The Gravity Field of the Papuan Fold Belt and its Geological Implications

by

David Harrison

Thesis Submitted for the Degree of PhD. Department of Geological Sciences University College London

University of London July 1991 ProQuest Number: 10611107

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Abstract

The structure of the Papuan Fold Belt at both local and regional scales has been deduced from the interpretation of gravity data from the fold belt and its foreland basin.

A digital terrain model of the fold belt region was created and used in the calculation of terrain corrections for the gravity stations in the fold belt. A computer program was developed to calculate such corrections for data from extensive areas of rugged terrain, and corrections were obtained for over 7000 existing gravity stations. In addition a field survey was undertaken and a further 300 gravity readings were taken in a previously unsurveyed area.

Digital filtering techniques, including upward continuation and spectral analysis, were used to separate the regional and residual gravity anomalies, which were then used to constrain the lateral extent of both regional and local geological structures. 2D forward modelling was carried out along five SW-NE and two NW-SE profiles and provided the third dimension to the structural interpretation.

In most areas the interpretation of the gravity data confirmed the existing models of the local structure based on geological fieldwork, though evidence of previously unidentified structures, notably in the Muller Anticline region, and beneath the northeastern fold belt, was discovered. On a regional scale it has been shown that the Mesozoic extensional structures of the Papuan Basin have a major controlling effect on the distribution of compressional structures within the fold belt, and on the isostatic compensation of the topographic load of the highlands. 3

Contents

Abstract 2

Contents 3

List of Figures 8

List of Equations 12

List of Plates 12

List of Tables 12

Acknowledgements 13

Aims and Introduction 14

Chapter 1 - The Regional Geology of Papua New Guinea 15

Introduction 15 Regional Setting 15 Major Geological Provinces 15 1.3.1 The Southern Plains 18 1,3.2 The Central Orogen 18 1.3.3 Intermontane Depressions 20 1.3.4 North New Guinea Province 21 1.3.5 The Papuan Peninsula 22 1.3.6 New Britain 23 1.3.7 The Melanesian Arcs 23 1.3.8 The Bismarck Sea 24 1.3.9 The Solomon Sea 24 4

1.4 The Geological History and Evolution of Papua New Guinea 25 1.4.1 Mesozoic Geological History 25 1.4.2 Mesozoic Evolution 26 1.4.3 Cenozoic Geological History 28 1.4.4 Cenozoic Evolution 30

Chapter 2 - The Geology of the Papuan Fold Belt 33

2.1 Introduction 33 2.2 The Stratigraphy of the Papuan Fold Belt 33 2.2.1 Basement Lithologies 37 2.2.2 Upper Palaeozoic and Mesozoic Units 38 2.2.3 Cenozoic Units 41 2.2.4 Quaternary Stratigraphy 45 2.3 The Structure of the Papuan Fold Belt 46 2.3.1 Zone 1 - The Muller-Kutubu Zone 46 2.3.2 Zone 2 - The Oksapmin-Tari-Mt Murray Zone 48 2.3.3 Zone 3 - The Porgera-Nipa-Poroma-KarimuiZone 48 2.3.4 Zone 4 - The Laiagam-Giluwe-Kubor Zone 48 2.3.5 Zone 5 - The Darai Plateau 48 2.3.6 Zone 6 - The Aure Tectonic Belt 49 2.3.7 Relationships between Stratigraphy and Structure 49 2.4 The Geological History of the Papuan Fold Belt 50 2.4.1 The Papuan Basin 50 2.4.2 The Foreland Basin 51 2.4.3 The Central Orogen 51 5

Chapter 3 - Gravity Data 52

3.1 The Papua New Guinea Gravity Database 52 3.2 The Lavani-Juha Regional Profile 52 3.3 The Tiengo Gravity Survey 56 3.3.1 Introduction 56 3.3.2 Field Survey 56 3.3.3 Data Reduction 64 3.3.4 Data Presentation and Interpretation 70

Chapter 4 - Gravity Data Reduction 73

4.1 Introduction 73 4.2 Production of the Digital Terrain Model 73 4.2.1 Digitising 74 4.2.2 Data Quality 74 4.2.3 Data Presentation 75 4.3 Reduction and Terrain Correction of Gravity Data - Theory 75 4.3.1 Initial Data Format 75 4.3.2 Latitude Correction 80 4.3.3 Free-Air Correction 81 4.3.4 Topographic Correction 81 4.4 Reduction of Gravity Data - Method 85 4.4.1 Introduction 85 4.4.2 Reduction to Simple Bouguer Anomaly 86 4.4.3 Topographic Correction 86 4.4.4 Application of the Topographic Correction 94 4.4.5 Overview 102 6

Chapter 5 - Qualitative Interpretation and Further Processing 103

5.1 Extended Bouguer Anomaly Maps 103 5.1.1 General Subdivision 103 5.1.2 Qualitative Interpretation 107 5.2 Enhancement of Terrain Corrected Gravity Data 109 5.2.1 The Fourier Transform 109 5.2.2 The Digital Filtering Program 110 5.2.3 Source Depth Constraints 112 5.2.4 Trend Analysis 116 5.3 Results of Data Enhancement 116 5.3.1 Upward Continuation 116 5.3.2 Power Spectrum Analysis 118 5.3.3 Wavelength (bandpass) Filtering 122 5.3.4 Trend Analysis 127 5.4 General Comments on Enhanced Results 127

Chapter 6 - Forward Modelling 130

6.1 Forward Modelling Theory 130 6.2 Forward Modelling of the Papuan Fold Belt Data 131 6.3 Constraints on the Models 135 6.3.1 Surface Geology 135 6.3.2 Formal Gravity Interpretation 137 6.3.3 Density Determinations from Borehole Gravimetry 139 6.3.4 Density Values used for Modelling the Fold Belt Rocks 142 6.4 Overall Density Structure of the Models 145 6.5 Description of the Profiles and Models 146 6.6 Overview of Regional Profiles 170 6.6.1 Foreland Region 170 6.6.2 Fold Belt Region 170 6.7 Detailed Modelling 173 7

Chapter 7 - Geological Implications and Principal Conclusions 181

7.1 Structures affecting the sedimentary cover in the fold belt 181 7.1.1 Shallow Structures 182 7.1.2 The Muller Anticline 183 7.1.3 The Darai Plateau 185 7.1.4 Significance of Hydrocarbon Accumulations 186 7.1.5 Thin v thick skinned tectonics 187 7.2 The Foreland Basin 188 7.3 Regional Tectonics and Isostasy 189

References 200

Appendix A - Gravity Base Station Descriptions Gravity Meter G-513 Calibration Table 207

