direction of core disking propagation (arrows) by forking.

FRACTOGRAPHY KTB CORE

(b)

45

46

e. B A N K W I T Z & E. B A N K W I T Z

Conditions of K T B investigation In the coarse-grained cores of the crystalline rocks of the KTB (Fig. l b) with a core diameter of 10cm well-developed fracture planes are found, but without good fractographic features, as are found in sedimentary rocks. Additionally, veins and veinlets disrupt propagation and, hence, cause irregular fracture surfaces. In many cases the coring-induced fracture surfaces contain only a rough pattern without detailed structures, but their overall shape reflects the origin, propagation direction and the ring-like fracture zones. At depths of 3000-8000 m, cores with a diameter of 23.5 cm were taken. Due to the large diameter of these cores the fracture surface structures are better developed. On cores at the KTB, orientation is marked by two colour lines rather than a scribe knife; therefore, no scribe knife fractures occur in the KTB core.

Fracture types There are several types of fractures in core from KTB, described below.

Drilling-induced centre-line fractures These fractures (parallel and oblique to the core axis) are rarely developed. Their drillinginduced origin is evident from fracture prolongation into the wall. These joints are not terminated in the core. In some cases the centre-line fractures represent fracture bifurcation, which implies a rapid fracture growth with early forking and subsequent propagation in two different parallel planes 2 m m apart in the centre of the core. In a few places, fractures seem to have started as petal lips, but their further propagation continued as core disking surfaces, not as petal-centre-line fractures.

Coring-induced fractures These fractures, initiated and terminated in the core, are dominant. Some of these fractures are oriented perpendicular to the core axis and show effects of torsion. Cores of 23.5cm diameter from 3000 to 8000m depths are intensely fractured but are sometimes without loss of cohesion (for example, Fig. 3). Therefore, not all fracture surfaces of such cores could be studied directly, otherwise the disks would have to be separated. In some cases, 0nly the traces of the fractures could be studied on the outside of the core. These traces indicate the fracture type and reflect twist hackles by the forking trace (Fig. 3). None the less, various disk fracture types are easy to distinguish because the borehole with a depth of 9100 m gives the opportunity to observe enough separated disk surfaces. Generally two main groups of disk fractures, which are dominant and occur nearly perpendicular to the core axis, can be distinguished. One group includes the core disking surfaces (CDS), characterized by saddle and trough structures, and is restricted to finegrained parts of the core. The second group contains disks with usual fractographic surface structures (UFS) similar to those found on natural joints, without the particular saddle and trough relief of the CDS. The second group is much more frequent than the first. Several subgroups are used to divide the groups depending on the differing stress conditions. The CDS on the KTB drill cores (Fig. 4) were investigated previously by R6ckel & Natau (1990) and Wolter et al. (1990). These authors have observed additional fractographic surface structures on CDS (Figs 3, 5 & 6). Sometimes there is a difference of

47

FRACTOGRAPHY KTB CORE

down-core

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Fig. 4. Core disking fracture with trough structure downcore (depth 7012 m, hornblende gneiss). The axis between the low points is SH.

several centimetres between high and low points, effecting a wavy trace on the core (Figs 3a & 6). The CDS-type is equivalent to the four-part fractographic type of natural joints (Fig. 2D). This type represents a special case of the more frequent UFS fractures, which are equivalent to type C natural joints (Fig. 2C). The CDS (Figs 4 & 6) appear in finegrained and quasi-isotropic rock, whereas the UFS (Fig. 2A-C) appear in cores of all kinds of rock. Besides the CDS, all other coring-induced fractures perpendicular to the core axis, with UFS, represent the second, and considerably more common, group (Fig. 2A-C). Their rough fractographic pattern is arranged on a surface (Figs 7-9) without the trough morphology of type D. They often initiate in the centre of the core and grow in regular fashion with plumes, hackles and arrest lines, terminating in twist-hackle fringes (Fig. 2A & B). Mostly, the fringe tilts from the opposite sides of the main plane in opposite directions, as can be seen in Fig. 2C. The core in Figs 7 & 8 shows such fracture surface structures as well, though roughly, developed. Barbed-type fractures similar to those described by Kulander et al. (1990) are sometimes formed, with the point of origin always located in the core, often near the centre and propagating downwards with the barb (Fig. 9). Transitional forms between CDS and UFS types have been recognized. In the literature opinion dominates that the CDS, with their saddle-trough structure, represent a special case of disk development. The assumption that CDS were initiated at the high-point areas of the core margin, and that the fracture propagated from these two points towards the centre of the core, seem to be supported (Dyke 1989) by experimental results concerning the strain distribution in the core, influenced by the drill equipment. From the physical

48

P. BANKWITZ & E. BANKWITZ

~t

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Fig. 5. Core disking surface with overall triangular structure (depth 6668 m, amphibolite). The typical trough structure, with two low (T) and two high (H) points, is superimposed by a central ridge in the trough which broadens towards the propagation direction (top of the right-hand figure). This shows three low points and is evidence for crack propagation in a preferred direction. It is a transitional fracture type between core disking and common fractographic joints. Arrows, direction of SH; + and - , positive and negative surface morphology.

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52

P. B A N K W I T Z & E. B A N K W I T Z

point of view, if the CDS initiates at two points of the margin there should be some evidence of overlap in the centre (T. Engelder, pers. comm.). Such CDS propagation could not be confirmed on the KTB cores. On CDS we have observed some hitherto unknown fractographic features and details of the shape of the core disking fracture which give additional information on the disking process. Such fractographic features, superimposed on the saddle structures of the CDS, are often very smooth and not easy to detect. In nearly all cases our observations indicate an initiation of the CDS in a low-point region and not at the marginal high-points. The fracture grew from this single point through the entire core to the opposite margin. No fractographic structures indicating a propagation from two marginal points towards the centre could be observed. From our observations it is apparent that the development of the CDS does not differ from the development of the UFS joints. The fractographic features superimposed on core disking saddle relief include fine hackle marks (rare) and small twist-hackle fringes (common). The fringe zones (1-2 cm) become broader in the high-point region and may disappear in the low-point area. Often core disking is characterized by a triangular surface structure (Fig. 5), caused by a ridge within the trough of the downcore curved plane. In this case three low points develop. Also of significance is a faint forking (Figs 3a, b & 6) and larger forking (Fig. 10a & b), which is also known from common tension fractures. The forking (bifurcation) marks the termination of CDS which are not developed through the whole core (Fig. 10). Concerning their shape and surface structures, a range of fractures was defined, differing from both CDS and UFS. These are both end members with many transitional forms in between.

Fracture process (origin and propagation) An origin of drilling-induced fractures at inhomogeneities (rock-internal boundaries or crossing veins) is rarely observed. The disk fracture growth is sometimes influenced by older shear planes parallel to the foliation, which dips c. 60 ~ (deviation of 30 ~ from the core axis). More commonly, thicker veins parallel to the layering have interrupted fracture growth. Sometimes the disks have started in the centre of the core and propagated towards the margin, terminating in small fringe zones with rough twist hackles. Plumes are often poorly developed. The propagation traces are marked by rough hackles. Dominantly, disk origin is initiated near or at the margin, and the disk fractures grow across the entire core, preferentially towards SH. This is counter-controlled by break-out and core relaxation analysis. Consideration of fine fractographic features reveals that CDS originate in the same way as the disks of the UFS. They initiate at the centre, or at one point, of the core margin, spreading across the whole core. The larger sections of the core provide good evidence for the relative timing of the recent joint succession (Fig. 11). In some cases slickenlines (a few centimetres) were observed in the area of origin of the brittle tensile fracture (Fig. 12). These suggest two stages of fracturing, separated by a brief hesitation of fracture growth. Probably, during propagation of an initial small fracture, sliding occurred on the initial fracture area which had developed before the fracture hesitation. Alternatively, the disk fracture may have propagated across the core, followed by a slight change in the stress field, which permitted slippage and slickenline development on the smooth section of the fracture. We prefer the second assumption because no arrest line can be observed.

FRACTOGRAPHY KTB CORE

53

(a)

Fig. 10. (a) Core disking surfaces (KTB core, 23.5cm in diameter, depth 6669m, metagabbro) strongly forking before termination; not growing across the whole core. Top, zone without traces of fractures; broken line, boundary of the fractured zone. (b) Schematic drawing from the core in (a). Broken lines, zone with veins parallel to the foliation. Arrows indicate the direction of fracture

Fig. 11. Relative timing of coring-induced fractures in KTB cores (23.5cm in diameter, depth 6434m, amphibolite). The oldest fractures are parallel O and oblique (~) to the core axis; the youngest fractures are the horizontal disk joints (~). Arrow, downcore direction.

54

P. BANKWITZ & E, BANKWITZ

SYMMETRY AXIS

.S J - ' ~ ~ " ~ - " \

?;1%

SHmi n

SHmox

/

=SHmox-150 ~ ist hackle steps

/ , /:/

~

forking SYM-AXIS 1

d

~

/

=SHmox=15@o

b C

///"C////

slickenlines Fig. 12. Two examples of coring-induced fractures with slickenlines in the area of origin, (a) Origin near the margin and preferred direction of growth through the whole core (depth 6668m, amphibolite). (b) Origin near the centre (~), followed by smooth areas (~) and forking (bifurcation) at opposite sites in S~ direction (depth 6359 m, amphibolite). (e) & (d) Interpretation of the core in (a) and in (b), schematic diagrams at the left: R, fringe; R1, inner fringe; R2, o u t e r fringe (both cores 23.5 cm in diameter); C, centre of primary initiation.

Recent stress field All joints reflect the stress state during their initiation and growth, Coring-induced fractures reflect the in situ stress field and the regularity of their fractographic geometry along the whole KTB cores (over a height of several kilometres), suggesting a relationship

55

F R A C T O G R A P H Y KTB CORE

Orientation of drillhole and drillcore instabilities

0

45

Sh 19o

SH la5 118o

Sh 225 127o

SH 31s 136o

lllililtlitilllillililiittliittillillllllllllllliittllitillliililililli

3000

V

V

3050 V

V

V

3100

V

VV

3150

V

3200

*

W vw

w wv

w

3250

w

3300 3350 v

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3400

a

-

V

t

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V

V

v

V

V

V

V

3500 V V

-

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~

. ~

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-v 3700 -

V

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~ ~

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~

o'1,-"

vvvv

V VV

w'vV ~

g"" 380O

V

w v, ---eJlr L'r_-

.... i vv I

,r

IIo

Fig. 13. Core disking high points 9 and centre4ine fractures ~' in cores of the pilot hole of the KTB, with SH at c. 1 6 5 ~ (Th. R6ckel, pers. comm.). X, dip of the centre-line fractures;., dip of the nonvertical boring induced fractures; V, trend of the non-vertical boring-induced fractures.

to the in situ stress orientation determined by other techniques, e.g. borehole break-outs (Brudy et al. 1993). The above mentioned method to infer the principal stress axis on the basis of fractographic symmetry led to the determination of the recent SH. Fractures suitable for such

P. BANKWlTZ & E. BANKWITZ

56

I

1

l

I

I

Symmetryaxis

M~nchberg

dBerneck "~

(~H ~

s0*00'-origin

' z~

dorf "45;--

?

1p

11'30'

1

b

2,o km 11~ '

12~ '

I

I

12115'

Symmetrya x i s (IH 165~

Fig. 14. (a) Maximum horizontal stress axes (small arrows) determined from disking surface symmetry of KTB cores in relation to the Franconian Line, which is the platform cover boundary between the Bohemian Massif and South German Block. Large arrows, supposed regional stress distribution. (b) The dominant type of fracture surface structures with the origin near the margin and a directed propagation through the core, mostly from south to north (depth 5778 m). The dominant NW-SE compression (b) changes below a depth of 6000 m (e.g. depth 6669 m), sometimes to a 30~ NE-SW compressional direction (e). O, Point of origin; arrows, direction of SH correlating with the symmetry axis. investigations suggest that SH is oriented at c. 165 ~ This direction coincides with the results from the core disking analysis of Wolter et al. (1990) and our break-out research (Fig. 13). At a depth below 6000 m a partial change of the SH direction to 30 ~ (NE-SW) is evident from our data. Such a change is possibly of local nature (Fig. 14). At greater depths SH seems to again have a persistent N W - S E direction. The recent main horizontal stress orientation is oblique to the fault boundary of the Bohemian Massif (130-145~ in some cases the stress orientation rotates to a perpendicular direction (30~ The dominant N W - S E compression coincides with the data in the World Stress Map. Agreement between stress directions derived from our method of analysing surface feature symmetry and other techniques is very close, so the former is suitable for both checking the reliability of other techniques and for use in its own with high level of confidence. Because it is applicable for all types of rocks, the method could become a routine method in core analysis for finding stress orientation in oriented cores.

Main results Most of the studied fractures from the core barrels are coring-induced fractures, restricted to the core itself. They are initiated by the coupling of the stress states related to in situ

FRACTOGRAPHY KTB CORE

57

stresses within the undisturbed rock, as well as a stress imbalance caused by removal of overburden pressure (Kulander et al. 1990). CDS, as well as UFS, are part of this group. 1. The direction of the maximum horizontal stress axis was determined by means of UFS features known from natural joints (Bahat 1991). In most cases the results were the same as from core disking data, indicating SH with a N W - S E direction (c. 165~ The Six orientation reflects compression oblique to the main boundary. 2. A sequence of joints is commonly developed (Fig. 11) and includes three groups (a) joints quasi-parallel to the drill core axis, as the oldest fractures; (b)joints oblique to the core axis (35 ~ deviation) formed next. Both (a) and (b) formed in response to drilling-induced stresses coupled with in situ stresses within the undisturbed rock. (c) Joints perpendicular to the core axis, these are the latest group formed primarily in response to unloading stresses coupled with in situ stresses, most likely assisted by torsional stresses (and vibrations accompanying drilling and coring). Frequently, only this last group (c) is present in the core. 3. According to their frequency, CDS with a saddle and trough structure (two high points and two low points) is the special case of UFS. Other UFS joints (Fig. 2A-C) represent the most common cases. These UFS joints are widespread and occur in different shapes (classes) in all rocks throughout the World, including highly anisotropic rocks. 4. Joint categories can be differentiated. Using data from fracture physics it was possible to differentiate between categories of fracture planes in KTB drill cores as: tension fractures (Six c. Sh), relaxation fractures (SH > Sh), compression-tensional fractures (Six > > Sh), compressional fractures (Srt > > > Sh, relatively large part of differential stress), torsion fractures. Joints approximately parallel to the core axis are rare and will not be discussed. UFS joints occur mainly in two classes: (1) as pure tensional fractures, initiating at the centre of the joints, reflecting a low amount of energy, proved by the equal dip direction of the twist-hackle steps on the entire marginal fringe; (2) as fractures formed under higher energy conditions, indicating differential stress. They originate near one border of the core and continue across the whole core from one side to the other, following the direction of the main stress. A N W directed polarity of the main horizontal compression Six seems probable.

Conclusions 1. CDS, and other joints with symmetric UFS structures, can occur together or separately. The high and low points of the CDS represent symmetry axes. CDS may also show fractographic surface elements, e.g. forking structures towards the margin, which reflect symmetry and reveal the mechanics of their growth. In many cases the CDS is the lower plane between the forking fractures, which are superimposed on the trough morphology (wavy trace on the core). 2. In many cases core disking fractures are characterized by a triangular structure with the point of origin at or near the margin of the fracture. The typical trough structure, with two low and two high points, is superimposed by a central ridge in the trough which broadens in the propagation direction. The shape of the CDS with three low points reveals crack propagation in a preferred direction. The high-point areas of this CDS type appear like fringes. 3. In a few cases slickenlines occur on small central parts of coring-induced KTB disk

58

P. BANKWITZ & E. BANKWITZ joints. The lines are positioned at the area of origin of these tensional joints with the central area being mostly tongue-shaped. It appears t h a t a small initial disk fracture experienced brief hesitation in p r o p a g a t i o n , followed by continued crack g r o w t h across the core. A slight change in the local compressional stress field caused m i n o r sliding along the initial fracture segment.

This investigation was supported by the Deutsche Forschungsgemeinschaft (DFG) since 1992. It is part of the programme 'Stresses in Rocks' with contributions from several teams, above all in the KTB-field laboratory, in several universities, and in the GeoForschungsZentrum Potsdam. We thank the D F G for financial support. We thank, for sharing their knowledge on various topics, H.G. Dietrich and Th. R6ckel, KTB-field laboratory. We are profoundly indebted to Terry Engelder and Stuart L. Dean for stimulating discussions. We thank Mohammed Ameen for helpful comments and correcting this manuscript. The paper has much benefited from the reviews.

References BAHAT, D. 1991. Tectonofractography. Springer, Berlin. BANKWlTZ, P. 1966. Uber Kliifte. II. Geologie, Berlin, 13, 896-941. & BANKWITZ, E. 1984. Die Symmetrie von Kluftoberflfichen und ihre Nutzung ftir eine Palfiospannungsanalyse. Zeitschrift fiir Geologische Wissenschaften, 12, 305-334. - &- 1993. Stress analysis on KTB drill cores derived from fractographic features. KTB Report 93-2, 213-218. - & - 1994. Event related jointing in rocks on Bornholm (Denmark). Zeitschrift flit Geologische Wissenschaften, 22, 97-114. BRUDY, M., FUCHS, K. & ZOBACK, M. D. 1993. Stress orientation profile to 6 km depth in the KTB main borehole. KTB Report 93-2, 195-197. DYKE, C. G. 1989. Core disking: Its potential as an indicator of principal in situ stress directions. In: MAURY, V. & FOURMAINTI~UX, O. (eds) Rock at Great Depth. Balkema, Rotterdam, 10571064. HIRSCHMANN, G. 1993. Seismischer Reflektor SE1 erbohrt - und was folgt darunter? 6. Schwerpunktkolloquium 'KTB', Abstracts, 1.-2.4.1993, GieBen, 63-64. HODGSON, R. A. 1961. Classification of structures on joint surfaces. American Journal of Science, 259, 493-502. KULANDER, B. R., DEAN, S. L. & WARD, JR., B. J. 1990. Fractured core analysis interpretation, logging, and use of natural and induced fractures in core. AAPG Methods in Exploration Series, 8.

ROBERTS, J. C. 1961. Feather-fracture, and the mechanics of rock-jointing. American Journal of Science, 259, 481492. R6CKEL, TH. & NATAU, O. 1990. Results from rock mechanical index tests of the pilot hole 'KTB Oberpfalz VB'. KTB Report 90-8, H 1-H 13. VERNIK, L., ZOBACK, M. D. & BRUDY, M. 1992. Methodology and application of the wellbore breakout analysis in estimating the maximum horizontal stress magnitude in the KTB pilot hole. Scientific Drilling, 3, 161-169. WEBER, K. (Coord.) 1985. Excursion Guide Oberpfalz. IUCL Coordinating Committee "Continental Drilling". WOLTER, K. E., ROCKEL, TH., BIJCKER, CH., DIETRICH, H. G. & BERCKHEMER, H. 1990. Core disking in KTB drill cores and the determination of the in situ stress orientation. KTB Report 90-8, G1-G13. WOODWORTH, J. B. 1896. On the fracture systems of joints, with remarks on certain great fractures. Proceedings of the Boston Society of Natural History, 27, 163-184. ZOBACK, M. L. 1992. First- and second-order patterns of stress in the lithosphere: the world stress map project. Journal of Geophysical Research, 97, 11 703-11 728.

From Ameen, M. S. (ed.), 1995, Fractography: fracture topography as a tool in fracture mechanics and stress analysis

Geological Society Special Publication No. 92, pp. 59-82

Observations on fractography with laboratory experiments for geologists B. R . K U L A N D E R

1 & STUART

L. D E A N 2

1Department of Geological Sciences, Wright State University, Dayton, OH 45435, USA 2Department of Geology, University of Toledo, Toledo, OH 43606, USA

Abstract: Brittle fracture growth proceeds through unique stages, each marked by

distinct fractographic features that can only develop during Mode I loading at the propagating crack tip. Fractographic features in any substance can be interpreted in terms of the location of the failure origin, as well as changes in propagation velocities, stress directions and stress magnitudes at the crack tip during failure. Joints in granular pervious rock, however, do not contain fractographic features, commonly formed in glass, that develop at unstable propagation rates under the influence of appreciable amounts of stored strain energy. Yet, features that develop at lower stable rates of propagation are present on fractures in glass and rock implying that the absence of certain fractographic features provides useful information. Simple laboratory experiments, primarily on glass, are discussed to provide geologists an opportunity to recognize and qualitatively interpret fractographic features diagnostic of brittle failure. The exercises demonstrate the changing propagation dynamics that control the morphological evolution of artificially-induced fractures and natural joints in rock.

The term fractography was first used in its modern sense by Zapffe & Clogg (1944) who described it as a new tool for studying fractures in metals. The current field of fractography investigates fractographic features (fracture surface structures) that comprise the topography of a fracture surface. The goal of these investigations is to interpret the dynamic history of a fracture from its origin to final termination. This paper compares fractographic features developed during simple laboratory experiments, utilizing primarily glass, to surface structures that are common on joints. The comparison provides a basis for the discussion section that focuses on selected aspects governing the dynamics of joint propagation. All experimental observations are qualitative in the sense that complex equipment and measurements are not necessary. Nevertheless, the experiments provide geologists an opportunity to gain experience in recognizing and interpreting fractographic features in all materials that undergo brittle failure. Primary fractographic features in any material can be classified as either transient or tendential (Frechette 1972). Transient structures are those that comprise the surface morphology of a fracture and commonly act together to form a complete plumose structure. Plumes with well-developed plume axes are particularly evident in uniform layers of contrasting lithology. Plumes in glass, however, can differ morphologically from those in rock.

60

B. R. K U L A N D E R & S. L. DEAN

(a)

(b)

61

(c)

(d) Fig. 1. Electron micrographs of a fracture surface on a glass rod subjected to three-point bending. (a) Fracture origin surrounded by mirror surface disrupted locally by two twist hackle faces. Lower face curves abruptly into upper face to form twist-hackle step. Twist hackle disappear as irregular stress distribution about the origin becomes uniform and mirror surfaces become a single plane. ~) Same fracture origin surrounded by mirror, then mist, region containing fine to coarse velocity hackle, then recurring mirror surface broken by pronounced twist-hackle faces and steps. (c) Mirror region. (d) Velocity hackle.

62

B . R . K U L A N D E R & S. L. D E A N

Fig. 2. Wallner lines on mirror surface in glass rod.

Tendential features are shown by the trace of a fracture plane on a free surface and may reflect specific transient structures, as well as gradual or abrupt changes in the propagation direction. Tendentially viewed features, including abutting patterns, may also result from interaction of multiple cracks (Kulander et al. 1979; Pollard & Aydin 1988). This paper focuses primarily on transient features. A brief description of terms commonly used in ensuing sections follows. A joint is a naturally occurring brittle fracture in rock that originates and propagates due to Mode I loading at the crack tip (Pollard & Aydin 1988). Plumose structures (Hodgson 1961) containing twist and inclusion hackle, and rib marks occur on joint faces (Kulander et al. 1979, 1990). Plumes on glass fractures can also contain Wallner lines and velocity hackle. An origin flaw that concentrates applied stress commonly serves as the location from which a single fracture develops. Origins have also been termed focal points or pits by Kies et al. (1950). An isolated inclusion that served as an origin in glass is shown in Fig. 1A. The mirror region is a flat featureless surface that surrounds the origin (Poncelet 1965; Frechette 1972). Extent of the mirror region ranges from microscopic to the entire fracture surface. A complete mirror developed in a glass rod is shown in Fig. lB. The featureless nature of the mirror is suggested by Fig. 1C. Development of velocity hackle terminates the mirror and gives the fracture surface a foggy appearance that typifies the mist region (Poncelet 1958). Preston (1939) referred to the mist region as a mat surface. The mirror to mist transition and the chaotic nature of velocity hackle is shown in Fig. 1B & D. Wallner lines (Wallner 1939; Quackenbush & Frechette 1978), or ripple marks (Poncelet

63

OBSERVATIONS ON FRACTOGRAPHY FOR GEOLOGISTS

2 CM

I ,

/I.

!

'~

~}

"

.

.

~.~

Fig. 3. Sketch of joint in siltstone. All parts of the joint developed during the same fracture event and propagation directions can be traced back to a single origin. Overall propagation direction from left to right is shown by plume geometry about plume axis 'A'. A twist-hackle band is marked 'B', twisthackle faces and steps in twist-hackle fringe marked 'C' and 'D', respectively. One rib mark surrounding the origin, and two constructed fracture front lines that are geometrically regular and parabolic, decorate the main joint face. Arrows show local propagation directions that can be determined at this scale.

1958), are common features in glass. They are generally subtle rounded features of low relief that are convex in the overall direction of fracture propagation (Fig. 2). Fracture forking occurs when a fracture, accelerated by increasing tip stresses in the mist region suddenly bifurcates into two or more divergent planes. The point where forking occurs is called the radient by Preston (1926). Twist hackle (Frechette 1972) has been discussed under the general term hackle or hackly fractures by Gash (1971), Murgatroyd (1942) and Preston (1926). Identical features have been termed striations by Poncelet (1958, 1965) and Preston (1939), and river lines by Pugh (1967). Geologists have discussed twist-hackle morphology as feather fracture border planes, cross fractures (Woodworth 1896), plumose structure, F-joints and C-joints (Hodgson 1961), fringe faces and steps (Bankwitz 1965), and dilatent echelon cracks (Pollard et al. 1982). Twist hackle develops when a fracture breaks into en echelon twisthackle faces and steps that are generally more closely spaced than adjacent main fracture planes (Fig. 3). Inclusion hackle, a form of twist hackle, can be a common constituent of hackle plumes in granular pervious rock (Kulander & Dean 1985; Kulander et al. 1990). It forms when a fracture plane is warped in passing around, or through, an inclusion and consequently does not maintain a single planar surface immediately on the inclusion's other side (Kulander et al. 1990, fig. A8). Rib marks (Preston 1926; Murgatroyd 1942) have been referred to as arrest lines

64

B. R . K U L A N D E R

& S. L . D E A N

"r; gi~)~,._! 9' .~.-.'r ~:..

'

L

~

~

"w"'~':'-~":" : ' : 1 / , ',:.'~,'22 ~

;

~

.

~

9 .:-:.\

Fig. 4. Common cross-sectional morphologies of rib marks. (A) In form of cuspate wave. (13)and (C) In form of line separating tilted panels. (D) Rounded form that may not mark points of arrest.

(Frechette 1972; Kulander et al. 1979, 1990), annular structures (Bankwitz 1965, 1966) and hesitation lines (Barton 1983). Rib marks represent a velocity discontinuity and commonly develop when propagation hesitates before or after abrupt deviation from the mean fracture plane (Fig. 4A-C). Rounded rib marks (Fig. 4D) may not mark points of hesitation. Fracture front lines that mimic rib marks can be constructed on fracture faces adequately marked by accompanying fractographic features (Fig. 3). Subsequent surface structures, caused by pressure solution or shear, include stylolites possessing cone and socket structure, undulating pressure solution seams, and slickenlines that may be marked by fibrous mineralization. These can be superposed on primary fractographic features formed initially by brittle failure. This sequence of events is geologically common in orogenic regions of lateral deformation. When these subsequent features are developed on a joint surface, they should be included in a complete interpretation of past dynamic events that led to final surface form. In this sense they are viewed as fractographic features.

Simple laboratory experiments in fractography - experimental methods The authors relied primarily upon the destruction of glass objects to provide surface structures that were compared to those found in rocks. Glass objects included large microscope slides, solid glass rods and capillary rods. One or both sides of several slides

OBSERVATIONS ON FRACTOGRAPHY FOR GEOLOGISTS

,~cribe mark--~ tapedfaceJ~ ~

65

, A

slide face under greatest tension Wallner lines

Fig. 5. (A) Glass lath prepared for cross-bending. (B) Broken lath folded back using tape as hinge.

were scoured, using a circular motion, with fine grit sandpaper and some slides were left unscoured. A number of the solid and capillary rods were also scoured. Transparent tape was then placed on one side of each slide, except those destined for thermal tests, to prevent scattering of broken pieces (Fig. 5A). The tape also acted as a hinge permitting exposure of fracture surfaces for examination (Fig. 5B). For slides broken by bending, origin flaws of varying size were scribed on some untaped slide surfaces. All slides were then broken by applying bending, torsion, point load and thermal stresses. Torsion and point load stresses were applied as shown in Fig. 6A & B. Due to lack of precise controls, several tests for a given experiment were necessary to produce results that most clearly show contrasting fractographic features related to scribe flaw size and vigour of scouring. The solid glass and capillary rods were induced to fail by stresses related to bending and torsion. Torsion stresses were applied to glass rods with pliers after wrapping their ends with layers of tape. Blackboard chalk is also an instructive medium. For example, failure was induced by hand applied torsion. Torsion was applied to chalk sticks while simultaneously adding varying amounts of compression and tension parallel to the cylindrical axis of the chalk. Failure was also caused in slides by stresses related to thermal expansion and contraction. For thermal fractures, a Bunsen burner was placed in a large open box to facilitate retrieval of broken fragments. For expansion fractures, scoured and unscoured slides, preferably square ones, were held over the flame so that the reducing flame tip was in contact with the slide centre. After rupture, that can be explosive, the slides were reassembled. For contraction fractures, a slide corner was heated until a yellow sodium flame was observed. The slide was then removed from the flame. Failure occurred within 10 s. If failure did not occur the slide edge was scoured and reheated. Next, the long edge of a slide was heated in similar fashion and cooled.

66

B. R. K U L A N D E R & S. L. D E A N

~Stress

~"'~ Tape ~-BB ~ B Fig. 6. (A) Glass lath prepared for torsion test. (B) Glass lath prepared for point load test.

Fracture surfaces were exposed for examination on slides by folding along the tape hinge. Slides and rods were then anchored in modelling clay for microscopic examination. Oblique illumination was adjusted to highlight subtle transient features observed on scoured and unscoured specimens, and those with scribed flaws were sketched. Local propagation directions were shown with arrows. Sketches of tendential features for fractures in glass slides, rods and chalk also proved instructive. Fracture development and development of fractographic features in any material selected for experimentation can be caused by diverse loading configurations applied singly or in mixed mode (Barton 1983, chapter 7). Slow (stable) propagation rates in brittle substances possessing lower tensile strength can be controlled by varying the intensity of applied stress, in such materials, rib marks commonly develop at points where fracture progress stopped. In addition, configuration of loading stress during slow fracture can be altered during tests to observe fracture reaction.

Observed experimental results Fracture surfaces in glass slides and solid glass rods All methods of stress application led to fractures that were each initiated at a discrete origin flaw. Furthermore, unscoured glass slides and rods with no scribed origin flaws, excluding fractures developed during thermal contraction, commonly show a progression of distinct fractographic features (Fig. 1B). These features, in order of development, include the mirror region containing Wallner lines that are most pronounced on slides with lightly scoured surfaces. The mirror is terminated by the mist region marked by development of velocity hackle that may become increasingly rough and chaotic in the direction of propagation. The mist region may revert to a mirror region that is commonly disrupted by smooth twist-hackle faces and steps, as shown in Fig. 1B. The mist region is also commonly terminated by forking of the initial fracture into diverging planar

67

OBSERVATIONS ON FRACTOGRAPHY FOR GEOLOGISTS

velocityrelated forking

no scribe flaw

IAI

(-~ torsion

both sides scoured

j

Y

~

i

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top origin and overall propagation direction

unscoured

bottom origin and overall propagation direction

B

4- 4~-4rr~-.~_ _' , ~ - - - ~ ~ , ~ . ] r~ Scoured I

[unscouredA)jLI Slide ~ "(~

I

side

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( D

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at Fracture

Fig. 7. (A) Glass laths broken by bending that show typical tendential view of fracture patterns. Slide bent concave towards the viewer and fractures led on reverse side. (B) Tendential fracture pattern in laths that failed under torsion. (C) Tendential fracture pattern in laths that failed under point load. View of slide face upon which pressure was applied. (D) Tendential fracture patterns on plates that failed explosively.

68

B. R. K U L A N D E R & S. L. D E A N

segments. After the inception of forking velocity hackle disappears, or its development is diminished. Fractures originating from scribed origin flaws in rods and slides required less applied stress to cause failure. This difference in failure stress was noticeable, even though stress magnitudes were not precisely measured in these qualitative tests. In every case, fractures from scribed flaws of sufficient dimension are all mirror surfaces. Furthermore, velocity hackle and forking did not develop.

Fracture forking in slides Fracture forking was produced by all stress applications, except thermal contraction, and is most commonly developed in unscoured slides with no scribed origin flaws. In every case, the point of bifurcation is immediately preceded by development of velocity hackle. After inception of forking, velocity hackle disappears or its development is commonly diminished. In this case, individual bifurcating segments may fork again. The initial angle between forking segments is c. 45 and 15~ respectively, on unscribed slides broken by bending and torsion stresses (Fig. 7A & B). On slides broken by point load and thermal stresses during expansion this forking angle develops over a short distance and is e. 180~ (Fig. 7C & D). These angles, predicted by a relationship between maximum and minimum failure stresses (Preston 1935), can be precise in controlled tests but are more variable in these qualitative experiments.

Fracture hooking in slides and rods Fracture hooking is most pronounced on slides and rods containing no scribed or scoured flaws. The hook develops towards the face that is under the least tension at failure and initially attempts to turn away from this surface (Fig. 5B). A twist-hackle fringe is commonly developed within the zone of hooking. There is an inverse relationship between stress to failure, related to origin size and distance from origin to initiation of the hook. This observation is highlighted qualitatively by the fact that hooking initiated closest to the point of fracture inception on unscoured rods with no scribed origin flaw (Fig. 8). Pronounced hooking of this nature requires rapidly accelerating propagation. Hooking of a different nature commonly develops when a fracture approaches a free surface at an oblique angle. The fracture tip will turn towards the free surface in an attempt to meet it orthogonally. Fracture hooking of this nature is particularly evident in the tendential view of themal contraction fractures that propagate slowly (Fig. 9).

Torsion fractures in slides and rods Unscoured slides, subjected solely to torsion, generally contain fractures that can be grouped into two sets orthogonal to each other (Fig. 7B). This angular relationship can be varied by a compression or tension added parallel to the slide faces during torsion. Slides scoured on one face commonly contain only one fracture set that originates on the scoured side. Slides with both faces lightly scoured may contain two fracture sets; however, fracture frequency is reduced. Chronology of development for fracture sets is shown by an abutting relationship that can be reversed on opposite ends of the slide (Fig. 7B). Furthermore, fractures of one set are initiated and lead on one slide face, whereas fractures within the other set originate and lead on the opposite side. This reversal occurs because

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69

large flaw Fig. 8. Tendential hooking in glass rods. Distance from origin to inception of hooking varies with origin size.

t__

E: 0 r "(3 (1)

v

v

(..t..-

(1) 13) "13 (1) "13

E: t,-

Thermal Contraction Fractures

Fig. 9. Tendential fracture patterns on laths that failed during cooling.

torsion fractures remain perpendicular and parallel to resultant tensile and compressive stress directions that are reversed on opposite ends of the slide (Fig. 7A). Some fractures within each set induced by torsion appear to mutually transect each other in a pattern commonly attributed to development of conjugate sets of shear joints. However, close examination reveals that one fracture is throughgoing (Fig. 10). The others, that can be seemingly continuous across the first-formed fracture, actually originate at adjacent origins on opposite sides of the throughgoing fracture. These origins are commonly twist-hackle faces and steps within a twist-hackle fringe. Torsion stresses applied to a glass rod produce a fracture surface seen tendentially as a

70

B.R. KULANDER & S. L. DEAN

m S~

> " " x J--

slideface origin /) toward ~,, t ,, (h) viewer '77 ~ "-"'"~, 9P '

"

Wallner lines Velocity hackle

,l~j t

j

origin ~ slidefacesunder greatesttension

Fig. 10. Transient morphology of abutting torsion fractures in glass laths.

helical trace. The trace ideally makes an angle of 45 ~ with the rod axis. As with slides, this geometry is attributed to the reversed sense of simple shear induced on diametrically opposite faces of the rod. Blackboard chalk, easily twisted by hand, demonstrates that this angle is dependent upon the amount of compression or tension applied parallel to the chalk axis during torsion. Axial tension and compression increase and decrease the angle, respectively. Similar fractures commonly develop in rock cores when the core jams in the barrel.

Point load fractures in glass slides Fractures produced in slides by point load pressure also commonly form at right angles to each other (Fig. 7C). Again, fractures that seemingly intersect are actually three different fractures. In this case, however, each fracture is parallel to a slide boundary. Subsequent fractures induced in previously broken segments also parallel previous fractures and slide boundaries. In any case, unscoured slides and those scoured on the loaded face require greater point load pressure to cause failure.

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Circumferential .__ tension

t

A

9r a d i a !

compression

B

Annealed"

Edge

Corner

Fig. 11. (h) Expansion stresses in glass plate caused by heat application at centre. (B) Contraction stresses in a glass lath subjected to corner and edge heating, then cooling.

Thermal fractures in glass slides Thermal expansion fractures develop at explosive rates and produce abundant velocity hackle and forking segments (Fig. 7D). Fracture is initiated immediately away from the heated area at a point of maximum circumferential tension. Circumferential tensile and radial compressive stresses causing fracture are shown in Fig. 11A. In contrast, thermal contraction fractures can grow at visibly slow rates and follow a sinuous propagation path (Fig. 9). Wallner lines are not developed. Failure is controlled by rapidly changing stress fields of low intensity caused when the heated and annealed edge or corner attempts to become circular upon cooling (Fig. 11B). Consequently, contraction fractures commonly arrest and then reinitiate propagation. Rib marks are commonly developed at these hesitation points. Thermal contraction fractures do not reach a propagation velocity necessary to develop velocity hackle and are consequently all mirror surfaces.

Bending fractures in capillariy rods Fracture surfaces transecting glass capillary rods contain two unique fractographic features that are related to the capillary bore. Both features originate at the inclusion on the side opposite that of initial fracture approach. Neither feature is developed on fractures in solid rods (Fig. 12A & B). One such feature is a set of paired Wallner lines,

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B.R. KULANDER & S. L. DEAN

h o o k ~ ~~ve!ocit](. ook

hack!e region

mist with velocity hackle

A

origin

...irror with Wallnerlines

Fig. 12. Transient features commonly develop on fracture surfaces in (A) glass rods and (B) glass capillary rods during three-point bending.

commonly called gull wings because of their combined shape. Gull wings are subtle and must be viewed with the microscope under proper oblique lighting. The second feature is an inclusion hackle tail that is an isolated twist-hackle step.

Discussion Each stage of fracture development in any brittle medium is marked by development of unique fractographic features. The features are produced when the fracture or joint deviates locally from its mean plane of propagation. Deviations are attributed to tensional crack-tip stresses that are affected by the amount of strain energy locally released, the passage of sonic waves, changes in propagation rate, changes in far-field stresses, and material discontinuities related to varying elastic moduli and textures. Even though these far-field stresses may be attributed to Mode II or III shear loading, resultant crack-tip stresses in a brittle material are tensile. Glass fractures, however, can develop velocity-related fractographic features that are absent from joints in granular pervious rock. Conversely, fractographic structures, common on joints, also form in glass. Appreciation of all features is vital for a wider understanding of factors that affect propagation dynamics. For example, realizing why certain fractographic features are absent in rock can benefit analysis. The sequential development of fracturing stages varies with mechanical and textural properties of the material and stress concentrated at the origin flaw that is required for failure. For example, glass can contain few and small flaws, high tensile strength, and a capability to store large amounts of elastic strain energy. In this case, fracture development can proceed through stages containing fractographic features that are dependent upon velocities that accelerate to higher levels than in rock. Such features include the mirror region, the mist stage containing velocity hackle, velocity-related forking and Wallner lines. Wallner lines are included in this category because propagation velocity must be elevated to the level where sonic waves do not immediately overtake and pass the fracture front. Joints in permeable clastic rocks propagate at lower rates then fractures in glass and, therefore, do not display fractographic stages related to propagation at high velocity.

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Fig. 13. Complex hackle plume geometry in siltstone marked by irregular and bifurcating rib marks and large irregular twist-hackle faces with distinct plume components. Development of subsequent twist-hackle steps has altered original extent of twist-hackle faces. Two origins shown at 'A', third origin off lower right. Irregular and bifurcating rib marks shown at 'B'. Arrows show local directions of joint advance determined from plume components at megascopic scale. Nevertheless, many joints do progress through distinct stages of development. For example, a relatively smooth portion of a joint surface may contain only low relief twisthackle, inclusion hackle, rib marks and minor undulations attributed to vagaries of the stress field. However, this relatively smooth area may break abruptly on a single joint into twist-hackle and rib marks, both consisting of adjoining panels, that stand in high relief (Bankwitz 1965).

Origins Failure origins in glass and rock occur at material flaws that allow local stress concentrations sufficient to cause failure (Fig. 1A). Flaws in glass, unless artificially induced, are generally small and widely dispersed. Whereas those in granular pervious rock are commonly larger and more numerous, including pore spaces, large grains, fossils, previously developed microcracks and joints, and irregularities in the bedding surface (Kulander et al. 1990, fig. A2). Externally applied stresses that cause brittle failure in both glass and rock generate local failure stresses that are magnified by such flaws. Glass fractures and individual joints in rock, developed during a single jointing event,

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& S. L. D E A N

progress from a single origin. Closely spaced multiple origins can, however, occur in glass when fracturing is violently induced. Hot glass quenched in water is an example. Here, thermal shock produces widespread stresses well above the minimum necessary to cause failure (Frechette 1972). In this case, velocity-induced hackle and forking, related to unstable propagation at rapid rates, are common. An exposed surface caused by jointing in rock may also contain multiple origins, with each marking a unique jointing event. In this case, the individual joint surfaces that coalesce can produce an irregular pattern (Fig. 13). However, fractographic features developed at high unstable velocities are absent.

Mirror-m&t regions and velocity hackle The featureless mirror region (Fig. 1B & C) surrounding the origin in glass fractures develops perpendicular to the principal tensile stress at the crack tip (Preston 1926; Frechette 1972, 1984). Poncelet (1958) states that this anatomy develops because fracture acceleration, caused by persistent stress at the crack tip, has not reached critical levels necessary to break a large number of bonds oblique to the mean fracture plane. Fractures in glass, initiated by larger stresses, have smaller mirror radii. Here, greater stress to failure and elevated stored strain energy accelerate the fracture tip. Consequently, the mirror region is terminated by development of velocity hackle that marks the inception of the mist region (Fig. 1B & D). At initial development, the roughened surface of the mist region appears frosted due to development of minute velocity hackle during uncontrolled (unstable) fracture propagation. A quantitative relationship between mist and mirror is described by Shand (1959). He determined from tests on glass rods, containing no residual stress that failure stress varies inversely with mirror radius (average distance from origin to inception of mist), if this radius is small when compared to rod diameter. Development of the mist region is generally attributed to sonic waves generated along the fracture front by parting bonds that may be oblique to the main fracture plane. At some critical propagation velocity, these waves begin to react with one another and with the stress field at the crack tip to cause local deviations from the main fracture plane. Individual velocity hackle so produced are irregular and chaotic in appearance. Furthermore, velocity hackle size increases with increasing angles of deviation from the main fracture plane at accelerating fracture velocities (Poncelet 1958). Shand (1959) and Poncelet (1965) state that, in glass under given test parameters, the propagation velocity critical to development of velocity hackle is approximately one-half the speed of transverse wave velocity. Schardin (1959) states that this velocity, under normal circumstances, is the approximate final fracture velocity. He found these velocities ranged from 750 to 2155ms -1 in glasses containing 40.7% PbO and 100% SiO2. The authors have not observed velocity hackle on joints in rocks, even when an electron microscope is used to examine individual grains cut by joints (Kulander et al. 1979, fig. 41A). This implies that natural joints propagate at stable rates and can be stopped at any point by decreasing the applied stress. In this case, the entire joint surface, including twisthackle and rib marks, would comprise the mirror. However, smooth mirrors, as in glass, would not develop because of granularity, porosity and fractographic features developed at low propagation velocities. Neither would rock mirrors, encompassing the entire joint, be planar because stress distributions in large natural systems are commonly more complex at all scales than in smaller glass samples that are more homogeneous and have regular dimensions.

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Fig. 14. (A) Diagram showing intersecting points of sonic wave, generated at a surface flaw, and crack front at selected times along a Wallner line. (B) Diagram showing how Wallner lines shaped like gull wings are developed. Displacement stresses of a sonic wave, generated on the lee side of an interior inclusion, catch up and couple with crack-tip stresses. Selected intersection points of sonic wave and crack front along the gull wings are shown. Two different fracture fronts (A and B) during propagation are shown to demonstrate this effect on Wallner line geometry.

Wallner lines Wallner lines are produced w h e n displacement stresses related to a passing sonic wave couple with principal stresses at the propagating crack tip. The sonic waves can be generated when a spreading fracture tip encounters a surface inhomogeneity. Successive

76

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~.B

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~

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,~

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Fig. 15. Transient features on a fracture that forked. elements of the sonic wave overtake the fracture tip and cause it to oscillate briefly out of the mean fracture plane in an attempt to remain perpendicular to the resultant tension (Fig. 14A). In a similar sense, Frechette (1984) mentions Wallner lines caused by sonic waves generated by vibration resulting from stress release during cracking. Wallner lines are convex in the overall direction of fracture propagation and can be used to infer the general shape of the fracture front. They do not show the exact shape at a particular instant because a sonic wave does not overtake the entire fracture instantaneously. Consequently, Wallner line distortion is most pronounced near the wavegenerating flaw where curvature of the wave front is most pronounced. A unique pattern of paired Wallner lines, commonly called gull wings (Frechette 1972), is generated by sonic waves as the fracture passes interior inhomogeneities (Fig. 14B). The geometry of all Wallner lines is dependent upon the shape of the fracture front and inclusion, as well as the velocity, of sonic waves and fracture propagation. It follows that analysis of Wallner line-shape in the mirror region can be used to calculate fracture velocity if thc sonic velocity is known (Poncelet 1965; Congleton & Petch 1967; Kulander et al. 1979). The distance between sections of adjacent Wallner lines qualitatively suggests variations of tensile stresses and propagation velocities along the crack tip. Away from the inception point of the causative sonic wave these quantities are low where Wallner lines are more closely spaced (Fig. 10). The authors have not observed Wallner lines on joints at a microscopic scale in individual grains or at larger intergranular scales. Their absence persists even though granular pervious rocks contain a myriad of inhomogeneities at which sonic waves are undoubtedly generated. Absence of Wallner lines in rock cannot be attributed, in every case, solely to a masking effect imposed by granularity. For example, Wallner lines, with maximum amplitude of several millimetres and maximum half-wavelengths e. 1 cm, have been observed on fractures in ceramic insulators (Kulander et al. 1979, fig. 41a). The authors, however, have never seen Wallner lines on rocks of similar fine-grained textures. In this case, absence of Wallner lines is attributed to low joint propagation velocities that permit sonic waves to overtake the entire joint tip over a short distance.

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Fork&g The propagation velocity at which fracture forking occurs is dependent on the material's elastic parameters. The velocity critical to development of velocity hackle and forking is similar, and is approximately one-half that of a transverse wave in a given material (Schardin 1959). Furthermore, if development of velocity hackle and subsequent forking are not artificially suppressed, the velocity at the inception of forking approximates the final propagation velocity. Fractographic details common to forking are depicted in a sketch of a fracture caused by bending (Fig. 15). The point of initiation for the bifurcation is on the slide surface under greatest tension. Forking is preceded by development of velocity hackle. After forking, fracture propagation slowed and then accelerated as shown by velocity hackle that disappears and then reappears. Such renewed acceleration may lead to additional forking. The number of individual bifurcations that are bounded by the outermost forking segments is dependent upon the intensity of applied stresses and the resulting amount of stored strain energy to be released. For example, Schardin (1959) concludes that an excess of elastic energy can exist within the fracturing material after forking. Consequently, surface energy, attributed to additional surface area of the bifurcating fractures, does not adequately slow propagation velocities, and individual bifurcations may fork again. In a similar fashion, excess tip stresses and elastic strain energy, that may exist immediately after fracture initiation, decreases the distance between the origin and the point of initial bifurcation. Bieniawski (1967) produced fracture forking at critical propagation velocities in crystalline rocks during controlled experiments. He concluded that forking and excess grain damage mark the boundary between stable and unstable propagation. Bieniawski's stable growth, at lower propagation rates, is controlled by the applied load and a deficiency in related stored strain energy. During this stage, fracture propagation can be stopped by decreasing the load stress. However, if unstable propagation is induced, marked by velocity hackle and forking, elastic energy, even though reduced, is sufficient to maintain propagation under a reduced load. Similarly, Irwin (1960) implies that the transition from stable to unstable propagation in brittle metals occurs when the energy released per unit of crack surface attains a critical value. Velocity-related forking of joint surfaces in sedimentary rocks, including fracture surfaces induced by drilling and coring, has not been observed by the authors. Joints studied in these observations include those developed regionally in horizontal rocks of diverse lithology within cratonic basins, and joints within foreland fold belts. In every case, joints of all sets contain fractographic features that must be attributed to Mode I loading and propagation. By adhering to this premise, the authors conclude that they have never seen a joint in sedimentary rocks within the described geological settings that propagated solely by Mode II or III shear stresses acting on the crack tip.

Twist and &clusion hackle Twist and inclusion hackle, unlike velocity hackle, can form at any propagation velocity and location on a crack surface. Twist hackle, common on joints, is initiated when a fracture progresses into a region where the principal tension rotates from perpendicular to

78

B. R. K U L A N D E R

& S. L. D E A N

I

I I

I I

/

A

leading rib segment bifurcation

point B Fig. 16. (A) Hypothetical plumose structure confined within an individual layer. Rib marks are shown by solid lines. Orthogonal plume components are shown by dashed lines. Straight and parallel rib segments are shown at A, circular rib mark at B, a parabolic section of rib mark is shown at C. See text for discussion of points D and E. (B) Plan view of bifurcating rib mark.

the crack tip line. Such a rotation implies a corresponding rotation of the crack plane. However, the entire tip cannot rotate by an instantaneous and equal rotation. Consequently, the fracture or joint breaks into en echelon, but disconnected, twist-hackle faces, with the tip lines on each face turning to become perpendicular to the new and rotated principal tension. Some, or all, twist-hackle faces may then curve abruptly, while maintaining a continuous surface, into an adjacent neighbour to form a twist-hackle step, thereby completing local separation (Kulander et al. 1990, fig. A-5). Alternatively, adjacent twist-hackle faces may be joined by later-formed twist-hackle steps that are not related to the initial fracture event. In any case, primary faces and steps can break into secondary, then tertiary, and even smaller, faces and steps at increasingly diminished scales (Kulander et al. 1990, fig. A-4). Development of such twist-hackle components attests to the complex stress distributions, at decreasing scales, along a propagating crack tip. Inclusion hackle, a form of twist hackle, can be a common constituent of hackle plumes in granular pervious rock (Kulander & Dean 1985). It is generated when a fracture plane is warped by local alterations of tip stresses near or within an inclusion. Inclusions are commonly pore spaces or material possessing elastic moduli in contrast to the matrix, as demonstrated in capillary rods. The planar warping and separation, during passage of a fracture through or around an inclusion, causes the initial plane to be at different levels on the inclusion's opposite side.

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To complete material separation, one segment of the original plane will hook abruptly into the other. This connecting segment is an isolated twist-hackle step that diminishes in relief away from the inclusion and is elongate in the direction of local fracture advance, However, on an even more local level, the direction of propagation within the step is oblique to this direction (Kulander et al. 1990, fig. A-8). In any case, the diverse shape and distribution of inclusions can cause related hackle to be irregularly shaped, producing a surface that is rough overall. Twist hackle commonly form distinct fringes near a stratum boundary where a joint commonly terminates. Twist hackle can also develop as bands where a single joint completely transects a layer of contrasting lithology and continues into adjacent layers where the twist hackle may disappear (Fig. 3), Twist hackle can also be irregularly distributed and form a complex collage of faces and steps (Fig. 13). This latter case is common in massive rock units In glass slides and layered rock, the most widespread twist-hackle fringe generally shows greatest relief between adjacent hackle faces and is located toward the stratum boundary furthest from the origin and plume axis. At the inception of this fringe, more closely spaced Wallner lines in glass (Fig. 11), as well as rib marks or constructed fracture front lines in rock (Fig. 3), show that fracturing velocities and related tip stresses were diminished. In this case, the hackle plume generally has a well-developed and fairly linear plume axis that reflects a regular distribution of crack-tip stresses. In a similar sense, the irregular collage of twist hackle faces (Fig. 13) reflects a highly irregular distribution of low differential stresses at the crack tip throughout the stratum. Consequently, no initial planar segment of the joint develops that contains a fairly linear and extensive plume axis (Fig. 3). Rib marks

Rib marks are produced at a velocity discontinuity where fracture propagation can hesitate in response to an abrupt change in the direction and magnitude of principal tension at the crack tip. If propagation continues this tension direction is commonly altered causing the new fracture surface to be abruptly inclined to that formed previously. Figure 4 depicts common morphologies of rib marks caused by fracture arrest. However, the rib mark in Fig. 4D is rounded in cross-section and may not reflect total arrest of fracture propagation. For this reason, the authors suggest that the term arrest line be replaced by rib mark. A geometrically regular rib mark in plan view, with sharp crest and no inflection points, or a constructed fracture front line drawn perpendicular to hackle components, traces points of instantaneous hesitation all along the fracture front. Both naturally developed rib marks and constructed fracture front lines with this geometry are convex everywhere in the direction of crack propagation. They may be nearly circular near the origin but approach a parabolic form as the fracture progresses in a layered medium (Fig. 16). Geometrically regular rib marks and front lines contain no sections that are locally concave in the direction of fracture propagation. Such regular features are useful for inferring and comparing, to the smallest feasible scale, stress configurations, relative stress magnitudes and relative propagation velocities during advance along a past fracture tip (Kulander & Dean 1985; Kulander et aL 1990). For example, a straight segment of a rib or front line marks points of instantaneous hesitation on a planar crack and indicates constant propagation velocity and stresses unchanged in magnitude and direction along its

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B.R. KULANDER & S. L. DEAN

length (Fig. 16A). A circular rib or rib segment on a planar crack also marks instantaneous hesitation, thereby indicating constant fracturing velocity and stresses that remained unchanged in magnitude. In this case, however, the greatest and intermediate principal stresses changed direction at a constant rate along its length (Fig. 16B). An entire rib or isolated rib segment with parabolic form (Fig. 16C) on a planar crack indicate that propagation velocities, stress magnitudes, and the directions of the greatest and intermediate stresses were altered at a regularly changing rate along its length. Adjacent rib marks and front lines that are geometrically repetitive with respect to a plume axis show maximum separation along the axis (Fig. 16A, point D). Here, tensile stresses (~r3) along a given rib or line acting perpendicular to the fracture plane at the tip, and related and propagation velocities, were greatest. Conversely, these values were lowest where adjacent rib marks were most closely spaced (Fig. 16A, point E). Maximum (Crl) and intermediate (a2) tip stresses within the fracture plane at the tip were perpendicular and parallel to rib mark tangents at any point. These stresses rotate and generally change magnitude from D to E as a3 decreases. If these stress parameters do not hold, the regular and repetitive geometries of propagating crack tips could not be maintained. Consequently, the crack tip would become geometrically irregular. A geometrically irregular rib mark and fracture front line possess sections that are alternately convex, then locally concave, in the direction of local propagation. Such features imply propagation rates that change irregularity and may not trace points of instant fracture hesitation. This observation is strengthened by geometrically irregular rib marks that bifurcate from a common point (Figs 13 & 16B). Here, the crack front delineated by the lone rib mark remained stationary on one side of the bifurcation point. Whereas, the crack front, on the other side of the bifurcation point, spreads to the position of the leading rib segment. Irregular geometry of rib marks mandates an irregular geometry for other fractographic features within a plume. These combined geometries reflect irregularly changing fracturing velocities, stress magnitudes and stress directions along the crack tip. Frechette (1972) states that, in glass, irregular and very closely-spaced arrest lines develop at stable propagation velocities below 10-3m s. Furthermore, rib marks are not developed on glass fractures that are accelerating. They may develop, however, on thermal contraction fractures.

Conclusions Simple fractography experiments utilizing easily obtained glass objects give geologists a greater appreciation of joint propagation dynamics than would be gained by rock study alone. For example, glass fractures are characterized by fractographic features, absent on joints, that develop at high (unstable) propagation velocities. This acceleration is attributed to mechanical and textural characteristics of glass that lead to levels of stored strain energy much elevated over that stored in rocks during jointing. Furthermore, fractographic features that develop at high propagation velocities are absent on joints in granular pervious rock. Fractographic knowledge provides a basis for conclusions centred upon something other than joint-trend geometry and postulated geometrical relationships with surrounding structures. For example, primary fractographic features at any scale are formed by brittle failure during Mode I loading at the crack tip. A shear origin for joints, caused by Mode II or III loading at the crack tip during propagation, is precluded. However, externally applied stresses sufficient to cause failure may contain a strong shear

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component. Secondary structures such as slickenlines or solution features, are superposed on joint surfaces by subsequent shear movement or compression. Rib marks and the related geometry of other plume components, as well as constructed fracture front lines, record the details of propagation dynamics. Geometrically regular rib marks show the location of the fracture tip at some past instant of time. Such rib marks show principal stress directions and relative magnitudes of tensile stress operating perpendicular to the propagating crack tip as well as relative propagation velocities. These observations can be compared throughout a hackle plume to a scale determined by surface structure development and preservation. Conversely, geometrically irregular rib marks may not mark points of instantaneous hesitation and are accompanied by an irregular geometry of hackle patterns within a plume. Geometrically irregular rib marks and those that are closely spaced, coupled with irregular geometry of plume components, imply stable joint propagation at low velocities. Furthermore, formation of velocity hackle and forking, commonly developed during unstable propagation, have not been found by the authors in geometrically regular or irregular plumes, on joints cutting granular pervious rock. Similarly, Wallner lines have not been observed on joint surfaces in sedimentary rocks. The authors appreciate reviews by M. Ameen, M. Cooke and an anonymous reviewer. M. Ameen's time and effort in seeing this collection of papers through to publication is especially noted.

References

BANKWITZ,P. 1965. Uber Klufte, beobachtungen in Thurinischen schiefergebirge. Geologie, 14, 242253. - 1966. Uber Klufte II. Geologie, 15, 896-941. BARTON, C. C. 1983. Systematic jointing in the Cardium Sandstone along the Bow River, Albena, Canada. PhD Thesis, Yale University, New Haven, CT, USA. BmNIAWSKI, Z. T. 1967. Mechanism of brittle fracture in rock, part II - experimental studies. International Journal of Rock Mechanics, Mineral Science, 4, 407-423. CONGLETON,J. & PETCH,N. J. 1967. Crack branching. Philosophical Magazine, 16, 749-760. FRECHETTE, V. D. 1972. The fractology of glass. In: PYE, D. L. (ed.) Introduction to Glass Science. Plenum Press, New York, 432-450. - 1984. Markings on crack surface in brittle materials: a suggested unified nomenclature. In: MECHOLSKY,JR, J. J. & POWELL,JR, S. R. (eds) American Society for Testing and Materials, STP 827, 104--109. GASH, P. J. 1971. A study of surface features relating to brittle and semibrittle fracture. Tectonophysics, 12, 349-391. HODCSON, R. A, 1961. Classification of structure on joint surfaces. American Journal of Science, 259, 293-502. IRWIN, G. R. 1960. Fracture mechanics. In: GOODIER,J. N. & HOFF, N. J. (eds) Structural Mechanics: Proceedings of the 1st Symposium~Naval Structural Mechanics, Pergamon, New York 557-591. KIES, J. A., SULLIVAN,A. M. & IRWIN, G. R. 1950. Interpretation of fracture markings. Journal of Applied Physics, 21, 716-720 KULANDER, B. R. & DEAN, S. L. 1985. Hackle plume geometry and joint propagation dynamics. In: STEPHANSSON,O. (ed.) Fundamentals of Rock Joints Proceedings. Cenetek Press, Lulea, Sweden, 85-94. , BARTON, C. C. & DEAN, S. L. 1979. The application offractography to core and outcrop fracture investigations. Morgantown Energy Technology Center, United States Department of Energy, METC/SP-79/3 , DEAN, S. L. & WARD, B. J, 1990. Fractured core analysis: interpretation, logging and use of natural and inducedfractures in core. American Association of Petroleum Geologists, Methods in Exploration Series, 8.

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MURGATROYD,J. B. 1942. The significance of surface marks on fractured glass. Journal of the Society of Glass Technology, 26, 155-171. POLLARD, D. D. ~I; AYDIN, A. 1988. Progress in understanding jointing over the past century. Geological Society of America Bulletin, 100, 1181-1204. --, SEGALL,P. • DELANEY,P. T. 1982. Formation and interpretation of dilatent echelon cracks. Geological Sociey of America Bulletin, 93, 1291-1303. PONCELET,E. F. 1958. The markings on fracture surfaces. Journal of the Society of Glass Technology, 279-288 1965. Modern concepts of fracture and flow. Poulter Research Laboratories, Stanford Research Institute, Menlo Park, Technical Report 002-65. PRESTON,F. W. 1926. A study of the rupture of glass. Journal of the Sociey of Glass Technology, 10, 234-269. 1935. Angle of forking of glass cracks as an indicator of the stress system. Journal American Ceramic Sociey, 18, 175-176. 1939. Bottle breakage - causes and types of fractures. Bulletin of the American Ceramics Sociey, 18, 35~50. PUGH, S. F. 1967. The fracture of brittle materials. British Journal of Applied Physics, 18, 129-162. QUACKENBtJSH, C. L. & FRECHE~E, V. D, 1978. Crack-front curvature and glass slow fracture. American Ceramic Society, 61,402-406. SCHARDIN,H. 1959. Velocity effects in fracture. In: AVERBACK,B. L., FELBECK,D. K., HAHN, G. T. & THOMAS, D. A. (eds) Proceedings, International Conference on Atomic Mechanisms of Fracture. John Wiley, New York. SI-IAND, E. B. 1959. Breaking strength of glass determined from dimensions of fracture mirrors. Journal American Ceramic Society, 42, 474-477. WALLNER, H. 1939. Linienstrukturen an Bruchflachen. Zeitschriftffir Physik, 114, 368-378. WOODWORTH, J. B. 1896. On the fracture system of joints, with remarks on certain great fractures. Boston Society of Natural History Proceedings, 27, 163-184. ZAPFFE, C. A. & CLOGG, JR, M. 1944. Fractography - a new tool for metallurgical research. Transactions, American Society of Metallurgy, 34, 71-107. 42,

From Ameen, M. S. (ed.), 1995, Fractography: fracture topography as a tool in fracture mechanics and stress analysis

Geological Society Special Publication No. 92, pp. 83-96

83

The character of rock surfaces formed in Mode I C. B U T E N U T H & M. H. de FREITAS

Centrefor Geological Engineering, Department of Geology, Imperial College of Science, Technology and Medicine, London S14/'72BP, UK Abstract: Samples failed in extension, i.e. in Mode I, are distinguishable from those failed in shear and mixed modes, i.e. in Modes II and III, by the macroscopic and microscopic character of their failure surface, i.e. by the fractography of their surfaces. This paper describes the use of fractography to discover whether the results of a rock test to failure have been influenced by both the properties of the rock tested and the test method used. The test in question (the hoop test) is relatively new and its operation has not yet been fully analysed; however, the use of fractography has confirmed the mode of failure operating in the test is tensile (Mode I). Further, fractography enables the results obtained from these tests on rock to be better interpreted and analysed with greater confidence than was the case before. Fractography has been used by metallurgists since the 1940s to study brittle failure but the availability of increasingly sophisticated imaging techniques, and the associated processing of data from them, has enabled the subject to be extended to cover both metallic and nonmetallic materials, including composite materials (Roulin-Moloney 1989). With rock it is possible to obtain very useful data relevant to brittle fracture using field-scale, as well as laboratory-scale, observations, as has been shown in a review of the subject by Bahat (1991). This paper describes the application of fractographic observations made in the laboratory to the analyses of a test of tensile strength (a hoop test): in doing this, the observations so made also provided data on the characters of rock surfaces failed in tension, i.e. in Mode I (Fig. 1).

Fig. 1. The basic modes of loading and their associated type of failure; (a) Mode I, opening; (b) Mode II, sliding; (c) Mode III, tearing.

84

C. B U T E N U T H & M. H. de FREITAS

MQrbte hoops

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! Z

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QJ t,,3 t,--

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Fig. 3. The ratios AF/AA and intercepts on the axis for area. The hoop test and the results obtained with it have been described in some detail by Xu et al. (1988), John et al. (1991) and Butenuth et al. (1993). In the hoop test an annular

hoop of rock is formed by coaxial drilling; two hemi-cylindrical loading platens are inserted into its central hole and expanded until the hoop fails. Two parameters are required to calculate the tensile strength of the material, namely the ultimate load line force at failure and the area of failure at that time.

R O C K S U R F A C E S F O R M E D IN M O D E I

85

Examples of results obtained from wet and dry hoop samples of Penrith sandstone and Italian marble are illustrated in Fig. 2, together with the strength of the material measured in direct tension (Butenuth et al. 1993). Note, that in the hoop test, the slope of the force v. area relationship describes a ratio AF/AA which is in very good agreement with the ratio F/A defined by tests in direct tension (for details see Butenuth et al. 1993). This indicates that there is a phenomenological correlation between the two tests which permits the ratio AF/AA defined by the hoop test to be accepted as a measure of tensile strength (Fig. 3). Hoop tensile strength, fiT, is thus defined as: tYT = Aft'/mA.

(1)

However, as can be seen from Fig. 2, the slope defining AF/AA for hoop tests sometimes has a positive intercept on the area axis: thus, it holds that for tests in direct tension (suffix

D): O'TD =

(FD--O)/(AD--O) ----FD/AD,

(2)

whilst for tests in hoop tension (suffix H): OrTn =

( F H - O ) / ( A - AO)H = FH/(A--AO)H.

(3)

Although the hoop test has proved to be a very convenient method for studying the strength of rock in extension there are two aspects of the results obtained from it which require clarification, namely (1) the deviation of some hoop test results from the line defined by the majority of tests (see Fig. 2); and (2) the cause of the A0 intercept (see Figs 2 &3).

Spread of results The results for saturated hoops of sandstone (see Fig. 2) show this spread of results most clearly; samples which had a failure area of (4.2 to 4.6) • 10-3 m 2 (arrowed) did not fall on the line defined by the results obtained for the same material, cut from the same block and tested in the same way. There was obviously something different about them and the samples were thus re-examined. Figure 4 illustrates one of these samples whose area, as calculated from the external dimensions of the hoop (the stippled area shown in Fig. 3), was 4.6 • 10-3 m 2. It can be seen that this is an overestimate of the actual area of failure because only the right-hand side of the hoop has completely failed in tension; tensile failure has not completely broken the left-hand side. Subtracting the area that has not broken in tension from the originally calculated area reduces the area from 4.6 • 10-3 to 3.57 • 10-3 m 2 and brings the result of this hoop on line with those of the other samples. That sample (Fig. 4) also showed the path of the fracture indicated by the heavy solid line on the right and left sides of the hoop differs from that shown by dotted lines on the left side of the hoop, the former having a more planar path than the latter. This difference in macroscopic features permitted other samples in the saturated sandstone set of Fig. 2, which had been completely broken in half during testing, to be re-examined on the basis that essentially planar surfaces represented the areas which had failed by the time F had reached its ultimate value: i.e. they represent the relevant A for the ultimate value of F. All the samples whose values deviated from the straight line could be remeasured using this principle: all values then fell on to the line.

86

C. BUTENUTH & M. H. de FREITAS

5 cm 9

I

I

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Fig. 4. Main failure path in a hoop, shown by the continuous line, and branching of the main failure on the left side of the hoop, shown by dotted lines (see text). This simple application of the most visible and obvious features of fractography thus resolved a problem with the results obtained that could not otherwise be answered. Smaller scale observations revealed even more about the dynamics of the test used and the manner of hoop failure.

Microscopic observations The failure surfaces formed slowly by extension and prior to fast fracture, and ultimate hoop failure, always started at the inner boundary of the hoop, opposite the area of platen separation. It is here that unequivocal evidence of the fractographic character of failure in tension can be found. Such evidence can be illustrated most clearly with reference to the samples of sandstone. The sandstone tested is of wind-blown origin having been accumulated as sand dunes: its constituent particles are rounded grains of quartz. These particles have subsequently been cemented by quartz which has formed an overgrowth to the original grains. The quartz cement adopts the crystal structure of the grain it coats and shrouds the grain in quartz that has a crystal form of rhombohedra and prisms when free to grow into pores. Under tension these crystal overgrowths separate from each other revealing the mineral cement bridges they have formed between grains and occasionally the faces of their crystal form which were originally in contact with their neighbour. Sometimes the overcoat of quartz is plucked cleanly away from the rounded grain it covers to reveal the grain itself. Figure 5a illustrates these features from a location < 0.5mm from the inside edge of the hoop opposite the area of platen separation. At the other end of the fracture, i.e. near the hoop boundary furthest from the platens in the area where the macroscopic planarity of the failure surface begins to be lost, or is lost in a bifurcating system of cracks, the microscopic character of the surface is similar to that

87

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Fig. 6. Schematic illustration of failure surfaces with essentially planar surfaces shown stippled.

Fig. 7. Schematic illustration of types of failure surface with Aplanar shown stippled. shown in Fig. 5a, with crystal faces of the overgrowth being visible and rounded grains exposed from beneath their overgrowths. Thus, despite the macroscopic differences referred to earlier (i.e. planar surfaces passing with distance from the initiation of failure to bifurcating surfaces) the microscopic evidence suggests that the surfaces always separated in tension. Figure 5b illustrates the failure surface developed in a hoop of marble: this photograph was also located within 0.5 mm of the inside edge of the hoop where failure commenced. The crystalline nature of the marble is clearly seen with fracturing occurring between and through the crystals of calcite. These microscopic characters are visible along the whole length of the failure surface, from where it starts at the inside edge of the hoop to where it ends at the outer boundary of the hoop. Here again the microscopic evidence suggests that the entire failure surface was formed under tension.

ROCK SURFACES FORMED IN MODE I

89

The intercept A o Reference to Fig. 4 and samples displaying similar incomplete failure across both sides of the hoop indicated that the total cross-sectional area (A) defined by the external dimensions of the hoop, as shown stippled in Fig. 3, did not correctly describe the area of failure. In addition, many hoops had a failure surface which departed from the plane of platen separation seen in plan and/or in elevation. This should not have occurred if they had failed in pure tension (see the path of failure on the left side of the hoop shown in Fig. 4). Such failure is schematically illustrated in Fig. 6: the stippled areas represent failure planes lying in the plane of platen separation and the unornamented areas represent failure surfaces which depart from that plane. It should be noted that in many cases the failure surfaces which depart from the plane of platen separation form part of a bifurcating system of fractures. In such cases only one side of the bifurcating fracture pattern forms a plane of sample separation, as can be seen in Fig. 4. Such bifurcations increase the total area of failure. The appearance of many hoop samples, once failed, is illustrated in Fig. 7, the stippled areas being the area of failure in the plane of platen separation. Summing the total failure area produces a larger area than that defined by the simple rectangular dimensions of the width of the hoop and its depth, A. Thus, for any sample: (4)

A=l+m+n+o+p, whereas: A p l a n a r ~-~ m

for A F / A A

=

+ n = A related to

Fultimate ,

(5)

o T.

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(6)

It will be noticed (Figs 2 & 3) that Ao is a constant regardless of A.

The constancy of Ao The areas of failure which have a consistent relationship with the load line force at failure (Fig. 3) i.e. the stippled surfaces of the type m and n in Fig. 7, are bounded, and thus have an area defined by the onset of bifurcation; the cause of this bifurcation is thus suspected to be related to the constancy of Ao. There are two possible causes for these bifurcations and they could be operating either independently or together. One cause is related to the presence of bending in the sample and the other is related to crack branching associated with dynamic crack growth.

Bending The hoop test is not as straightforward a test as may first appear - its operation is the combination of direct tension and pure bending, as shown in Fig. 8a (John et al. 1991). This combination changes the magnitude of the hoop stress across the sample (Fig. 8b). Photoelastic modelling and the strain gauging of elastic samples confirms the existence of an axially oriented neutral axis (na in Fig. 8b) and an outer zone of compression (Gentier

90

c. BUTENUTH & M. H. de FREITAS

Pure bending +

(a)

Dmect tension

(b)

Nill

Fig. 8. (a) Super position principle and its application to the analyses of a hoop test; (b) hoop stress distribution in a curved beam subjected to pure bending: maximum negative hoop stress (i.e. tensile stress) is at the inside edge of the hoop and decreases with an increase in radial distance becoming positive (i.e. compressive) beyond the neutral axis (na) (from John et al. 1991).

et al. 1991; John et al. 1991), even though the microscopic characters of the surfaces

indicate that failure has always been in tension (Fig. 5). The position of this neutral axis on either side of the hoop will only be symmetrical about the plane defined by the load line if the platens and the sample are both coaxial and concentric. Unless extreme measures are taken with sample preparation, some departure from such symmetry must be expected. Platens can move in one or more of three ways; 1. They may separate as they are supposed to, so that each pair of points on opposite sides of the plane of platen separation moves apart by the same amount as measured in the direction of the force line. This assumes the platens expand uniformly in a sample whose inner hole has perfectly parallel sides coaxial with the vertical axis which passes through the centre of expansion of the platens. 2. They may tilt by opening about a vertical axis which is not concentric to the sample so that one side of the platens, seen in plan, opens more than the other (Fig. 9a). 3. They may tilt about a horizontal axis so that one end of the platens, seen in vertical section, opens more than the other (Fig. 9b). As nothing other than ordinary care in good workshop practice is taken with machining the platens, which are made from steel, and drilling the samples, which are cored with good quality diamond impregnated bits, it is inevitable that condition (1) cannot always be guaranteed. Conditions (2) and (3) commonly occur and at present are considered to be the most likely reasons for failures of the type shown in Fig. 4. The cause of tilting need not always be due to the geometrical mismatch between platens and sample. A perfect match may be upset if one side of the sample begins to fail before the other and this can be thought of as occurring either in plan or in section, or in both. This has been observed using reflecting holographic interferometry (which is capable of detecting displacements measured in nanometres); this has shown that although both sides of a sample deform, and that the deformation can be almost symmetrical, one side invariably reaches yield and ultimate failure before the other (A1-Samahiji 1992). Whatever the cause, axes are created about which bending occurs (neutral axes).

ROCK SURFACES FORMED IN MODE I

91

(a)

onQ (b)

"~149o9 '1 ~176176 ~176176

Fig. 9. (a) Asymmetric opening of the loading platens about a vertical neutral axis (na) at time of maximum load line force, and appearance of failed sample; (b) asymmetric opening about a horizontal neutral axis (na) and appearance of failed specimen.

Figure 10 illustrates the consequence of having such axes within a specimen; they divide the sample into zones of tension and compression. One such axis, e.g. either vertically or horizontally oriented (Fig. 10a or b) will divide the total potential failure area into two and if both axes co-exist the potential failure area is quartered (Fig. 10c). Figure l la is a photograph of the macroscopic character of a sample where fractography suggests such axes existed, and Fig. 1 l b illustrates its basic features 9 The sample is sandstone from the same batch as shown in Fig. 2 and has a failure similar to that on the right side of the specimen shown in Fig. 4 (although it is not a photograph of that specimen). Tensile failure commenced at the left-side margin of the sample, which was the inside of the hoop adjacent to the platens, and progressed to the right. The surface roughness increases towards the top-right quadrant where it is hilly and rough 9 A more detailed inspection of the surface further revealed that escarpments shadow two areas of origin (Fig. 11c), so that by using morphological evidence on the surface (see, for example, Kulander & Dean 1985) the progression of failure can be retraced (Fig. 1ld). If bending is accepted as the explanation for the surface topography shown in Fig. 11, then it can be argued that this sample commenced its failure with one, vertically oriented, neutral surface, as in Fig. 10a, but whose more rapid failure in its lower half (Fig. 1lc & d) resulted in the later development of a horizontally oriented neutral axis, as in Fig. 10b, culminating in a final region of compression in the top-right quadrant, shown stippled in Fig. 1lb.