Appendix B - Tiengo Gravity Survey Local Terrain Corrections 212

Appendix C - Tiengo Gravity Survey Extended Bouguer Anomaly Data 215

Appendix D - FORTRAN program listings 224

Enclosure - PNG Gravity Study, extended Bouguer anomaly map (1:500000) 8

List of Figures

Chapter 1

1.1 The Location of Papua New Guinea. 16 1.2 Plate tectonic steeing of Papua New Guinea. 16 1.3 Geological Provinces of Papua New Guinea. 17

Chapter 2

2.1 Geological summary map of the Papuan fold belt. 34 2.2 Diagrammatic relationship of the major stratigraphic units. 36 2.3 Major structural subdivisions of the fold belt. 47

Chapter 3

3.1 Gravity survey location map. 53 3.2 Tiengo survey location map. 57 3.3 Road quality map. 60 3.4 Tiengo survey station location map. 62 3.5 Tiengo and Kandep daily pressure readings. 66 3.6 Leru, Nipa and Mendi daily pressure readings. 67 3.7 Humidity correction to aneroid heights. 68 3.8 Tiengo extended Bouguer anomaly map. 71 3.9 Southern Highlands extended Bouguer anomaly map. 72 9

Chapter 4

4.1 Areas covered by the Digital Terrain Model. 76 4.2 Typical 1:50000 contour check plot. 77 4.3 The thin rod approximation. 84 4.4 Calculation of the angle subtended at the centre of the Earth. 84 4.5a Flowchart illustrating the logic of the one-stage correction program (pi). 88 4.5b Flowchart illustrating the logic of the one-stage correction program (p2). 89 4.6 Location of the Komo (8641) survey lines. 91 4.7 Profiles along Komo survey lines 1 and 2, showing both the one and two stage correction results using a standard crustal density of 2670kgm*3. 92 4.8 Difference between the results of the one and two stage corrections. 93 4.9a Plot of the difference between the one and two stage results against station elevation. 93 4.9b Plot of the difference between the one and two stage results against the terrain correction. 94 4.9c Plot of the difference between the one and two stage results against the total terrain correction. 94 4.10 Profiles along Komo survey lines 1 and 2 using variable density and uniform density of 2670kgm'3. 96 4.11 Profiles along Komo survey lines 1 and 2 using variable density and uniform density of 2500kgm'3. 97 4.12 Profiles along Komo survey lines 1 and 2 using variable density and constant density multiplied by a depth factor. 97

Chapter 5

5.1 Extended Bouguer anomaly map showing general subdivisions. 104 5.2 Extended Bouguer anomaly map showing the location of the major geological features. 105 5.3 Complex wavforms resulting from the summation of two sine wave components of frequency f and 2 f 111 5.4 Frequency domain representation of the waveforms shown in Fig 5.3. I ll 5.5 Typical logarithmic power spectrum plot. 115 10

5.6 Extended Bouguer gravity field upward continued by 10km. 117 5.7 Residual gravity map having removed the upward continued field (Fig 5.6) from the extended Bouguer anomaly field (Fig 5.1). 117 5.8 Logarithmic power spectrum of the entire region. 119 5.9 Logarithmic power spectrum of the foreland region. 119 5.10 Logarithmic power spectrum for the eastern Muller Anticline area. 121 5.11 Logarithmic power spectrum for the western Muller Anticline area. 121 5.12 Wavelength filtered extended Bouguer anomaly field - 100km+ passed. 123 5.13 Wavelength filtered extended Bouguer anomaly field - 50km+ passed. 123 5.14 Wavelength filtered extended Bouguer anomaly field - 0-100km passed. 125 5.15 Trend filtered extended Bouguer anomaly field - trends from 15° to 75° passed. 128

Chapter 6

6.1 Calculation of the gravitational attraction of an infinite body. 132 6.2 Location of gravity profiles and models. 133 6.3 Model of the Darai Escarpment showing synthetic and actual Bouguer anomaly values for a range of topographic densities. 138 6.4 Results of the Hides BHGM survey. 140 6.5 Effect of karstification on BHGM results. 143 6.6 Effect of dipping beds on BHGM results. 144 6.7 Profile 1.1 148 6.8 Profile 1.2, Model a 149 6.9 Profile 1.2, Model b 150 6.10 Profile 1.3, Model a 152 6.11 Profile 1.3, Model b 153 6.12 Profile 1.3, Model a, detail. 154 6.13 Profile 1.3, Model b, detail. 154 6.14 Profile 2.1 156 6.15 Profile 2.2 158 6.16 Profile 3 160 6.17 Profile 4 161 6.18 Profile 5 163 11

6.19 Profile 6.1 165 6.20 Profile 6.2 167 6.21 Profile 7 168 6.22 Map showing extend of lower crustal bulge. 171 6.23 Location of Hides profiles. 174 6.24 Hides profile 1, Model a 175 6.25 Hides profile 1, Model b 175 6.26 Hides profile 2, Model a 177 6.27 Hides profile 2, Model b 177 6.28 Location of the Kutubu/Nembi profile. 179 6.29 Kutubu/Nembi profile. 180

Chapter 7

7.1 Geological summary map of the Muller Anticline region. 184 7.2a Free-air anomaly map 191 7.2b Regions showing no isostatic response to crustal loading. 191 7.3 Density structure of the eastern fold belt. 193 7.4 Mesozoic isopach map. 193 7.5 Diagram showing the Mesozoic crustal structure of the Papuan Basin. 197 7.6 Major volcanic trends of the Papuan fold belt. 198 12

List of Equations

4.1 Observed Gravity datum correction formula 80 4.2 Latitude Correction 80 4.3 Free-Air Correction 81 4.4 Bouguer Correction 81 4.5 Line-Mass Approximation 83 4.6 Thin Rod Approximation 83 4.7 Vertical Gravitational Attraction of a Prism 83

List of Plates

4.1 Isometric view from the SW of the high data density region of the DTM. 78 4.2 Isometric view from the SSW of the area covered by theKaroma 1:100000 topographic map. 78 4.3 Isometric view from the SSE of the area covered by theKoroba 1:100000 topographic map. 79 4.4 Close-up isometric view of the Lavani Valley. 79 5.1 Wavelength filtered extended Bouguer anomaly map - 0-100km passed. 125 5.2 Trend filtered extended Bouguer anomaly map - trends 15° to 75° passed. 128

List of Tables

1.1 Summary of the Geological Provinces. 19 3.1 Papuan fold belt - gravity survey details. 54 6.1 Typical rock density values. 136 6.2 Density values of the rock units of the Papuan fold belt. 136 6.3 Effect of terrain correction on BHGM apparent densities. 141 6.4 Density contrasts used for modelling. 146 6.5 Distances along profiles between fold belt structures and deep crustal features. 172 13

Acknowledgements

This project was sponsored by British Petroleum (Australia) Ltd and I am grateful to them for their financial support, for their logistical support of the fieldwork in PNG, especially the help of Mike Nosiara in Port Moresby and Marcus Parrish and Peter Fenessey at Tiengo, and for their provision of data. I also acknowledge the valuable help of Peter Morris and Andy Billings at BP in London, and John Bennett in Melbourne.