Crack branching The topography shown in Fig. 11 could also be attributed to crack branching during dynamic crack growth. Fracture mechanics has demonstrated that branching occurs when the release of stored strain energy at a propagating crack tip is so great that continued

92

C. BUTENUTH & M. H. de FREITAS

)

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(b)

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no

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t tx ICztx I (c) Fig. 10. Schematic illustration of (a) vertical and (b) horizontal neutral axes (na), (c) quartering a failure surface created by their coexistence: tz, Cz, tension and compression about a vertical neutral axis; tx, Cx, tension and compression about a horizontal neutral axis, respectively.

growth of the existing crack produces insufficient new surface area to consume the energy being released. Shock waves reflected back from the boundaries of the specimen can cause new cracks to form close to the existing (or main) crack (Ewalds & Wanhill 1984). The failure on the left side of the sample shown in Fig. 4 could be a result of this process. The question thus arises: is Ao = constant defined by bending or branching or by both. Fractography does not answer this directly but provides a powerful piece of evidence which unites all the observations made. Of the many samples tested none have exhibited slickensides or damage that can be clearly attributed to shear: all the surfaces appear to have failed totally in tension. For example, the surface in the top-right quadrant of Fig. 1 lb is not slickensided but is rough, even though this area could well have been under compression and susceptible to shear failure under compression. Such roughness as seen here is a fractographic characteristic of branching. Branching is most likely to occur when the tip of the main crack is 'blunt' as blunt ends to cracks tend to cause branching whereas sharp ends do not (Ewalds & Wanhill 1984). The coincidence of the onset of branching

ROCK SURFACES FORMED IN MODE I

93

with possible positions of neutral axes suggests that crack branching and specimen bending are in some way related. Although the nature of any such relationship has yet to be defined the evidence implies the following. As A = Ao + Aplanar, it follows that Aplanar (A is related to Fultimate for A F / A A = a.r: equation 5) has to increase as A increases, given that Ao proves to be a constant. Thus, for any ratio of inner and outer diameter of the hoop, the neutral axis or axes move with progress of the test. However, the position of the axes is dictated by a balance between the tensile strength of the material and the unfailed distance of hoop ahead of the major propagating crack tip. As the hoop is progressively deformed prior to ultimate failure, and as the neutral axes can be thought of as moving away from the boundary at which tensile failure commenced, so the aspect ratio of the crack decreases, its tip bluntens and the conditions conducive for branching commence. The failure path of the sample in Fig. 4 could be interpreted as showing how this tried to happen. Failure commences on the inside of the hoop and developed a reasonably planar surface for c. 2.5 cm on the right-hand side and 1.0cm on the left-hand side, i.e. the right side appears to have been failing in preference to the left. The character of the surface on the right side then changes to a rougher form (possibly reflecting incipient branching as the tip of the propagating crack is progressively bluntened by the bending illustrated in Fig. 8a). This continues for the remaining 2.0 cm of its length until it intersects the outer boundary of the hoop. The failure surface on the left side also changes in character, it too having a rougher character for its remaining 2.0 cm of propagation, possibly for the same reason as that suggested for the roughness of the crack on the right. Branching had clearly begun to develop on the left-hand side at c. 1.5cm from the inside edge of the hoop and was the sole form of fracture after 3.0 cm. If branching is associated with blunting of the propagating crack tip and linked to bending of the sample (Fig. 8a) then the onset of branching on the left side could have occurred at the time when failure had completely broken the right side, enabling the hoop to open like the letter C. Such movement would tend to blunten any crack tip on the left side of the hoop. Confirmation of such a model requires a linkage between the load-displacement plots for such tests and the onset of bifurcation, which is most likely to occur when the loaddisplacement curve has a negative gradient; future testing will attempt to reveal the nature of this linkage.

Conclusion Fractography has proved to be an extremely useful tool in studying the behaviour of rock samples tested to ultimate failure. It provides data that cannot easily be generated from numerical models founded on concepts of continuum mechanics, elasticity and statics. It also enables the dynamic aspects of fracture propagation and development in rock to be studied and offers considerable insight to the interpretation of results based on statics. The surfaces formed in porous and crystalline rocks by Mode I type failure have distinctive marks of tensile fractures. At microscopic level, at magnifications of as little as x 30, clear evidence of grains being pulled apart can be seen in the porous sandstone. In the marble, failure of complete crystals and crystal mosaics along cleavage and crystal surfaces gives witness to the presence of Mode I failure. A striking feature in the specimens of both sandstone and marble is the cleanness of the fractures - they are free from slickensides and debris associated with shear displacements. At higher magnifications

94

c. B U T E N U T H & M. H. de FREITAS

0

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~ compression ~escarpmenf

ROCK SURFACES FORMED IN MODE I

95

"*%.

~176176

(c)

,

(d)

Fig. 11. Macroscopic appearance of a failure surface formed in Mode I. (a) Photograph of failure; (b) general appearance of the surface, tensile failure progressed from left; (e) escarpments on the surface indicating the progress of failure in this plane; (d) reconstructed directions of fracture propagation from the trace of the escarpments on the failure surface. Note this surface is not that shown in Fig. 4 but was formed in the same way.

(• 60 + ) the microscopic facets associated with brittle failure can be seen on grains and crystals cut by failure surfaces. N o t only does fractography enable the results of tests to failure to be analysed, it further permits the relationship of the sample to its test method to be studied, thus enabling properties which are a function of the material to be separated from those which are a function of the test (e.g. the Ao v. A problem for the tests described in this paper). Fractography also has the great advantage of enabling new ideas and observations on the failure of samples to be applied to old results for force and displacement from samples failed during earlier experiments, provided the fracture surfaces are carefully preserved and are still available for study.

References AL-SAMAHIJI,D. K. 1992. Experimental investigation of the tensile failure of rock in extension. PhD Thesis, Imperial College of Science, Technology & Medicine, London University, UK. BAHAT, D. 1991. Tectono-fractography. Springer-Verlag, Berlin. BUTENUTH, C., DE FREITAS, M. H., AL-SAMAHIJI, D. K., PARK, H. D., COSGROVE, J. W. & SCHETELIG, K. 1993. Observations on the measurement of tensile strength using the hoop test. International Journal of Rock Mechanics and Mining Sciences and Geomechanics Abstracts, 130, 157-162. EWALDS, H. L. & WANHILL, R. J. H. 1984. Fracture Mechanics. Edward Arnold, London. GENTIER,S., POINCLOU,C. & BERTRAUD,L. 1991. Essai de traction sur anneux (Hoop Tensile Test), application aux schistes ardoisiers. R 31682.BRGM. Orleans, France. JOHN, S., AL-SAMAHIJI,D. K., DE FREITAS, M. H., COSGROVE, J. W., CLARKE, B., LOE,N. & TANG, H. 1991. Stress analysis of a unidirectionally loaded hoop specimen. Proceedings of the 7th International Congress on Rock Mechanics, Aachen, Germany, 513-518. KULANDER,B. R. & DEAN, S. L. 1985. Hackle plume geometry and propagation dynamics. In:

96

C. BUTENUTH & M. H. de FREITAS

STEPHANSSON,O. (ed.) Fundamentals of Rock Joints. Proceedings of Lulea University of Technology, Sweden, 85-94. ROULIN-MOLONEY, A. C. (ed.) 1989. Fractography and Failure Mechanisms of Polymers and Composites. Elsevier Applied Science, London. Xu, S., DE FREITAS, M. H. • CLARKE, B. A. 1988. The measurement of tensile strength of rock. International Symposium, International Society of Rock Mechanics and Power Plants, Madrid, September 1988, 1, 125-133.

From Ameen, M. S, (ed.), 1995, Fractography: fracture topography as a tool in fracture mechanics and stress analysis Geological Society Special Publication No. 92, pp. 97-147

Fractography and fracture characterization in the Permo-Triassic sandstones and the Lower Palaeozoic Basement, West Cumbria, UK MOHAMMED

S. A M E E N

GeoScience Ltd, Silwood Park, Buckhurst Road, Ascot SL5 7Q IV, UK Present address: Independent Consultant, 122 Tachbrook Street, Pimlico, London SW1 V 2ND, UK

Abstract: Three distinctive phases of fracturing (mesofractures and macrofaults) are

recognized in the Permo-Triassic part of the cover rocks of West Cumbria. The earliest phase caused syn-depositional growth macrofaults with northerly strikes in the lower part of the cover rocks. This phase is not evident in any of the macrofaults and mesofractures which are exposed in the study area. The second phase led to the occurrence of the easterly striking macrofaults and mesofractures and is clearly evident from surface exposures. The latest phase caused the exposed northerly striking macrofaulting and mesofracturing. Some of the latter macrofaults most probably propagated upwards from the tips of the pre-existing (phase 1) northerly macrofaults. Others resulted from the coalescence of mesofractures and propagated downwards. The fractographic features on both the second and third phase mesofractures suggest that they occurred as Mode I (tensile) fractures and were later modified by shear or mixed-mode loading. The geometry and degree of development of fracture surface morphology are controlled by a combination of factors. The Lower Palaeozoic basement is characterized by complex meso- and macrofracture patterns, including a large number of local sets with a wide range of orientations. Although 'statistically' prominent regional sets are recognized in the basement rocks, these are purely geometrical and include various types of fractures that developed under different stresses. Each local set of mesofractures shows varied attitudes of fractographic features, indicating different dynamic implications of fractures within the same set. Some macrofaults in the cover may have propagated from the basement, or are related to movements on basement faults which controlled the evolution of the sedimentary basin and its subsequent structures.

The study area is a part of the Lake District which is dominated by two distinctive stratigraphic/tectonic divisions: the Lower Palaeozoic basement rocks and the Upper Palaeozoic to Mesozoic sedimentary cover rocks. Each of these divisions includes several lithostratigraphic units (Fig. 1). The area covered in this work includes parts of the Lake District Block and the onshore parts of the northeastern edge of the East Irish Sea Basin, and is flanked to the north by the Solway Basin (Fig. l a). The Lake District Block is a structural high composed mainly of Palaeozoic igneous rocks (granitic batholiths and volcanic rocks) which has received a thin sedimentary cover during the Carboniferous to the Cenozoic. The detailed stratigraphic and tectonic evolution of the region (particularly in the Palaeozoic) is complex, and has been the subject of ongoing debate which has been described in a voluminous literature (see the bibliography given in Smith 1974; Jackson et al. 1987; Mosely 1990; Chadwick et al. 1993 a, b). The post-Palaeozoic tectonic regime is relatively better defined in terms of the regional structures (Fig. l a). However, the

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Fig. 1. (a) Location (insert) and general geology map of the English Lake District (modified after Chadwick et al. 1993a). (b) A schematic map showing the distribution of the studied localities (stars).

FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

99

temporal evolution of some of the tectonic events is poorly constrained and the overall effect of these events is therefore equivocal. The summary of the post-Palaeozoic tectonic events given below is mainly based on Chadwick et al. (1993a, b). The East Irish Sea and the Solway Basin developed as extensional/transtensional basins due to an E-W tension which dominated NW-Europe, in the Permian to Early Jurassic. In the Late Jurassic to the Early Cretaceous these basins were affected by N-S tension. The Mid to Late Cretaceous was characterized by regional post-extensional subsidence. In the Palaeogene to Neogene a regional uplift was accompanied by some degree of reverse movement on the pre-existing extensional faults in the East Irish Sea and the Solway Basins. The latter movement is the result of N-S to N W - S E compression related to the convergence of Africa and Europe which was culminated in the Oligocene to the Miocene. The Lake District block was affected by a considerable uplift due to the crustal shortening during this phase. In addition, both the Lake District Block and the surrounding basins suffered a regional flexural uplift as a peripheral effect of doming associated with the onset of seafioor spreading between Scotland and Greenland. The maximum depth of burial of the Permo-Triassic St Bees and Calder Sandstones in the study area are estimated to vary from 1400 m over the Lake District Block to c. 2000 m in the eastern flank of the East Irish Sea Basin. The maximum horizontal in-situ stresses in the region are oriented N-S to N W SE.

Objectives of the present work This work forms a part of the ongoing geological and hydrogeological investigation into the suitability of the Sellafield area for a nuclear waste repository. The present study focuses on the fracture patterns and their evolution in selected parts of the cover rocks and the basement rocks. The study uses field observations of fractures (including joints, faults and veins), particularly meso- and macrofractures, to achieve its objectives. It incorporates fractography of mesofractures (see Definitions) as a tool to determine the dynamic implications of the fractures and the tectonic evolution of the study area. Fracture data were collected from coastal and inland exposures of the Ordovician basement (the Borrowdale Volcanic Group and the Ennerdale and Eskdale Granites) and the PermoTriassic cover rocks (the Calder and St Bees Sandstones within the Sherwood Sandstone Group). In the following sections of the paper these will be referred to as 'basement rocks' and 'cover rocks' respectively. A total of 452 planar and 69 linear data were collected in ten localities in the basement rocks, and a total of 500 planar and 31 linear data were measured in eight localities in the cover rocks (Fig. 1). Individual localities cover an area of 400 m wherever possible.

Definitions of terms used in the study Fractures

The term 'fracture' is used here in the broad sense defined by Kulander et aL (1990) as a break or physical discontinuity in a rock caused by stresses exceeding the rock's strength. Therefore, it covers faults, joints, stylolites and veins. Faults are fractures which exhibit appreciable shear displacements, i.e. displacement parallel to the fracture surface. Joints are fractures which show little or no displacement parallel to the fracture surface. The limit of displacement (parallel to the fracture plane) beyond which a fracture is classified as a

100

M.S. AMEEN

fault is arguable. The scale of observation should be considered, i.e. the presence of an observable amount of offset along the fracture surface on the scale of observations should be used as a criterion for that distinction. Stylolites are pre-existing fractures or bedding plane surfaces along which pressure solution or dissolution has taken place leaving insoluble residues (e.g. clay minerals) giving the surface a dark colour. They tend to show various degrees of waveforms in profile section. Veins are fractures infilled with minerals such as quartz and calcite. Fractures develop in rocks on various scales and can accordingly be classified into four scale groups (Turner & Weiss 1963). Microfractures occur on microscopic scale and include those fractures which affect individual crystals or grains. Mesofractures can be effectively observed in 3D with or without a hand lens in hand specimens or large continuous exposures. Macrofractures are too large or too poorly exposed to be examined directly in their entirety in one exposure. Megafractures are hundreds or thousands of kilometres in length, and may occur on continental scale, e.g. the Dead Sea fault. Fractures can be classified into two main groups according to their geometrical consistency in terms of their orientation and shape (Ramsay & Huber 1987). Systematic fractures are planar and develop into well defined sets, each include parallel or sub-parallel fractures within the field of observation. Non-systematic fractures are non-planar and tend to show no consistent orientation within the field of observation. The present paper deals with systematic fractures.

Fractography and fractographic features Fractography is defined here as the discipline of diagnosing, describing and interpreting the fracture surface morphology or topography (fractographic features) which are related to the initiation and propagation of the fractures. The fractographic features include two main groups. The first are referred to here as 'surface structures' or 'surface patterns' and cover all features which can be observed directly on the exposed fracture surface. They occur as perturbations or undulations forming distinctive geometrical patterns. Such features were first described by Woodworth (1896), and their anatomy is illustrated in Fig. 2. Surface features related to shear movement (faulting) on fractures (slickenlines) are not included in Woodworth's, or later, classifications (e.g. Kulander et al. 1979, 1990). These are often related to later rejuvenation of fractures, and are therefore not included in the present definition. The second group of fractographic features includes the forms of the trace of the fractures as observed on a free surface (e.g. bedding surface or cliff face). They are referred to here as fracture traces (Fig. 2). The fractographic features are the result of the interaction between a complex combination of factors including remote stresses, lithological and mechanical heterogeneity in the rocks, and pore-water pressure. The dynamic implications of fractographic features are discussed in a separate section.

Mesofracture patterns in the cover rocks

Systematic fractures The systematic fractures in the cover rocks of the study area include joints, veins and faults. They are classified according to their orientation (and sense of slip of faults) into sets. The fractures pattern is dominated by three distinctive sets, a southerly striking set

FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

101

2a

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I 1. Main joint face 2a. Abrupt twist-hackle fringe 2b. Gradual twist-hackle fringe 3. Origin 4. Hackle plumes 5. Plume axis

6. Twist-hackle face 7. Twist-hackle step 8. Rib marks ( front lines of the fracture ) 9. Hooking 10. En echelon fractures

Fig. 2. A block diagram illustrating the different types of surface structures (patterns) and fracture trace architecture (partly based on Kulander et al. 1990).

(referred to here as set 1, which includes subsets la and lb), an easterly-striking set (referred to here as set 2, which includes subsets 2a and 2b), and bedding-parallel fractures (referred to here as set 3, or bedding fractures). A brief description of these fractures is given in Table 1 and their particular features in each locality are summarized in Table 2 and plotted in Fig. 3.

Non-systematic fractures These are a group of fractures which develop nearly at a right angle to the systematic fractures, as irregular, often curved, relatively short and open fractures (Table 1). The present study concentrates on the systematic fractures.

Surface features of the systematic mesofractures

Fracture traces Most of the fracture traces are observed on gently dipping bedding surfaces and to a lesser extent on coastal cliff faces. The observed characteristics of sets 1 and 2 are generally similar and are summarized below. On bedding surfaces the traces of the fractures form discrete discontinuities (Fig. 4a) or fracture zones (Fig. 4b-d). The latter consist of smaller fractures parallel to the zone (Fig. 4b) or oblique to it (Fig. 4c), with or without a consistent en echelon pattern (cf. Fig. 4b with 4c). In thick, cross-bedded sandstones some fractures show anastomosing traces (Fig. 4e). In poorly lithified sandstones the fractures occur as zones of anastomosing microfaults (granulation seams or microgouges) which distinctly protrude above the wave-cut patchy

102

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108

M.S. AMEEN

exposures of gently dipping beds (Fig. 4f). Similar features were recognized by Aydin & Johnson (1978) who called them deformation bands. Some fractures develop splay patterns and pinnate joints (Fig. 4g). Subsets la and lb, and 2a and 2b do not occur together in each of the studied localities. However, when subsets l a and l b and/or 2a and 2b occur together their traces appear to cross-cut each other forming an 'X' shape or abut against each other forming a 'Y' shape (Fig. 4h & i). In some places a ' l a ' or '2a' fracture zone is found to consist of smaller lb or 2b fractures, respectively, and/or vice versa with an en echelon arrangement (Fig. 4j) similar to those described by Beach (1975). On cliff faces, the traces of set 1 and set 2 fractures occur as discrete discontinuities or fracture zones which either cut across bedding discontinuities or abut against them (Fig. 5a & b). Some of the fractures are offset by bedding fractures (Fig. 5c), though this is rare and is characterized by small displacement. In some cases a fracture is limited to one or a few beds, depending on the bedding-joints spacing. Some fractures continue upsection as vertical or steeply dipping discontinuities (Fig. 5b), others tend to become gentler, branch off and may join bedding discontinuities (Fig. 5d). The interaction of set 1 and 2 fractures is relatively simple. The filled or sealed set 2 fractures are consistently cut across, and in some places displaced by, set 1 fractures (Figs 4 & 6a, b). However, the latter abut the open set 2 fractures. The interaction of the traces of the set 1 and set 2 fractures in each of the studied localities is schematically shown in Fig. 7.

Surface patterns (structures) No identifiable surface patterns occur on the bedding fractures, probably due to weathering and the limited exposure. However, both set 1 and 2 fractures show surface structures which are summarized in Table 3 and illustrated in Figs 8-11 (although these are better developed and/or preserved on set 1). The fact that coastal cliff faces (which mostly trend N-S) form the main source of data has undoubtedly limited the exposure and thence observations of set 2 fractures. The interpretation of the surface patterns is discussed in a later section.

Is the layering control on surface patterns universally pertinent? The configuration and degree of development of the surface patterns on both sets 1 and 2 fractures are constrained by lithological heterogeneity, syn-sedimentary structures and the occurrence of free surfaces (open fractures), such as bedding parallel joints. The shape of the rib marks and associated plume axes are considered by some authors (e.g. Kulander et al. 1979 ) to be constrained by the existence of layering, where fractures located far from bedding discontinuities may develop circular rib marks, whereas those developed between more closely spaced layer boundaries tend to show parabolic or elliptical rib marks and layer-parallel plume axes. Such a constraint is unequivocal when a bedding discontinuity is a free surface and is used here to date the fracture (chronology). However, the constraints of bedding or layering, defined by lithological changes rather than free surfaces, on the fracture surface patterns is equivocal, as is evident from this study. The systematic set 1 fractures commonly contain plume components that cross sedimentary layering, thereby ruling out exclusive control of 'non-free surface' layering on the development and geometry of surface features. Lithological heterogeneity and syn-sedimentary structures

FRACTOGRAPHY AND F R A C T U R E CHARACTERIZATION

109

110

M.S. AMEEN

Fig. 4. (a-k) Photographs illustrating the different patterns of fracture traces on bedding surfaces. For details see text. Hammer shaft, compass and pencil are 0.3m, 10 and 14cm in length respectively, rucksack is 0.5 m in height.

FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

111

112

M.S. AMEEN

Fig. 5. (a--d) Photographs illustrating the architecture of the fracture traces on cliff faces, bedding traces (B) are nearly horizontal. For details see text. Hammer shaft and measuring-tape box are 0.3 m and 5 cm in length, respectively.

(e.g. cross bedding) seem to affect the degree of development rather than the geometry of surface features (Fig. 8). This may be related to the stress intensity at the propagating fracture tip being high enough for the fracture to propagate across coherent sedimentary layering or lithological heterogeneity. In addition, asymmetry in the development of the fracture surface structures may reflect a corresponding asymmetry in the intensity of fracturing stresses (Fig. 9a & b).

FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

1 13

Fig. 6. An overall view (a) and a close-up view (b) showing an older fault of set 2 (F2) with two sets of slickenlines (arrows); horizontal (older) and oblique (younger) and later calcite seal (c) which has been cut and displaced by younger fracture of set 1 (F1).

114

M.S. AMEEN

od I

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FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

115

Table 3. A summary of the surface patterns (surface features) on the systematic sets 1 and 2 Feature

Set 1

Set 2

Origins

Mostly located in the beds; in the middle or at top or bottom third of the fractured bed of horizon (Figs 8 & 9). Voids, coarse grains or lithological heterogeneity form the nuclei of the fractures (O; Fig. 8a). Multi-origins are common on large fractures (Fig. 8a).

Mostly located in the beds in the middle or at top or bottom third of the fractured bed or horizon (Fig. 11). Voids, coarse grain or lithologlcal heterogeneity form the nuclei of the fractures.

Rib marks

Well developed, mostly parabolic in shape with their axes mostly parallel or nearly parallel to the bedding (Figs 8 & 9). The spacing of the rib marks, their geometrical symmetry and degree of development vary greatly (Figs 8 & 9).

Weakly developed, circular concentric around the origin (Fig. 11) or parabolic in shape. The axes of the latter are mostly parallel or nearly parallel to the bedding. The spacing and degree of development of rib marks vary greatly.

Hackle plumes Well developed, chevron shaped with their axes mostly parallel or nearly parallel to the bedding trace on the fracture. Their geometrical symmetry and degree of development vary greatly (Figs 8 & 9).

Weakly developed, either radial associated with concentric rib marks, or chevron shaped with their axes gently inclined to the west at ~< 15~ (Fig. 11).

Twist-hackle fringes

Both gradual and abrupt twist-hackle fringes occur with average twist angles of 30~ for subset la and 15~ for subset lb. The sense of rotation of twist faces is clockwise in subset la and counterclockwise in subset lb. The degree of development, geometry and symmetry of the occurrence of the twist hackles on the opposing fringes of a fracture face vary greatly (Figs 9b & 10).

Mostly sharp with average twist angles of 11~ for subset 2a and 13~ for subset 2b. The sense of rotation of twist faces is clockwise in subset 2a and counter-clockwise in subset 2b. The degree of development, geometry and symmetry of the occurrence of the twists on the opposing fringes of a fracture vary widely

Degree of development, symmetry and continuity of surface features

The surface features of well-exposed fractures continue across many bedding parallel fractures and sealed fractures of set 2 (Figs 9 & 10). However, the degree of development, symmetry and continuity of the surface features appear to be affected by lithological heterogeneity, synsedimentary structures, and the magnitude and orientation of local effective principal stresses across the fractured horizon, in addition to the existence of open fractures (free surfaces).

Characterized by relatively faint surface features, partly due to obliteration by later shear movements (faulting), and masking by calcite/ haematite coating and filling. However, similar factors to those quoted for set 1 have most probably affected the evolution of the fracture surface features for set 2.

C h r o n o l o g y o f the mesofractures Cross-cutting and abutting relationships of the fracture sets described above suggest that set 2 developed first, and were then sealed by calcite and/or iron oxide. Some mineralization occurred during subsequent faulting, as evident from fibrous mineral filling (Fig. 12a & b). This phase was followed by set 1 fractures which cut and displaced all sealed set 2 fractures. However, where the seal never existed, or had been partly or fully removed, set 1 fractures abut set 2 fractures. The continuation of surface structures on set 1 across the traces of sealed set 2 fractures support this conclusion (Figs 8b & 12c).

116

M . S. A M E E N

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FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

117

The majority of bedding-parallel fractures occurred later than set 1 and 2 fractures. However, some bedding fractures seem to have developed throughout the different phases of deformation. Some set 1 and 2 fractures are terminated against bedding fractures or are displaced by them. However, most of set 1 and 2 fractures cut across and displace the bedding fractures (Fig. 5). In addition, the continuation of surface patterns on many systematic fractures, across existing bedding fractures, indicates that the latter developed after the former (Figs 8 & 12c). Alternatively, at least some bedding fractures could have developed first and then healed before being cut by the systematic fractures. However, the lack of any mineralization along bedding discontinuities makes the latter explanation less likely. The non-systematic fractures always end at the open systematic fractures or cut across them, if they are healed (Fig. 13d), and are most probably the result of late-release stresses which post-date the systematic fractures.

Dynamic implications of the mesofractures The configuration of surface patterns (structures) and traces of fractures are considered to be a direct result of the magnitudes and directions of the local principal stresses at the tips of the propagating fractures, within a fractured stratum, as illustrated in Fig. 13a. The dynamic implications of each of the systematic fractures (sets 1 and 2) are discussed according to their chronology, starting with the older set (set 2). In this context, the author uses local principal stresses to mean the net or effective maximum, intermediate and minimum principal stresses (al eft, a2 eft, a3 eft, respectively).

Set 2 mesofractures The majority of the set 2 fractures lack visible surface patterns, but show slickenlines and/or are mineralized. Visible surface structures (Table 3) indicate that the fractures originated within the beds. The fractures then propagated either radially and cyclically (evident from rib marks, e.g. Fig. 1la), or steadily (evident from the lack of rib marks) with their leading edge (defined by the hackle-plume axis) inclining at 15~ to the west. These configurations suggest that the fractures originated as Mode I (tensile) fractures, with the local minimum principal stress (cr3 eft) oriented at right angles to the main fracture surface (nearly horizontal). The local maximum and intermediate principal stresses (al eft and a2 eft) lie within the fracture surface. At the leading edge of individual fractures al eff is inclined at 15~ to the west and a2 eft is at right angles to it (Fig. 11). Fractures which are not mineralized, and show no visible surface patterns, can be considered as Mode I (tensile) fractures which have propagated steadily and therefore lack rib marks. However, fractures which show evidence of shear movements (e.g. slickenlines) most probably originated as Mode I (tensile) fractures and were subsequently subjected to shear loading. This explanation is evident from the occurrence of rib marks and hackle plumes on some fractures which show indications of 'later' shear movement (faulting). In addition, as shown in Fig. 7 (Loc. 5), some 2a faults are formed by the coalescence of small en echelon 2b fractures. The latter, in turn, consist of smaller, right stepping en echelon fractures (counter-clockwise twist faces). The occurrence of some 2a and 2b fracture zones in which the constituent small en echelon 2a and 2b fractures enclose a small acute angle (< 20~ as shown in Fig. 7 (Loc. 4), offers a good constraint on the directions of the

118

M.S. AMEEN

F R A C T O G R A P H Y AND FRACTURE CHARACTERIZATION

119

principal stresses during the mixed-mode loading. The sense of stepping and twist on these zones indicate an overall N N W - N N E horizontal, 0-3 and E N E - W S W vertical 0-1/0-2plane. The sense of slip on the fractures indicates that 0-1 was gently dipping to the west during the final stages, which resulted in the sinistral normal and dextral normal faulting on the 2a and 2b fractures, respectively. Therefore, subset 2a and 2b fractures indicate a consistent overall remote stress regime characterized by a horizontal northerly 0-3 which varied in orientation temporarily and/or spatially between a N N E and a N N W direction. The O'1/O"2 plane is perpendicular to 0"3, with 0"1 > 0"2 and 0-1 gently inclined to the west during the latest phase of faulting (Fig. 13). However, early (burial and/or tectonic) vertical or near vertical 0-1 is very likely. The dynamic implications of the set 2 fractures are used to construct stress maps in the study area associated with the evolution of these fractures in individual localities (Fig. 15a-c). The occurrence of out-of-plane propagation of both subset 2a and 2b fractures is manifested by the occurrence of both gradual and abrupt twist-hackle fringes. These are the results of gradual or abrupt changes in the orientation and/or the magnitude of the local principal stresses. Assuming a linear-elastic behaviour, the angle of rotation of a twist-hackle face (twist angle) is a function of the rock's rheology and the ratio of shear to normal effective stresses. Such a relationship was first analysed and quantified (in terms of Poisson's ratio and the ratio of Mode I to Mode II stresses) by Lawn & Wilshaw (1975) for brittle solids. Pollard et al. (1982) and Cruikshank et al. (1991) applied the latter analysis to certain sandstones. This method for estimating stresses is useful, but the results have to be considered cautiously, because it is based on the linear-elastic fracture mechanics, a simplified theory of a complex phenomenon (Lawn & Wilshaw 1975). The average twist angles on subset 2a and 2b fractures of 11 and 13~ respectively (Table 3), can be accordingly used to estimate the ratio of shear to tensile stresses at the time of fracture propagation. Assuming a Poisson's ratio of 0.25 for the St Bees and the Calder Sandstones the tensile/ shear stress ratios for subset 2a and 2b fractures are 0.11 and 0.13, respectively. Due to the lack of information about the pore-fluid pressure at the time of fracturing it is impossible to estimate the regional stresses from these ratios. However, it can be concluded that the net normal stresses were considerably higher than the shear stresses. The common occurrence of calcite mineralization along the set 2 fractures suggests an important role of pore-fluid pressure in the initiation and/or propagation of the fractures. Detailed fluid inclusion analysis may help in estimating the pore pressure (Engelder & Lacazette 1990) Fig. 9. Photographs (left) and maps (right) showing: (a) A set 1 fracture propagated with a maximum rate along arrow 1 to rib mark R1 then downward (arrow 2), as evident from the larger spacing of the rib marks in these directions. Note the convergence of several rib marks between R1 and R2 at I, indicating (slower) propagation than along arrows 1 and 2. This drop in the propagation rate coincides with the occurrence of a smaller, parallel joint which originated at O' and propagated along arrow 4 (whose projection on the main fracture lies on the same line as I). Rib mark deflections at II coincide with the intersection of the rib marks and a carbonate-rich lens and are accompanied by a zone (E) of smaller, interfering rib marks and hackle plumes (small arrows indicate the local direction of propagation). Note the continuation of the surface features across a sealed set 2 fracture (F). (b) Asymmetrically developed surface patterns on a set 1 fracture. (c) A partly exposed set 1 fracture surface showing three main phases of propagation history: an early, cyclic, often interrupted, phase evident from closely spaced rib marks (phase 1); a transitional, cyclic but less interrupted phase, evident from wider spaced rib mark (phase 2); and a final non-interrupted phase evident from the absence of rib marks (phase 3). Note the localized occurrence of twist hackles in phase 2 near its border with phase 1.

120

M.S. AMEEN

Fig. 10. Set 1 fractures showing: gradual (a & b) and sharp (e) twist-hackle fringes, both with a counter-clockwise sense of rotation of the hackle faces (1 and 2) relative to the main fractures (M). Large and small arrows indicate the directions of propagation of the main fracture and twist faces, respectively. The dashed line in (a) is a rib mark. The photograph in (b) shows the trace of the upper fringe zone of the fracture in (a) on a bedding plane.

FRACTOGRAPHY

AND

FRACTURE

121

CHARACTERIZATION

A,.o

0

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122

M.S. AMEEN

Fig. 12. (a) A dextral-normal fault (F) of sub-set 2a with calcite filling exposed partly on a cliff face (C) and partly on a bedding plane (B). (b) Fibrous calcite filling on the face of the fault in (a) indicating an oblique normal-dextral movement (parallel to arrow). (c) A dextral fault (F2) of subset 2a with incongruous step structures (S) and fibrous calcite coating which are cut by later set 1 fractures (F1) and bedding joints (B). (d) The traces (on bedding surface) of mineralized systematic set 2 fractures (F2) cut by later non-systematic fractures (F).

FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

123

124

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FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

125

and hence the regional stresses. Such studies are currently being carried out by UK Nirex Ltd, as a part of their ongoing geological/hydrogeological investigation in the study area (Andrew Bowden, pers. comm.).

Set 1 mesofractures Visible surface structures on the fractures (e.g. twist-hackle plumes) indicate that they originated within the beds as Mode I (tensile) fractures. The local minimum principal stress (a3 eft) is nearly horizontal, at right angles to the main fracture surface. The local a~ eft and cr2 eft lie within the fracture plane and are horizontal and vertical, respectively, along the leading edge of the majority of the fractures (Fig. 13). Some fractures lack rib marks and therefore indicate a steady propagation. However, the majority of fractures show rib marks, indicating incremental interrupted growth. Such interrupted growth is linked to local fluctuation in the minimum stresses (tr3 eft), possibly due to changes in pore-fluid pressure. The common hooking of lb fractures on to la fractures indicates that the former developed later. Therefore, it can be concluded that la and lb fractures developed due to NE-SW to E-W tensions (~r3), respectively. Both subsets were later affected by a stress system characterized by an ENE-WSW trending a3. Therefore, they suffered mixed mode loading which resulted in the breakdown of the fractures into twist faces. The consistent clockwise and counter-clockwise twist of the l a and l b fractures, respectively (Table 3 & Fig. 7, Loc. 2a & b), is indicative of dextral and sinistral sense of shear, respectively. The average twist angles on la and lb fractures are 15 and 30 ~ respectively (Table 3). These indicate ratios of shear to tensile stresses of 0.15 and 0.36 for subsets la and lb, respectively. The former ratio is similar to those obtained for subset 2a and 2b fractures, and indicates that the net normal stresses were considerably higher than the shear stresses during fracture development. However, the ratio for subset lb fractures is considerably higher, indicating that they developed under considerably lower normal stresses and higher shear stresses than those of set 2 and subset 1a fractures. This may be the result of sub-set l b developing later, and therefore at shallower depths than subset la. The coalescence of the la and lb mesofractures under such shear loading resulted in the development of faults. Stress pattern maps in the study area based on the temporal evolution and dynamic implications of the set 1 fractures are shown in Fig. 15d-f.

The relationship of the mesofractures to the macrofaults in the cover rocks

Macrofaults pattern Macrofaults in the St Bees and the Calder Sandstones crop out in limited coastal exposures between Seascale in the south and Whitehaven in the north. The best exposures of these macrofaults occur on the coastal cliff faces between South Head and North Head. The observed faults dip steeply (> 60 ~ and strike at N-S to NW-SE and ENE-WSW to E-W. The lower hemisphere, equal area projection of these faults is dominated by these two sets which are referred to as set 1 (northerly set) and set 2 (easterly set), respectively (Fig. 14a). The observed exposed width of the macrofaults ranges from a fraction of a metre to several metres. They show various degrees of development of fault rocks (cataclasites and gouges), as illustrated in Fig. 14b.

126

M.S. AMEEN

The 3D geometry and size of the macrofaults are difficult to determine from the small field exposures. However, seismic and borehole data indicate the presence of faults up to several kilometres in depth and several tens of kilometres in length. In addition, a wide range of geometries are evident with linear, wavy and anastomosing traces of the faults on horizontal plan, and linear or curvilinear (concave upward) profiles of individual faults (Chadwick et al. 1993ab). The net slip on all the exposed macrofaults have a consistent extensional component and vary from oblique-slip (dextral-normal or sinistral-normal)to strike-slip. There lacks field evidence of syn-depositional movements on any of the exposed macrofaults of sets 1 and 2. However, subsurface data indicate the occurrence of such movements on some of the northerly faults at depths within the lower part of St Bees Sandstones (Jon Gutmanis, pers. comm.). The implications of this are discussed in the section on the regional stresses and the evolution of macrofaults.

Chronology of the exposed macrofaults At the immediate vicinity of some of the macrofaults there occurs a localized group of synthetic and antithetic fractures (Fig. 14c). Such a relationship suggests that the mesofractures are related to the macrofaulting. However, all the exposed macrofaults are consistently parallel or nearly parallel to the regionally diffused, systematic mesofractures, i.e. set 1 and 2 macrofaults are parallel to set 1 and 2 mesofractures, respectively (cf. Fig. 14a with Figs 3c &13d). In addition, set 1 macrofaults consistently cut and displace the set 2 mesofractures, whereas set 1 mesofractures persistently abut against open segments of the set 2 macrofaults and cut across healed segments of the same macrofaults. Therefore, it can be concluded that the majority of the exposed set 1 and 2 macrofaults resulted from the coalescence of the set 1 and 2 mesofractures, respectively, during the same phases of deformation which caused these mesofractures. Hence, the exposed parts of the northerly striking macrofaults (set 1) developed after the exposed parts of the easterly macrofaults (set 2).