At UCL, the advice and encouragement I received from Steve Kaye, who spent many long periods helping with problems of all sorts, and developing geophysical computer software, was a major benefit to the study. Paul Schooling, computer manager in the Geography Deparment at UCL, also provided invaluable help with the reading of magnetic tapes.

The assistance of all the technical and ancillary staff in the Geology Department at UCL has been greatly appreciated, particularly that of Ron Dudman, who has always been willing to help with equipment and administrative difficulties.

Finally I must thank John Milsom, who initiated this project and gained the financial backing of BP. John has always provided useful and practical advice on all aspects of the study, and has spent a considerable amount of time proof reading and criticising this thesis. 14

Aims and Introduction

Papua New Guinea is situated at the boundary between the Australain and Pacific Plates, one of the few areas in which active arc-continent collision is currently taking place. The Papuan Fold Belt, which lies along the southern (continental) edge of the resulting deformation zone and comprises folded and thrust Mesozoic and Cenozoic sediments forming mountains which rise to in excess of 3000m, is therefore an ideal place to study the effects of accretionary processes on a continental margin. These effects may then be used to help unravel the histories of other, older collision events elsewhere.

Geological and geophysical investigations have been carried out in the fold belt region since the early part of this century, primarily by oil and mineral exploration companies searching for hydrocarbons and gold. Large amounts of data have been collected, but the evolution of the region is still not fully understood.

This study aimed to bring together all the existing gravity data from the fold belt region, and to collect additional data where necessary, and to process it all to a uniformly high degree of accuracy, including the application of detailed terrain corrections. This required the representation of the rugged fold belt topography in digital form, and the production of a digital terrain model was an important part of this study.

Interpretation of the resulting extended Bouguer anomaly data using 2D forward modelling and digital filtering to enhance specific features within the data has enabled new ideas on both the near-surface structures of the fold belt and the deeper, regional structure of the region, to be developed. These will be of use to the hydrocarbon exploration industry and have also improved our understanding of the processes involved in the formation and evolution of the fold belt. 15

1. The Regional Geology of Papua New Guinea.

1.1 Introduction

The island of New Guinea in situated to the north of Australia, at the eastern end of the Indonesian archipelago (Fig 1.1). Politically it is divided into two parts along the 141°E meridian, with the Indonesian province of Irian Jaya to the west, and the state of Papua New Guinea to the east. This discussion considers the geology of mainland Papua New Guinea (PNG), and the islands of the Bismarck Archipelago, including New Britain and New Ireland. This is an area of some three million square kilometres divided almost equally between land and sea, which includes a great diversity of physiographic and geological provinces, ranging from the flat, swampy plains of southern mainland PNG, to the 4500m peaks of the central ranges, and the volcanic islands of the Melanesian arcs.

1.2 Regional Setting

Papua New Guinea lies at the boundary of the Australian and Pacific plates, and includes one of the few areas where arc-continent collision is currently taking place. Convergence across New Guinea has been calculated to be 122mm/yr in a WSW-ENE direction (Abers and McCaffrey, 1988), though the mode of accommodation of this shortening is not fully understood. The regional tectonics are complicated somewhat by the presence of the Melanesian microplates, the Caroline Sea, Solomon Sea, and North and South Bismarck plates (Fig 1.2). However, with the possible exception of the convergence of the South Bismarck and Solomon Sea plates, which is absorbed at the deep New Britain Trench, the relative motions of these plates are small when compared with the overall Australia-Pacific movement.

1.3 Major Geological Provinces

Nine geological provinces have been identified (Milsom, 1985), and are shown in Figure 1.3. The two major upland areas, the Central Orogen, and the North New Guinea Province, are separated by an area of low lying, swampy valleys, the Intermontane Depressions, which approximately delineate the boundary between the island arc units to the north, and the 16

A rafura

1 20 1 30 1 40 150 Fig 1.1 Location of Papua New Guinea

120 140 160

y < i ‘

Philippine Sea Plate

Pacific Plate

Eurasian Plate Caroline Plate

• o •Bismarck Rates

Solomon Sea Rate

Indo-Australian Plate

120 140 160

Fig 1.2 Plate tectonic setting of Papua New Guinea 145 150

Melanesian Arcs

‘X JntermorrtaneW\X. Bismarck Sea

. ' x v\ 1 "'X^Centrai OrogenN^

Southern Plains Solomon Sea

Papuan Peninsula 0

145 150

Fig 1.3 Major geological provinces. 18

Australian continental margin sequences of the Central Orogen and its foreland basin, the Southern Plains, to the south. The Papuan Peninsula is considered separately because it has a significantly different geological history to the surrounding provinces.

The islands of western Melanesia, and their intervening seas, have been subdivided on the basis of their geological affinities. These are the island arc units of New Britain, and the geologically similar part of northern New Guinea; the volcanic islands of the Melanesian Arcs, and the Bismarck and Solomon Seas.

The geology of each of these regions is described in the remainder of this chapter, and summarised in Table 1.1.

1.3.1 The Southern Plains

The Southern Plains are an extensive area of gentle, jungle-covered terrain, generally not more than 75m above sea level. Fluvial plains covering an area in excess of 4000km2 are characterised by broad meandering rivers with extensive floodplains, which deposit sediment in the form of point bars, levees, floodplain muds and crevasse splays. The plains are bounded to the south by a coastal belt of shifting deltas, lagoons, beaches, and bars (Brown, 1977).

Geologically, the Southern Plains are underlain by slightly deformed north Australian continental margin. Late Permian crystalline basement is overlain by a sequence of Late Triassic to Pliocene shallow marine sediments deposited on the north Australian continental shelf. Following the Neogene uplift of the highlands to the north, the Southern Plains now form a foreland basin in which clastic sediments derived from the highlands are deposited (Pigram et al., 1989).