Mesofractures and the evolution of the macrofaults The stress regimes which caused the exposed macrofaults are similar to those which caused the mesofaults (cf. Figs 14a & 13b-d). The 3D geometry of the exposed macrofaults is not clear due to their limited outcrops. However, the geometrically parallel, spatially and genetically related mesofractures (faults and joints) form zones which are better exposed and can therefore be used to shed some light on the macrofaults and the early stages in their evolution.

Styles of mesofaults and mesofracture zones. There are four styles of smaller faults and fracture zones which constitute part of the systematic fractures (Fig. 16): (1) Discrete discontinuities which occur as discrete fracture planes (Fig. 16a); (2) Zones of granulation seams (deformation bands), as shown in Fig 4f; (3) Zones of closely spaced fractures which are curved in 3D, i.e. have curved vertical and horizontal profiles and occur as 'duplexes' enclosing blocks of rocks up to a few metres in length (Fig. 16b). The shape of these fractures will necessitate ploughing and friction large enough to produce substantial amounts of fault rocks at relatively small displacements of a few metres; (4) Localized zones of closely spaced fractures of a systematic set. These often include two subsets (e.g. 2a and 2b), one of which tends to end at the other dominant subset. The junction zone

FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

127

Fig. 14. (a) Lower-hemisphere, equal-area projection of poles to macrofaults [dextral (O); sinistral (A)] in the St Bees Sandstone and their slickenlines (corresponding solid symbols) showing two main sets striking northerly (~ set 1) and easterly (, set 2) and are the ineffective principal stresses (cf. Figs 3b & c and 13b-d). (b) Cataclasite and gouge (G) developed along the macrofault shown in (c) & (d). (e & tl) A photograph and its interpretation of a sinistral-normal macrofault dipping at 70-85 ~ towards 7~ with fibrous haematite growth (slickenlines) dipping at 38~ towards 285 ~ and vertical offset of 6m exposed on a cliff face at South Head (NX9513 1195).

128

M.S. AMEEN

between the two subsets is highly susceptible to fault rocks development (Fig. 16c); (5) Flower structures (Fig. 16d & e). These different styles can be considered as small-scale models of the macrofaults. Furthermore, they can be used to understand the embryonic stages of the macrofaults. Therefore, any model for the evolution and mutual relationship of the macrofaults should take into consideration the field evidence discussed here.

Regional stresses and the evolution of the mesofractures and the macrofaults This study shows that there is a consistent mesofracturing, macrofracturing and related stress regime in the exposed cover rock masses. An early N-S tension, evident from set 2 fractures (meso- and macrofractures) is most probably linked to the regional N-S tension (Late Jurassic to the Early Cretaceous). Later E-W and NE-SW tensions which caused the set 1 fractures (meso- and macrofractures) are in agreement with the Palaeogene to Neogene N-S to NW-SE compression and uplifting which resulted from the convergence of Africa and Europe. In addition there exists another phase of E-W and NE-SW tension (Permian to Early Jurassic), evident only from subsurface data (Chadwick et al. 1993a, b; J. Gutmanis, pers. comm.) in the lower part of the St Bees Sandstones and in older rocks (e.g. growth macrofaults parallel to the set 1 of the exposed macrofaults). It is also evident in exposed Carboniferous rocks where the northerly-striking mesofractures pre-date the easterly-striking fractures (e.g. at Saltom Bay, NX 95801593). Therefore, it can be concluded that the upper part of the St Bees Formation and the Calder Sandstone in the study area were mostly not affected by and/or post-date this early phase of the E-W tension. However, the lower part of the St Bees Sandstones was unlithified and thus adapted itself by differential compaction and consolidation, and a limited degree of syndepositional macrofaulting. Therefore, there are three main phases of fracture development (including macrofaulting) in the cover rocks, which resulted from the three post-Palaeozoic regional stress regimes (outlined above): (1) Phase 1 (Permian to Early Jurassic), is the earliest phase during which syn-depositional growth macrofaults (parallel to the set 1 of the exposed macrofaults) occurred. This was mostly restricted to the lower part of the cover rocks and was accompanied by little or no mesofracturing. Therefore, this phase is not evident in any of the exposed macrofaults which are covered in this study. (2) Phase 2 (Late Jurassic to the Early Cretaceous) occurred after the deposition of the cover rocks and led to the occurrence of set 2 macrofaults and mesofractures, and is clearly evident from surface exposures. (3) Phase 3 (Palaeogene to Neogene) is the latest phase during which the exposed set 1 macrofaulting and mesofracturing took place. Some of these macrofaults most probably propagated upwards from the tips of the pre-existing (phase 1) macrofaults. Others resulted from the coalescence of the set 1 mesofractures and propagated downwards. Some of the movements on the macrofaults during this phase were most probably also accommodated by the pre-existing set 2 macrofaults. The occurrence of a few slickenline sets on some of the set 2 faults may support this suggestion (Fig. 6). It is difficult to put a depth constraint on the evolution of the set 1 and 2 fractures. Fluid inclusion analysis of the calcite flling may be helpful. Such analyses are at an early stage and their results are not available yet (Andrew Bowden, pers. comm.). However, an upper limit of depth equivalent to the maximum estimated depth of burial of the rocks can be applied. Such a maximum depth is estimated 1400-2000 m (Chadwick et al. 1993a) and can be invoked for set 2 fractures. However, at least some of the set 1 fractures

129

FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

-

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--295

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(particularly subset 1b) are likely to have occurred at shallower depths due to the uplifting which accompanied their causative tectonic stresses. Although bedding fractures (set 3) seem to have been active throughout the tectonic history of the region most of them post-date set 1 and 2 fractures. Later stress release associated with uplifting, and removal of overburden, was probably the main contributing factor in the development of bedding-parallel fractures. This is evident from the continuation of fracture surface structures, e.g. hackle plumes, on individual fractures of sets 1 and 2 across bedding fractures (Figs 8a & 12c), and the occurrence of faults which cross-cut bedding fractures (Fig. 5b ).

130

M.S. AMEEN

FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

131

Table 4. Classification of the two statistically dominant sets of fractures in the Borrowdale Volcanics Group (Fig. 18a & b) according to their types (joints, veins etc.), local set number, location and direction of propagation Regional set

Corresponding local fractures type/local set no. (Locality), direction of propagation Joints

Veins Strike-slip

Faults Dip-slip

Oblique-slip

Unknown

1

2(14), 6(9), 6(9), 4(11)~=~, 3(9)~, 3(11), 2(15), 2(15), 3, 4(9), 2(10), 2, 3, 4(13) 4(12)yD 2(11)~=~, 3(15)/'/,," 3(14)

2(12)TR

2(14)D-R 3(10)S-R 3(15)$D-N

1(15) 2(10) 3(14) 5(12)

2

2(9) B.J., 1(11) 1(12) B.J. 2(9), 5(9)~ 1(10), 3(10)~ 1(13), 1(14) 4(14), 6(9) 3(11)

2(12)1R 1(13)

3(9)S-R

5(12) 1(14) 1(11)

1(13) 4(14) 8(9)T

5(9)S 1(13)S l(13)D 6(9)TD

Same as those used in the observations given in the Appendices (e.g. 1(9): Local set no. 1 in Loc. No. 9). Arrows indicate direction of propagation; ~, top to bottom; T, bottom to top; r h o r i z a o n t a l ; / / , oblique and top to bottom or bottom to top (if the direction is not known none of these symbols is shown). B.J., Bedding joint; D, dextral fault; S, sinistral fault; R, reverse fault; D-R, dextral-reverse fault; S-R, sinistral-reverse fault; D-N, dextral-normal fault (if sense of slip is not known none of these symbols is shown).

Mesofracture patterns in the basement rocks Several sets of fractures occur in each of the studied localities. The summary of the fracture observations in individual localities is reproduced in Figs 17 & 18 and in the Appendices. Individual sets generally show a wide range of dip and strike, and evidence of multiphase rejuvenation. Therefore, it is difficult to recognize clearly defined regionally prevalent fracture sets according to orientation, type and relative chronology. However, Fig. 16. Schematic block diagrams and photographs illustrating the different styles of smaller faults and fracture zones in St Bees Sandstone. (a) Discrete fractures, illustrated by a set 1, dextral fault dipping at 80-90 ~ towards 53-81 ~ although the fault's length and height are large (> 70, 20m, respectively) its displacement is c. 0.1 m (St Bees Head, NX9445 1333). ~) Zones of closely spaced fractures which are curved in 3D. The photographs illustrate profile and bottom views of dextralnormal fault zones consisting of such fractures of set 1 at St Bees Head with width, height and length of 1, > 20 and >30 m, respectively. The dip angle varies from 60 to 80~ and dip direction varies from 76 to 92 ~ The net slip on this zone is very small (c. 1 cm) and therefore there are no fault rocks. (e) Localized zones of closely spaced systematic fractures. The photograph shows such a zone developed between subsets 2a and 2b. Figure 5b shows the highly fractured profile section of the junction zone. (d) Negative flower structures (transtensional faults), illustrated by a set 1 fault with a width ranging from 2 m at the bottom of the cliff to a minimum of 12 m at the top of the cliff, and a minimum length of 40 m. (e) Negative flower structure (transtensional fault), illustrated by a set 1 fault dipping at 78~ towards 20 ~ with slickenlines dipping at 25~ towards 300~ associated with antithetic faults dipping at 85~ towards the west (St Bees Bay).

132

M.S. AMEEN

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FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

133

synoptic pole diagrams indicate the existence of two 'geometrical' sets in the Borrowdale Volcanics. These are generally steep dipping (> 45~ striking NE-SW to ESE-WNW and NW-SE to NNE-SSW, respectively (Fig. 18a & b). As shown in Table 4, each of the two regional geometrical sets includes a wide spectrum of fractures and has no consistent type, origin or fractographic features. In the plutonic rocks, synoptic pole diagrams of the fractures are characterized by more diffused data which may be vaguely divided into five geometrical sets. These include one nearly horizontal and four steep dipping sets striking NW, NE, NNE and WNW, respectively (Fig. 17c & d).

Surface features of the mesofractures in the basement The surface features for the individual sets in each locality are summarized in the Appendices. These include both the features described in Fig. 2 and features related to shear movement (e.g. slickenlines). The latter vary from one set to another and generally include fibrous mineral growths and frictional wear. No clear evidence of pressure solution was noticed in the field on any of the fractures. Fractographic features are often obliterated or masked by slickenlines, mineral filling and coating, and are described below.

Fracture traces The exposed traces of the fractures are either discrete discontinuities (Fig. 19a) or zones of closely spaced discontinuities. Fracture zones vary in width from a fraction of a centimetre to a few tens of metres (Fig. 19b--d) and are made up of parallel fractures with or without a consistent sense of overlapping. Some zones are defined by en echelon tension veins (Fig. 20). The different fractures in each locality may show consistent or varied cutting/abutting relationships. This seems to depend on the presence or the absence of fillings or seals and the number of phases of deformation (see Appendices). However, on a regional scale, it is difficult to find a clearly defined, consistent cutting/abutting relationship between the regional geometrical sets. Examples of relatively well established cutting/abutting relationships between local sets of fractures are given in the next section.

Surface patterns (structures) These mainly include hackle plumes and twist hackles. The pitch of the plume axes varies from 0 to 90 ~ (Fig. 21 and Appendices). Some individual sets show both horizontal and steeply dipping plume axes (Table 4).

Chronology and dynamic implications of the mesofractures in the basement As mentioned above, it is difficult to define regionally consistent sets of fractures in the basement rocks with clear relative chronology and origin. Instead, local sets are commonly observed that show no clear evidence for origin and chronology. The fractures in the plutonic rocks are the least defined due to their complex pattern, and/or the relatively small number of data. In the Borrowdale Volcanic Group each of the two statistically dominant geometrical

134

M.S. AMEEN

Fig. 18. Synoptic pole and great circle diagrams of all the fractures in: (a & b) the Borrowdale Volcanics; (c & d) the Eskdale Granite, Granodiorite and the Ennerdale Granite (e & f) all the studied basement rocks shown in (a-d)

FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

13 5

Fig. 19. Typical examples of fracture traces in the Basement: (a) Borrowdale Volcanics; feldsparphyric andesite, at Wasdale (Water Crag, Loc. 13, NY1553 0610); (b & d) Eskdale Granite, at Hooker Crag (Loc. 17, SD1120 9830); (e) Ennerdale Granophyre, at Wasdale (Loc. 18, NY 1510 0543). Hammer shaft is 0.3 m in length and the coin is 2 cm in diameter. See text for details.

sets (Fig. 18) consists of several overlapping local sets of diverse types of fractures (e.g. dextral, sinistral, reverse and normal faults). These fractures do not show consistent local or regional dynamic implications. For example, some fractures within set 1 (Table 4) have propagated from top to bottom, or vice versa, others horizontally. Such a difference reflects differences in the orientation and/or relative amount of the causative effective principal stresses. In addition, an individual fracture zone may show evidence of

136

M.S. AMEEN

Fig. 20. A nearly-vertical fracture zone in feldspar-phyric andesite of the Borrowdale Volcanics at Wasdale (Water Crag, Loc. 13, NY1553 0610) defined by en echelon tension veins and showing evidence of multiphase deformation: (a) early phase dextral movement; (b) late phase sinistral movement; (e) a close-up view of the bottom vein in (b). The coin is 2.8 cm in diameter. For details see text.

multiphase movements. An excellent example is shown in Fig. 20 where a fracture zone with a N - S trend consists of sigmoidal tension veins with an en echelon arrangement, which indicates that the zone was originally developed as a dextral shear zone. However, a later sinistral fault parallel to, and along, the fracture zone cuts and displaces the veins (Fig. 20). Further movement phases are indicated by fibrous quartz filling along the fault.

FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

137

A minimum of four phases of fibrous growth can be recognized in the field. Earlier stages are all perpendicular to the fault walls indicating extension. The latest stage is oblique to the fault walls indicating oblique (dip-slip and dilational movement) on the fault. The fracturing history in the Borrowdale Volcanic Group is complex. Fracture initiation in the the group has most likely occurred since the emplacement of the volcanic rocks, due to cooling (e.g. columnar joints), volcano-tectonics (caldera-collapse; Branney & Soper 1988), and later tectonic movements. Thus, a purely tectonic origin for all the fractures would be an oversimplification. The release of tectonic strain would have occurred at least partly by rejuvenating the suitably oriented pre-existing discontinuities (including mineralization zones) irrespective of their origin.

The relationship of the mesofractures and macrofaults in the basement rocks No macrofaults were observed in the plutonic rocks. This is most probably related to the limited exposure and small number of localities covered in the study. Many macrofaults occur in the Borrowdale Volcanic Group (Fig. 22). These include two sets which strike NNW and ENE, respectively, marked by gouges and cataclasites. In addition, most of the macrofaults are parallel or sub-parallel to immediately adjacent mesofractures. However, the frequency of the latter does not always seem to be affected by the presence of macrofaults. Multiphase movement on individual macrofaults is evident from the occurrence of a few sets of slickenlines, as shown in Fig. 22.

Conclusions The present study shows a clear contrast between the fracture characterization of the Lower Palaeozoic basement and the Permo-Triassic cover rocks. The latter is characterized by clearly defined regional systematic sets with consistent relative chronology, whereas the former shows a complex fracture pattern with no identifiable consistent chronology and origin. Such a contrast is the result of the long and complicated history of the basement rocks compared to that of the cover rocks. The cover rocks were systematically fractured during burial, uplift and unloading, in addition to tectonic deformation. Three distinctive phases of fracture development are recognized: (1) Phase 1 (Permian to Early Jurassic) is the earliest phase during which syn-depositional growth macrofaults with northerly strikes occurred. This was mostly restricted to the lower part of the cover rocks and was accompanied by little or no mesofracturing. Therefore, this phase is not evident in any of the exposed macrofaults which are covered in this study. This phase resulted from an E-W tension which dominated NW Europe and caused the development of the East Irish Sea and the Solway Basins as extensional/transtensional basins. (2) Phase 2 (Late Jurassic to the Early Cretaceous) occurred after the deposition of the cover rocks and led to the occurrence of the easterly striking (set 2) macrofaults and mesofractures and is clearly evident from surface exposures. They resulted from burial stresses (vertical al) and tectonic stresses characterized by E-W, horizontal to moderately, westerly inclined O"1 and N-S horizontal a3. (3) Phase 3 (Palaeogene to Neogene) is the latest phase during which the exposed northerly striking (set 1) macrofaulting and mesofracturing took place. Some of these macrofaults most probably propagated upwards from the tips of the pre-existing (phase 1) northerly, macrofaults. Others resulted from the coalescence of the set 1 mesofractures and propagated downwards. Some of the movements on the macrofaults during this phase

138

M.S. AMEEN

Fig. 21. Fracture surface patterns and their interpretation in the Borrowdale Volcanics. For details see text. Long arrow indicates the direction of propagation of the main fracture, small arrows indicate the direction of propagation of twist faces. Stress directions are for the leading edge of the fracture near its origin. Hammer shaft and pencil are 0.3 m and 14cm in length, respectively.

were most probably also accommodated by the pre-existing set 2 macrofaults. The occurrence of a few slickenline sets on some of the set 2 faults may support this suggestion. This phase resulted from N - S to N W - S E compression related to the convergence of Africa and Europe which was culminated in the Oligocene to the Miocene (NNW-SSE horizontal to gently, northerly dipping O"1 and E N E - W S W horizontal era) and a regional flexural uplifting (vertical crl) associated with the onset of seafloor spreading between Scotland and Greenland. The fractographic features on both sets of mesofractures (sets 1 and 2) suggest that they occurred as Mode I (tensile) fractures and were later modified by shear or mixed-mode (tensile-shear) loading. The geometry and degree of development of fracture surface morphology are controlled by sedimentary layering, lithological heterogeneity, the

FRACTOGRAPHY AND FRACTURE CHARACTERIZATION

139

Fig. 22. Lower-hemisphere, equal-area projection of poles to macrofaults in the Borrowdale Volcanics (O), and their slickenlines (O) showing two main sets striking northerly (set 1) and easterly (set 2). Note that slickenlines 1-3 occur on the same fault (F).

presence of open fractures, the level of pore-fluid pressure, the orientation and relative amounts of the principal stresses, and their temporal and spatial variations across the fractured horizon. The conventionally accepted control of sedimentary layering on the size and shape of fractures and their fractographic features is found to be equivocal, unless the layering forms free surfaces at the time of fracturing. The normal stresses were considerably higher than the shear stresses during the development of set 2 and subset 1a fractures. However, the ratio of the normal to the shear stresses was considerably lower during the development of the subset lb fractures. This difference is probably linked to the fact that the l b fractures were the youngest systematic fractures, which could have developed at shallower depths than set 2 and l a fractures. It is difficult to recognize any consistent regional phases of deformation from the fractures in the basement. A combination of weakness zones [e.g. mineralization bands in plutonic rocks, flow surfaces, cooling shrinkage cracks (columnar joints), bedding (in tufts) and volcano-tectonic discontinuities (in volcanic rocks)] are most likely to have contributed to the wide spectrum of fracture orientations. In addition, during each phase of tectonic deformation some of the pre-existing fractures are likely to have adjusted to dissipate any suitably orientated stresses. Furthermore, new fractures were most likely initiated during ailphases of tectonism where pre-existing anisotropies were locally absent or ineffective. The most important feature to fracture characterization of the basement and the cover rocks is the abundance of dilational movements and vein development and the lack of clear evidence of pressure solution from field evidence. Although the exposed macrofaults in the cover rocks are parallel, or nearly parallel, to those occurring in the basement, they are not necessarily related. Those in the basement do not show a clear and consistent sense of slip like those in the cover rocks. The majority of faults in the basement are most probably older than those in the cover rocks (e.g. of volcano-tectonic origin), and therefore suffered a complex movement history. However,

140

M.S. AMEEN

the macrofaults exposed in the cover rocks are probably the result of direct physical propagation from those in the basement and/or they are merely controlled by the strain pattern induced into the cover rocks by movements on the basement faults (Ameen 1988, 1990, 1992). The existence of the evaporite units in the Permo-Triassic, though patchy (Chadwick et al. 1993b), may limit the propagation of macrofaults from the basement into the cover rocks and hinder the connectivity of faults in the two units in some parts of the study area. The 3D geometry and evolution of the small-scale faults and fracture zones in the cover rocks, described here, should be used to constrain any model for the evolution of macro faults. The author is grateful to Terry Engelder, Jon Gutmanis and Byron Kulander for their review and comments which helped to improve an earlier version of this paper. This paper is partly based on work carried out on behalf of UK Nirex Ltd as a part of their ongoing geological investigation into the suitability of the Sellafield area for a nuclear waste repository. Thanks are due to UK Nirex Ltd and GeoScience Ltd for their permission to publish this paper. All interpretations and conclusions presented here are solely the author's; they do not necessarily reflect the ideas of UK Nirex Ltd or Geoscience Ltd.

References AMEEN, M. S. 1988. Folding of layered cover due to dip-slip basement faulting. PhD Thesis, Imperial College, London, UK. 1990. Macrofaulting in the Purbeck-Isle of Wight monocline. Proceedings of the Geologists' Association, 101, 31-46. 1992. Strain pattern in the Purbeck-Isle of Wight monocline: A case study of folding due to dip-slip fault in the basement. In: BARTHOLOMEW,M. J., HYNDMAN,D. W., MOGK, D. W. & MASON, R. (eds) Basement tectonics 8: characterization and comparison of ancient and Mesozoic continental margins (Proceedings of the 8th International Conference on Basement Tectonics). Kluwer Academic Publishers, Dordrecht, The Netherlands, 559-578. AYDIN, A. & JOHNSON,A. M. 1978. Development of faults as zones of deformation bands and as slip surfaces in sandstone. Pure and Applied Geophysics, 116, 931-942. BEACH, A. 1975. The geometry of en echelon vein arrays. Tectonophysics, 28, 245-264. BRANNEY, M. J. & lOPER, N. J. 1988. Ordovician volcanotectonics in the English Lake District. Journal of the Geological Society, 145, 367-376. CHADWICK, R. A., BAILY,H. E., KIRBY, G. A. & ROWLEY,W. J. 1993a. Jurassic to Neogene structural evolution of the Sellafield area. UK Nirex Report No. 518, UK Nirex Ltd, UK. --, EVANS,D. J. & HOLLIDAY,D. J. 1993b. The Maryport fault: the post-Caledonian tectonic history of southern Britain in microcosm. Journal of the Geological Society, 150, 247-250. CRUIKSHANK,K. M., ZHOA,G. t~ JOHNSON,A. M. 1991. Analysis of minor fractures associated with joints and faulted joints. Journal of the Structural Geology, 13, 865-886. ENGELDER, T. & LACAZETTE, A. 1990. Natural hydraulic fracturing. In: BARTON, N. & STEPHANSSON,O. (eds) Rock Joints. A.A. Balkema, Rotterdam, 35-44. JACKSON, D. I., MULHOLLAND,P., JONES, S. M. & WARRINGTON,G. 1987. The geological framework of the East Irish Sea Basin. In : BROOKS,J. 8s GLENNIE,K. (eds) Petroleum Geology of North West Europe. Graham & Trotman, London, 191-203. JOHNSON, M. R. W., SANDERSON,D. J. & lOPER, N. J. 1979. Deformation of the Caledonides of Ireland and Scotland. In: HARRIS,A. L., HOLLAND,C. H. & LEAKE,B. E. (eds) The Caledonides of the British Isles. Geological Society, London, Special Publication, 8, 165-186. KULA~ER, B. R., BARTON,C. C. & DEAN, S. L. 1979. The application offractography to core and outcrop fracture investigation. US Department of Commerce Technical Report METC/SP-79/3, National Technical Information Service, Springfield, Virginia, USA. --, DEAN, S. L. & WARD, B. J. 1990. Fractured Core Analysis: Interpretation, Logging, and Use of Natural and Induced Fractures in Core: Methods in Exploration Series, No. 8. American Association of Petroleum Geologists, Tulsa, Oklahoma, USA. LAWN, B. R. & WlLSHAW, T. R. 1975. Fracture of Brittle Solids. Cambridge University Press, -

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POLLARD, D. O., SEGALL, P. • DELANEY, P. T. 1982. Formation and interpretation of dilatant en echelon cracks. Geological Society of America Bulletin, 93, 1291-1303. RAMSAY,J. G. & HUBER, M. I. 1987. The Techniques of Modern Structural Geology. Academic Press, London. TURNER, F. J. d~ WEISS, L. E. 1963. Structural Analysis of Metamorphic Tectonites. McGraw-Hill, New York. SMITH, R. A. 1974. A bibliography of the geology and geomorphology of Cumbria. Cumberland Geological Society, 1-32. WOODWORTH, J. E. 1896. On the fracture system of joints, with remarks on certain great fractures. Proceedings of the Boston Society of Natural History, 27, 63-183.

From Ameen, M. S. (ed.), 1995, Fractography: fracture topography as a tool in fracture mechanics and stress analysis

Geological Society Special Publication No. 92, pp. 149-174

Fracture characterization in the Chalk and the evolution of the Thanet monocline, Kent, southern England M O H A M M E D S. A M E E N

Independent Consultant, 122 Tachbrook Street, Pimlico, London S W 1 V 2ND, UK Abstract: Systematic studies of fractures in the exposed Upper Cretaceous Chalk and

from seismic data are used to establish the evolution of the regional structure and the related fractures on the Isle of Thanet, Kent. The results show that the main structure developed as an easterly trending, south facing, monoclinal forced fold with a WNWESE trending cross forced fold along a basement grain (Pegwell Bay-Minnis Bay transtension/transpression zone). The folding started during the deposition of the Chalk in the Upper Cretaceous due to an extensional tectonic phase accompanied by differential sedimentation and compaction [vertical (0.1) and horizontal (0-3)] which induced vertical and steep macrofractures (mainly joints and some extensional faults) in the Chalk. This was followed late in the Upper Cretaceous by a horizontal, N-S compression (0-1)which continued in the Tertiary and caused inversion of movement on the basement faults and the occurrence of the final stages in the forced folding. The latter phase led to the rejuvenation of the pre-existing fractures in the cover (mainly by faulting) and the development of new fractures (mostly faults and joints). An uplifting/ unloading phase started late in the deformation history and caused the occurrence of near horizontal (bedding-parallel) fractures (joints) and bedding-perpendicular or steep mesofractures parallel to, and perpendicular to, the burial/tectonic macrofractures. These have particularly developed at the top several metres of the exposed Chalk where a rapid increase in frequency of fractures is clear. The fracture characterization of the Chalk varies significantly in different parts of the monocline. Based on such variation a minimum of four domains (zones) have been identified. The heterogeneity in the fracture development is most probably controlled by the variation in orientation and type of the causative basement faults and the amount and type of displacement on t~em.

The study area covers the Isle of Thanet in Kent, an E - W trending peninsula protruding into the southern part of the N o r t h Sea. The peninsula constitutes the upthrown gentle limb and the outer anticlinal hinge zone of a south-facing monoclinal structure in the Upper Cretaceous to Tertiary sedimentary rocks, referred to here as the Thanet monocline (Fig. 1). Although the presence of the monocline is evident from a structure contour map published by the Geological Survey of Great Britain (1980), the fold was not described until 1988 when a very brief account of the structure of the country around Ramsgate and Dover was given by Shephard-Thorn (1988), who called the fold 'Thanet anticline'. There have been no detailed systematic fracture studies in the monocline. Cawsey (1977) gave a brief account of the fracture sets in Kent, including the Isle of Thanet (according to their strikes only), and considered all fractures as joints. Shephard-Thorn (1988) schematically indicated the traces of some of the macrofaults on coastal cliff sections of the Upper Chalk, and called them tension faults associated with the 'Thanet anticline'; however, neither their orientation nor their chronology or origin was mentioned.

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AMEEN

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152

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The present study discusses the evolution of the Thanet monocline and the associated fracture characterization in the Upper Chalk. The term 'fracture' is used here in the broad sense defined by Kulander et al. (1990) as a break or physical discontinuity in a rock caused by stresses exceeding the rock's strength. Therefore, it covers faults, joints, veins and stylolites. According to their scale the fractures covered in the present study include mesofractures and macrofractures. Mesofractures can be effectively observed in 3D with or without a hand lens in hand specimens or large continuous exposures. Macrofractures are too large or too poorly exposed to be examined directly in their entirety in one exposure. Fractography is defined here as the discipline of diagnosing, describing and interpreting the fracture surface morphology or topography (fractographic features), which are related to the initiation, propagation and subsequent history of the fractures. Details concerning the definitions of fractures and fractographic features (e.g. origins, hackle plumes, hackle faces, hackle steps and rib marks) and their interpretation and methods used in this study are the same as those outlined in Ameen 1995. A total of 431 fractures (faults and joints) was studied in 11 field localities, mostly on the coastal parts of the Isle of Thanet. In addition, an offshore seismic section is interpreted and used to study the profile of the regional structure and understand its origin and development, in particular the basement/cover interaction.

Regional structure The Thanet monocline The Thanet monocline is a relatively gentle monoclinal fold, facing south to southwest. The steepest exposed part of the monocline (the outer anticlinal hinge zone) occurs in Pegwell Bay and dips at 10~ to the south-southwest. The fold has a relatively small amplitude (c. 50 m, measured on the base of the Chalk) and its axial trace strikes at an E-W direction along most of its length onshore; however, it swings to a N W - S E direction at Pegwell Bay (Fig. 1a, b & c). The onshore length of the monocline is c. 20 km; however, the structure is expected to continue offshore for a minimum of a few kilometres to the southeast of Pegwell Bay. The fold affects the relatively thin sedimentary cover (a few hundred metres to 1 km in thickness) which is dominated by Upper Cretaceous and Tertiary rocks (Fig. l a, c & d). The underlying basement rocks are of Lower Palaeozoic and Precambrian age, and form the southern part of the London-Brabant Massif (defined below). The monoclinal flexure evolved as a result of the differential compaction 'draping' of the Upper Cretaceous to Tertiary sediments over the topographic high of the LondonBrabant Massif and the late Cretaceous to Pliocene basement-block faulting. The origin of the monocline will be discussed further in a later section.

The London-Brabant Massif This covers a wedge-shaped area with an E - W to N W - S E trend, and extends from east England (where it has its maximum width of nearly 200 km) across the southernmost part of the North Sea into Belgium, where it is considerably narrower. The massif has a length of c. 550 km. It is bounded by faults which separate it from surrounding Mesozoic basins (Fig. 2). Geophysical evidence suggests the presence of Precambrian cratonic rocks beneath the London-Brabant Massif. The massif, which received some Lower Palaeozoic

CHALK FRACTURE CHARACTERIZATION, THANET

153

Fig. 2. Regional tectonic framework of southern England (based on Hamblin et al. 1992 and Cameron et al. 1992) showing the location of the London-Brabant Massif.

sediments, became established as a relatively stable upland area early in the Dinantian and formed the southern boundary of the southern North Sea Basin and the northern boundary of the Weald Basin until mid-Cretaceous times. The southern boundary of the massif, which is close to the study area, is defined by a Variscan thrust and related transpression zones (Variscan Front in Fig. 2). The massif received very little and patchy sediments during the Upper Palaeozoic to mid-Cretaceous. This is reflected by the shallow depth to the top of the Lower Palaeozoic (c. 1 km), and to the top of Carboniferous rocks (Cameron et al. 1992, figs 11 & 13). In the Cretaceous the London-Brabant Massif was submerged for the first time since the early Palaeozoic. This led to the deposition of the Wealden, Greensand, Gault and the Chalk Group (Lower, Middle and Upper Chalk). However, the thickness of these rocks over the London-Brabant Massif is considerably thinner than that in the North Sea Basin and the Weald Basin (Cameron et al. 1992, fig. 82). The submergence of the massif is probably due to a combination of a gentle (thermal?) crustal downwarping and a contemporary global rise in sea level, and was accompanied by basement-block faulting. During the Tertiary both shallow marine and continental deposits accumulated over the

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M.S. AMEEN

Fig. 3. (a) A map outlining the 4 different domains (zones) of the Thanet Monocline and showing the distribution of the studied localities (a-i); ~) Lower-hemisphere, equal-area stereographic projection of the poles to the fractures and the slickenlines in the different localities (a-i).

London-Brabant Massif. The contact between the Upper Cretaceous Chalk and the Tertiary rocks is marked by a time gap and erosional surface (Fig. l d). It is arguable whether the surface is the result of subaerial or shallow marine erosion.

Fracture characterization of the Upper Chalk The fracture patterns in the exposed parts of the Thanet monocline can be divided (according to their host structural habitat) into two main groups, those in the upthrown nearly horizontal limb of the monocline and those in the outer anticlinal hinge zone of the monocline.

Fractures in the upthrown limb of the monocline Three different zones (domains) have been recognized in this part of the monocline, according to their fracture patterns. They occur in the northeastern, northwestern and western parts of the upthrown limb of the monocline, respectively, referred to here as zones 1, 2 and 3 (Fig. 3), respectively, and are described below.

Fractures in the northeastern part of the limb (zone 1). These fractures occur in the area east of Margate (Walpole Bay, Palm Bay, Botany Bay and Kingsgate Bay). They include two sets, which are nearly orthogonal, and dip steeply (>~ 70 ~ and strike at ESE to E - W and NNE to N-S respectively. The former set is the most dominant. The fractures are mostly joints - faults are rare and include extensional faults only (Fig. 3). The joints include two distinctive groups according to their dimensions, fractography and their relationship to the bedding. These are referred to as the macro- and mesojoints. The macrojoints cut across the coastal cliff faces from top to bottom and continue across bedding parallel joints, except some hard ground and marl horizons (Fig. 4). They have a

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155

156

M.S. AMEEN

Fig. 4. Easterly striking macrojoints affecting the nearly-horizontal beds exposed in Walpole Bay (a & b) and Kingsgate Bay (c & d) and a northerly striking macrojoint exposed in Palm Bay (e). For location see Fig. 3.

CHALK FRACTURE CHARACTERIZATION,THANET

157

Fig. 5. Fractographic features on: (aid) easterly striking macrojoints, indicating dominantly horizontal, overall propagation (a--e) and less common vertical propagation (d); and (e) on a northerly striking macrojoint indicating vertical, top to bottom propagation direction (a, in Walpole Bay, b-e, Botany Bay). The 3D interpretation of the fracture shown in (c) is illustrated in the line drawing to the right. For location see Fig. 3. Long arrow indicates direction of propagation of the leading edge of the joints. Small arrows are local direction of propagation. Stresses are for the leading edge of the joint near its origin.

158

M.S. AMEEN

minimum height of several to a few tens of metres (about cliff height) and trace lengths, on bedding surfaces, of tens of metres (up to 50 m). They are mostly tight or slightly open (up to a few millimetres). Surface structures of these joints are characterized by origins, hackle plumes, hackle faces, hackle steps and rib marks. Rib marks are rare and twist angles are very small (_<5~ The surface features indicate an overall horizontal direction of propagation of most of the easterly striking joints, some show evidence of vertical propagation from top to bottom (Fig. 5). The main differences between the easterly and northerly trending macrojoints are the apparent greater frequency of the former (see Fig. 3 and cf. Fig. 4b & d with e), and the tendency of the northerly set to abut against the easterly set and to propagate from top to bottom (Fig. 5d). The mesojoints of both N-S and E-W striking sets are constrained by adjacent beddingparallel joints and tend to originate at such joints. In addition, the N-S joints are also constrained by the E-W set. Hackle plumes and rib marks on the mesojoints indicate horizontal or vertical propagation from top to bottom or vice versa (Fig. 6).

Fractures & the northwestern part of the limb (zone 2). These occur in the areas extending from Margate westwards to Minnis Bay (Westgate Bay and Epple Bay) and include two systematic, regionally persistent, vertical or steeply dipping sets, striking NNW-SSE to NW-SE and N-S, respectively, and a third non-systematic set of fractures which show flint fillings (Fig. 7a). The northwesterly set is the most prominent in this zone (Fig. 3). As for the eastern part of the limb, the systematic fractures are dominated by macroand mesojoints with relatively rare extensional faults. The macrojoints occur as discrete discontinuities (Fig. 7b & c), or discontinuity zones (Fig 7d) which cut across the beddingparallel joints (except some hardgrounds and marl horizons) from top to bottom of the cliff faces. Most of the fractures are tight or slightly open (_< 1 mm) with no filling; however, iron oxide staining is common on the NNW-SSE to NW-SE set. (Fig. 7c). The exact dimension of each joint is difficult to determine due to the limited exposure; however, the exposed lengths of the macrofractures is up to 20 m and their heights are equal to or greater than those of the cliff faces, i.e. several to tens of metres (Fig 7c). Fractographic features on these surfaces indicate that they mostly originate within a bed in the middle or upper part, at a void, fossil or flint nodule. Hackle plumes, twist hackles and gradual twisthackle fringes occur and indicate an overall horizontal propagation of most joints. The faults are dip-slip, extensional, and belong to the northwesterly striking set. Early diagenetic, compaction joints and faults are mostly flint-filled and non-systematic, i.e. vary in attitude (e.g. Westgate Bay, East, Fig. 3). The mesojoints are parallel to the macrofractures and show similar fractographic features and fracture/bedding relationship to those for the mesofractures in the northeastern part of the limb (zone 1, described above). Fracture patterns in the southeastern part of the limb (zone 3). This zone covers the part east of the Pegwell-Minnis Bay Transpression Zone and south of Joss Bay and is characterized by steeply dipping fractures, mostly macrofaults, striking NNE-SSW to NNW-SSE, E-W to ESE-WNW and NW-SE, respectively (Fig. 3). The northerly fractures are very rare. The fractures in this zone are well exposed on the coastal cliffs between Joss Bay and Ramsgate, particularly in the Broadstairs area (North Cliff and Dumpton Gap to East Cliff). The dimensions of the macrofractures are similar to those in other parts of the monocline (Fig. 8).

CHALK

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Fig. 7. Macrofractures exposed in zone 2. (a) Flint-filled, early diagenetic fault zone (F), Westgate Bay. (b) Traces of N-S (f 1) and N W - S E (f2) fractures on bedding-parallel wave cut platform, Epple Bay. (e) Steeply dipping N W - S E macrojoints (J) stained with iron oxide, Westgate Bay. (d) A N W SE fracture zone (f) dipping at 80 ~ towards 230 ~ defined by smaller fractures (sf) dipping at 60-70 ~ towards 32 ~ Epple Bay. For locations of (a)-(d) see Fig. 3.

CHALK FRACTURE CHARACTERIZATION, THANET

161

The faults are mostly strike-slip, although some show multi-phase movements with strike-slip, dip-slip and oblique-slip components as evident from the occurrence of a few slickenline sets on some faults (Figs 8 & 9). The slip indicators are pressure solution and frictional wear grooves. It is difficult to analyse the slip history, i.e. chronology and amount of each movement phase. However, it is possible to determine the net-slip on the faults which mostly have an extensional component. Fault rocks (cataclasites and gouges) have evolved to various degrees on the faults and tend to vary in thickness along individual faults (Fig. 10a & b). The cross-cutting relationship indicates that the N W - S E set were the latest active faults which, together with the N - S fractures, terminate against the older E - W fractures. Fractographic features in this zone are not clear, they have mostly been obliterated by frictional wear due to shear movement on the majority of the fractures (Fig. 10c).

Fractures in the outer anticlinal hinge of the monocline (zone 4). This part of the monocline is exposed at Pegwell Bay and inland in Monkton Quarry. The former exhibits a spectacular group of steeply dipping fractures dominated by joint zones and multi-phase faults. These include four sets, striking N-S, N W - S E , E - W and N E - S W , respectively. The former two dominate the fracture pattern. In M o n k t o n Quarry, which is relatively overgrown, the N E - S W and E - W fractures do not occur (Fig. 3). The two dominant sets occur as macro- and mesofractures, the former cut across bedding-parallel fractures (except some hard grounds and marl horizons) and extend from top to bottom of the cliff faces. The exposed height of most of the macrofractures is equal to or greater than that of the cliff face (i.e. several to tens of metres) and their exposed length is equal to or greater than a few to several tens of metres (Fig. 11 a-d). The northerly striking fractures occur mostly as discrete faults or narrow fault zones up to 1 m in width. The northwesterly striking fractures include both faults and joints, and occur mostly as zones up to several metres in width (Fig. 1 l a & c). The northwesterly fractures tend to hook on to the northerly striking fracture (Fig. lid). Macrofaults of both sets show anastomosing, wavy, stylolitized traces on the bedding-parallel, wave-cut platform (Fig. 1 le). Slickenlines are common on fault surfaces of both sets and are particularly well preserved inside coastal caves associated with differential weathering and erosion along the fracture zones. Up to four sets of slickenlines occur on some faults and it is not always possible to draw a firm conclusion as to their overprinting relationship, and thence their relative age (Fig. 12). In addition, it is impossible to separate the complex movement history on individual faults (which show multiphase movements) into its increments. However, it is possible to determine the net (resultant) slip on each fault from the orientation of the fault and the offset of key horizons, particularly flint bands. These are consistently oblique-slip with extensional component (i.e. normal--dextral or normalsinistral). Such an extensional component to the net-slip is spectacularly displayed on the nearly E - W trending cliff face where the Whitaker's 3 inch flint band steps downward progressively from the eastern end to the western end of the bay along the traces of the macrofaults with a net vertical offset of c. 10 m over a distance of 870 m. Various degrees of cataclasis occur along the faults; these are evident from w e d g e shaped bodies of cataclasites and/or gouges which change in thickness drastically along the exposed faults length and are similar to those shown in Fig. 10a & b. Fractographic features (particularly hackle plumes) on the macrofractures of both major sets are obliterated by the shear movements on them. However, the profile sections

162

M.S. AMEEN

CHALK FRACTURECHARACTERIZATION,THANET

163

of some fractures show a breakup of the main fracture in a fashion characteristic of major twist hackles, similar to that shown in Fig. 5c, indicating that the fracture originates at the top of the fractured horizon with an overall horizontal propagation, individual segments (twist faces) have propagated mainly vertically from top to bottom. There is a wide variety of hackle plumes and rib marks on the mesojoints of all sets. They indicate cyclic or steady growth of the joints, evident from the occurrence or the lack of rib marks, respectively. Most of the joints originate at flint nodules, fossils or voids within the beds or at their borders. The use of blasting prior to concrete injection into some of the coastal caves (to stabilize the cliffs) might have inevitably induced some joints with clear hackle plumes and rib marks. Therefore, careful consideration should be given to the analysis of fractographic features on the joints. The presence or lack of iron oxide staining (common on the tectonic joints in the area), the closeness of the joints to the blast locations and the local variation in their frequency are used as guidance for recognizing tectonic from blast-induced joints.

Origin of the Thanet monocline and related fractures Origin of the Thanet monocline In order to understand the fractures, it is essential to throw some light on their host regional structure, i.e. the Thanet monocline, which was described in the section on the regional structure above. It is possible to evoke several hypotheses for the origin of the Thanet monocline, these are: (1) Halokinesis (salt diaperism) is very unlikely due to the absence of the Upper Permian Salts underneath the monocline; such salts caused halokinetic activity in the North Sea Basin, to the north of the London-Brabant Massif; (2) buckle folding requires the detachment of the mainly Upper Cretaceous sedimentary cover from the early Palaeozoic and Precambrian basement - both stratigraphic and geophysical evidence discredit this hypothesis (Fig. 13); (3) drape or forced folding, this includes non-tectonic and tectonic forced folding. The non-tectonicforcedfolcling can be the result of differential compaction of sediments over rigid topographic highs, or the result of the growth of igneous intrusions, particularly laccoliths and sills. The latter is ruled out due to the lack of evidence. However, the former is apparent from stratigraphic and geophysical data. These reveal that the sedimentary cover (mainly Upper Cretaceous Chalk Group) is thinner over the London-Brabant Massif compared to the Weald Basin and the southern North Sea Basin, to the south and the north respectively (Cameron et al. 1992, fig. 82). Locally, in the Thanet monocline, the

Fig. 8. Macrofractures affecting horizontal beds exposed on the cliff faces in zone 3. (a) An older easterly striking fault zone (F) dipping at 73~ towards the north abutted by later NW-SE joints (J). Note that both sets are stained by iron oxide (North Cliff). (b) A NW-SE striking fault (F) dipping at 85~ towards the southwest with three sets of slickenlines which pitch at 20~ NW, 20~ SE and 90~ SW. The fault is stained by iron oxide (South Cliff). (c) A NW-SE striking fracture zone (F) dipping at 77~ towards SE. The zone is partly defined by en echelon fractures (EF) dipping at 85~ to the NE (South Cliff). Notebook is 21 cm in length. (d) Northerly striking faults (F1) dipping at 60~ to the east with an extensional slip (South Cliff) evident also from frictional-wear slickenlines. Note that the NW-SE fractures (F2) abut against F1 (South Cliff). (e) A NW-SE striking macrofault (F) dipping at 52~ to the southeast with c. 2 m dip-slip component (South Cliff). For locations see Fig. 3.

164

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CHALK FRACTURE CHARACTERIZATION, THANET ~

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Fig. 11. (a) A NW-striking nearly-vertical fracture zone (> 7 m in width) consisting of closely spaced fractures (Pegwell Bay). (b) A N-striking fault at Pegwell Bay dipping at 72 ~ to the east with dextral-normal and later sinistralnormal slips. The estimated net dip-slip is 0.21 m. (e, d & e) The traces of: (c) a northwesterly fracture zone; (d) northwesterly fractures (F1) hooking on to the older northerly fractures (F2). (d) A northerly striking fault, shown in (b), the dark colour of the trace is due to pressure solution seams on the fault, all exposed on a beddingparallel wave-cut platform, Pegwell Bay. See Fig. 3 for location.

CHALK FRACTURE CHARACTERIZATION, THANET

167

Fig. 12. Slickenlines (frictional wear/pressure solution grooves), parallel to arrows, on macrofault surfaces in zone 4. (a) Horizontal slickenlines on a northerly striking sinistral fault dipping at 85 ~ to the east, Pegwell Bay. (b) Four sets of slickenlines on a northerly striking fault dipping at 78 ~ to the west, Pegwell Bay. (c) Gently dipping slickenlines on a dextral-normal, northwesterly striking fault in Monkton Quarry. Note the iron oxide staining of the surface and the relatively feeble development of the slickenlines, compared to those in (a) & (b). For location see Fig. 3.

168

M.S. AMEEN

cover rock (mainly Upper Cretaceous Chalk) is considerably thinner in the upthrown limb compared with the external, anticlinal hinge zone and the downthrown limb, respectively. Furthermore, the present study discovered local differences in thickness within the upthrown limb. Both variations are controlled by extensional and transtensional fault escarpments in the basement which can be seen on an offshore seismic section of the monocline interpreted in this study (Fig. 13). The tectonic forced folding is caused by the draping of sedimentary layers over a growing fault escarpment in the basement. The fault can be dip-slip (normal or reverse), obliqueslip, positive or negative flower structures (i.e. transpression or transtension zones). The strain patterns and hence the fracture patterns in the cover rocks, are strongly affected by the type, angle of dip and amount of displacement of the basement fault/fault zones (Ameen 1988, 1990). Such a type of folding has affected the evolution of the Thanet monocline, which seems to be underlined and controlled by at least two groups of faults/fault zones striking E-W and WNW-SSE, respectively. The former are mainly dip-slip and control the dominantly it E-W trend of the monocline, west of Pegwell Bay (Fig. 1a,b). The second group forms a transpressional fault zone (Pegwell-Minnis Bay Zone) detected in the present study from seismic sections (Fig. 13), and is expressed in the cover rocks as deflections of the otherwise E-W trending structure contours to a WNW direction along the zone (Fig. l b), which defines an upward bulge in these rocks to accommodate the flower structure in the basement. This is also supported by the change in the fracture patterns from zone 1, which is far from the transpression zone, to those in the immediate vicinity (zones 2 and 3) and those within the transpression zone (zone 4), respectively (Fig. 3). Both groups of basement faults were most probably active during the Cretaceous and the Tertiary N-S compression. Such faults would have affected the evolution of the monocline in two different, though related, ways; differential deposition and compaction of sediments on existing fault escarpments (non-tectonic forced folding) and the draping of the sedimentary layers over active faults (tectonic forced folding). The evolution of the forced fold occurred in two main phases; an early extensional phase and a late compressional phase. The extensional phase included non-tectonic and tectonic forced folding and took place during the deposition of the Chalk where a transtensional movement on the Pegwell-Minnis Bay Zone resulted in thicker accumulation of the Chalk in the zone. Similarly, extensional movements on the easterly striking faults caused thicker Chalk to be deposited on the downthrown side of the faults (Fig. 13). The second (compressional) phase started late in the Cretaceous and caused inversion of the movement on the basement faults and was accompanied by mainly tectonic forced folding (including cross-forced folding above the transpressed Pegwell-Minnis Bay zone), uplifting and erosion evident from the angular erosional unconformity between the Chalk and the Tertiary rocks. Further compression took place during the Tertiary when the monocline evolved to its present shape.

Origin o f the fractures The fractures in the Upper Chalk of the Thanet monocline resulted from a few different phases, at least some of which are related to the evolution of the fold.

Phase 1 - burial compaction. Unlike limestone, chalk does not usually lithify during early diagenesis at, or very close to the surface, therefore early jointing occurs mainly after a

169

CHALK FRACTURE CHARACTERIZATION, THANET

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considerable burial has taken place. Burial stress regime is characterized by vertical maximum principal stress (or1) and horizontal minimum principal stress (or3). Under such circumstances the chalk will tend to consolidate, by expulsion of pore water, and lithify by pressure solution and cementation (Jones et al. 1984). However, the intact chalk has an exceptionally low permeability (13.9-37) • 10-9m s-1 for a rock with high porosity; 29.645.6 (Bell 1977; Bell et al. 1990). Therefore, a buildup of pore-water pressure with burialinduced consolidation and lithification is expected to facilitate the initiation and propagation of fractures. The orientation of the effective least principal stress is affected by any shallow, buried topographic features underneath the Chalk. On the Isle of Thanet effective least principal stress is expected to be in a direction perpendicular to the local trend (and parallel to the maximum slope) of the local topographic highs in the basement and the associated boundary faults. Therefore, this resulted in the development of vertical or steep macrofractures (joints and normal faults) striking easterly (zone 1), northwesterly (zone 2), northerly and easterly (zone 3) and N-S (zone 4). Some fractures propagated rhythmically others propagated smoothly with little hesitation. This is evident from the occurrence of rib marks and rhythmic hackle plumes on the former and their lack on the latter. Stress release after the development of these fractures led to a change in the relative amounts of the principal stresses and the development of vertical or steep macrofractures (mostly joints with some normal faults) orthogonal to the first set. These late occurring burial fractures are generally rare, form a few to several per cent of the total number of fractures, except in zone 3 (Fig. 3), and are constrained between the older burial fractures. The majority of the macrofractures show fractographic features indicative of a horizontal, bedding-parallel propagation. This resulted from local reorientation of the vertical maximum effective remote stresses to horizontal due to the existence of horizontal bedding-parallel lithological discontinuities. The Upper Chalk was deposited as a relatively homogenous lithological unit which lacks real bedding discontinuities, except for some widely spaced hardgrounds, thin marl zones and flint bands, which acted as 'stress deflectors'. Before the onset of uplifting there were no closely spaced horizontal (bedding) joints (see below) to constrain the burial phase fractures to smaller dimensions. This is evident from the fact that the majority of the macrofractures and their fractographic features continue across existing bedding discontinuities (e.g. Fig. 5) but tend to abut against hardgrounds and marl horizons. The joints which occurred during this phase and the consequent phases (described below) originated as tensile (Mode I) fractures with the local (net or effective) minimum principal stress (cr3 eft) oriented at right angles to the main fracture surface. The local (net or effective) maximum and intermediate principal stresses (crl eft and ~r2 eft) lie within the fracture surface. At the leading edge of each individual fracture ~rl eft and or2 eft are parallel and perpendicular to the direction of propagation of the fracture, respectively (Figs 5, 6 & 10c). However, these and consequent faults resulted from the coalescence and shear or mixed-mode (tensile-shear) loading of microscopic and/or mesoscopic tensile fractures and/or shear or mixed-mode loading of macrojoints. P h a s e 2 - tectonic p h a s e . This includes an early extensional (Upper Cretaceous) and a late compressional stage (late Upper Cretaceous to Pliocene) associated with the development of the monocline. The extensional period was caused by a vertical maximum principal stress (Crl) and horizontal minimum principal stress (or3). This is similar to, and difficult to separate (in terms of the resulting fractures) from, the burial phase. The two were acting penecontemporaneously as discussed in the section on the origin of the Thanet monocline

CHALK FRACTURE CHARACTERIZATION, THANET

171

(above). Therefore, many of the macrofractures developed in the burial phase are also most probably linked to the extensional tectonic stage. The second (compressional) stage was associated with the Alpine lateral N-S compression (Crl) which affected the Weald Basin to the immediate south of the Thanet monocline. Such a N-S compression is expected to have reactivated the suitably oriented fractures (both in the basement and the sedimentary cover) and initiated new sets of fractures, where necessary, to take up the strain. The reactivation of fractures is evident from extensive faulting in zones 3 and 4, which was accompanied by the development of new northwesterly striking fractures (mostly faults). These fractures displaced or abut the older burial/extensional phase fractures. The majority of the northwesterly striking faults are strike-slip or oblique-slip faults, with an extensional movement component (Fig. 3). The existence of such faults is also evident from seismic sections a few kilometres offshore. The section, interpreted in the present work, indicates the existence of a major fault zone in the basement along the Pegwell-Minnis Bay grain shown on Fig. lb. The sense of slip and geometry of the zone are characteristic of a convergent strike slip-fault (positive flower structure) indicative of transpression (Fig. 13). Such a major fault zone is expected to cause the development of a suite of fractures and the rejuvenation of pre-existing fractures to accommodate the associated strain pattern. Transpression faults (Harland 1971) have a component of horizontal shortening across the strike-slip fault zones which is compensated by cataclastic deformation and uplift of the rock wedge enclosed in the fault zone (Lowell 1972; Wilcox et al. 1973). The sedimentary cover above such a bulging basement wedge is expected to act passively to accommodate itself to the uplift of the wedge and the strike-slip movement of the main bounding blocks (Fig. 14). Such an accommodation in the Thanet monocline has taken place by gentle, cross-forced folding along the WNW-SSE basement grain, outlined on Fig. lb and also evident from the pattern of exposure of the Tertiary and Cretaceous rocks and the structure contour map of the base of the Chalk (Fig. l b). The amplitude of this flexure decreases dramatically from a few tens of metres near Pegwell Bay to several metres near Minnis Bay. Such a decrease is noticeably accompanied by an equally significant decrease in the intensity of faulting in the exposed Upper Chalk (Fig. 3). Therefore, the causative faulting in the basement seems to diminish significantly northwestwards. The fracture patterns associated with this phase is expected to include a combination of those typical of a strike-slip zone and those associated with mild drape folding (Fig. 14), i.e. a combination of strike-slip, oblique-slip and normal faults (and joints parallel to them). The complete suite of fractures shown in Fig. 14 do not always occur, especially when pre-existing fractures exist which are orientated suitably to take up the strain, as is the case on the Isle of Thanet. Phase 3 - uplifting~unloading phase. This phase started late in the deformation history and was accommodated by the occurrence of near-horizontal (bedding-parallel) fractures (joints) and bedding-perpendicular or steep mesofractures parallel to, and perpendicular to, the burial/tectonic macrofractures described above. These have developed particularly at the top several metres of the exposed Chalk where a rapid increase in frequency of fractures is clear. It is difficult to assess the effect of the Quaternary, relatively brief, glacial burial which took place during the final stages of the tectonic uplifting. Such a burial is expected to be taken up by movements on pre-existing fractures. Detailed studies, including fault rocks dating, may help in recognizing neotectonic/glaciation related faulting.

172

M.S. AMEEN

Fig. 14. A schematic block diagram illustrating the potential fracture patterns in a drape fold associated with a transpression fault zone (positive flower structure) in the basement. For details see text.

Conclusions and discussion The Thanet monocline and related fractures resulted from a complex combination of nontectonic and tectonic (extensional and compressional) forced folding largely controlled by dip-slip and oblique-slip faulting in the Precambrian to Lower Palaeozoic basement. The non-tectonic and extensional-tectonic forced folding occurred throughout most of the Upper Cretaceous and was characterized by a stress regime with a vertical maximum principal stress (trl) and a horizontal minimum principal stress (a3). The largely Chalk cover rocks responded to these stresses by vertical and steep macrofractures (mostly joints and some extensional faults). The orientation of the fractures vary in the different parts of the monocline depending on variation in the direction of or3, which vary with the type and orientation, of the underlying fault escarpment in the basement. Accordingly, a minimum of four fracture domains (zones) have been recognized in the monocline. The tectonic forced folding and fracturing of the cover rocks continued during the compressional phase which started late in the Upper Cretaceous and continued during the Tertiary. This phase was characterized by a N-S compression (trl) which inverted the movement on the basement faults and rejuvenated the macrofractures in the cover rocks (mostly by faulting) and initiated new fractures (faults and joints). The joints which occurred during the different phases of folding and fracturing originated as tensile (Mode I) fractures with the local (net or effective) minimum principal

CHALK FRACTURE CHARACTERIZATION, THANET

173

stress (cr3 eft) perpendicular to the main fracture surface. The local (net or effective) maximum and intermediate principal stresses (Ol eft and cr2 eft) lie within the fracture surface. At the leading edge of each individual fracture crl eft and cr2 eft are parallel and perpendicular to the direction of propagation of the fracture, respectively. However, the faults resulted from the coalescence and shear or mixed-mode (tensile-shear) loading of microscopic and/or mesoscopic tensile fractures and/or shear or mixed-mode loading of macro joints. An uplifting/unloading phase started late in the deformation history and was accommodated by the occurrence of near-horizontal (bedding-parallel) fractures (joints) and bedding-perpendicular or steep mesofractures parallel to, and perpendicular to, the burial/tectonic macrofractures described above. These have developed particularly at the top several metres of the exposed Chalk where a rapid increase in frequency of fractures is clear. It is difficult to assess the effect of the Quaternary, relatively brief, glacial burial which took place during the final stages of the tectonic uplifting. Such a burial is expected to be taken up by movements on pre-existing fractures. Detailed studies including fault rocks dating may help in recognizing neotectonic/glaciation related faulting. In addition, some joints are induced by blasting during remedial measures to protect the coastal cliff from sea erosion. Other joints are probably related to, or opened by, episodes of sudden subaerial erosion and sudden cliff retreats, common during the winter, and tend to coincide with and follow periods of maxima of air frost and water surplus (rain). The present results can be used as a guidance to review faults and folds in the immediate vicinity of the Thanet monocline. The dominantly N W - S E to W N W - S S E striking faults in the Kent Coalfield of the Richborough syncline, to the south of the study area, and in the English Channel (described by Shephard-Thorn et al. 1972) are most probably related to flower structures similar to the Pegwell-Minnis Bay transtension/transpression zone in the Precambrian-Lower Palaeozoic basement rather than strike-slip or dip-slip faults. In addition, the assumption that joints in the cover rocks, including the Chalk, in other parts of east Kent are propagated from, or related to, basement faults of the same strike (Middlemiss 1983) is an oversimplification. As the current study shows, the fractures are caused by movements on basement faults and/or differential burial compaction over basement fault escarpments. The latter and dip-slip faulting in the basement result in fractures (faults and joints) in the cover rocks striking parallel to the causative fault in the basement. However, strike-slip and oblique-slip (transpression and transtension) faults result in fractures in the cover rocks striking at a range of angles to the causative fault/ fault zone in the basement. The present studies show clearly the occurrence of extensional/transtensional movements on the basement faults during the deposition of the Upper Cretaceous rocks, including the Chalk. This raises serious doubts about the commonly held view that the extensional movement in southern England ceased before the onset of the deposition of the Chalk (e.g. Stoneley 1982) and hence the assumption that the Chalk thickness does not vary drastically across basement faults. Recent surface and subsurface observations support the occurrence of extensional movement on basement faults as evident from significant and abrupt difference in the Chalk thickness across the basement faults in the Dorset and the Isle of Wight area (Andrew Gale, pers. comm.). The author is grateful to Dr J. Cosgrove for reviewing the paper, Dr E. R. Shephard-Thorn for reviewing the manuscript and his helpful comments, Dr Brian D'olier for supplying the raw seismic data, and Dr Andrew Gale for the helpful discussions during the course of this work.

174

M.S. AMEEN

References AMEEN, M. S. 1988. Folding of layered cover due to dip-slip basement faulting. PhD Thesis, Imperial College, London, UK. 1990. Macrofaulting in the Purbeck-Isle of Wight Monocline. Proceedings of the Geologists' Association, 101 (1), 31-46. 1995. Fractography and fracture characterization in the Permo-Triassic sandstones and the Lower Palaeozoic Basement, West Cumbria, England. This volume. BELL, F. G. 1977. A note on the physical properties of the Chalk. Engineering Geology, 11, 217-115 , CRIPPS, J. C., EDMONDS, C. N. & CULSHAW, M. G. 1990. Chalk fabrics and its relation to certain geotechnical properties. In: Chalk: Proceedings of the International Chalk Symposium. Thomas Telford, London, 187-194. CAMERON,T. n . J., CROSBY,A., BALSON,P. S., JEFFERY,D. J., LoTr, J. K., BULAT D. J. & HARRISON, D. J. 1992. The geology of the southern North Sea, British Geological Survey, UK Offshore Regional Report. HMSO, London. CAWSEY, D. C. 1977. The measurement of fracture patterns in the Chalk of southern England. Engineering Geology, 11, 201-215. GEOLOGICAL SURVEYOF GREAT BRITAIN. 1980. Ramsgate Sheet, 1:50,000. HMSO, London. HAMBLIN,R. J. E., CROSBY,A., BALSON,P. S., JONES,S. M., CHADWICK,R. A., PENN,I. E. & ARTHUR, M. J. 1992. The geology of the English Channel, British Geological Survey, UK Offshore Regional Report. HMSO, London. HANCOCK, J. M. 1976. The petrology of the Chalk. Proceedings of the Geologists" Association, 86, 499-535. HARLAND,W. I . 1971. Tectonic transpression in Caledonian Spitsbergen. Geological Magazine, 108, 27-42. JONES, M. F., BEDFORD, J. & CLAYTON, C. 1984. On natural deformation mechanisms in the Chalk. Journal of the Geological Society, London, 141, 675-676. LOWELL,J. D. 1972. Spitsbergen Tertiary orogenic belt and Spitsbergen fracture zone. Geological Society of America Bulletin, 83, 3091-3102. MIDDLESMISS,F. A. 1983. Instability of Chalk cliffs between the South Foreland and Kingsdown, Kent, in relation to geological structure. Proceedings of the Geologists" Association, 94, 115-122. SHEPHARD-THORN,E. R. 1988. Geology of the country around Ramsgate and Dover, British Geological Survey Memoir for 1:50 000 geological sheets 274 and 290, England and Wales. HMSO, London. , LAKE, R. D. & ATITULLAH,E. A. 1972. Basement control of structures in the Mesozoic rocks in the Strait of Dover region and its reflexion in certain features of the present land and submarine topography. Philosophical Transactions of the Royal Society, London, A272, 99-113. STONELEY, R. 1982. The structural development of the Wessex Basin. Journal of the Geological Society, London, 131, 543-554. WILCOX, R. E., HARDING, T. P. • SEELY, D. R. 1973. Basic wrench tectonic. American Association of Petroleum Geologists, Bulletin, 57, 74-96. -

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From Ameen, M. S. (ed.), 1995, Fractography: fracture topography as a tool in fracture mechanics and stress analysis

Geological Society Special Publication No. 92, pp. 175-186

Fracture surface markings in Liassic limestone at Lavernock Point, South Wales JOHN

C. R O B E R T S

University of Ulster, Coleraine Campus, Coleraine, Northern Ireland, BT52 1SA, UK

Abstract: A study of fracture surface markings (FSM) on three joint sets, striking at 190,

335 and 280~ cutting a 2.5 m thick slice of interbedded limestones and mudstones of lower Liassic age, yields evidence which enables determination of the orientation of the palaeostress field at the time of joint formation. A specific element of the FSM, namely the border planes, or the twist hackle, has been used to determine the orientation of a~, in that these twist-hackle structures are considered to be developed normal to the remote tensile stress ~r3.In the case of the joint sets striking at 335 and 190~ the sense of rotation of the twist plane would suggest that for both sets Crl was aligned N-S and almost parallel to bedding. Other FSM features suggest that the 190~ set propagated rapidly with S-type plumose markings being the most common. The small angle of border plane rotation (10-15 ~) suggests that all of the joint sets propagated under relatively low confining pressures. The sense of border plane rotation on the 335 and 190~ sets suggests that there is a possible conjugate relationship between them. However, the other main joint set which strikes at 280~ cannot be accounted for by simultaneous development in a N-S aligned ~rl which might have caused the development of the 335-190 ~ joint sets.

In recent years there has been renewed interest in various aspects of rock jointing. These studies have ranged from the analyses of interacting joint sets (Reches 1976; Hancock 1985; Bahat 1987), joint interpretation investigated regarding joint frequency (Rives et al. 1992), and the concept of fracture saturation and the sequential development of joint patterns (Gross 1993). There has been a movement away from the investigation of regional joint systems and attempted interpretations of regional joint patterns in terms of palaeostress fields. The emphasis has moved towards an understanding of the factors that produce fracture and propagation mechanics of individual joint planes. The nature of this change is documented clearly in the synoptic paper by Pollard & Aydin (1988). The relevance of the fracture surface markings (FSM) and what they mean in terms of the mechanics of joint initiation and propagation has been made possible by using fractography as a geological tool. Kerkof (in Bahat 1991) states, 'while in photography the light fixes the picture of an event, in fractography it is the fracture that plays the corresponding part'. FSM have been observed for almost a century in many different kinds of rocks and in a variety of tectonic settings. Woodworth (1896) produced the fundamental article on the subject in his study of the Mystic River Slates. However, the adoption of fractographic principles by Kulander et al. (1979) to a theoretical and applied approach to rock fracturing laid the path to future interpretative studies. The synoptic study by Bahat (1991) provided an updated synthesis of the subject. As a result of fractographic studies what has to be called into question are the former interpretations of joints as kinematic indicators of rock deformation. The classification of

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Fig. 1. The components of fracture surface markings; descriptive classification from Hodgson (1961b) and genetic classification from Kulander

et al.

(1979).

joints as being either 'shear' or 'tension' in origin is no longer tenable. There would be broad general agreement that the bulk of joints result from Mode I tensile fracturing, but where the total surface area of the joint is displayed (and in suitable lithologies) the joint can be seen to be the result of mixed-mode failure. The result of mixed-mode failure, i.e. Modes I and III, can be seen clearly in Fig. 1, where in broad terms the hackle plume area represents Mode I failure (opening) and the fringe (border planes-twist hackle) the out-ofplane propagation of the main joint plane in response to a twist imposed by the influence of the remote stress field (a3). However, even in the absence of twist hackle most workers would regard a3 to be normal to the joint plane.

The study area The basal succession of the lower Liassic crops out west of Penarth, South Wales, between Lavernock Point and Saint Mary's Well Bay (Fig. 2). The succession has been described as the Saint Mary's Well Bay Formation (Waters & Lawrence 1987), consisting of alternating blue-black shales, mudstones and impure muddy limestones. This formation comprises the lowest 16m of the Liassic succession and passes upwards into the Lavernock Formation, a predominantly shaley sequence c. 15 m thick. The studied horizon is located in the upper part of the Saint Mary's Well Bay Formation (Fig. 3). The succession is well

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Fig. 2. Location and general geology of the study area.

exposed in a low cliff line and also in a series of gently dipping limestone pavements. The pavements provide information regarding the geometry of the joint patterns and their interactive elements, whilst the FSM were studied in the cliff outcrops. The structural geology of the area is that of a gently SSE plunging syncline (Fig. 2) and one fault subparallel to the synclinal axis. The fault strikes at c. 330 ~ dips 65~ with a dip-slip displacement of c. 2 m. Where the fault trace intersects the pavement exposures there is a 50 cm zone of calcite mineralization, usually in the form of parallel-sided veins, 3-5mm in width and filled with sparry calcite. This is the only manifestation of mineralization in the study area.

The fractured horizon The best developed examples of FSM are stratigraphically located towards the top of the Saint Mary's Well Bay Formation. Three limestone beds, i.e. 32-34, display abundant FSM (Fig. 3). However, for descriptive purposes, it is from Bed 22 that most of the observations are derived; the FSM geometries are similar in Beds 33 and 34, and with particular reference to the sense of twist-hackle rotation, identical. Some FSM geometries are recorded in Fig. 4a-d, all relating to Bed 32, whilst one example of joint pattern and interactions is given in Fig. 4e. The distinctive lithology of this horizon had already been identified by Waters & Lawrence (1987), who named it the laminated horizon. The most

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J.c. ROBERTS

L.S.

Lavernoc 34

I I

to

34 } ES.M. 33 32 Horizon

E 0 la_

1 m

33 1 32

I

I

I lOcm

m m

31

c

1

m

D

m

Fig. 3. Simplified stratigraphic column with the location of the studied horizon.

FRACTURE SURFACE MARKINGS

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Fig. 4. (a) A point source in Bed 32 with an elliptical termination. (b) Point source of an S-type chevron in Bed 32. Note the influence of the distinct parting on the location of the plumes. (e) A double point source in Bed 32. (d) An S-type chevron in Bed 32 with a deep fringe and twist-hackle planes. (e) The joint trace pattern in Bed 32, looking towards 10~ Chalk marks are 1 m apart and are parallel to 190~ joint traces and perpendicular to 280 ~ joint traces.

highly developed l a m i n a t i o n occurs in Bed 32, a 10cm thick limestone with at least 25 laminations, some of these can be seen clearly in Figs 4 a - c & 5a & b. Therefore, the study is concentrated within a 2.5 m slice of interbedded limestones and shales, with the F S M relating m a i n l y to observations within the limestone layers. However, their influence u p o n the interbedded shales can be seen clearly in Fig. 5 b - d (deep twist-hackle structures).

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Fig. 5. (a) Parallel 190~ joint faces displaying localization of the chevron axis along a well-defined lamination. (b) Parallel 190~ joint faces displaying asymmetric barbs. Note that the chevron at the left of the figure displays within the limestone layer and exhibits breakdown into broader planes. (e) An asymmetrically positioned point source in Bed 32. Note the strongly developed border planes above the joint plane. These are rotated counter-clockwise in relation to the joint surface. (d) An asymmetric chevron breaking down into deep fringe with en echelon border planes. The joint system The joint system within the study area has been described by Roberts (1974). The systematic joint system is represented in the rose diagram of Fig. 2. However, the synoptic nature of the diagram does not convey accurately the joint pattern, Beds 32-34 are cut by three main joint sets (better referred to as spectra as there is a spread of azimuth), i.e. 335, 280 and 190 ~ (Fig. 4e, sets 280 and 190~ As is evident in Fig. 4e, the 190 ~ set displays a high frequency and a short joint trace length, but is restricted to Beds 32-34.

Method of study During the previous study of the systematic joint patterns (Roberts 1974) FSM were noted at Lavernock, and many other localities along the coast. However, the best examples occur within the described location. The attributes of FSM were recorded subsequent to the investigation of the joint system, when ground truth data were recorded for the joints. Details of FSM have been derived using oriented photographs, however, ease of obtaining

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detailed surface morphologies was reduced by the general E - W alignment of the cliff face which precluded extensive development of the 190~ joint plane surfaces. In addition, the low relief of FSM induced a further constraint in that incident light (sun) angles could only be optimized over short time intervals.

Fracture surface markings (FSM) The morphological elements of FSM are illustrated in Fig. 1. The current descriptive terminology is due to Hodgson (1961b), but this has been amplified by Kulander et al. (1979) into a genetic scheme. The main departures in the respective terminologies relate to structures in the fringe, and the recognition by Kulander et al. (1979) of this structure as representing a particular mode of fracturing.

FSM

on t h e 1 9 0 ~ j o i n t s e t

As stated above, excellent FSM occur in Beds 32-34; however, the bulk of the descriptions relate to observations on Bed 32. Hackle plume markings. These are ubiquitous features on the surfaces of the 190~ joint set, and the bulk of the observations indicate that the common direction of propagation of the joint plane was layer parallel and controlled by lamination within the bed. This control exerted by lamination is easily seen in Figs 4a-c & 5a & b. The plumes, in a geometric sense, can be described as chevron type, in that the plume barbs have a straight linear relationship to the plume axis, in contrast to some of the wavy plumes which are frequently associated with fracture initiation points on the base of a bed (Bahat & Engelder 1984). Details of the variations in the location of the plume axis and its asymmetry with respect to the bedding surfaces are given in Figs 4c & d & 5d. In Fig. 5d the plume axis is inclined at an angle of c. 25 ~ to the left of the photograph, but it becomes layer parallel within 10 cm of the cross joint striking at 280 ~ A parallel joint plane at the same exposure displays identical plume axis orientation, indicating that the principal effective stresses were uniform over this volume of rock (see also Fig. 5a). Divergence in propagation direction is seen clearly in Fig. 5c, where there is a fan arrangement of the plume structure indicating that in the initial stage of propagation the joint surface propagated vertically as well as parallel to bedding. However, the majority of the hackle plume marks on the 190~ joint set were of the S-type (Bahat & Engelder 1984). Point sources. The reality that joints initiate at points (small discrete areas of failure) can be demonstrated in many places (Roberts 1961). Point sources are seen clearly in Figs 4a-c & 5b, where the hackle plumes diverge from a central point. Figure 4a shows a point source that lies mid bed, a lamination-parallel plume axis and a rib mark 20 cm to the left of the initiation point. A most uncommon situation is seen in Fig. 5b where two in-plane point sources have given rise to a rhomboid-shaped interference pattern. Rib marks. These structures are not common on the 190~ joint set, but where they appear they indicate a deceleration of the fracture front (Fig. 4a), and in some cases actual termination of the joint plane (Kulander et al. 1979, fig. 77).

182

J . c . ROBERTS

~'~'~~ 010~

335~

l

Twist //! j hackle plane

oo

E

ccw Twist c% cw Twist 0 Twist angle Greatest principle driving stress Remote stress field

~ ....,,

I

\ 155~

1900 Fig. 6. Relationships of joint traces and twist-hackle planes to the stress field.

Twist-hackle planes of the fringe. The recognition of twist-hackle planes is important because according to Kulander et al. (1979) they form when 'a propagating fracture abruptly enters a region of different stress orientation. The fracture breaks (twists) into a series of en echelon individual lance-like twist-hackle faces, each perpendicular to a resultant tension'. In this investigation the salient feature, as evidenced by Beds 32 and 33, is that the twist-hackle component is not found in the limestone layer but in the

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183

interbedded shale layers, i.e. above and below the limestone; this relationship is seen clearly in Figs 4d & 5c & d. What should also be noted is the development of twist hackle from the hackle plume (Figs 4d & 5d); in Fig. 5d the twist-hackle zone is seen to have developed to a depth of 15 cm below the hackle-plume axis. In Fig. 5d the discrete planar elements of the twist hackle are evident, also the distinct out-of-plane orientation of the twist-hackle planes, this is recorded as the twist angle (Fig. 6), the values of which have implications for the intensity of confining pressures at the time of fracture propagation (see below). In addition, the recognition of a zone of twist hackle is important in that if the twist hackle were to develop on a larger scale, then it could be mistaken for a zone of close jointing (Kulander et al. 1979). Indeed, such a process of large-scale twist-hackle formation leading to the production of close joint zones has been noticed in more quasiisotropic rocks, e.g. granites and volcanic rocks (P. Bankwitz, pers. comm.), and indicates high velocity propagation of the joint planes in the close joint zone.

F S M on other j o i n t sets FSM are not confined to Beds 32-34, it is that they attain their acme at this horizon, and in particular on the 190~ joint set. The 335 ~ joint set frequently displays S-type plumes mainly of a chevron geometry. In contrast, the 280 ~ joint sets are characterized by FSM of a more delicate plumose pattern with many elliptical rib markings.