1.3.2 The Central Orogen

The central orogenic belt consists of mountain ranges up to 4500m high, separated by deep valleys, often only 1000m above sea level. Geologically it separates the relatively undeformed Australian continental margin sediments to the south from the accreted island arc units to the north. The region has a very varied physiography, including (Brown and Robinson, 1982, 19

Geological Province Geology

Upper Palaeozoic continental basement overlain 1 Southern Plains by relatively unreformed Mesozoic-Recent sediments. Upper Palaeozoic continental basement overlain 2 Central Orogen by Mesozoic-Late Neogene sediments. Extensive Pliocene-Recent deformation. Metamorphic and ultramafic basement overlain 3 Inteimontane Depressions by Miocene-Recent sediments. Igneous and Metamorphic rocks including 4 North New Guinea Province ultramafics and acid to intermediate intrusives, dating from Jurassic to Earliest Miocene. Metamorphosed sedimentary basement, with 5 Papuan Peninsula Cretaceous-Eocene submarine basalt and Mesozoic ultramafics. Eocene-Lower Miocene and Pliocene-Recent 6 New Britain island arc volcanics with Miocene sediments. Oligocene and Pliocene-Recent volcanics with 7 Melanesian Arcs Miocene sediments. Consists of two smaller basins, separated by a 8 Bismarck Sea seismically active zone. Southern basin contains thick sediments. Oligocene-Miocene age, overlain in the west by 9 Solomon Sea thick deltaic sediments.

T able 1.1 Summary of the Regional Geology of Papua New Guinea. 20

Davies, 1983): (i) glacial landforms of the higher peaks (ii) extensive karstified limestone plateaux with no surface drainage (iii) high relief areas of karstified limestone with deeply incised river channels forming spectacular gorges (iv) volcanic centres showing characteristic radial drainage (v) deep valleys bounded by cliffs formed of volcanic and clastic material (vi) broad, shallow valleys, infilled by recent sediments, separated by rounded ridges of limestone or clastic sediments. which reflect the complex geology and variety of structural styles.

Although Lower Palaeozoic rocks have been reported in the Central Orogen of Irian Jaya, the oldest known rocks of the Papuan Central Orogen belong to the crystalline basement, thought to be Permian in age. These are overlain by remnants of Upper Permian to Lower Triassic rocks, and up to 7km of terrigenous sediment which was deposited in the Mesozoic Papuan Basin (Jenkins, 1974). This is in turn overlain in the SW by a Tertiary sequence of shallow water limestone and shale, and to the NE by trough tuibidites and volcanigenic sediments, which are in fault contact to the NE with a belt of uplifted continental basement and ophiolitic units, known as the Mobile Belt (Bain, 1973; Dow, 1977). The Tertiary sequences are capped in places by late Tertiary-Quaternary volcanoes.

Structural style varies across the orogen from wrench faulting in the New Guinea Mobile Belt (NGMB) to the northeast, to thrusting and folding in the Papuan Fold Belt to the southwest.

133 Intermontane Depressions

Separating the Central Orogen from the North New Guinea Province to the north is an area of low lying, swampy depressions, which vary in width from around 100km in the case of the Sepik Plains to the west, to only 5km in parts of the Ramu-Markham divide to the east.

Beneath the Sepik Plains, the metamorphic basement of upper greenschist and blueschist facies, in places interleaved with cumulate and tectonite ultramafics, is regarded by some (eg Pigram 21 and Davies, 1987) as being formed from accreted terrenes which had been rifted from the Australian continental margin during the Mesozoic. The basement is overlain by up to 3km of Miocene to Recent marine and terrigenous clastic sediment, which includes volcanic detritus. Interpretation of aeromagnetic data postulates a major fault, the Turu-Keram Fault, running beneath the Sepik Plains, separating the rocks of oceanic affinity to the north from metamorphosed continental sediments to the south (Davies and Hutchison, 1982).

The Ramu-Markham valley, which separates the Central Orogen from the Sarawaged and Adelbert Ranges of north eastern PNG, is flanked on both sides by steep mountain ranges. Alluvial fans shed detritus into the valley from both south and north, those on the northern flank currently being more developed, indicating the continuing uplift and erosion of the northern ranges (Tingey and Grainger, 1976; Crook, 1989). Along with the recently folded Quaternary fluvial and lacustrine sediments of the valley floor, the alluvial fans cover the Ramu-Markham fault (Tingey and Grainger, 1976), assumed to be a major left-lateral wrench fault probably formed in the Pliocene (Pigott et al., 1985). Together, the Turu-Keram and Ramu-Markham faults represent the northern edge of the suture between the post-Cretaceous island arc units to the north, and the Mesozoic and Tertiary continental margin sediments of the Central Orogen.

1.3.4 The North New Guinea Province

The north coast of PNG to the west of the Sepik River is flanked by a narrow belt of mountains, only 25km wide, rising up to 1800m. The Torricelli Mountains, comprising of narrow ridges separated by deeply incised valleys, lie to the north and west of the lower and less rugged Prince Alexander Mountains. In places, the Torricelli Mountains are separated from the coast by a narrow, low lying, coastal plain; elsewhere they descend steeply to the coast (Hutchison and Norvick, 1978; Norvick and Hutchison, 1980).

The Prince Alexander Mountains are formed of Middle Jurassic to Miocene ultramafic rocks, with acid to intermediate intrusives and metamorphics. These contrast with the Late Cretaceous to Miocene intrusives (Torricelli Intrusive Complex) and the Palaeocene? to Miocene basic to intermediate Bliri Volcanics (Pigram and Davies, 1987) of the Torricelli Mountains, which comprise hypersthene-normative rocks with both calc-alkaline and tholeiitic island arc affinities (Hutchison and Norvick, 1978). The Prince Alexander Mountains also contain faulted blocks of 22 the metamorphic basement of upper greenschist facies inferred to exist under the Sepik Plains (Davies and Hutchison, 1982). Both regions are covered by Miocene to recent clastic sediments.

The interpreted history of the area is that prior to an eariy Miocene arc-continent collision, late Cretaceous to Eocene sediments were metamoiphosed. Collision resulted in uplift of these metamorphics, and emplacement of the mafic and ultramafic rocks (Milsom, 1985).

It should be noted that although the boundary between the North New Guinea Province and the Intermontane Depressions is distinct at the surface, the complex structure of the basement, which includes up to eight proposed terranes of both continental and oceanic affinities (Pigram and Davies, 1987), and the presence of metamoiphic rocks of continental affinity in the Prince Alexander Ranges, preclude precise separation at basement level.

1.3.5 The Papuan Peninsula

Forming the eastern part of Papua New Guinea, the Papuan Peninsula is dominated by the 1800- 3500m peaks of the Owen-Stanley Range, which are flanked by lower ranges of volcanic and ophiolitic origin.

The core of the mountain ranges comprises metamoiphosed continental slope sediments, which suggests that the Papuan Peninsula is a part of the Australian continental margin which was rifted away during the opening of the Coral Sea (Milsom, 1985). Pieters (1978) identifies a separate basic metamorphic group on the northern flank of the Owen Stanley Range. Two other pre-Miocene units have been identified in the region: a sequence of Cretaceous and Eocene submarine basalts and gabbroic intrusives which are found in the extreme east of the peninsula, and the Papuan Ultramafic Belt, a piece of Mesozoic ocean floor, which was probably emplaced during the Oligocene or Early Miocene, and which forms the northeastern mountain ranges (Davies and Smith, 1971, 1974).

Terrestrial and marine volcanism has accompanied uplift since the Late Oligocene, and Recent outpourings of shoshonitic lavas are indicative of the later phases of stabilisation and elevation of the orogen (Davies and Smith, 1974; Pieters, 1978). 23

1.3.6 New Britain

The island of New Britain, together with the Sarawaged, Adelbert and Finisterre Ranges of northern PNG, forms a continuous line of uplifted pre-Pliocene volcanics and sediments, rising to 4100m on mainland New Guinea, and 1800m on New Britain.

The oldest rocks, both on the mainland and on New Britain, are Eocene island arc volcanics intermixed with argillaceous sediments. These grade upwards into a Lower Oligocene to Lower Miocene volcanic sequence, known as the Finisterre Volcanics on the mainland, and the Merai and Kapuluk Volcanics in New Britain. These are overlain by a thick sequence of Mid Miocene to Early Pliocene limestones, representing a period of relatively stable conditions before island arc volcanism returned in the Pliocene (Robinson et al., 1976; Rybum, 1975).

Initial uplift of the area had occurred by the Late Miocene, and collision of the western end of the arc-terrane with the continental block followed soon after the onset of the Pliocene island arc volcanism (Milsom, 1985). Erosion has been more vigourous in the western part of the province, where uplift is greater (2.0-7.6m/1000yrs, Crook, 1989), and where there are many Pleistocene fluvial, fluviodeltaic and lacustrine deposits.

1.3.7 The Melanesian Arcs

The Melanesian Arcs Province is bounded to the north by the Manus Trench (eg Exon and Marlow, 1988), also known as the Outer Melanesian Trench (Milsom, 1985), a 6km deep trench which extends from the north coast of New Guinea close to the Irian Jaya border in the west to the northern end of the Solomon Islands. Two arcs have been identified in the region, an outer arc (the Mussau-Feni Arc), and an inner arc, formed by the islands of New Ireland, New Hanover and Manus.

The outer arc, though not volcanically active, exhibits thermal activity indicative of cooling magma, and a seismically defined magma chamber is thought to lie beneath the arc near Mussau (Dadisman and Marlow, 1988). All the islands show evidence of a mid Tertiary period of igneous activity, a second late Cenozoic alkaline phase, and raised platforms of Miocene and Pliocene limestones. The lack of seismic activity in the outer arc suggests that subduction along 24

the Manus trench is intermittent or has recently commenced (Dadisman and Marlow, 1988; Johnson, 1979).

The inner arc has a basement of Oligocene island arc volcanics similar to those of New Britain (Milsom, 1985), which are overlain in the west (Manus) by Upper Miocene to Pliocene volcanigenic sandstones and Pliocene to recent volcanics, and in New Ireland by Miocene limestones similar to those in New Britain, and by Pliocene volcanics.

1.3.8 The Bismarck Sea

Lying between the island arc of New Britain and the Manus-New Ireland chain, the Bismarck Sea consists of two smaller basins, the Manus Basin to the north, and the New Guinea basin to the south, separated by the NW-SE Willaumez-Manus rise.

The most striking geophysical feature of the sea is a narrow zone of earthquake foci crossing it from west to east (Taylor, 1979). Fault plane solutions have been interpreted as showing both left-lateral shear and spreading offset by left-lateral transforms. Whichever process is dominant, the Bismarck Sea is cut into two parts by this major structure. If the Manus Trench is active, then a small plate, the North Bismarck Plate, must exist to the north of the fault zone; but if the trench is inactive, the fault zone marks the present southern boundary of the Caroline Sea Plate (Ryan and Marlow, 1988).

There are thick sediments on the southern plate, close to the Ramu and Sepik deltas and close to the New Britain volcanoes. In contrast, sediment in the Manus basin is thin, with little input from the north. The lack of sediment in the Manus basin has been used as evidence of recent spreading (Milsom, 1985).

1.3.9 The Solomon Sea

Bounded to the north by the New Britain and North Solomons trenches, and to the south by the Trobriand Trough, the Solomon Sea is thought to be of Oligocene or Miocene age (Milsom, 1985). Geophysical evidence suggests the crust is 10km thick in the centre and east, and up to 25

33km thick in the west, where it is overlain by thick deltaic sediments. This crustal thickening has been explained in terms of a triple junction at the western end of the Solomon Sea, where the two upper plates converge above the doubly-subducting Solomon Sea Plate (Davies et al., 1987).

1.4 The Geological History and Evolution of Papua New Guinea

This section describes the geological history of Papua New Guinea using the evidence from the geological provinces described in the previous section. The geological evolution of the region is then explained using evidence from this geological history.

1.4.1 Mesozoic Geological History

In early Mesozoic times, Papua New Guinea formed part of the stable Australian continental margin. Subsidence of this margin started in the Triassic, but deposition did not become widespread until the Upper Jurassic. To the northeast, the Late? Triassic agglomerates and volcanolithic sandstones of the Kana Volcanics overlie crystalline basement, and are indicative of andesitic to dacitic volcanic activity of unknown tectonic affinity (Brown et al., 1980). To the south and west, Late Triassic? to Mid Jurassic coarse clastic sediments, such as the Bol Arkose, which were apparently deposited locally in troughs, overlie Mid Triassic (236Ma) acid igneous basement in the Southern Plains region (Harding, 1969). In the western fold belt basement comprises Lower Triassic granite (214+/-3Ma, Jenkins, 1974) and granodiorite (minimum 222+/- 4Ma, Page, 1976).

During the Mid to Late Jurassic, continental fluviodeltaic and marginal marine sediments progressively transgressed over the continental margin (Brown et al., 1980) forming a sequence of alluvial fan and braided river deposits, grading northeastwards into a sequence of deltaic and marginal marine sediments with coal beds, and marine shales in the basinal areas further to the northeast By the Oxfordian, marine deposition was present throughout the region, with marginal marine deposition on the continental platform to the southwest, and the deposition of marine shales on the platform margins. 26

Other Mesozoic rocks, referred to as being of Sepik association, and consisting mostly of deep water marine sediments and volcanic detritus, are found to the north of the Central Orogenic Belt and are inferred beneath the intermontane depressions. They have been interpreted as having been deposited in a deep trough to the north of the continental margin (Brown et al., 1980).