Discussion FSM were studied in three limestone beds on both flanks of an open syncline- dips < 10~ - which is crossed by a single dip-slip fault with a downthrow of 2 m to the west (Fig. 2). Both limbs of the syncline are cut by three sets of joints (Rose diagram, Fig. 2). This diagram, a cumulative frequency plot, gives an impression of the geometry of the joint system which pervades the Saint Mary's Well Bay Formation of some 36 limestone beds. There is nothing in the morphology of the joint sets to suggest that they can be differentiated in terms of their origin, e.g. uplift as opposed to burial (Bahat 1991), however, it was suggested by Roberts (1974) that the joint system resulted from a weak tectonic deformation, with O"1 aligned approximately N-S. The somewhat tenuous kinematic relationship of the joint sets to the local synclinal structure requires substantiation, some of which might be provided by the evidence of FSM. However, whatever conclusions are arrived at they must be considered solely with respect to the nature of the fracture pattern and the associated FSM. The pertinent observations are that all of the FSM in Beds 32-34 which are developed on the 190 ~ joint set are of the S-type and show a chevron hackle plume, and that the joint set is of a multilayer nature, i.e. they cut more than one bed. The jointing architecture of Beds 32-34 suggests that the 190~ set formed last in the jointing sequence, as evidenced by their high frequencies, large number of intersections upon other joint sets and selfinteraction. In addition, the high frequency of cross-cutting (intersections) indicates that the 335 and 280 ~ joint sets were closed at the time the 190 ~ set propagated. The development of clear FSM in this rock succession is not unexpected since the role of lithology as a controlling factor has been demonstrated by many workers, and is summarized by Syme-Gash (1971), who states that FSM are 'commonly developed on joints with inter-laminated strata such as occurs in the Lias or flysch and on fractures in metal plates', he also recognized the fundamental control of grain size on inhibiting FSM,

184

J.C. ROBERTS

their development being favoured by fine-grain sizes. In contrast, it must be noted that large-scale FSM have been observed in coarse-grained igneous rocks, and therefore grain size alone cannot be regarded as an inhibiting factor for their development. The confinement of the hackle plume to the limestone layers, and its initiation within the limestone layer, must result from differences in the elastic module of the interbeds. This is borne out by the development of the twist hackle in the overlying and underlying shale beds (Fig. 5c & d) where there is a reduction in the crack-tip stresses due to local plastic flow (Kulander et al. 1979). The most important observation with respect to the twist hackle is that the sense of twist remains constant along the length of the joint plane, and that the angle of twist is usually in the order of 10-15 ~. This study has shown that the sense of twist on the 190~ joint set is always counter-clockwise. Recent work by Cruikshank et al. (1991) has shown that a relationship exists between the size of the twist angle and ratio of Mode III to Mode I driving stresses. Although their calculations were based upon a sandstone with a Poisson's ratio of 0.25, their values would indicate that the twist angles in the shales suggest a stress ratio in the region of 0.2-0.3. The significance of the value of the angle of twist has been discussed by Bahat (1987) who has shown that it can also be related to the angle of internal friction of the rock. In the case of Bahat's study of Eocene chalks the characteristic twist angle was found to be 17~ The 335 and 280 ~ joint sets display clockwise twists of 10-15 ~. These observations suggest that all the joint sets formed under similar stress conditions, but not necessarily at the same time. The salient point is that certain conditions are conducive to the production of the twist-hackle component of the joint plane. The presence of the twist hackle suggests low confining pressures, Bahat (1991) stated, 'It appears that en echelon cracking is inhibited beyond certain confining pressures'. In terms of the orientation of the palaeostress field, the twist-hackle plane is an indicator of the direction of maximum tension immediately ahead of the crack front (Cruikshank et al. 1991). Bahat (1991) states that the size of the twist angle is influenced by the magnitude of the remote stress field. These conclusions were arrived at by Kulander et al. (1979) who stated that, 'Twist hackle steps increase in relief towards stratum boundaries. In this direction principal effective stresses decrease, superposed stresses have greater effect, and the resultant stress perpendicular to the twist-hackle face may alleviate more from the principal stress'. The observations on the magnitude of the twist angle may have a bearing upon speculations regarding possible stratigraphic thickness of the Jurassic succession in the study area. The remnant thickness of the Liassic is in the order of 150 m (Wilson et al. 1990), but Wobber (1966) and Owen (1967) have suggested that Mesozoic strata covered the South Wales Coalfield, a depth of burial that would have required a much greater thickness of post-Triassic strata. Evidence of a thicker Jurassic succession is found in faultbounded synclines in the Bristol Channel (Wilson et al. 1990). Therefore, with regard to the small value of the twist angle, either the joint set in question evolved early during the post-lithification stage of sedimentation or much later under conditions of low overburden (confining pressure), which would suggest an attenuated Jurassic succession. Finally, some of the lithological controls exerted by the succession should be considered. The fractured succession consists of alternations of mud and lime rocks in almost equal proportions, and has responded in two different ways with respect to the development of the joint sets (Hodgson 1961a). The limestone layers respond to the stress as relatively rigid members, and it would appear that limestones 32-34 acted either as stress concentrators or occupied a zone of high stress concentration within the succession. The role of the limestone beds as stress concentrators might have been influenced by the relative

FRACTURE SURFACE MARKINGS

185

rates of diagenesis of the lime and mud sediments. The lime sediments lithify earlier and would therefore be capable of taking stress earlier than the mud sediments, and so would act as accumulators of early stresses. There is no evidence to suggest that if joints formed early in the limestone layers that they propagated upwards or downwards at a later stage into the mud rocks, as stated above where hackle plumes are exposed to the principal direction of joint propagation which was layer parallel with joint limits terminating in the mud rocks. What is evident is that the limestone layers acted as the stress concentrators and provided sites for fracture initiation.

Conclusions Within the basal part of the Liassic succession a 2.5m thick horizon comprising three limestone layers with shale interbeds displays abundant, well-developed FSM. These represent an episode of rapid, brittle fracturing during one phase of the evolution of the joint system. The fracture pattern includes three joint sets striking at 335, 280 and 190 ~. The F S M are best developed on the 190 ~ joint set which is characterized by having a high joint frequency and short joint trace length. The other two sets also display F S M but not as frequently as the 190 ~ set. The most common geometry of F S M on the 190 ~ set is Stype, chevron hackle plumes. The joints terminate in the overlying and underlying mud rocks with deep borders of twist hackle. The sense of rotation of the twist-hackle planes is consistent for a given joint set, albeit the degree of twist suggests low confining pressures at the time of joint propagation. The clockwise rotation of the twist hackle on the 335 ~ joint and the counter-clockwise rotation on the 190 ~ set suggests that or1 was oriented approximately N - S at the time of propagation of both joints. This, however, does not imply that the joint sets are of a conjugate nature, but rather that if conjugate relationships were being sought then this pair set could be so considered, whereas the clockwise twist on the 280 ~ set excludes any conjugate relationship with either the 335 ~ or the 190 ~ set.

Many thanks are due to Nigel McDowell for the photographic work, Killian McDaid for the graphics, and to Mary McCamphill and Anne Hayes for typing the manuscript. Also, in the earlier years of the investigation, to the Research fund of the University. The manuscript was improved greatly by the constructive comments of the referees - M. Cooke, A. Aydin and P. Bankwitz. I thank them all.

References BAHAT, D. 1987. Correlation between surface morphology and orientation of cross-fold joints in

Eocene chalks around Beer Sheva, Israel. Tectonophysics, 136, 323-333. 1991. Tectono-fractography. Springer-Verlag, Berlin. - 8s ENGELDER, T. 1984. Surface morphology on cross-fold joints of the Appalachian Plateau, New York and Pennsylvania. Techonophysics, 104, 299-313. BANKWITZ,P. 1965. Uber Klufte, beobachtungen in Thurinischen Schiefergebirge. Geologie, 14, 242253. - 1966. Uber Klufte II. Geologie, 15, 896-941. CRUIKSHANK, K.M., ZHOA, G. 8s JOHNSON, A. M. 1991. Analysis of minor fractures associated with joints and faulted joints. Journal of Structural Geology, 13, 865-886. GROSS, M. R. 1993. The origin and spacing of cross joints: examples from the Monterey formation, Santa Barbara Coastline, California. Journal of Structural Geology, 15, 737-752. HODGSON, R. A. 1961a. Regional study ofjointing in Comb Ridge-Navajo Mountain Area, Arizona and Utah. Bulletin of the American Association of Petroleum Geologists, 45, 1-38.

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1961b. Classification of structures on joint surfaces. American Journal of Science, 259, 493502. HANCOCK, P. L. 1985. Brittle microtectonics: Principles and practice. Journal of Structural Geology, 7, 437-457. KULANDER, B. R., BARTON, C. C. & DEAN, S. L. 1979. The application offractography to core and outcrop fracture investigations. Morgantown Energy Technology Centre, METC/SP-79/3. OWEN, T. R. 1967. 'From the South': a discussion of two recent papers in the proceedings. Proceedings of the Geologists" Association, 78, 595-599. POLLARD, D. D. & AYDIN, A. 1988. Progress in understanding jointing over the past century. Bulletin of the Geological Society of America, 100, 1181-1204. RECHES, Z. 1976. Analysis of joints in two monoclines in Israel. Bulletin of the Geological Society of America, 87, 1654-1662. RIVES, T., RAZACK, M., PETIT, J. P. & RAWNSLEY, K. D. 1992. Joint spacing: analogue and numerical simulations. Journal of Structural Geology, 14, 925-938. ROBERTS, J. C. 1961. Feather-fracture, and the mechanics of rock-jointing. American Journal of Science, 259, 481-492. 1974. Jointing and minor tectonics of the vale of Glamorgan between Ogmore-by-Sea and Lavernock Point, South Wales. Geological Journal, 9, 97-114. SYME-GASH, P. J. 1971. Surface features relating to brittle fracture. Tectonophysics, 12, 34-391. WATERS, R. A. & LAWRENCE, D. D. 1987. Geology of South Wales Coalfield, Part III, the country around Cardiff, 3rd Edition. Memoirs of the British Geological Survey, sheet 263 (England and Wales). WILSON, D., DAVIES, J. R., FETCHER, C. J. N. & SMITH, M. 1990. Geology of the country around Bridgend. Explanation of the 1:50 000 sheet. British Geological Survey, sheet 262, England and Wales. WOBBER, F. J. 1966. A Study of the depositional area of the Glamorgan Lias. Proceedings of the Geologists' Association, 77, 127-137. WOODWORTH, J. B. 1896. On the fracture system of joints, with remarks on certain great fractures. Boston Society of Natural Historical Proceedings, 27, 163-183. -

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From Ameen, M. S. (ed.), 1995, Fractography: fracture topography as a tool in fracture mechanics and stress analysis Geological Society Special Publication No. 92, pp. 187-196

The expression of hydraulic fracturing in rocks and sediments J. W. C O S G R O V E

Department of Geology, Imperial College of Science, Technology & Medicine, London SW7 2BP, UK Abstract: The orientation and spatial organization of hydraulic fractures in rocks and sediments is determined by the magnitude of the differential stress, the orientation of the principal stress axes and the intrinsic properties of the rock, particularly its cohesive strength. Depending on the values of these various parameters the expression of hydraulic fracturing in a rock or sediment can be extremely varied ranging from conjugate shear fractures through parallel tension fractures to randomly oriented tension fractures and fluidized sediments. The influence of a rock's properties on its response to hydraulic fracturing is illustrated by considering the mechanics of formation of sedimentary dykes.

The theory of brittle failure and hydraulic fracturing are discussed in most texts on structural geology (e.g. Price 1966, Phillips 1972; Price & Cosgrove 1990) and the reader is referred to them. Only a brief review of aspects of the topic which bear directly on the ideas discussed in this article is given here. The stress conditions necessary for the formation of shear and tensile failures are shown in Fig. 1. In this figure the Navier-Coulomb criteria of shear failure and the Griffith criteria of tensile failure are expressed graphically, and the stress states are represented by M o h r circles. In order for tensile failure to occur the Mohr circle must touch the brittle failure envelope at the point 7- = 0, 0"3 = T, where T is the tensile strength of the rock. It follows from the geometry of the failure envelope (Fig. 1) that this can only occur if the diameter 0"1

a)

0"1

0

~

x2+4Tcrn - 4T =

Fig. 1. The graphical expression of the two brittle failure criteria and Mohr circles representing stress states capable of causing tensile (a) and shear (b) failures.

188

J.w. COSGROVE

"cj i I

D,

0

0-1

0"3 ~

O'1

~ 0"3

0"3 ~

/ ~ /

(3"1

0"1

0"1

0"1

t

t

0"3 ~

~ 0"3

b 0"1

0"3 ~

F?, 1

~-0"3

~ 0"3

0"1

Fig. 2. (a) Mohr circles (i-iv) representing a range of stress states which will lead to tensile failure. The Mohr 'circle' iv that represents hydrostatic stress is a point. (b) (i-iv) Patterns of tensile fractures generated, respectively, by the four stress states shown in (a).

HYDRAULIC FRACTURING

189

Fig. 3. (a) A regular array of tensile fractures exposed on a bedding plane in Carboniferous sandstone at MiUook, North Cornwall. The alignment of the veins indicates that a significant differential stress existed during fracturing. The diameter of the coin is 3 cm. (lo) Devonian sandstone cut by randomly orientated tension fractures which indicate a hydrostatic stress state during fracturing, at Combe Martin, North Devon. The diameter of the coin is 3 era.

190

J.W. COSGROVE

of the Mohr circle (the differential stress 0 - 1 - 0"3) is small, i.e. < 4T. Mohr circles representing four stress states that satisfy these conditions are shown in Fig. 2, and in the following section the orientations of the fractures that form in response to these stresses are discussed.

The orientation of tensile fractures Consider the orientation of fractures that form in response to the range of differential stress states represented by the Mohr circles shown in Fig. 2, all of which satisfy the conditions for tensile failure. The differential stresses vary from just < 4T (circle i, Fig. 2) to zero ('circle' iv, Fig. 2). Note that when the differential stress is zero, the stress state is hydrostatic and the Mohr circle collapses to a point. Tensile fractures form parallel to the maximum principal compressive stress, 0-1, i.e. they open in the direction of the minimum principal stress, 0-3 (Fig. 1a). Thus, in the stress state represented by Mohr circle i in Fig. 2, which has a relatively large differential stress, there is a definite direction of easy opening for the tensile fractures, i.e. parallel to 0-3. The fractures would therefore exhibit a marked alignment normal to this direction (Fig. 2b, i). However, for the stress states represented by the Mohr circles ii-iv, the differential stress becomes progressively smaller until, for the hydrostatic stress represented by Mohr 'circle' iv, the differential stress is zero. In a hydrostatic stress field the normal stress across all planes is the same and there is therefore no direction of relatively easy opening for the tensile fractures. Thus, the fractures will show no preferred orientation and, if they are sufficiently closely spaced and well developed, will produce a brecciation of the rock (Fig. 2b, iv). It is to be expected, therefore, that as the differential stress becomes progressively lower (stress states i-iv, Fig. 2), so the tendency for the resulting tensile fractures to form a regular array normal to a3 decreases. Tensile fracture systems ranging from well aligned fractures to randomly oriented fracture arrays are to be expected in rocks, and field observations (Fig. 3a & b), support this idea.

The role of fluid pressure in tensile failure It is generally acknowledged that the state of stress in the Earth's crust tends to be compressional and that true tensile stresses are uncommon. For example, in a tectonically relaxed region of the crust, i.e. a region where the overburden is the main source of stress, both the vertical and horizontal stresses are compressional. The relationship between the two depends on the boundary conditions and the rock type. If, for example, the boundary conditions are such that horizontal strains are prevented by the constraints of the rock mass surrounding the area of interest, then it can be shown (see Price 1966, p. 69) that the relationship between the vertical and horizontal stresses is given by: 0-H =

0-v/(m- 1),

(1)

where m is Poisson's number - the reciprocal of Poisson's ratio. Under these boundary conditions the vertical stress will be 0-1 and the principal stresses will be compressive at all depths in the crust. It can be seen from equation 1 that the differential stress (0-v - o-n) will depend on the depth of burial and the material properties of the rock.

HYDRAULICFRACTURING

"s SJ'~

191

J"

s t S tI . _ ...t

-

', s S

2

a"

/ (0"3-PH20)

I~3

(l~l_PH2-~O)

0"1 O'v

Fig. 4. Mohr circle representation of four stress states (solid line circles 1-4) which will not cause failure. The effect of fluid pressure is to move the stress state to the left by an amount PH20. If the new circle (dashed) touches the failure envelope hydraulic fracturing occurs. Whether this failure is by shear or tension depends on the differential stress.

The stress states in four different rocks at a particular depth are represented in Fig. 4 by four Mohr circles. They have the same overburden stress, el, and their horizontal stresses are determined by their material properties. Because the principal stresses are compressional none of these stress fields can cause tensile failure. However, field and borehole core observations show that tensile failure is common in sediments subjected to the 'zero horizontal strain' boundary conditions mentioned above and this dilemma has been resolved by arguing that the stress states shown in Fig. 4 (and represented by the solid-line Mohr circles 1 4 ) can be modified by fluid pressure. The build up of fluid pressure during burial acts so as to oppose the compressive stresses generated by the overburden and causes them to be reduced to the effective stresses (0-1 - - P n 2 o ) and (0-3 PH20). Thus, the Mohr circles are moved to the left by an amount equivalent to PH20 and by this means, if the differential stress is < 4T, can be brought into contact with the tensile failure envelope at the point 0-3 = - T , r = 0. Failure caused in this way is known as hydraulic fracturing and the type of fracture that develops depends on the differential stress. For example, if the fluid pressures in rocks with the stress states represented by the solid line Mohr circles in Fig. 4 were sufficiently large to cause failure, rock 1 would fail by shear failure and rocks 2-4 by tensile failure, as indicated by the dashed Mohr circles. As mentioned earlier, as the differential stress is reduced, so the tendency for the tensile fractures to form normal to 0-3 decreases until in the limit when 0-1 -- 03, the fractures are randomly oriented (Fig. 2b, iv). It is a common misconception that the result of hydraulic fracturing in rock is the formation of randomly oriented tensile fractures and the generation of breccia textures (Fig. 3b). The above discusion shows that, depending on the differential stress, the expression of hydraulic fracturing can range from randomly oriented tensile fractures through aligned tensile fractures to shear fractures.

192

J.W. COSGROVE

Fig. 5. Diagrammatic representation of fluidization. The build up of fluid pressure in the pores of an uncemented granular sediment (a) reduces the normal stress acting across the grain contacts, thus facilitating flow (b). The normal stress may be reduced to zero and the grains may move apart (e).

Fluidization Now consider the effect of hydraulic fracturing on uncemented sediments with no intrinsic cohesion, such as an uncemented sandstone lens in a shale. A material such as this has no strength and therefore cannot support a differential stress. Thus, the stress state within it will be close to hydrostatic. If the fluid pressure is such that hydraulic 'fracturing' can occur then the grains of the sediment will simply move apart slightly and the sediment will fluidize. Thus, an additional diagram should be added to Fig. 2b showing the effect of hydraulic fracturing under conditions of hydrostatic stress in a material with zero cohesion such as an uncemented sediment (Fig. 5). Evidence for two different modes of tensile failure occurring at the same depth and at the same time in different 'rocks' is provided by sandstone dykes frequently found cross cutting shales (Fig. 6a). Because of the electrostatic charges between clay particles making up a shale there is an intrinsic cohesion in the sediments from the moment they are deposited. This cohesion will increase with compaction and cementation. However, sand

HYDRAULIC

(a)

]I.2

193

FRACTURING

' Sandstonel

I I

i

I

I I

,

m m

I

I

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(b)

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T Fig. 6. (a) Diagrammatic representation of a folded sandstone dyke cutting the Mercia mudstone at Watchet, North Somerset. (b) Unfolding of the dyke to determine the thickness ll, of the bed at the time of dyke formation. (e) Original thickness of the sediment, 10.

lenses interbedded with the shale will remain cohesionless until the processes of cementation are initiated. Thus, during the early stage of burial the shales will possess a cohesion and the sandstones will not. If, during this stage, the fluid pressures become high enough to cause hydraulic fracturing in both sediments, the shales, because of their intrinsic cohesion, will be able to sustain a differential stress. If the differential stress is < 4T of the shale, tensile failure will occur and the resulting fractures will be vertical. At the same time the effect of hydraulic fracturing on the sand lenses which have no cohesion will be to cause fluidization. Fluidized sand will then be injected into the vertical fractures in the shales wherever these fractures intersect the sandbodies. Subsequent compaction of the shales will cause these originally vertical sand dykes to become buckled (Fig. 6a). The thickness, ll, of the shale when the dykes were intruded can be calculated by unfolding the dykes (Fig. 6b). If the present state of compaction of the shale (Fig. 6a) can be determined then its original thickness, 10, can be calculated together with the stage in burial when hydraulic fracturing occurred. Tensile failure in anisotropic

rocks

The general condition for tensile failure by hydraulic fracturing along any plane is that the fluid pressure, PH20 must be equal to or greater than T, the tensile strength of the rock

194

(a)

(b)

J . W . COSGROVE

HYDRAULIC FRACTURING

195

(c)

Fig. 7. Bedding-parallel veins indicating hydraulic fracturing along bedding with fluid pressures in excess of the lithostatic load, (a) calcite (beef) vein from Lulworth Cove (Cosgrove 1993, fig. 2a) and (b) satin spar veins from the Mercia mudstone at Watchet Harbour, North Somerset. The bedding is weakly developed and forms anastomosing, sub-horizontal planes of weakness which are exploited by the veins. The diameter of the coin is 2 cm. (e) Vertical tension veins with small offshoots along bedding cutting a Carboniferous sandstone in the Rusey Fault zone, Boscastle, North Cornwall. normal to that plane and the normal stress acting across it, i.e. PH~O > T + 0".

(2)

In an isotropic rock this condition is first met along the plane normal to the least compressive principal stress, 0"3 (Fig. la). However, in an anisotropic rock, i.e. one in which the tensile strength varies in different directions, this is not necessarily so. Consider a well-laminated shale with a marked planar anisotropy parallel to bedding. The tensile strength in the direction normal to bedding, Tn, will be considerably less than that parallel to it, Tp. From equation 2, the conditions for tensile failure parallel and normal to bedding by hydraulic fracturing are: PH20 > Tp + O"H,

(3)

PH20 > T. + o"v,

(4)

respectively, recalling that 0-v = 0-1 and 0-H = 0"3.Which of these two conditions is satisfied first depends upon the relative values of the differential stress (0"1 - 0"3) and the difference between the two tensile strengths (Tp - 7",). If (0"2 - 0"3) > ( T p - T n ) then, as the fluid

196

J.W. COSGROVE

pressure increases, equation 3 is satisfied first and the resulting hydraulic tensile fractures form normal to a3, i.e. are vertical. If the differential stress is less than the difference between the two tensile strengths then the condition represented by equation 4 will be satisfied first and hydraulic tensile fractures will form parallel to bedding, i.e. normal to al. Field evidence for these fractures can be found in the form of bedding parallel, horizontal veins of fibrous calcite (beef) and gypsum (satin spar) (Fig. 7a & b). The fibres are vertical and indicate the direction of opening of the veins. These veins show that fluid pressures equal to or in excess of lithostatic pressure can develop (Cosgrove 1993). If the condition (a3 + Tp) = (al + Tn) occurs then the fluid pressure required to form hydraulic fractures parallel and normal to the bedding will be the same and there would be an equal likelihood that both sets of fractures will form. An example of this is illustrated in Fig. 7c, which shows vertical viens of quartz with small horizontal offshoots forming along the bedding.

Conclusions Because the state of stress in the Earth's crust is generally compressive it is usually necessary for high fluid pressures to be present in the rock in order for tensile failure to occur. The orientation of the tensile fractures can vary between parallel arrays of fractures which form normal to the least principal stress, a3, when the differential stress is relatively large (but still < 4T of the rock) to randomly oriented fractures which are generated when the state of stress in the rock is approximately hydrostatic, i.e. when the differential stress approaches zero. Hydraulic failure of a rock with a very low cohesive strength such as an uncemented sandstone, may lead to fluidization. Once fluidized the sediment may be injected along hydraulic fractures in the surrounding, more coherent rocks, resulting in the formation of vertical sedimentary dykes which buckle during subsequent compaction and dewatering. The stage in the burial history at which the dyke formed (i.e. when hydraulic fracturing occurred) can be determined by 'unfolding' the dykes to obtain the thickness of the injected bed at the time of dyke formation. Although tensile failure generally occurs along a plane normal to the least principal stress, a3, it is possible for such failure to occur normal to al if the rock is mechanically anisotropic. For example, in a sedimentary succession in which the main source of stress is the overburden, failure will occur normal to O"1 when the differential stress is less than the difference between the tensile strength of the rocks parallel and normal to bedding.

References COSGROVE,J. W. 1993. The interplay between fluids, folds and thrusts during the deformation of a sedimentary succession. Journal of Structural Geology, 15, 3-5, 491-500. PHILLIPS, W. J. 1972. Hydraulic fracturing and mineralisation. Journal of the Geological Society of London, 128, 337-359. PRICE, N. J. 1966. Fault and Joint development in Brittle and Semi-Brittle Rock. Pergamon Press, Oxford. - & COSGROVE,J. W. 1990. Analysis of Geological Structures. Cambridge University Press, Cambridge, UK.

From Ameen, M. S. (ed.), 1995, Fractography: fracture topography as a tool in fracture mechanics and stress analysis

Geological Society Special Publication No. 92, pp. 197-213

Spatial change in joint geometry in the Chalk of eastern England A. M . G O O D W I N

Department of Earth Sciences, Cambridge University, Downing Street, Cambridge, CB2 3EQ, UK Abstract: Prominent NE and NW striking joint sets previously recognized in southern

England have been traced northwards in the Chalk of eastern England. Both sets change their geometry northwards, defining two joint continua, from steep conjugate hybrid joints in Norfolk to a sub-vertical single joint set in Humberside. This spatial change is consistent both with the observed northward increase in rock strength of the Chalk and a possible northward decrease in differential stress across eastern England. Fractures, particularly joints, form one of the most common rock structures. Although studied for over a century (Pollard & Aydin 1988) it is only recently that their usefulness within the field of brittle microtectonics has been recognized (Hancock 1985). Despite their sometimes kinematically ambiguous characteristics many studies have highlighted the valuable nature of fractures in microtectonic investigations, particularly for inferring regionally significant stresses and their relative magnitudes. Such studies show the contribution that systematic joints can make in establishing past and contemporary stress regimes, especially in stable platform regions of mild deformation (Engelder & Geiser 1980; Engelder 1982; Eyal & Reches 1983; Hancock et al. 1984; Bevan 1985; Caputo 1991; Zoback 1992). Despite the increase in use of brittle-microtectonic features the term 'joint' has been defined in a variety of ways (e.g. Dunne & Hancock 1994) and the genetic classification of joints and fractures is still somewhat disputed. For the purposes of this paper, the following definition of the term 'joint' will be used - a planar discontinuity which at the scale of observation possible in the field shows no discernible evidence for offset related to shear, dilation, or shortening (altered from the definition given by Hancock 1985). Previous work has demonstrated the predominance of two joint strikes, NW-SE and NE-SW, within the Mesozoic and Cenozoic lithologies of eastern and southern England (Norfolk, Essex, Kent and Dorset), northwest France, and Belgium (Toynton 1983; Bevan 1985; Balson & Humphreys 1986; Bevan & Hancock 1986; Vandyke et al. 1991). Joints have also been documented in the Chalk of Yorkshire (Patsoules & Cripps 1990) in a similar NE and NW pattern, but no such study has been undertaken in Lincolnshire. Bevan (1985) regarded the two identified joint sets as coeval (questionably the NW set was the older) and due to a Miocene-Pliocene deformation phase associated with Alpine compression. Balson & Humphreys (1986) assigned a Pleistocene age to their joint systems whilst Vandycke (pers. comm.) believes the joints to be of a late Cretaceous to early Tertiary age. Toynton (1983) makes no comment on the age and tectonic relationships of the joints observed in Norfolk. There are several possible explanations for the formation of the two joints sets within the tectonic and intra-plate stress regimes of northwest Europe since the end of the Cretaceous: Palaeocene-Eocene uplift of NW Britain, Tertiary

198

A.M. GOODWIN

(a)

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CHANGING JOINT GEOMETRIES IN CHALK

199

tectonic activity in the North Sea, a 'neotectonic' origin related to the present day stress field (Hancock & Engelder 1989). Structural and stratigraphic controls are known to occur in the chalk in southern England (Ameen 1990; Clayton 1990) affecting the joint patterns formed. Toynton did not address this point in his study of Norfolk. Recent re-analyses of existing gravity and magnetic data by the British Geological Survey have also shown the existence of deep seated lineaments (NW-SE, NE-SW, ENE-WSW, E-W) within the concealed Caledonide basement of eastern England (Pharaoh et aI. 1987; Lee et al. 1990, 1991). No manifestation of these structures is evident, superficially, in the Mesozoic and Cenozoic units. The fracture study reported here covers the Upper Cretaceous Chalk of eastern England, the counties of Norfolk, Lincolnshire, north Cambridgeshire and south Humberside (Fig. 1). Prolific joints are visible in an essentially undeformed Mesozoic sedimentary sequence that is fiat lying or gently dipping (< 10~ to the NE. The study reviews the occurrence of joint patterns in eastern England extending the existing studies from the adjacent areas. The collection of outcrop data has allowed a classification of identified systematic joints into individual joint sets (Hancock 1985). The analysis of these data have allowed the parallelism and continuity to be tested between the eastern England joints and those further to the south and any previously unrecognized controls and variations on fracturing to be defined. Insights can also be made into the formation environment and age of the joint sets and their previous regional tectonic interpretations.

Methods and techniques The field data collected from the study area comprise joint orientation, spacing, geometry, architecture, dihedral angle, continuity, dilation and surface morphological features. The orientation data of Toynton (1983) have augmented these new data particularly for recently infilled quarries. Bias in orientation readings due to sampling procedure was minimized by collecting readings from faces of varying orientation. Toynton has shown that 50 readings are sufficient to identify the significant trends within the data set. All orientation data have been analysed using the 'STEREOPLOT' software (Mancktelow 1989) and cluster analysis software (SPSS 1990). Based on the relationships with other planes, and the similarity of the criteria stipulated above, systematic joint sets were identified and individual joint planes were assigned to a joint set whilst at a field locality. Equal area stereonet scatter plots (dip and dip azimuth) of both a standard lower hemisphere projection and a rotated lower hemisphere projection (rotated about a horizontal axis of minimum point density) confirmed the field-based observations. Single-linkage cluster analysis producing a heirarchy of cluster solutions based on the squared Euclidean distance between variables (data points) enabled the significant cluster(s) to be identified from any background noise occurring in each set at every locality. This process also identified a mean plane for each joint set. Once the set was identified, statistical parameters such as the standard distribution could be calculated. Distribution analyses related to the point density on the sphere, primarily the Bingham analysis (Cheeney 1983; Fisher et al. 1987) confirmed the mean plane. This paper concentrates on presenting the variations seen in the field criteria and characteristics of the joint sets identified, not the secondary statistical differentiation of the data and joint sets. Average planes and joint set data are presented for representative localities.

200

A.M. GOODWIN

Grasby i (65)

Fox Coven

(83)

(44)

(77) Tefford Hill

20 I

0

" ~ (73)

I

Candlesby Hill

Km

(58)

Hunstanton _ r .

North Ormsby'

(68)

~

BurnhamOvery (88)

Wells-Next-the-Sea

(122) Mill Hilt

(72)

Lower Hcllesdon

(41)

Snettisham

(108) #

Gayton

(64)

1

Whittingto

(40)

Non wo#:'" (40)

") .~/

Soo Pieken m (41)

Caistor

(118)

Fig. 2. Lower hemisphere equal area projections of collected data. For each locality scatter plots of poles to joint planes are shown with an accompanying plot of the mean plane for the dominant joint sets. The NW joint set is highlighted in each case, the number represents the sample size (N). Joint patterns Joint sets, represented by their mean plane, are shown in Fig. 2. The existence of two clusters on the stereonets is clear, correlating with the division of joint planes into two similar sets in the field (Fig. 3). The summary plots mostly show steeply dipping N W and NE striking sets of similar orientation to those in southern England. This relationship is clearest on a horizontal surface perpendicular to the joint planes effectively eliminating the varying geometric styles of set members. Both joint sets comprise steep conjugate and vertical joints, everywhere consistent with bedding. The N W set is generally, but not consistently, the more prominent of the two (Table 1). Fractures showing small amounts of dip-slip displacement (< 15 cm) form a minor division of the N W set at only a few localities. N o surface morphological features are seen on any joint planes. Cross-cutting and abutting relationships between (and within) the two sets are ambiguous. Little variation in jointing is seen between the different stratigraphic horizons of the Chalk

CHANGING JOINT GEOMETRIES IN CHALK

201

202

A.M. GOODWlN

CHANGING J O I ~ F GEOMETRIES IN CHALK

203

Fig. 3. Photographs showing jointing in vertical sections from Humberside, Lincolnshire and Norfolk. Examples of structures are shown in each photograph where: j, joint plane; b, bedding plane; c, conjugate joint plane (steeply inclined); e, extension joint plane (vertical). (A) Dominance of vertical joints (dip > 85~ on a vertical SW-NE face in Burnham, south Humberside (hammer is 40 cm). (B) Vertical SW-NE cliff face showing an increase in the number of steeply inclined joints at Mill Hill, south Lincolnshire (pen is 15 cm). (C) Vertical SW-NE cliff face showing the dominance of steeply inclined joints at East Harling, Norfolk. (D) Close-up of the conjugate nature of steeply inclined joints at East Harling, Norfolk (compass is 10cm). (E) Dominant joint traces on a horizontal surface are the NW trending set, West Runton, North Norfolk coast. Scale bar graded in 1 and 0.5cm. (Upper, Middle and Lower). Minor discrepancies in the density of jointing can be seen between different localities, higher intensities of jointing being seen in 'harder' chalk horizons (in the present study mainly in localities in the Middle Chalk). Likewise, digressions in the nature of jointing - orientation, architecture and geometry - are also apparent but again are only slight. A more significant difference in joint set characteristics can be observed on a regional scale. Although both the N W and NE striking joint sets are apparently continuous through the eastern England Chalk, this continuity hides a subtle, but equally important, change in their detailed geometry across the area. This change is illustrated with respect to the more prevalent N W striking joint set (Fig. 4a). The N E striking set follows an equivalent pattern. 1. The southern area. In Norfolk, joint traces on a vertical N E - S W section typically

show a conjugate, steeply inclined form with a dihedral angle of between 25 ~ and 45 ~ Vertical joints occur less commonly. Displacement on joint planes is not evident in

204

A . M . GOODWIN

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A.M. GOODWIN the field (excepting the minority of recognizable fractures) and microscope analysis has not proved effective in detecting any shear displacement. Joint traces terminate at horizontal joints, with an overall length of c. 20-40 cm. Ten per cent of the vertical joints have a continuous trace across these horizontal joints of up to several metres. Dilation across joint planes is small ( < 2 mm), except for the more continuous traces where it can reach 5-6 cm. . The northern area. By contrast, vertical N E - S W sections across the joint sets in south Humberside/north Lincolnshire show mainly vertical joints ('H' and T architectural traces of Hancock 1985) with few steeply inclined planes occurring. As in Norfolk, the vertical joints are truncated at the horizontal joints, although c. 10% are more continuous and these joint traces show the greatest dilation.

The origin and stress implications o f the joint patterns The joints of Norfolk (the southern area) are interpreted as the conjugate hybrid joints of Hancock (1985) showing typical 'x' and 'y' architectural traces. They would have formed in a stress field with the maximum principal stress (0-1) vertical and the minimum principal stress (0-3) horizontal and perpendicular to joint strike (principal stresses are assumed to be compressive when positive). The effective stress component acting normal to the failure plane (when the dihedral angle is < 45 ~ would be < 0 ( Price & Cosgrove 1990), allowing for an element of shear along the fracture plane as well an element of extension across the plane - hence the term 'hybrid'. The lack of detection of any component of shear may simply be a factor of the absence of any suitable attribute to record any occurrence of shear (fractured grains, displaced microscopic veins). The Humberside joints are interpreted as extension joints (dihedral angle < 10 ~ formed perpendicular to the minimum principal stress (0-3). The effective minimum stress would, necessarily, have to be < 0, possibly achieved through the effect of pore-fluid pressure counteracting a compressive 0-3. Any fluid present in a sediment will, on compression, counteract the applied stress because of its incompressible nature. This will decrease both 0-1 and 0"3 but not the differential stress (0-z-o3), effectively weakening the Chalk, instigating failure (Fyfe et aL 1978). If 0-3 is lOW or the hydrostatic pressure is large enough, the effective 0-3 can become zero or negative giving rise to tensile brittle failure. There appears to be a transition between the conjugate hybrid joints in the south and the extension joints in the north, with localities in the intervening south Lincolnshire area showing both types of trace (Fig. 4a). In both south Humberside and Norfolk, the joints form a joint continuum rather than a singular joint set. Hancock (1986) suggests that systematic joints can form either a well-defined fracture set, or they may form a coaxial spectrum of joints (including both hybrid and extension joints) describing an angular continuum with a maximum dihedral angle of c. 45 ~ The dominantly conjugate joint set of Norfolk and the singular set of Humberside are therefore seen as end members of a transition from south to north through a single joint continuum. The minor variation observed in jointing within different stratigraphic horizons of the Chalk reflects the similar N - S control of joint geometry - 'harder' horizons show the same characteristics as the northern area. Similarly, a control on fracturing by local features and the reactivation of basement features is not clearly seen. No major structural features, folds and faults, occur in this area to act as a controlling influence. Other than a superficial correlation of trend to identified basement lineaments (NW and SE), a lack of information

CHANGING JOINT GEOMETRIES IN CHALK

207

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208

A.M. GOODWlN

on the deep structures and a paucity of surface exposure coverage does not allow conclusive suggestion of a link between the two.