The thick band of sandstones, such as the Toro Sandstone, regarded as a sheet of coalesced barrier-bar deposits, and dated as Late Tithonian to Neocomian, are evidence of a marine regression close to the Jurassic-Cretaceous boundary (Brown and Robinson, 1982). This regression also accounts for the occurrence of sandstones within the marine shale sequences of the deeper water areas to the northeast. However, by the Late Neocomian, a transgressive regime had been re-established, resulting in the deposition of marginal marine sediments over much of the region, the exception being the Kubor region, which was again emergent, and acting as a local source of volcanic detritus (Brown et al., 1980). Volcanic detritus was also deposited seawards of the shelf, where it is found interbedded with siltstones and shales.

Deposition appears to have ceased in the Late Cretaceous, probably Maastrichtian, and uplift, block faulting and erosion of the southern parts of the region occurred as a result of thermal updoming (Brown et al., 1980). Today the Cretaceous is represented on the platform by 750m of pre-Turonian (mostly Early Cretaceous) sediments, as compared with 4500m of sediment representing most of the Cretaceous period, at least 2750m of which is Early Cretaceous, in the Southern Highlands Basin (Jenkins, 1974).

Elsewhere in New Guinea, Cretaceous oceanic crust is inferred beneath the arcs of the North New Guinea Province (NNGP), and Cretaceous volcanic detritus is found with the trough sediments of Sepik association (Brown et al., 1980).

1.4.2 Mesozoic Evolution

The Mesozoic geological history of the region has been explained by several authors (eg Brown et al., 1980; Cullen and Pigott, 1989; Pigram and Davies, 1987; Pigram and Panggabean, 1984), with varying degrees of agreement. It is generally accepted that the Early Mesozoic setting of the region was that of the subsiding Australian continental margin. The localised Lower Jurassic trough sediments are interpreted as having been deposited in rifted grabens, which formed as a 27 result of the subsidence and break-up of the continental margin, and the uplifted areas, such as the Kubor Anticline and Pasca Ridge represent the intervening horst blocks. However, by the Late Jurassic, the basement topography had been buried, and deposition was influenced by regional sea-level changes and episodes of volcanic activity, with only occasional associated uplift.

The Late Cretaceous updoming, which was responsible for the major base-Tertiary unconformity, is thought to have developed as a result of the opening of the Coral Sea to the southeast, and was accompanied by extensional faulting (Brown et al., 1980).

Pigram and Panggabean (1984) relate the Mesozoic sequence of Papua New Guinea to a rift-drift sequence, in that the New Guinea margin formed by rifting in the Triassic, and the rifted micro-continents are thought to have been accreted northern New Guinea and/or to eastern Indonesia (Hill, 1989b). Pigram (in press) suggests that only fragments derived from the Late Cretaceous rifting events could have been accreted to eastern Indonesia or northern New Guinea, and those fragments resulting from the Jurassic rifting must have been carried into the eastern Tethyan region and possibly accreted to the Laurasian landmass.

The mode of evolution of the northern parts of New Guinea during the Mesozoic is less clear. Some authors regard the rocks of Sepik association as having formed in a trough to the north east of the continental shelf, sourced, like the shelf sediments, from the southwest The inclusion of volcaniclastic sediments has been interpreted as showing the presence of a volcanic arc to the northeast which formed in the Late Jurassic-Early Cretaceous and was particularly active during the Cretaceous (Brown et al., 1980). To further complicate the situation, the rapid facies changes from the shelf to the trough sequences suggests that the two areas may have been juxtaposed more by recent faulting.

Brown et al. (1980) suggest that the Pacific plate was subducted to the southwest beneath an island arc to the north of New Guinea, and that the arc was separated from the continental shelf by a back-arc basin, in which the trough sediments were deposited. It would seem likely that such a subduction zone would initially form at the continental margin (Andean style) and would later migrate northwards, but there is no volcanic evidence of this (Milsom, 1985). It is also unclear what happened to the intervening oceanic crust following the later collision of the arc with the continental margin. Initially, northwards subduction does seem to present fewer 28 difficulties (Milsom, 1985), but it appears to necessitate the transport of volcanic detritus across the trench to be deposited on the Australian margin, which is difficult to envisage (Cullen and Piggot, 1989). However, it has been suggested that the source of volcanic rocks may not be arc- related at all, but a result of intra-cratonic rifting, the break-up of the continental margin, from the Early Cretacoeus initial rifting of the Coral Sea (Hill, 1989b; Pigram and Panggabean, 1984). This perhaps more likely situation requires neither the transportation of detrital material across a trench, nor the "disappearance" of a stretch of oceanic crust.

1.4.3 Cenozoic Geological History

The base Tertiary unconformity indicates that the Cenozoic commenced with a period of erosion, which was more dominant in the south where most of the Late Cretaceous succession was eroded (Jenkins, 1974). This erosion was accompanied by extensional faulting, with downthrows generally to the north (Hill, 1989b), preserving the Late Cretaceous succession in the north, and allowing the deposition of Palaeocene calcareous mudstones (Moogli Mudstones).

Further to the north, the Early Cenozoic (Palaeocene?) volcanics of the North New Guinea Province are thought to represent the early developments of a south-facing volcanic arc to the north of the Australian continental margin, which formed to accommodate the increasing northwards movement of the Australian plate.

By the Late Eocene, deposition had returned to the western parts of the region, with up to 1000m of Mendi Group limestones being deposited on the northern edge of the continental margin; to the south a thinner carbonate sequence was formed, and clastic sediments were deposited in the Morehead sub-basin (Durkee et al., 1986). This change from the Mesozoic predominantly clastic sedimentation to carbonate depositional regimes marks the movement of the continent into more tropical climatic zones.

To the east, the Eocene volcanics (including pillow lavas) of the Aure Trough, separated the continental margin from the Papuan Peninsula. Davies (1980) suggests that metamorphism, accompanied by emplacement of the ophiolite of the Papuan Ultramafic belt, also occurred during the Eocene, though other writers (Pigram, in press; Pigram and Davies, 1987) suggest that the Eocene event in the Ultramafic Belt was a result of inter-terrane collision, and accretion did 29 not take place until the Late Oligocene or Early Miocene. The Eocene event in the Ultramafic Belt was followed by the Late Eocene to Oligocene emplacement of the ophiolites of the New Guinea Mobile Belt (NGMB) to the west (Marum Ophiolite and April Ultramafics), and associated metamorphism. Whether this obduction was a result of continent-arc collision, terrane accretion, or just an accelerated convergence between the Australian and Pacific plates (Hill, 1989b) is the subject of much discussion.

Slow subsidence, accompanied by patchy deposition and some erosion, continued to the south and west throughout the Eocene and Oligocene, and it was not until the Early Miocene that sedimentation commenced at any significant rate, possibly as a result of foreland basin formation and the development of the Central Highlands (Pigram et al., 1989). A broad carbonate shelf, flanked by barrier reefs (Borabi Reef trend) became established on the Fly Platform to the south, whilst in the deeper waters to the northeast, deep water limestones and their turbiditic clastic equivalents were deposited. Pinnacle reefs (eg Pasca reefs) formed on highs on the continental slope (Duikee et al., 1986).

Carbonate sedimentation ceased to the northeast in the late Early to early Middle Miocene (Durkee et al, 1986), as elastics and volcanic detritus (Pynyang Fm, Orubadi Fm and Burgers Fm) derived from an island arc immediately to the northeast of the continental margin inhibited carbonate deposition. Throughout the Late Miocene the locus of carbonate deposition was pushed further south (Pigram et al, 1989), though deposition of clastic sediments on the continental slope and in deeper water areas to the east, including the Aure Trough, continued until the Late Miocene.

The mid-Miocene onset of volcanism observed in the NGMB followed the abrupt Early Miocene cessation of volcanism in the NNGP, and resulted in uplift and major strike-slip faulting. Deformation in the Papuan fold belt commenced in the latest Miocene, and with the continuing deformation and uplift during the Pliocene, molasse deposition took place in depressions within the orogenic belt and the NNGP. Plio-Pleistocene sedimentation continued on the platform to the south, with detritus derived from the orogen to the north, and terrestrial strato-volcanoes forming high mountains developed in areas of the fold belt and NGMB during the Pleistocene. 30

1.4.4 Cenozoic Evolution

The first major event to affect the region during the Cenozoic was the onset of active spreading in the Coral Sea during the Early Palaeocene (62Ma, Weissel and Watts, 1979). This phase of extensional tectonism was responsible for the uplift of the southern parts of the continental shelf, now represented by the sub-Terdary unconformity, the block faulting of the continental margin, and the formation of the Aure Trough, which can be regarded as the northwestern limit of the Coral Sea basin. Spreading is thought to have ceased by the Late Palaeocene (56Ma).

Increased convergence between the Australian continent and the Pacific plate commenced in the Eocene as Australia separated from Antarctica. This may have been responsible for the heating and metamorphism of the Mobile Belt, though it is also possible that orogenesis associated with ophiolite emplacement was the cause (Hill, 1989b). However, only the ophiolites of the Papuan Peninsula, and not those of the Central Orogen, are thought to have been emplaced at this time (Cullen and Pigott, 1989). In addition, Pigram and Davies (1987) suggest that the deformation of the Papuan Peninsula ophiolites at this time was a result of terrane amalgamation, and emplacement of the Papuan Ultramafic Belt followed in the Late Oligocene or Early Miocene.

Another consequence of the increased convergence of the Australian and Pacific plates may have been the development of a northwards dipping subduction zone to the north of the continental margin, its associated volcanics now incorporated into the NNGP (Hill, 1989b). The relationship of this newly-formed arc with the arc which may have been active during the Cretaceous is not clear. If Cretaceous subduction was to the south, and back-arc spreading was active, the Cretaceous arc could have migrated several thousands of km to the north of the region by this time. However, if Cretaceous subduction was also to the north, then collision would have occurred, and Cretaceous volcanic units should have been accreted to the northern New Guinea margin. It is possible that these volcanic units could nave been incoiporated into those of the Mobile Belt in an event associated with the Eocene-Oligocene ophiolite emplacement, or form part of the "scrapland terranes" (Silver, 1988) which lie between the Ramu-Markham and Bismarck fault zones and have suspect stratigraphy with respect to the Australian continental margin. However, there is no evidence of a major tectonic event during the Late Cretaceous or Eariy Palaeogene to support this. Of course, if there was no Cretaceous arc, or if the "newly- formed" arc was simply regeneration of the old arc, these problems do not arise. 31

As Australia continued its rapid northward movement throughout the Eocene into more tropical climatic zones, carbonate sedimentation became dominant. Though much of the southern part of the shelf was close to sea-level at this time, and little sedimentation took place, up to 1000m of carbonate sediments were accommodated on the rifted shelf edge, and pelagic sedimentation continued on the continental rise and abyssal plain (Pigram et al., 1989).

Throughout the Oligocene, subduction continued to the north of the New Guinea margin, and the Australian continent continued its migration northwards towards the arc. The relative positions of the continent and arc at this time are the subject of debate. Milsom (1985) invokes arc-continent collision during the Oligocene. Pigram et al. (1989) cite the emplacement of the ultramafic rocks of the Mobile Belt, and the lack of Mid-Late Oligocene sediments as evidence of a mid-Oligocene collision and Pigram (in press) suggests the mid Oligocene (30Ma) accretion of the Sepik Terrane was the first in a series of terrane docking events which increased lithospheric loading and provided a structural downwarp in which Miocene carbonates accumulated in the starved areas to the south, and elastics derived from the developing orogenic belt were deposited in the deeper northern parts of the basin.

Hill (1989b) suggests that the increased convergence during the Oligoceneresulted only in ophiolite emplacement. This was sufficient to cause crustal subsidence, but did not involve collision of the continent with the arc, which was still several thousand km to the north.

Many authors (eg Cooper and Taylor, 1987; Hill, 1989b; Milsom, 1985) suggest that by the Middle Miocene, subduction to the southwest beneath New Guinea was active. As a result of the arc-continent collision, active volcanism in the NNGP appears to have ceased early in the Miocene, and the southwest-dipping subduction zone developed to take-up the continued convergence of the Australian and Pacific plates. This subduction, which produced the arc volcanoes of mainland New Guinea, known as the Maramuni Arc (Dow, 1977), is regarded as the onset of subduction in the New Guinea Trench and Trobriand Trough. This cannot be reconciled with those who believe the arc of the NNGP still to be | some distance to the north of New Guinea, and Hill (1989b) suggests that subduction beneath the Maramuni Arc was related only to the southwestwards subduction in the now inactive Trobriand Trough, the Solomon Sea plate having been doubly subducting beneath the continental margin to the south, and the arc of the NNGP and New Britain to the north. Pigram (in press) suggests that the 32 volcanoes of the Maramuni Arc may not be subduction related, but may be a result of the increased loading of the continental margin.