Causes of the spatial change in joint geometry It is suggested here that there are two possible mechanical explanations for the change from conjugate to extension joint geometry from Norfolk to south Humberside. These are explained with reference to Mohr stress diagrams (Fig. 5a). They are not mutually exclusive and the natural explanation may involve components of both mechanisms.

Variable rock strength hypothesis The first possibility is that the differential stress field (0-1- 03) remained constant across the region but that the rock strength increased from south to north. The weaker rock would have a Mohr failure envelope which lay within that of the stronger rock with a less steep gradient (smaller angle of internal friction; Fig. 5a). For a range of differential effective stresses, the resulting stress circle would touch the failure envelope of the weak rock at two points, giving conjugate hybrid joints, yet touch the envelope for the strong rock at only one point, giving extension joints. The minimum effective principal stress (0-3) would need to be < 0. The northward increase in rock strength and hardness in the Chalk is well known (e.g. Scholle 1974; Bell et al. 1990). In the mechanically variable Chalk of England induration, stylolitization and density all increase northwards, whilst permeability and porosity decrease (Table 2). The Chalk of southern Britain is thus considered to be 'softer' than that occurring further north in Lincolnshire and Yorkshire. A definition of strength terms can be found in Mortimore & Fielding (1990). The observed northwards hardening is thought to be primarily due to a variety of diagenetic factors (Clayton 1983, 1990; Clayton & Matthews 1987), influenced locally by tectonically induced deformation (Mimran 1975). Representative failure envelopes for high porosity (soft) and low porosity (hard) English Chalks are shown, for comparison, in Fig. 5b. Scholle (1974) and Carter & Mallard (1974) have suggested that the changes in geotechnical properties of the Chalk are due to the controlling processes of either consolidation ('compression associated with over pressuring or positive excess pore pressures') or drained compression ('compression controlled by the rate of application of an overburden in the absence of excess pore pressures') associated with an increasing burial depth (increasing northwards). Clayton & Matthews (1987), however, suggest that these processes were insignificant contributors to the confined deformation and densification of the Chalk and that excess pore fluids were unimportant (Jones et al. 1984). If excess pore pressure was not a significant factor at burial and consolidation of the Chalk the minimum compressive stress (0-3) would ;need to be low or < 0, therefore requiring only a minimal effect of pore pressure to promote extensional failure.

Variable stress field hypothesis The second hypothesis to explain the observed joint variation is that the rock strength remained constant across the region, but that the differential stress (0-1-0-3) decreased from south to north, Fig. 5a. A large stress circle would touch any Mohr envelope at two points, giving conjugate joints, whereas a small stress circle would touch the same envelope

CHANGING JOINT GEOMETRIESIN CHALK

209

Table 2. Density, porosity and strength values for the Chalk of eastern England showing the N - S variation highlighted in Fig. 4b

Yorkshire

Norfolk

Kent

Dry density (mg m-a) Bell et al. 1990 Clayton 1983

2.06-2.14 1.97-2.25

1.76-2.17 1.54-2.19

1.44-1.61 -

Absolute porosity (%) Bell et al. 1990 Scholle 1974

24.4-26.8 1-10

26.6-35.1 -

40.6-47.0 43

Compressive strength (MPa) Bell et al. 1990

25.6-30.7

13.0-29.2

5.5-9.5

1.7-2.2

0.8-2.1

0,5-0.7

Tensile strength (MPa) Bell et al. 1990

at only one point in the tensile zone. A greater stress difference in the south would require a larger effective tr~ than in the north. Because crl is vertical, inferred from the geometry and architecture of the joint planes, this implies a greater overburden in the south. There is no obvious geological evidence for this greater overburden. Indeed, the higher porosity and lower density of the soft 'southern' Chalk has been taken to indicate a lower overburden to the south. A high differential stress also implies that the rock mass has a similarly high degree of internal strength, soft rocks flowing to dissipate stress evenly into a hydrostatic state and to minimize the difference between crland a3 (Engelder 1993). Increasing the differential stress to the south would then suggest an associated increase in Chalk strength rather than the observed decrease.

Discussion It follows that the most likely explanation for the spatial variation in the geometry of jointing in eastern England is variable mechanical properties of the Chalk, namely an increase in strength and hardness northwards (Table 2, Fig. 4b). The strength variation seen is directly correlated to changes in density and porosity and accompanying textural differences (Fig. 4b) (Clayton 1983; Jones et al. 1984). Variations in initial porosity and matrix behaviour have been shown to affect the stress-strain path of a sample significantly (Clayton & Matthews 1987; Clayton 1990). This sedimentological difference between the northern and southern Chalks may be the underlying factor in the changing hardness/ strength and therefore the joint geometry. Increasing the tensile and cohesive strength of the Chalk will increase the internal angle of friction and so delimit a broader Mohr failure envelope (Bell et al. 1990; Clayton 1990) influencing the mechanisms of deformation of the rock. Despite a paucity of evidence for a northward increase in overburden, however, the diagenetic gradient within the Chalk from south to north alone infers a greater depth of burial northwards (Scholle 1974). Variations noted in the lithological horizons of the chalk reflect the broader scale N-S change and is similarly linked to the hardness of the individual horizon. Chalk is a marine deposit, the originally deposited sediment having a very high porosity,

210

A.M. GOODWIN

possibly c. 80%, and presumably very fluid-rich. Although the differential stress (0.1- 03) might have remained constant across the area, the effective stresses would have been reduced due to an increase in pore pressure with burial, however small a contributing factor. Extension failure in the north implies an effective 0"3 that is negative or equal to zero with the pore pressure exceeding the minimum compressive stress by the tensile strength of the rock. Excess hydrostatic pressures would have been feasible if depositional and diagenetic fluids were partially trapped by a low permeability layer. Two possible explanations could account for the increase in pore-fluid pressure within the Chalk: 1. There existed above the Chalk an unknown amount of Upper Cretaceous and early Tertiary overburden. The Chalk is unconformably overlain in the southern North Sea by argillaceous Palaeocene sediments. Following early Palaeocene deposition, a major regression and widespread uplift led to extensive erosion (possibly in the region of 200m or more) of these and Upper Cretaceous sediments (Ziegler 1987; Cameron et al. 1992; Cope et al. 1992). The precise amount of Cretaceous and early Palaeocene sediment that was removed is uncertain. On this hypothesis, the joint sets would be due to the stress regime before the early Palaeocene uplift of the North Sea, and therefore be of latest Cretaceous or early Palaeocene age. 2. An effective overburden was provided by later Palaeocene and Eocene sediments unconformably overlying the Chalk, equivalent, for example, to the late Palaeocene Lista Formation (Cameron et al. 1992, figs 87 & 88). Post-early Palaeocene subsidence deposited upper Palaeocene and Eocene basinal sediments up to 2000 m thick, averaging c. 1000m (Ziegler 1987; Cameron et al. 1992; Cope et al. 1992). Much of northern, central and western Britain was at a structural high during this period, shedding sediment eastwards into the North Sea Basin. However, the extent of onlap of late Palaeocene and Eocene sediments on to onshore Britain is unclear (Cope et al. 1992). On the hypothesis of significant post-Palaeocene overburden, the joint sets would then have formed prior to or during the uplift associated with the Alpine orogeny. An additional factor in either of the above hypotheses is the relatively late Cretaceous impermeability of the Chalk, its present porosity and permeability being mostly fracture related. This impermeability may have been sufficient to form an effective seal without the necessity of any overlying impervious unit. However, flow within the Chalk could have been appreciable due to the deposits high original porosity possibly allowing only small hydrostatic pressures to arise in the Chalk. Even if increased pore pressure was not a significant factor in the brittle failure of the Chalk a small 0.3, as postulated by Bevan (1985), may have been low enough to fall into the tensile field of the Mohr diagram by only a minimal amount of hydrostatic pressure. The joint geometries in eastern England suggest that the Chalk there was overlain by an overburden, probably of missing Upper Cretaceous and earliest Tertiary sediments, of substantial thickness. The vertical application of an overburden would have had the effect of increasing the vertical principle stress to become 0.1, as implied by the joint geometries. Bevan (1985) postulated that 150 m of unconsolidated sediment would have been sufficient to achieve this. Clayton & Matthews (1987)estimated a possible past maximum vertical effective stress of 2000 k N m -2, equivalent to 200 m of overburden (based on the average thickness of post-Cretaceous sediments plus 150 m of ice). They note that this estimate would be greater than all past in situ vertical effective stresses, tectonic and diagenetic

CHANGING JOINT GEOMETRIES IN CHALK

211

processes hardening even the softest Chalk by some degree. Clayton & Matthews (1987) suggest that the most likely possibility for the maximum past overburden would be the deposition of c. 240m of early Tertiary sediments (Thanet sands, Reading Beds and London Clay). Although it is possible that enough overburden existed prior to the uplift and erosion during the early Palaeocene, it is more likely that a greater thickness of overburden of late Palaeocene-Eocene sediments prior to uplift associated with Alpine compression was responsible for the main episode of jointing in the Chalk. The effect of the Pleistocene glaciations cannot be ignored, as Bevan (1985) noted, a thickness of < 150 m of (denser) ice would have the same effect as a sediment pile of 150 m. As can be seen from Fig. 4b, the southern limit of the hard Chalk could correlate to a glacial ice front(s). It would be unreasonable to assume, especially in eastern England, that the sequence of glacial overburdens exacted no toll on the underlying solid rock. However, distinguishing such an effect from that of an earlier sedimentary overburden would prove difficult. The fact that both joint sets also occur in 'unglaciated' southern England suggests that their formation probably pre-dates the Pleistocene glaciations, with possibly no effect other than re-activation being due to any ice overburden.

Conclusions 1. The N E - S W and N W - S E striking joint sets previously recognized in southern England have been traced northwards in the Chalk of eastern England as far as the Humber Estuary. 2. Both sets exhibit a change in their joint plane geometry northwards from dominantly steep conjugate hybrid joints in Norfolk to sub-vertical extension joints in Humberside. 3. This change might be due to a decreasing differential stress towards the north, but more likely is an association with the observed northwards increase in rock strength and hardness in the Chalk. 4. The higher strength, lower porosity and higher density of the northern Chalk may record a greater overburden here, which may have included impermeable clays to seal the Chalk. This would generate the high pore-fluid pressures necessary for extensional failure. Alternatively, if o3 was small only a minimal effect of hydrostatic pressure would be necessary to cause the formation of the extension joints seen. 5. Stratigraphic and structural controls on jointing are acknowledged as possible influences on joint geometry but considered to play a minor role superimposed on the broader regional variation.

This work has been completed whilst in receipt of a three-year postgraduate Natural Environment Research Council award. I would like to thank Nigel Woodcock and Ben Goodwin (amongst others) for their thoughts and help towards completion of this manuscript.

References AMEEN,

M. S. 1990. Macro-folding in the Purbeck-Isle of Wight monocline. Proceedings of the

Geologists" Association, 101, 31-46. BALSON, P. S. & HUMPHREYS, B. 1986. The nature and origin of fissures in the East Anglian Coralline and Red Crags. Journal of Quaternary Science, 1, 13-19. BELL, E. G., CRIPPS,J. C., EDMONDS,C. N. & CULSHAW,M. G. 1990. Chalk fabric and its relation to

212

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certain geotechnical properties. In: Chalk: Proceedings of the International Chalk Symposium. Thomas Telford, London, 187-194. BEVAN, T. G. 1985. A Cenozoic tectonic history of Southern England inferred from mesofractures. PhD Thesis, University of Bristol, UK. & HANCOCK,P. L. 1986. A late Cenozoic regional mesofracture system in southern England and northern France. Journal of the Geological Society, London, 143, 335-362. CAMERON, T. D. J., CROCBY, A., BALSON, P. S., JEFFERY, D. H., LOTT, G. K., BULAT, J. & HARRISON, D. J. 1992. The geology of the southern North Sea. United Kingdom Offshore Regional Report, No. 7. British Geological Survey. CAPUTO, R. 1991. A comparison between joints and faults as brittle structures used for evaluation of the stress field. Annales Tectonicae, 5, 74-84. CARTER, P. S. • MALLARD, D. J. 1974. A study of strength, compressibility and density trends within the Chalk of south-eastern England. Quarterly Journal of Engineering Geology, 7, 43. CHEENEY, R. F. 1983. Statistical Methods in Geology for Field and Laboratory Decisions. Allen & Unwin, London. CLAYTON, C. R. I. 1983. The influence of diagenesis on some index properties of Chalk in England. G~otechnique, 33, 225--241. 1990. The mechanical properties of the Chalk. In: Chalk: Proceedings of the International Chalk Symposium. Thomas Telford, London, 213-232. & MATTHEWS, M. C. 1987. Deformation, diagenesis and the mechanical behaviour of Chalk. In: JONES, M. E. & PRESTON, R. M. F. (eds) Deformation of Sediments and Sedimentary Rocks. Geological Society, London, Special Publication, 29, 55-62. COPE, J. W., INGHAM, J. K. & RAWSON, P. F. 1992. Atlas of Palaeogeography and Lithofacies. Memoir of the Geological Society, London, 13. DUNNE, W. M. & HANCOCK, P. L. 1994. Palaeostress analysis of small-scale brittle structures. In: HANCOCK, P. L. (ed.) Continental Deformation. Pergamon Press, Oxford, 101-120. ENGELDER, T. 1982. Is there a genetic relationship between selected regional joints and contemporary stress within the liothosphere of North America? Tectonics, 1, 161-177. - 1993. Stress Regimes in the Lithosphere. Princeton University Press, N J, USA. & GEISER, P. A. 1980. On the use of regional joint sets as trajectories of palaeostress fields during development of the Appalachian Plateau, New York. Journal of Geophysical Research, 85, 6319-6341. EYAL, Y. & RECHES, Z. 1983. Tectonic analysis of the Dead Sea Rift regions since the lateCretaceous based on meso-structures. Tectonics, 2, 167-185. FISHER, N. I., LEWIS, T. L. & EMBELTON, B. J. J. 1987. Statistical Analysis of Spherical Data. Cambridge University Press, Cambridge, UK. FYFE, W. S., PRICE, N. J. & THOMPSON, A. B. 1978. Fluids in the Earth's Crust. Elsevier, Amsterdam. HANCOCK, P. L. 1985. Brittle micro-tectonics: principles and practice. Journal of Structural Geology, 7, 437-457. - 1986. Joint spectra. In: NICHOL, I. & NESBITT, R. W. (eds) Geology in the Real World- The Kingsley Dunham Volume. Institution of Mining and Metallurgy, London, 155-164. - d~ ENGELDER,T. 1989. Neotectonic joints. Bulletin of the Geological Society of America, 101, 1197-1208. --, AL KADHI, A. & SHA'AT, N. A. 1984. Regional joint sets in the Arabian platform as indicators of intra-plate processes. Tectonics, 3, 27-43. JONES, M. E., BEDFORD,J. d~ CLAYTON,C. 1984. On natural deformation mechanisms in the Chalk. Journal of the Geological Society, London, 141, 675-683. LEE, M. K., PHARAOH, T. C. & GREEN, C. A. 1991. Structural trends in the concealed basement of eastern England from images of regional potential fields. Annales Socidtd G~ologique Belgique, 114, 45-62. , PHARAOH, Z. C. & lOPER, N. J. 1990. Structural trends in central Britain from images of gravity and aeromagnetic fields. Journal of the Geological Society, London, 147, 241-258. MANCKTELOW, N. 1989. STEREOPLOT, Version 1.2 Faceware Software, FaceIt REsources. MIMRAN, Y. 1975. Fabric deformation induced in Cretaceous Chalks by tectonic stresses. Tectonophysics, 26, 309. MORTIMORE, R. N. & FIELDING, P. M. 1990. Relationships between texture, density, and strength of Chalk. In: Chalk: Proceedings of the International Chalk Symposium. Thomas Telford, London, -

-

-

-

-

-

CHANGING JOINT GEOMETRIES IN CHALK

213

109-132. PATSOULES, M. G. & CRIPPS, J. C. 1990. Survey of macro- and micro-fracturing in Yorkshire chalk. In: Chalk: Proceedings of the International Chalk Symposium. Thomas Telford, London, 87-94. PHARAOH, T. C., MERRIMAN, R. J., WEBB, P. C. & BECKINSDALE, R. D. 1987. The concealed Caledonides of eastern England: preliminary results of a multi-disciplinary study. Proceedings of the Yorkshire Geologists Society, 46, 355-369. POLLARD, D. n. & AYDIN, A. 1988. Progress in understartding jointing over the past century. Bulletin of the Geological Society of America, 100, 1181-1204. PRICE, N. J. & COSGROVE, J. W. 1990. Analysis of Geological Structures. Cambridge University Press, Cambridge, UK. SCHOLLE,P. A. 1974. Diagenesis of upper Cretaceous Chalks from England, Ireland and the North Sea. In: Hsu, K. J. & JENKYNS,H. C. (eds) Pelagic Sediments: On Land and Under the Sea. Special Publication of International Association of Sedimentologists, 1, 177-210. SPSS 1990. SPSS for the Macintosh, Version 4.0 SPSS Inc., Language Systems Corporation, Chicago. TOYNTON, R. 1983. The relationship between fracture patterns and hydraulic anisotropy in the Norfolk Chalk, England. Quarterly Journal of Engineering Geology, 16, 169-185. VANDYCKE, S., BERGERAT, F. dk DUPUIS, Ch. 1991. Meso-Cenozoic faulting and inferred palaeostresses in the Mons Basin, Belgium. Tectonophysics, 192, 261-271. ZIEGLER, P. A. 1987. Late Cretaceous and Cenozoic intra-plate compressional deformtion in the Alpine foreland - a geodynamic model. Tectonophysics, 137, 389-420. ZOBACK,M. L. 1992. First- and second-order patterns of stress in the lithosphere; the world stress map project. Journal of Geophysical Research, 97B, 11 703-11 728.

From Ameen, M. S. (ed.), 1995, Fractography." fracture topography as a tool in fracture mechanics and stress analysis

Geological Society Special Publication No. 92, pp. 215-233

Factors controlling joint spacing in interbedded sedimentary rocks: integrating numerical models with field observations from the Monterey Formation, USA MICHAEL

R. GROSS, 1 M A R K P. F I S C H E R , 2 T E R R Y E N G E L D E R 2 & R O Y J. G R E E N F I E L D 2

1Department of Geology, Florida International University, Miami, FL 33199, USA 2Department of Geosciences, Pennsylvania State University, University Park, PA 16802, USA Abstract: Local tensile stress normal to a joint is reduced in the vicinity of the joint because such stresses are not transmitted across free surfaces. This stress reduction prevents the formation of new joints in the vicinity of existing joints, and thus influences joint spacing. Lateral extent of this stress reduction shadow increases with joint height, which corresponds to bed thickness for many sedimentary rocks. The linear correlation between joint spacing and bed thickness commonly observed in outcrop is a direct result of this relationship. However, other factors in addition to bed thickness influence joint spacing. We evaluate these factors through both a review of the Hobbs model for joint spacing and a 2D finite element simulation of a crack confined to a lithology-controlled mechanical unit. The stress reduction shadow increases in length with increasing Young's modulus of the jointing bed, though fracture stress, flaw size, flaw distribution and extensional strain all interact with bed thickness and elastic properties ultimately to control joint spacing. One explanation for the observed decrease in joint spacing with increasing Young's modulus in outcrops of the Monterey Formation is that beds with higher Young's moduli fail at lower magnitudes of extensional strain.

In a 1967 issue of Geological Magazine, D. W. Hobbs published a paper entitled 'The formation of tension joints in sedimentary rocks: an explanation', which is commonly cited as a theoretical explanation for joint spacing in rock masses composed of layers, each with a different set of elastic properties. Based on the original derivation of Cox (1952), the Hobbs model provides one theoretical basis for the well-documented linear relationship between joint spacing and bed thickness in sedimentary rocks (e.g. Price 1966; McQuillan 1973; Ladeira & Price 1981; Huang & Angelier 1989; Narr & Suppe 1991; Gross 1993a). By relating joint spacing to the Young's modulus of the jointing bed and shear modulus of the bounding non-jointing beds, Hobbs implies that joint spacing is influenced by the elastic properties of each bed. Though both ingenious and remarkably accurate in its prediction of stress distribution, the Hobbs paper is often misunderstood because it is compact and lacks illustrations. We present a brief explanation of Hobbs' analysis in an effort to clarify the assumptions and implications of his model. In addition, we use a 2D finite element numerical simulation of a crack confined within a bed to describe the stress and displacement fields around the profile section of a crack in an interbedded sedimentary rock mass. Results of the finite element modelling are compared with predictions of the Hobbs model. Finally, we discuss

216

M.R. GROSS E T

AL.

Fig. 1. Sketch of a lithology-controlled mechanical unit, with systematic joints confined to discrete beds, and lithologic contacts serving as mechanical layer boundaries. Note that joint height equals mechanical layer thickness (MLT), and is proportional to joint spacing (from Gross 1993a).

Fig. 2. Photograph of steeply dipping limestone beds in the Monterey Formation exposed at Lion's Head along the Santa Maria coastline of California, demonstrating the relationship between joint spacing and bed thickness. Tall joints, in thicker beds, are more widely spaced than short joints, in thinner beds. Note the consistent joint height for each individual mechanical layer.

FACTORS CONTROLLING JOINT SPACING

217

Fig. 3. Photograph of interbedded siliceous (jointing) and mudstone (non-jointing) layers in the Monterey Formation exposed at Purisima Point along the Santa Maria coastline. The light coloured siliceous layers contain statistically regularly-spaced joints that terminate at contacts with adjacent mudstone layers.

the major factors controlling joint spacing and attempt to reconcile these factors with field observations that some stiffer beds in the Monterey Formation of California have more closely spaced joints than beds with lower Young's moduli.

Stress reduction near joints Joint spacing models for rocks commonly focus on an analysis of stress reduction in the vicinity of joints because the criterion for joint propagation is a threshold stress value, derived either from the tensile strength test (e.g. Jaeger & Cook 1976) or a measure of fracture toughness (e.g. Irwin 1957) of the rock bed. The mechanism for stress build up in these models is layer-parallel stretching, which can occur locally in the extensional region of a fold or regionally upon uplift and erosion. Stress perturbations occur in the vicinity of an open joint because the joint is a free surface, and hence acts as a barrier to the transmission of tensile stress. Consequently,

218

M.R. GROSS E T A L .

Fig. 4. Photomicrograph of a thin section from core taken from Point Arguello Oilfield, offshore Santa Maria Basin, showing a tensile fracture confined to a quartz chert layer. See text for details.

local crack-normal tensile stress at the joint surface is zero, and increases to the level of far-field tensile stress at an infinite distance from the joint. New joints form where the local tensile stress exceeds the fracture stress of the rock. Thus, models for joint spacing commonly predict the location, and hence spacing, of new joints by evaluating the ratio of local tensile stress to remote tensile stress as a function of distance from the stress-free joint (e.g. Lachenbruch 1961; Pollard & Segall 1987).

Bed thickness - joint spacing relationship and sequential infilling Joints in sedimentary rocks generally fall into two categories, those that terminate randomly within the rock mass and those that terminate at discrete mechanical layer boundaries. Lithologic contacts, as well as pre-existing fractures, can serve as mechanical layer boundaries, thereby dividing the rock mass into discrete mechanical units (Gross 1993a). The analysis presented in this study focuses on joints confined to individual lithologic beds, which are frequently observed in thin to medium bedded sedimentary

FACTORS CONTROLLING JOINT SPACING

219

Fig. 5. Location map of the southern Santa Maria coastline showing outcrops where fracture data were collected along with the photographs from Figs 2-4 (adapted from Sylvester & Darrow 1979). Sites include Lion's Head (LIO), Purisima Point (PUR) and Lompoc Landing (LOM). Towns are Santa Maria (SM) and Lompoc (L). SYRF refers to the Santa Ynez River Fault.

sequences (i.e. beds ranging in thickness from c. 1 to 100 cm) composed of diverse rock types. In such lithology-controlled mechanical units joints extend from the bottom of the bed to its top so that joint height equals bed thickness, which in turn is proportional to joint spacing, measured as the orthogonal distance between adjacent parallel joints of a particular set (Fig. 1). These relationships are demonstrated in the photographs in Figs 24, taken from the Monterey Formation along the central California coastline (Fig. 5). In the interbedded limestone and shale of Fig. 2, joint height in each mechanical layer is uniform, and tall joints are more widely spaced than short joints. The outcrop shown in Fig. 3 consists of a series of well-defined jointing and non-jointing lithologies, in which joints are confined solely to the light coloured siliceous layers and display a regular spacing. In contrast, the thin recessive shaley layers act as mechanical layer boundaries and remain unjointed. An example of a joint confined to a thin quartz chert layer is shown

220

M . R . GROSS E T AL.

in Fig. 4. Note the joint terminates abruptly at the contact with adjacent shale layers. Although slip has occurred in the immediate vicinity of the crack, in general the widest contact does not serve as a detachment. Though numerous statistical functions, ranging from negative exponential to normal, have been proposed to describe the general distribution of joint spacing (e.g. Priest & Hudson 1976; Rouleau & Gale 1985; Rives et al. 1992), the specific case of lithologycontrolled joints leads to a typically skewed log normal or gamma distribution for joints belonging to an individual systematic set (e.g. Huang & Angelier 1989; Narr & Suppe 1991). The regular spacing may be due to the stress reduction shadow created by preexisting joints (e.g. Hobbs 1967) and the sequential joint infilling process, whereby new joints form in a sequential fashion between adjacent pre-existing joints (Gross 1993a). The non-normal distribution may result from the random distribution of flaws, which serve as initiation points for joints. While other studies focus on interactive joint growth (e.g. Pollard et al. 1982; Olson & Pollard 1989), we are concerned with identifying where new joints will form with respect to pre-existing joints along the profile section of a layer. Therefore, throughout this paper we consider the actual joint propagation process as instantaneous relative to the time between sequential infilling stages of a systematic joint set. Material science studies provide evidence in favour of a sequential infilling mechanism for joints confined to lithology-controlled mechanical layers. Tension tests performed on layered ceramics demonstrate that in addition to an increase in crack spacing with layer thickness, average crack spacing decreases systematically with increasing applied stress (Garrett & Bailey 1977; Parvizi & Bailey 1978). The applied stress in turn scales linearly with extensional strain for the elastic laminates. The decrease in crack spacing with increasing load occurs by a process of sequential infilling. Geological evidence supporting sequential infilling includes fractured piedmontite grains (Masuda & Kuriyama 1988) and curving cross joints (Engelder & Gross 1993).

The Hobbs model for joint spacing The 1D Hobbs model consists of three discrete beds and predicts variations in the cracknormal tensile stress as a function of differences in elastic properties between a middle jointing bed and two non-jointing bounding beds. The beds behave elastically and are bonded together to prevent slip along the interfaces. Unlike the non-layered models of Lachenbruch (1961) and Pollard & Segall (1987), stress perturbations in a bedded medium are influenced by both elastic properties and shear stresses that develop along bed interfaces. According to our interpretation, the Hobbs model consists of a middle 'jointing bed' (Bed A) sandwiched between 'non-jointing beds' (Beds B). Hobbs assumes both bounding beds have the same elastic moduli, and that the thickness of bounding beds is much greater than the thickness (d) of the jointed bed. After initial extension all of the beds remain unjointed, however, as extensional strain increases, the tensile stress in the middle bed eventually exceeds the tensile strength of that bed, and a joint propagates. At this point Hobbs assumes that tensile stress cannot be transmitted across the joint surface. Shear stresses are produced within the bounding beds as a direct result of differences in elastic displacement within the jointed bed and bounding beds. The key parameter controlling joint spacing is the distribution of longitudinal tensile stress (~rxx) within Bed A after jointing, which increases from zero at the joint surface to a value equivalent to the remote stress at an infnite distance away from the joint.

FACTORS CONTROLLINGJOINT SPACING

221

Substituting Hobbs' equation 13 into equation 6 and rearranging the terms produces an equation that evaluates normalized crack-normal tensile stress after joint propagation (i.e. local stress divided by remote tensile stress) as a function of distance from the joint; P E A de1

ff local - -

1 + sinh ~

~

(1)

- cosh--d-- -~A

t7 r e m o t e

where P is the applied load, s is the strain at which the first joint forms, d is the jointing bed thickness (i.e. joint height), and EA and GB are the Young's modulus and shear modulus of the jointing and non-jointing beds, respectively. We prefer to rewrite equation (1) in the form: P - - = E A del

f(x/6),

(2)

,

(3)

where, = ~

~

and f ( x / 5 ) gives the dependence of the local stress on x. The function f ( x / 6 ) is 0 for x = 0 and increases to 1 for large x. The quantity 6 has units of length and can be thought of as a decay length, or the distance over which the stress shadow decays back to a fixed fraction of the remote stress. The distance 5 is proportional to d and to the square roof of (EA/GB).

1.21

..I

. . . . .

, , ,

I

I

I

I

I

I

I

I

1.0 -! .........EA/ GB= 1 .............. I.M

ml_ N

~

.~ o

0.8 0.6

EA/GB=20

"~ ,,.J o r162 o 0 . 4

i

rr" O =,

lEA/GB=(1, 2, 215, 5, 10, 20) I 0.2

0.0 0

1

2

3

4

5

NORMALIZED DISTANCE (x/d) Fig. 6. Predicted plots of stress reduction shadows based on the Hobbs model for various EA/Ga ratios.

222

M . R . GROSS E T A L .

Fig. 7. Finite element mesh and boundary conditions. Models were subjected to uniform longitudinal extension, c. Inset shows an enlarged portion of mesh around joint in Bed A. Bed B is shale in all models whereas the material properties of Bed A are varied. Joint shown slightly open for illustration. Cartesian coordinate system with (0, 0) at the centre of the joint is used for all related graphs and figures (after Fischer 1994).

A plot of normalized stress v. normalized distance from a joint for a series of EA/GB ratios is presented in Fig. 6. Because Hobbs' analysis predicts that for a fixed shear modulus of the bounding beds (i.e. holding GB constant), 6 will increase as the square root of EA, a high EA will produce a long stress reduction shadow, whereas a low EA creates a short perturbed zone.

Table 1. Boundary conditions and elastic properties assigned to finite element models

Model #

egg

1 2 3 4

5• 5• 5• 5•

EB (GPa) 10-4 10-4 10--4 10-4

16 16 16 16

EA (GPa)

uB

UA

GB (GPa)

GA (GPa)

56 17.5 35 70

0.14 0.14 0.14 0.14

0.26 0.25 0.25 0.25

7 7 7 7

22 7 14 28

Finite element modelling We conducted a series of 2D finite element models to determine the distribution of stresses and displacements in a jointed, interbedded rock body, and to investigate the effects of contrasting elastic moduli on stress perturbations near joints. Finite element numerical simulation was accomplished using the interactive fracture analysis program F R A N C (Wawrzynek & Ingraffea 1987). F R A N C simulates the r - ~ singularity in the elastic cracktip stress field by surrounding each crack tip with a rosette of eight quadratic, triangular, isoparametric quarter-point elements (Barsoum 1976); the remainder of the model mesh is composed of quadratic, isoparametric elements. The boundary conditions and geometry of the model are shown in Fig. 7. Model material properties (Table 1) are median values compiled from several published sources (Blair 1955, 1956; Clark 1966; Senseny & Pfeifle 1984; Atkinson & Meridith 1987). In the first model, the 2D distribution of crack-normal tensile stress (crxx) was analysed in detail for a model consisting of a dolostone bed (Bed A; E = 56 GPa; G = 22 GPa;

FACTORS CONTROLLING JOINT SPACING

223

Fig. 8. Contour diagram showing the 2D distribution of crack-normal tensile stress (trxx) in MPa for a portion of the dolostone and shale model to the right of the joint. Model subjected to a strain of 5 • 10--4 (after Fischer 1994).

u = 0.26) bounded by shale beds (Bed B; E = 16GPa; G = 7GPa; u = 0.14). A uniform longitudinal extensional strain of 5 x 10-4 was applied, resulting in theoretical precracking stresses of 28 MPa in the dolostone and 8 MPa in the shale beds. Because of edge effects in the finite element model, however, Crxxvaried longitudinally within each bed prior to cracking. This variation was minor ( < 2.0% mean error in each bed), and in the model the mean axx is 28.5 MPa in dolostone and 7.8 MPa in shale.

L o c a l c r a c k - n o r m a l tensile stress Local crack-normal tensile stress (trxx) in the dolostone is reduced near the joint whereas in the shale there is an increase in Crxxnear the joint (Fig. 8). At the joint wall ~rxx is zero and near each end of the joint the stress is elevated due to the stress concentrating effect of the sharp crack tip (Lawn & Wilshaw 1975; Brock 1978). Because the joint is elliptical, the greatest amount of elastic strain (and hence stress) is relieved by crack opening along the middle of Bed A. Consequently, the stress reduction is greatest and extends farthest along the centre of Bed A. This combination of elliptical crack shape, high stress near the joint tip and zero stress along the length of the joint is responsible for the concave-inwards shape of the stress reduction shadow in the dolostone bed. Because there is no crack opening in the shale beds, large elastic strains developed in shale adjacent to the joint tips result in a localized increase in stress. An integral part of linear elastic fracture mechanics (Irwin 1957; Paris & Sih 1965; Broek 1978), this stress concentration is maximum at the joint tips and decreases rapidly away from the joint. At horizontal distances > 0.3 m from the joint the stress in the shale beds is the same as the pre-cracking stress.

224

M . R . GROSS E T AL.

-o

1.2 I . . . . . . . . . . . . . . . . . . . . ~g

1 ............. EA/ Ga 0.8

............................ i

#: g o.~ . . . . . . . . . . . . . . . . . . . . . .

I ,~,

0.5 ~ [ ~ ~ = - - ~ = ~ ,

m

0

.~

-o.25

~,

-o.5 o.5

== 02

~-

-o25

o

~

-0.5

o

(a)

1 2 3 4 NORMALIZED DISTANCE (x/d)

5 ~ -~

N

0.25

(a) EA=17.5 GPa

0.25

t~ 0.4 .............

0

0.9

(b)1 EA=35GPal

.75

o

:~ -0.25 ~,

z

-o.5

3 0 1 2 Normalized Horizontal Distance from Joint (x/d)

(b)

Fig. 9 (a) Graph of normalized axx v. normalized horizontal distance from the joint for various EA/G8 values calculated using FRANC. GB for all models is 7 GPa. Stress values were recorded along centre of Bed A. ~) Contours of normalized axx in a middle bed of v = 0.25 and variable EA. Bounding beds are shale in all three models and EA/GBratios are 2.5, 5 and 10 in (a), (b) and (c), respectively. Contours are in percentages of the pre-cracking stress in each layer (i.e. a value of 0.5 indicates the stress at that position is 50% of the pre-cracking stress value) (after Fischer 1994).

Elastic moduli effects In order to investigate the effects of varying elastic moduli (i.e. EA/Ga ratios) on joint spacing we conducted a series of models in which a bed with v = 0.25 and variable EA is sandwiched between shale beds and subjected to a uniform longitudinal extensional strain of 5 • 10-5. The ratio of EA/GB controls the lateral extent of the stress reduction shadow around a joint, as shown graphically in Fig. 9a. To illustrate this effect better, Fig. 9b shows contours of normalized stress in the middle bed for three geologically reasonable values of EA/GB. For values of normalized Crxx < 0.75, the stress reduction shadows for various EA/GBratios are nearly identical. However, for values of normalized stress > 0.75 the size of the stress reduction shadow increases significantly with increasing EA.

Comparison with Hobbs' analysis In general there is good agreement between Hobbs' analysis and our finite element model results. Although Hobbs' prediction of the dependence of stress reduction shadow size on contrasts in elastic moduli is in qualitative agreement with the finite element model results (i.e. the stress shadow increases in extent with increasing Young's modulus), the dependence of the stress shadow decay length, d, on EA exhibited by F R A N C is weaker than the square root dependence predicted by the Hobbs model. To show this we consider the x position of the finite element contours of normalized axx at the centre of Bed A,

FACTORS CONTROLLING JOINT SPACING

225

denoted by x (m) (Fig. 9b). These positions are compared to those predicted by the Hobbs model, denoted by x (p). The x (p) values for each contour are computed as:

x(p)

(EA)

= x(m) , / EA

(17.5)V 17.5 '

(4)

and are given in Table 2. This definition of x (p) gives Hobbs' prediction of the x position of each contour for the case in which EA is 35 and 70GPa (i.e. xlP~), xlP~)) based on the position of the corresponding contour for the finite element modelinwhich EA is 17.5 GPa (i.e. "~(17.5V" ,.(m) a The x (p) values calculated for the models in which EA is 35 and 70 GPa are in all cases greater than the corresponding x (m) values, showing that dependence of the stress reduction shadow size on EA is weaker for the finite element model than the square root dependence predicted by Hobbs. Model geometry and method of stress determination may contribute in part to the differences in stress reduction shadow size predicted by the Hobbs and finite element models. In the finite element model the joint terminates at the shale boundaries and is entirely confined to the middle layer, thus reflecting geometry commonly observed in interbedded rocks with sharp lithologic contacts (Figs 3 & 4). In contrast, the joint in the Hobbs model extends significant distances into the neighbouring beds due to requirements imposed by shear stress assumptions. Furthermore, as opposed to Hobbs' derivation, the numerical model utilizes linear elastic fracture mechanics to describe the stress and displacement fields in the vicinity of the crack.

Table2. Comparison of finite element stress reduction contours to the (EA/GB) '/2 dependence predicted by the Hobbs model :

Contour level 0.25

EA(GPa) 17.5 35 70

0.50

0.75

0.90

x(m)

x~)

x(m)

x~)

x(m)

x@)

x(m)

x~)

0.403 0.420 0.470

0.403 0.570 0.806

0.655 0.682 0.706

0.655 0.927 1.310

1.042 1.075 1.176

1.042 1.473 2.083

1.613 1.848 2.318

1.613 2.281 3.226

In all models GB is 7 GPa and distances are normalized to a bed thickness of 1. x(m) refers to the contour position in the numerical model and x(p) is the position predicted by Hobbs.

Joint spacing predicted from numerical models Joint spacing models combine a rock fracture criterion with an analysis of the stress distribution around a pre-existing joint to predict the location where subsequent joints will form (e.g. Hobbs 1967; Pollard & Segall 1987; Narr & Suppe 1991; Gross 1993a). Hobbs employs a tensile strength fracture criterion wherein joints form when the remote tensile stress (O'remote) in a bed exceeds the tensile strength. We use a fracture criterion based on the principles of linear elastic fracture mechanics (Lawn & Wilshaw 1975).

226

M.R.

3 0

"" 2 =E

,

l

GROSS

,

I

,

,

9, \ + ' '

", i ~ . , i " ~ . . . .~. . . - - ~

25.

E T AL.

I

~ r

o

i'.i

i?:i,

1 5 .............. i.............. !

5-

I

...... ~..............;..........,t - - - - " - ....

.............. ~,. ........ %.+.....i ............. i .......................

m

,

,

li,,+

!

,

I

I

sandstone

t .....

shale

1/

i .............. i ............. ~-'............ :.

~i

i

~

i

+

"4 .

=

"~

,

i

,

dolostone

.........

i

........... cf = K m c . ( n c ) "I!~........... i............................. i...........................................

0

0

0.1

0.2 FLAW

0.3

HALF

SIZE

0.4

0.5

r (r

Fig. 10. Flaw size dependent fracture criterion. Fracture stress for dolostone, sandstone and shale illustrated for various flaw sizes. Fracture toughness values for sandstone (Kic = 1.2 MPa mW), shale (Kic= 0.9 MPa m v2) and dolostone (Kic= 1.7 MPa m v') are the mean of Kic values compiled in the literature (Senseny & Pfeifle 1984; Atkinson & Meredith 1987).

Fracture criterion Inglis (1913) and Griffith (1920) showed that brittle fracture is often induced by a tensile stress concentration at flaws (e.g. microcracks) in a material. Linear elastic fracture mechanics describes the stress concentration at the tip of a Mode I opening (i.e. tensile) crack by means of the stress intensity factor (Broek 1978): gI

~-

Yo'remote

v/-~,

(5)

where ~rr~moteis the remote crack-normal tensile stress, c is half the crack length, and Y is a geometric constant, commonly taken as 1.0 for an infinite fiat blade crack (Lawn & Wilshaw 1975). The fracture criterion states that a crack will propagate when stress intensity at the crack tip exceeds some critical value known as the fracture toughness (Kit) of the material (Irwin 1957). Substituting Kic for KI in equation (5), setting Y = 1, and solving for the fracture stress (af), the crack-normal tensile stress required to cause a flaw to propagate, it is found that: gxe o: -

v/ft.

Using published values of K~c for a variety of rock types, af can be calculated for a given flaw size (Fig. 10). However, because the distribution of flaw sizes in rocks is complex, it is unlikely that the fracture strength of a rock can be defined by a single value of af.

Flaw size distribution in two dimensions The remote tensile stress required to drive a crack decreases with initial flaw size (Fig. 10). As a consequence, for a constant applied extensional strain, the minimum allowable distance between two adjacent joints decreases with increasing flaw size. The dependence

FACTORS CONTROLLING JOINT SPACING

227

Fig. 11. Stress reduction shadow in the dolostone bed of the dolostone and shale model. Tensile stress contours are in MPa, and applied remote stress is 28.5 MPa. The closest distance at which a new joint will form is dependent on the size of the flaw from which it initiates. Flaws ranging in size from 0.3 to 0.5 cm will activate in the shaded region, with 0.5 cm flaws initiating closer to the joint. Joint spacing therefore decreases with increasing initial flaw size.

of joint spacing on flaw size is shown schematically in Fig. 11, which displays stress contours around a joint in the model dolostone bed (extensional strain of 5 x 10-4; applied remote stress of 28.5 MPa). As flaw size increases, the corresponding fracture stress at which point the flaw grows into a joint is attained at distances closer to an existing joint, resulting in a decrease in joint spacing. Joint propagation occurs in the shaded region of Fig. 11 for flaws ranging in size from 0.3 to 0.5cm. In the case of joints initiating in the middle of the bed (i.e. y = 0) the minimum allowable joint spacing ranges from (0.82)d for 0.5 cm flaws to (1.25)d for 0.3 cm flaws, where d is the bed thickness. In addition to flaw size distribution along the x-axis, another factor that determines joint spacing is the position relative to the bed boundary (i.e. along the y-axis) from where the joint initiates. Contour shapes in Fig. 11 predict that for a given flaw size, joints initiating near bed boundaries will be more closely spaced than joints propagating from the central region,of beds. The curvature of stress contours is less pronounced for large distances from the joint, so that for the 90% contour level the dependence of joint spacing on y-axis-initiation position is somewhat reduced. Furthermore, a WeibuU approach to the strength of materials (e.g. Weibull 1951; D avidge 1979) implies that joint initiation is more likely to occur near the centre of the bed rather than along its edges. This is due to the greater activation area in the central portion of the bed as indicated by the width of the shaded zone in Fig. 11.

Joint spacing determination Several approaches have been proposed for analysing joint spacing in 1D models, including the maximum tensile stress criterion that predicts precise midpoint fracturing regardless of initial joint spacing (Hobbs 1967), the introduction of random flaws into stress reduction models (Narr & Suppe 1991; Rives et al. 1992), and the minimum

228

M . R . GROSS E T AL.

Fig. 12. (a) Stress shadows representing 90% of the remote tensile stress for beds with Young's modulus of 17.5GPa (Bed #1) and 70GPa (Bed #2). ~) Expected joint spacing distributions for Beds #1 and #2 after an equal number of infiUing events. Note the larger stress reduction shadow leads to a larger median joint spacing.

allowable spacing criterion (Gross 1993a). Regardless of the approach, two boundary conditions always apply to models of joint spacing in lighology-controlled mechanical layers. First, a new joint will form between two pre-existing joints when the fracture stress is attained. Second, the stress reduction shadow inhibits the formation of new joints in the vicinity of pre-existing joints. Our numerical results from 2D models indicate that the extent of the stress reduction shadow increases with increasing Young's modulus. The effect of elastic properties on joint spacing is illustrated in an example of two jointing layers, #1 and #2, with Young's moduli of 17.5 and 70GPa, respectively (Fig. 12a). Consider the case whereby each bed has undergone an initial fracturing episode. A subsequent infilling event will result in the joint spacing distributions depicted in Fig. 12b. Random flaw distributions along with the 2D shape of the activation zone will lead to a wide range of joint spacing values and a characteristic skewed distribution. For example, joints will initiate preferentially from the largest flaws, and joints initiating near bed boundaries will be more closely spaced than joints propagating from the central portion of the bed. Nevertheless, one would initially predict that for a population of mechanically confined fractures, the bed with the higher Young's modulus will exhibit a larger median joint spacing because of the larger stress reduction shadow.

229

FACTORS CONTROLLING JOINT SPACING

Table 3. Fracture spacing measurements at Lompoc Landing along with physical properties of siliceous mechanical units in the Monterey Formation

Number of beds Number of joint spacings Matrix porosity Young's modulus Shear modulus Fracture spacing index (FSI) Fracture spacing ratio (FSR)

Chert

Diatomite

6 247 0.5% 70 GPa 33 GPa 2.41 2.29

7 315 37% 5 GPa 2 GPa 1.09 1.11

Porosity and elastic moduli measurements are preliminary and were determined from similar units exposed along the Santa Maria coastline. Elastic properties were measured dynamically using the resonance technique. From Gross 1993b.

Why stiffer beds have smaller joint spacings; a field example from the Monterey Formation, California The concept that differences in elastic properties among jointing lithologies will result in different joint spacings for a given bed thickness is intuitive. Our analysis thus far suggests the stress shadow increases in lateral extent with increasing Young's modulus, implying that joint spacing should be greater in beds with higher Young's moduli (i.e. stiffer beds) (Fig. 9b). However, joint spacing is often closer in stiffer beds. For example, the siliceous Monterey Formation exposed at Lompoc Landing along the Santa Maria coastline of California (Fig. 5) consists of a series of interbedded cherts and dolomitic diatomites, both jointing lithologies, along with non-jointing clay-rich mudstones. Typical physical properties of the two siliceous rock types, exposed along the Santa Maria coastline, are markedly different from each other, with a Young's modulus of 70 GPa for the quartz chert and 5 GPa for the impure diatomite (Table 3). Several beds at Lompoc Landing are welded together along their boundaries, and consequently were subjected to equal amounts of extensional strain. The vertical systematic joint set investigated trends N N E SSW and post-dates silica diagenesis, and therefore developed after differences in physical properties were established among the mechanical units. Joint spacing data for the diatomite and chert beds at Lompoc Landing are presented as plots of median joint spacing v. mechanical layer thickness (Fig. 13). The slope of the bestfit line, calculated with median joint spacing as the dependent variable, is defined as the fracture spacing index (FSI) of Narr & Suppe (1991). The fracture spacing ratio (FSR) equals the mechanical layer thickness divided by the median joint spacing, and represents the mean for all beds of a given lithology. Chert beds at Lompoc Landing display significantly higher FSI and FSR, and thus more closely spaced joints relative to bed thickness, than adjacent diatomites. This spacing relationship exists despite the much larger chert Young's modulus. How does such a joint spacing distribution develop in light of the numerical models and theoretical considerations discussed in this paper? One explanation is that despite the larger stress reduction shadow, a stiffer bed will contain more joints relative to other beds because jointing in the stiffer bed occurs at lower strain levels. Consider a rock mass composed of a sequence of dolostone and sandstone beds interbedded with non-jointing mudstone beds. The dolostone has a Young's modulus of

230

M.R. GROSS ,,

70

,,

I

. . . .

I

. . . .

I

. . . .

I

,,

,,

I,

,,

,I

,

,,

ET AL.

,

[I

I

L

I

,

I

I

E

v

o

~

l

,

l

,

l

,

I

E 39

A

Dolomitic Diatomite

60-

w w

,j18

~ 8_

,""

:

t"

Chert

f

eI

ir 5 0 ,~r

9 nS

r

/- 40I--

,"

-

I

=6~ "

a

67 m,,

t~

s

=,, 3 0 -

36 9 ,,

."

: ~ _

,_1

9

20-

,'"

82

FSI = 1.09

FSR

m'm 3 6

=E

26

"

9

0

....

0

I''''I

10

....

20

"~

30

.~ 4 2

''''I

....

40

=

I'''

50

1.11

'I

60

Median Joint Spacing (cm)

o

' ' ~''r''

70

FSI = 2.41 FSR = 2.29

6

: .~-r 2 -

o

10-

146

i

L " .C '~ o

o

s

1117

.j

9 53

111.64

r

>,4-

tl

9

m

,

s e

138 "

0

0

I

'

I

2

'

I

'

I

4

'

I

'

i

6

'

I

'

1

8

Median Joint Spacing (cm)

Fig. 13. Median joint spacing v. mechanical layer thickness for diatomite and chert beds of the Monterey Formation exposed at Lompoc Landing. Numbers next to data points refer to total number of joint spacing measurements taken in that particular bed.

5 6 G P a and a Kit of 1 . 7 M P a m ~, whereas the sandstone has a Young's modulus of 12GPa and a Kic of 1 . 2 M P a m ~. Hooke's Law states that for a given value of strain, stress will be higher in the bed with the higher Young's modulus (Fig. 14). Because our fracture criterion is a threshold stress value (i.e. the fracture stress, crf), jointing will occur in a given bed at the point where the stress-strain curve intersects cry for that bed. Assuming a range in flaw size from 0.3 to 0.5 cm (i.e. crack half-lengths of 0.15-0.25 cm) for both jointing lithologies, the fracture stress is 19.2-24.8 M P a for the dolostone and 13.5-17.5MPa for the sandstone (Fig. 14). Fracture stress ranges derived from the range in flaw size correspond in t u m to a range of strain values for the first jointing event in each lithology. The intersections of stressstrain curves with horizontal fracture stress lines in Fig. 14 indicate that joints will propagate in the dolostone bed for strain values within the range of 3.43 x 10-4 to 4.42 • 10-4. Likewise, joints will initiate in the sandstone unit between strains of 1.12 x 10-3 and 1.46 x 10-3. In the case whereby all beds in the rock mass are welded together, and thus remote longitudinal strain remains uniform throughout, the first series of joints will propagate in the dolostone bed once a strain of c. 4 x 10-4 is reached. At this juncture the sandstone bed is unjointed because stress within the sandstone is less than the minimum fracture stress of 13.5MPa. As longitudinal extension continues, the sandstone bed remains unjointed until strain values of 1.12 x 10-3-1.46 • -3 are exceeded, at which point the first set of joints propagates in the sandstone. As a result of the relationship between Hooke's Law and the fracture stress criterion, jointing occurs in the dolostone at lower strains, and hence earlier, than jointing in the sandstone. In fact, the situation may arise where a stiff bed undergoes two or more jointing events prior to initial jointing in a low modulus bed. Therefore, in spite of the larger stress reduction shadows predicted by both Hobbs and finite element modelling, beds with relatively high Young's moduli can actually contain relatively closely-spaced joints by virtue of the small strain increments required for infilling jointing events.

FACTORS CONTROLLING JOINT SPACING

231

Fig. 14. Fracture stress ranges superimposed on stress-strain plots for typical dolostone and sandstone beds. Ranges in fracture stress correspond to a range in flaw size from 0.3 to 0.5 cm. Note the dolostone bed will fracture at lower strain levels than the sandstone. See text for details.

In comparing two jointing beds with different elastic moduli (for Beds 1 and 2, E1 > E2), the factor that determines which bed will fracture at lower strains is the ratio of Young's modulus to the ratio of fracture stress. Assuming the same range of flaw sizes for both beds, the stiffer bed will fracture at lower strains when: E1 > K~__.)_)

(7)

The inequality in equation (7) holds true for a majority of sedimentary rocks, which may explain the higher joint densities often observed in stiffer beds despite the larger stress reduction shadows.

Conclusions Joint spacing models focus on the distribution of tensile stress about a joint in an effort to predict the location of subsequent joints. Stress distributions derived from finite element simulations compare favourably with Hobbs' theoretical predictions. Both the Hobbs model and finite element modelling predict longer stress reduction shadows, and hence wider joint spacings, for beds with higher Young's moduli. Finite element simulations,

232

M.R. GROSS ET AL.

though, show the dependence of the stress reduction shadow on Young's modulus in the jointing bed is weaker than the square root dependence predicted by Hobbs. Though bed thickness is a primary factor, elastic moduli, extensional strain and flaw size are all important parameters determining joint spacing in interbedded sedimentary rocks. A high Young's modulus contributes to an increase in stress reduction shadow length, yet tends to lower the strain level required for joint propagation due to its relationship with the fracture stress criterion. Small flaws situated in the middle of the jointing bed produce widely spaced joints, whereas those same flaws located near bed interfaces result in more closely spaced joints. Despite the complex interaction among these factors, creating a wide range of joint spacings for an individual mechanical bed, the linear correlation between median joint spacing and bed thickness for a given lithology, and hence given mechanical properties, is remarkably strong.

Funding for this project was provided by a grant from Texaco (TE, MRG) a Shell Doctoral Fellowship (MPF) and GRI Contract 5088-260-1746 (TE). We thank T. Ingraffea for providing the finite element program, T. Bittencourt and D. Swenson for technical assistance and Wendy Bartlett for helpful field information. We appreciate the thoughtful reviews provided by Wayne Narr, Atilla Aydin, Byron Kulander, Ernie Rutter and Mohammed Ameen.

References ATKINSON,B. K. & MEREDITH,P. G. 1987. Experimental fracture mechanics data for rocks and minerals, ln: ATKINSON, B. K. (ed.) Fracture Mechanics of Rock. Academic Press, London, 477525. BARSOUM,R. S. 1976. On the use of isoparametric finite elements in linear fracture mechanics. International Journal of Numerical Methods in Engineering, 10, 25-37. BLAIR, B. E. 1955. Physical properties of mine rock, Part 3. US Bureau of Mines Report of Investigations 5130. - 1956. Physicalproperties of mine rock, Part 4. US Bureau of Mines Report of Investigations 5244. BROEK, D. 1978. Elementary Engineering Fracture Mechanics. Sijthoff & Noordhoff, Dordrecht, the Netherlands. CLARK, S. P. 1966. Handbook of Physical Constants. Geological Society of America Memoir, 97. Cox, H. L. 1952. The elasticity and strength of paper and other fibrous materials. British Journal of Applied Physics, 3, 72-79. DAVIDGE,R. W. 1979. Mechanical Behavior of Ceramics. Cambridge University Press, Cambridge, UK. ENGELDER,T. & GROSS,M. R. 1993. Curving cross joints and the lithospheric stress field in eastern North America. Geology, 21, 817-820. FISCHER, M. P. 1994. Application of linear elastic fracture mechanics to some problems of fracture propagation in rock and ice. PhD Thesis, Pennsylvania State University, USA. GARRETr, K. W. & BAILEY, J. E. 1977. Multiple transverse fracture in 90~ cross-ply laminates of a glass fibre-reinforced polyester. Journal of Materials Science, 12, 157-168. GRIFFITH, A. A. 1920. The phenomena of rupture and flow in solids. Philosophical Transactions of the Royal Society of London, A221, 63-197. GROSS, M. R. 1993a. The origin and spacing of cross joints. Journal of Structural Geology, 15, 737751. 1993b. The effects of mechanical stratigraphy on failure mode and fracture spacing in the Monterey Formation of coastal California. PhD Thesis, Pennsylvania State University, USA. HOBBS, D. W. 1967. The formation of tension joints in sedimentary rocks: an explanation. Geological Magazine, 104, 550-556. HUAN6, Q. & ANGELIER, J. 1989. Fracture spacing and its relation to bed thickness. Geological Magazine, 126, 355-362.

FACTORS CONTROLLING JOINT SPACING

233

INGLIS, C. E. 1913. Stresses in a plate due to the presence of cracks and sharp corners. Transactions of the Institute of Naval Architecture, 55, 219-230. IRWIN, G. R. 1957. Analysis of stresses and strains near the end of a crack traversing a plate. Journal of Applied Mechanics, 24, 361-364. JAEGER, J. C. & COOK, N. G. W. 1976. Fundamentals of Rock Mechanics. Chapman & Hall, London. LACHENBRtJCH, A. H. 1961. Depth and spacing of tension cracks. Journal of Geophysical Research, 66, 4273-4292. LADEIRA, F. L. & PRICE, N. J. 1981. Relationship between fracture spacing and bed thickness. Journal of Structural Geology, 3, 179-183. LAWN, B. R. & WILSHAW, T. R. 1975. Fracture of Brittle Solids. Cambridge University Press, Cambridge, UK. MASUDA, T. & KURIYAMA, M. 1988. Successive 'mid-point' fracturing during microboudinage: an estimae of the stress-strain relation during a natural deformation. Tectonophysics, 147, 171-177. McQUILLAN, H. 1973. Small-scale fracture density in Asmari Formation of southwest Iran and its relation to bed thickness and structural setting. American Association of Petroleum Geologists Bulletin, 57, 2367-2385. NARR, W. & SUPPE, J. 1991. Joint spacing in sedimentary rocks. Journal of Structural Geology, 13, 1037-1048. OLSON, J. & POLLARD, D. D. 1989. Inferring paleostresses from natural fracture patterns: A new method. Geology, 17, 345--348. PARIS, P. C. & SIH, G. C. 1965. Stress Analysis of Cracks. American Society of Testing and Materials, Special Publication, 381, 30-76. PARVIZI, A. & BAILEY, J. E. 1978. On multiple transverse cracking in glass fibre epoxy cross-ply laminates. Journal of Materials Science, 13, 2131-2136. POLLARD, D. D. & SEGALL,P. 1987. Theoretical displacements and stresses near fracture in rock: with applications to faults, joints, veins, dikes, and solution surfaces. In: ATKINSON, B. K. (ed.) Fracture Mechanics of Rock. Academic Press, London, 277-349. - - , SEGALL, P. & DELANEu P. T. 1982. Formation and interpretation of dilatant echelon cracks. Geological Society of America Bulletin, 93, 1291-1303. PRICE, N. J. 1966. Fault and Joint Development in Brittle and Semi-brittle Rocks. Pergamon Press, Oxford. PRIEST, D. S. & HUDSON, J. A. 1976. Discontinuity spacings in rock. International Journal of Rock Mechanics, Mining Sciences, and Geomechanics Abstracts, 13, 135-148. RIVES, T., RAZACK, M., PETIT, J.-P. & RAWNSLEY, K. D. 1992. Joint spacing: analogue and numerical simulations. Journal of Structural Geology, 14, 925-937. ROULEAU, A. & GALE, J. E. 1985. Statistical characterization of the fracture system in the Stripa Granite, Sweden. International Journal of Rock Mechanics, Mining Sciences, and Geomechanics Abstracts, 22, 353-367. SENSENY,P. E. & PFEIFLE,Z. W. 1984. Fracture toughness of sandstones and shales. Proceedings of the 25th US Symposium on Rock Mechanics. Society of Mining Engineers of the American Institute of Mining, Metallurgical and Petroleum Engineers, Inc., New York, 390-397. SYLVESTER,A. G. & DARROW,A. C. 1979. Structure and neotectonics of the western Santa Ynez fault system in southern California. Tectonophysics, 52, 389-405. WAWRZYNEK, P. A. & INGRAFFEA, A. R. 1987. Interactive finite element analysis of fracture processes: an integrated approach. Theoretical and Applied Fracture Mechanics, 8, 137-150. WEIBULL, W. 1951. A statistical distribution function of wide applicability. Journal of Applied Mechanics, 18, 293-297.

Index

Page numbers in italics refer to Figures and Tables Alpine Orogeny 171, 210 andesite Japan 15, 17 West Cumbria 144 annular structure 64 see also rib mark arrest line 42, 63, 79, 176 see also rib mark Australia granodiorite roughness study 31 auto correlation function (ACF) defined 13 B-plane 176 basalt discontinuity measurements 15, 17 Bavaria 40, 41 bedding planes 12 discontinuity profile 16, 18, 19, 20 beef 196 bending, role in failure of 89-91 bending fracture, experimental production of 71-2 Bohemian Massif 40, 41 Borrowdale Volcanic Group 100, 141-4 see also West Cumbria basement rock Botany Bay 154, 157 burial/compaction and associated jointing 16870 Burnham chalk study 201 C-fracture 176 C-joint 63 calcite mineralization 115, 117, 122 veins 196 Calder Sandstone 99 cataclasite 127, 161,165 centre-line average height (CLAH) defined 12 effect of profile length on 20 effect of sampling interval on 18, 19 relation to ratio of profile length 24 relation to weighted asperity inclination 23 Chalk joint features 168-70 method of analysis 199 results 200-6 results discussed 209-11 significance of spatial changes 208-9 stress implications 206-8 setting 197-9 Chausuyama hornfels discontinuity study 15, 17

chert discontinuity measurements 15, 17 jointing study 229, 230 Chita mudstone discontinuity study 15, 17 compaction/burial and associated jointing 16870 compressional tectonics 171 compressive strength in Chalk 209 Continental Superdeep Borehole see KTB cooling discontinuity 12 core disking surface (CDS) 46-7, 48, 53 transitional forms 47 crack branching, role in failure of 91-3 cross fracture 63 Cumbria see West Cumbria density of Chalk 209 desiccation discontinuity 12 diatomite jointing study 229, 230 dilatent echelon crack 63 discontinuities classification 12 Japan case study effect of profile length 20, 21 effect of sampling interval 18-19 equipment 16-17 location of samples 15 primary v. secondary 22-4 role of eigen direction 21, 22 linear profile measurement 15-17 linear profile parameters 12-15 summary of observations 24--5 disparity, use in roughness measurement of 28 diverging plumes 41 drape folds 163, 168, 172 Dumpton Gap 158 dyke, sandstone 192-3 East Cliff 158 East Harling chalk study 202 East Irish Sea Basin 98, 99 Echizen tuff discontinuity study 15, 17 eigen direction 22 effect on discontinuity study 21, 22 eigen discontinuity 12 elastic moduli, effect on joint spacing of 224 en echelon fracture 101 Ennerdale granite and granodiorite 99, 134, 146 Epple Bay 158, 160

236

INDEX

Eskdale granite 99, 134, 145 experimental production of fractographic features methods 64-6 results bending fracture 71-2 fracture forking 68 fracture hooking 68 origin of fracture 66-8 point load fracture 70 thermal fracture 71 torsion fracture 68-70 results discussed forking 77 inclusion hackle 77-9 origins 73-4 rib marks 79-80, 81 speed of propagation effect 72-3 twist hackle 77-9 Wallner lines 75-7 extensional tectonics 170-1 F-joint 63, 176 failure mode 83 experimental testing 84-5 failure area analysis 89 bending 89-91 crack branching 91-3 microscope analysis 86-8 fast Fourier transform (FFT) images 35, 36 faulting, role in discontinuity of 12 faults 100 West Cumbria study 104-7, 125-6 feather fracture border plane 63 flow plane 12 fluidization 192-3 focal point 62 see also origin folding, role in discontinuity of 12 forced folds 163, 168, 172 forking 48, 54 see also fracture forking fractal defined 27 fractal dimension defined 13 fractographic features classified 100 fractography defined 1, 99-100, 152 evolution of studies 1-2 first use of term 59 relation to fracture analysis 2-3 research future 9 past 3-8 fracture defined 99-100 fracture face 176 fracture forking 63, 66, 68, 74, 77 fracture hooking 68

fracture plume 176 fracture processes 52-4 fracture propagation 52 fracture spacing index (FSI) 229 fracture spacing ratio (FSR) 229 fracture surface studies (markings, morphology and patterns) classification 27, 41-2, 176 interpretation of stress patterns 43 photogrammetric measurement principles 28 quantitation 31-2 regional studies KTB analysis core disking surfaces (CDS) 46-7, 48, 53, 54 drilling-induced centre-line fractures 46 transitional forms 47, 52 usual fractographic surface structures (UFS) 46-7, 49, 50, 51 Mt Alexandra Quarry 32-6 St Mary's Well Bay 176--7 methods of analysis 180--1 results 181-3 results discussed 183-5 setting 177-9 West Cumbria 108, 115, 133 symmetry 42 fracture trace defined 100 West Cumbria study 101, 108, 109, 110, 111, 112, 133 Franconian Line 39, 40 freezing discontinuity 12 fringe 176 fringe face 63 fringe step 63 Germany 40, 41 glaciation, overburden effects of 211 gouge 127, 161,165 grain plane, role in discontinuities of 12 granite Japan 15, 17 West Cumbria 100, 133, 145, 146 granodiorite Mt Alexandra 31 West Cumbria 145 Griffith criteria 187 gull wing 72, 75, 76 gypsum veins 196 hackle face 158 hackle (hackly) fractures 63 hackle plume 79, 101 KTB study 41

INDEX Lavernock Point 181, 183, 184 Thanet monocline 158, 159, 163 West Cumbria study 118, 121 hackle step 158 haematite mineralization 127 Harcourt roughness measurements 31 hardway plane, role in discontinuity of 12 hesitation line 64 see also rib mark hooking 3, 101, 125 hoop test for failure 84-5 hornblende, role in discontinuity of 12 hornfels discontinuity measurements 15, 17 hydraulic fracturing 191 igneous rocks, discontinuities in 12 inclusion hackle 63, 77-9 inclusion hackle tail 72 iron oxide mineralization 115, 117, 163 Italian marble, failure testing of 85, 87, 88 Japan discontinuity morphology study effect of profile length 20, 21 effect of sampling interval 18, 19 equipment 16-17 location of samples 15 primary v. secondary features 22-4 role of eigen direction 21, 22 joint defined 62, 99, 197 regional studies Eastern England 168-70, 197-9 method of analysis 199 results 200-6 results discussed 209-11 significance of spatial changes 208-9 stress implications 206-8 West Cumbria 104-7 joint face 101, 176 joint spacing bed thickness effects 218-20 history of research 215-17 mathematical modelling 1-D Hobbs model 220-2 finite element model 222-4 models compared 224-5 testing of models 225-8 Monterey Formation study 229-31 stress field effects 217-18 Kingsgate Bay 154, 156 Kitamatado shale discontinuity study 15, 17 KTB study drill core fractures fracture processes 52

237 fracture types core disking surfaces (CDS) 46-7, 48, 53, 54 drilling-induced centre-line fractures 46 transitional forms 47, 52 usual fractographic surface structures (UFS) 46-7, 49, 50, 51 stress field analysis methods 54-6 results 56-7 location 39, 40, 41

Lake District Block 98, 99 Lavernock Point 176, 177 fracture surface marking study methods of analysis 180-1 results 181-3 results discussed 183-5 setting 177-9 limestone Japan discontinuity measurement 15, 17 St Mary's Well Bay Formation 176-7 methods of analysis 180--1 results 181-3 results discussed 183-5 setting 177-9 lithology, effect on joint spacing of 229 London-Brabant Massif 152--4 macrofracture and macrofault defined 100 West Cumbria study 125-8, 137 marble, failure testing of 85, 87, 88 mathematical modelling of joint spacing 1-D Hobbs model 220-2 finite element model 222-4 models compared 224-5 testing of models 225-8 mean standard variation of height (MSVH) defined 13 mean standard variation of inclination (MSVI) defined 13 megafracture defined 100 Mercia Mudstone 193, 194 mesofracture defined 100 West Cumbria study mesofracture study in basement rock 131-3, 141-6 dynamic significance 133-7 surface features 133 fracture traces 133 surface patterns 133 mesofracture study in cover rock chronology 115-7 layering control 117

238

INDEX

non-systematic fractures 108, 112 systematic fractures 100-1 dynamic significance 117, 125 fracture traces 101, 108, 109, 110, 111, 112 surface patterns 108, 112 mesofracture/macrofault relations in basement rock 137 mesofracture/macrofault relations in cover rocks 126-8 regional stress patterns 128 metamorphic rocks, discontinuities in 12 micas, role in discontinuity of 12 microfracture defined 100 Mill Hill chalk study 201 mineralization 3 West Cumbria 115-7, 122 mineralogy, effect on discontinuities of 12 Minnis Bay 158 mirror region defined 62 experimental production of 74, 66, 71, 72 mist region characterized 62 experimental production of 66, 72, 74 mode I (tensile) fractures 188, 190 Thanet monocline 170, 172 West Cumbria 117-9 modes of failure 83 experimental testing 84-5 failure area analysis 89 bending 89-91 crack branching 91-3 microscope analysis 86--8 mode I (tensile) failure 83, 176 effect of tensile strength 193-4 relation to stress 187 role of fluid pressure 190-1 mode II (sliding) failure 83 mode III (tearing) failure 83, 176 Mohr circles 187, 188 chalk stress field 207 Monterey Formation joint spacing study 216, 217, 219-20, 229--31 Mt Alexandra Quarry roughness measurements 31 mudstone discontinuity measurements 15, 17

Nakatsukawa granite discontinuity study 15, 17 Navier-Coulomb criteria 187 non-systematic fractures defined 100 West Cumbria study occurrences 101, 117, 123 North Cliff 158, 165

Okumino discontinuity studies 15, 17 origin 62, 70, 72, 101 action in experiments of 66 role in chalk failure 158 Otake andesite and basalt discontinuity study 15, 17 overburden, stress effects of 196, 210-11 Palm Bay 154, 156, 159 Pegwell Bay 166, 167 Pegwell Bay-Minnis Bay Transpression Zone 158, 168, 171 Penrith Sandstone failure testing 85, 86, 87 photogrammetry Mt Alexandra Quarry tests 32-6 use in roughness measurement methods 31-2 principles 28 pit 62 see also origin plume axis 101 plumose structure 63, 176 plumose-coarse twist hackle boundary 176 point load fracture 70 point sources 181 porosity in Chalk 209 power spectral density (PSD) plots 36 pressure solution features 64 profile length, effect on discontinuity study of 20, 21 profilometer 14, 17, 31 radient 63 ratio of profile length (RPL) defined 13 effect of profile length on 20 effect of sampling interval on 18, 19 relation to centre-line average height 24 relation to weighted asperity inclination 25 rhyolite discontinuity measurements 15, 17 rib mark 63, 71, 78, 79-80, 81,101 Lavernock Point 181 Thanet monocline 158, 159, 163 West Cumbria study 118, 120, 121 Richborough syncline 173 ripple mark 62-3 see also Wallner line river line 63 root mean-square of height (RMSH) defined 13 root mean-square of inclination (RMSI) defined 13 roughness measurement 27-8 method 28-30 quantitation 31-2 regional study 32-6

INDEX St Bees Sandstone 99 fracture pole diagrams 102, 103 macro faults 127, 130 St Mary's Well Bay Formation 176-7 fracture surface marking study methods of analysis 180-1 results 181-3 results discussed 183-5 setting 177-9 sampling intervals, effect on discontinuity study of 18, 19 sandstone discontinuity measurements 15, 17 failure testing 85, 86, 87 see also West Cumbria study cover rock satin spar 196 schistosity planes 12 sedimentary rocks, discontinuities in 12 shale, discontinuity measurements in 15, 17 shear failure and relation to stress 187 shear planes discontinuity profile 16, 18, 19, 20 role in discontinuity of 12 shearing 3 sheeting planes discontinuity profile 21, 22 role in discontinuities of 12 Sherwood Sandstone Group 100 see also West Cumbria study cover rock shoulder 176 slickenlines 3, 52, 54 KTB study 52, 54 Thanet monocline 155, 161, 164, 167 West Cumbria study 113, 127 sliding (mode II) failure 83 Solway Basin 98, 99 South Cliff 165 spatial frequency (SF) plots 36 stereographic projection 155 stress 187 KTB stress field analysis 39, 40, 41 drill core fracture evidence 46 core disking surfaces (CDS) 46--7, 48, 53, 54 drilling-induced centre-line fractures 46 transitional forms 47, 52 usual fractographic structures (UFS) 467, 49, 50, 51 West Cumbria study 112, 124, 125, 128-9 striations 63 structure function defined 13 stylolites defined 100 stylolitization 3 surface features see fracture surface studies systematic fractures 99-100 West Cumbria study dynamic significance 117, 125 fracture traces 101, 108, 109, 110, 111, 112

239 surface patterns 108, 115

tearing (mode III) failure 83, 176 tectonic discontinuity 12 tectonic features of Thanet monocline 170-1 tectonic stress field study drill core fracture evidence 46 fracture types core disking surfaces (CDS) 46-7, 48, 53, 54 drilling-induced centre-line fractures 46 transitional forms 47, 52 usual fractographic structures (UFS) 467, 49, 50, 51 location 39, 40, 41 tendential fractographic features 62, 67, 69 tendential penetration 176 tensile (mode I) failure 83, 176 effect of tensile strength 193--4 failure area analysis 89 bending 89-91 crack branching 91-3 hoop test 84-5 microscope analysis 86-8 relation to stress 187 role of fluid pressure 190-1 tensile (mode I) fracture 188, 190 Thanet monocline 170, 172 West Cumbria 117-9 tensile strength 209 tensile stress analysis application to joints finite element model 222-4 Hobbs model 220--2 tension planes discontinuity profile 16, 18, 19, 20 role in discontinuity 12 texture, effect on fracture of 72, 73 Thanet monocline fracture characterization hinge area 161-3 NE area 154-8 NW area 158 SE area 158-61 fracture origin 168-71 origin 163-8 seismic section 169 setting 149, 150 stratigraphy 151 structure 152 synthesis 172-3 thermal fractures, experimental production of 71, 74 torsion fractures 68-70 experimental production of 68-70 tourmaline, role in discontinuity of 12 trace 176

240

INDEX

transient fractographic features 59, 70, 72 transpression in Thanet monocline 171 Tsuge granite discontinuity study 15, 17 tuff discontinuity measurements 15, 17 twist angle 125 twist hackle 63, 66, 77-9, 158 twist-hackle face 78 defined 101, 176 twist-hackle fringe 68, 69, 70, 79 defined 101, 176 KTB study 41-2 Thanet monocline 158 West Cumbria study 118, 120, 125 twist-hackle plane 182-3 twist-hackle step 54, 72, 78 defined 101, 176 uplift/unloading, associated jointing 171 usual fractographic surface (UFS) structures 46--7 transitional forms 47 veins defined 100 West Cumbria study 104-7 velocity hackle 62, 70, 72 experimental production of 66, 68, 71, 74 Wallner line 62, 62-3, 70, 72 experimental production of 66, 71, 71-2, 75-7 Walpole Bay 154, 156, 157 Watchet 193, 194

weighted asperity inclination defined 13 effect of profile length on 20 effect of sampling interval on 18, 19 relation to centre-line average height 23 relation to ratio of profile length 25 West Cumbria basement and cover rocks compare d 137-~9 macrofault study in cover rock 125-6 mesofracture study in basement rock 131-3, 141-6 dynamic significance 133-7 surface features 133 fracture traces 133 surface patterns 133 mesofracture study in cover rock chronology 115-7 layering control 108-12 non-systematic fractures 101, 117 systematic fractures 100-7 dynamic significance 117, 125 fracture traces 101, 108, 109, 110, 111, 112 surface patterns 108, 115 mesofracture/macrofault relations in basement rock 137 mesofracture/macrofault relations in cover rocks 126--8 regional stress patterns 128 setting 97, 98 West Runton chalk study 203 Westgate Bay 158, 160 Young's modulus 229

Fractography: fracture topography as a tool in fracture mechanics and stress analysis editedby

M.S. Ameen (Independent Consultant, London, UK) This volume introduces the principles and procedures involved in the description and analysis of fracture surface topographies. It contaim a wide range of case studies from core and field samples, as well as from experimentally produced fractures, which illustrate the application of fractography to fracture studies. Fractography is a powerful tool that has been largely neglected by those who deal with fracture characterization. A greater understanding of the subject is essential to the development of oil and gas fields, particularly deep and marginal fields. In eddWon, incorporating fractographic data into models of rod( masses, such as the permeability of hydrocarbon reservoirs and the stability of tunnels or underground nuclear waste repositories, will undoubtedly enhance their feasibility. The book also pmsents new mathods of quantitative meesummem of fracture topography, and outlines areas needing further research and development. The volume will be of interest to a broad range of researche~ and professionals in frKture mechanics,as well as ,bdctural, engineering and petroleum geology, material sciencesand mining engineedng. • • • •

248 pages over 200 Illustrations 12 papelrs index

ISBN

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1-897799-32-2