Following a volcanically inactive period, the collision of the western end of the New Britain arc with the continental margin to the south occurred in the Late Miocene. This collision is regarded as resulting in the onset of deformation in the fold belt Although collision occurred in a relatively small area, it is suggested that the collision restricted subduction in both the New Guinea (Milsom, pers comm) and Trobriand Trench systems, greatly increasing the convergent stresses across mainland New Guinea. Pigott et al. (1985) show that the Middle-Miocene to recent convergence rate between the Australian and Pacific plates across New Guinea is almost double that in the Eocene to Early Miocene. The observed increase in volcanic activity in New Britain during the Pliocene could result from increased subduction rates in the New Britain trench would almost certainally follow the decline of subduction in the Trobriand Trough.

The origin of the Quaternary volcanic centres in central mainland PNG are regarded as being the result of either decompressional melting of the upper mantle resulting from Cretaceous subduction (Cullen and Pigott, 1989), or the southwards subduction of the Solomon Sea plate beneath northern New Guinea, or the ongoing deformation of the fold belt (Hill, 1989b; Ripper and McCue, 1982).

* 33 2. The Geology of the Papuan Fold Belt

2.1 Introduction

The Papuan fold belt extends for about 500km southeast from the Irian Jaya border to the Papuan Gulf. It forms a belt up to 100km across, in which sub-parallel, southward-verging thrust faulted anticlines dominate the geological structure.

During the Mesozoic and Early Cenozoic the area was part of the Papuan basin, in which littoral, deltaic and shelf sedimentation took place on the subsiding north Australian continental margin. The area has been affected by two tectonic episodes, the first of which occurred during the latest Cretaceous to Palaeocene, when the breakup of the Australian continental margin and the opening of the Coral Sea resulted in uplift and tilting of the basin to the northwest (Durkee et al, 1986). The second episode resulted from the convergence of the Australian and Pacific plates, and led to the progressive southwestwards migration of a deformation front from the northern and eastern margins of the Papuan Basin. The resulting fold and thrust structures have been recognised as potential hydrocarbon traps, and exploration of the fold belt has been continuous since the early parts of this century.

2.2 Stratigraphy of the Papuan Fold Belt

The Papuan fold belt is covered by the 1:250000 series of geological maps published by the Australian Government Publishing Service on behalf of the Geological Survey of Papua New Guinea, and the Bureau of Mineral Resources (Australia). These maps, along with their accompanying notes, particularly the work of Davies and Norvick (1974), Davies (1983), Brown and Robinson (1982), and Bain and MacKenzie (1974), form the basis of this discussion, and a geological summary map (Fig 2.1) illustrates the location and extent of outcrop of the geological units described in this section.

Although Lower Palaeozoic rocks have been reported in Irian Jaya (Visser and Hermes, 1962), the oldest known rocks of the Papuan fold belt are the Upper Palaeozoic Omung Metamoiphics, which are known to be older than the intruding Kubor Granodiorite, dated isotopically as 215-244Ma (Page, 1976). These Permo-Triassic rocks are exposed in the core of the Kubor 34

l r I 85 Si *1 i I » I H § £§ &5 8* * 5s s. CO £ $ 3 <5 S |S 0 <5 0 !□□□□□■ Fig Fig 2.1 Geological summary map of the Papuan fold belt. 35

anticline. To the south and west, basement is represented in outcrop by an acidic intrusive, dated at 214 +/-3Ma, known as the Strickland Granite.

Immediately oveiiying the basement are remnants of Upper Permian to Lower Triassic formations, between which correlation is difficult. To the south-east (Karimui Sheet) the Omung Metamorphics and Kubor Granodiorite are unconformably overlain by the Permo-Triassic limestone reef remnants of the Kuta Formation, which are in turn oveiiain by the late Triassic Kana Volcanics to the south-west and their metamoiphic equivalents, the Bena-Bena and Goroka Formations, to the north-east. Further west (Kutubu Sheet) the Kubor Granodiorite is directly overlain by the Kana Volcanics, whereas to the north (Wabag Sheet) the volcanics overlie silts and shales of the Yuat Formation.

Overlying the Strickland Granite and Kana Volcanics in the Kutubu, Wabag, and Blucher Range areas are Triassic (un-named) and Jurassic (Kuabgen Group) fluvial and marginal marine elastics, finer Jurassic outer-shelf sediments (Om Formation), and their metamoiphic equivalent (Om Metamorphics). Elsewhere, the Om Formation is found directly overlying the Kana Volcanics.

The Om formation is overlain by the mid Jurassic to early Palaeocene fine marine elastics and minor volcanics of the Wahgi Group. Both the Om formation and the Wahgi Group are time equivalents of the shallow marine sediments of the Kuabgen Group, and the overlying late Jurassic to late Cretaceous Feing Group (Fig 2.2).

Throughout the south and west of the area, the Mesozoic clastic sediments are overlain by the late Palaeocene shale and fine carbonates of the Mendi Group, and the late Oligocene to Miocene Darai Limestone Formation. These grade north-eastwards through the micritic limestones of the Nipa Group into the clastic trough sediments of the Aure Beds. The Darai Limestone forms the impressive and inhospitable karst terrain which is typical of much of the south-western part of the fold belt.

In the north and east, the Mesozoic elastics are overlain by the suite of trench and arc sediments known as the Salumei Synthem, which are regarded as basinward equivalents of the Mendi and Feing Groups.

Overlying the Miocene sediments in the south and west are sequences of Miocene to Pleistocene fluvial, deltaic and shallow marine sediments. In the north and east, the Miocene elastics are 36

Cl

O- Fig Fig 2.2 Diagrammatic relationship of the major stratigraphic units

T3 Not to Not to scale 37

overlain by Plio-Pleistocene to Quaternary volcanics and ashes. In the higher mountain areas Pleistocene glacial moraines are found.

The relationship between all the units mentioned above is summarised in Figure 2.2.

Evidence of volcanic activity is found throughout the stratigraphic column. The more recent events (Miocene-Pliocene) dominate, though there is evidence of early Cretaceous, early Jurassic and late Triassic volcanic activity. Quaternary events are also represented in certain areas.

The following sections describe the lithologies of the various associations of rock of the fold belt in more detail. This is not an exhaustive list, and only contains details of those lithologies which are relevant to the gravity studies.

2.2.1 Basement Lithologies

Three principal basement lithologies have been identified. The Upper Palaeozoic Omung Metamorphics comprise slate, phyllite, and sericite schist, partially recrystallised shale and siltstone, and less commonly, metagreywacke and basic metavolcanics. They are found in outcrop in the core of the Kubor anticline, where they are in places tighly folded, and are inferred beneath much of the northern and eastern fold belt areas.

The Omung Metamorphics are intruded by the Upper Permian Kubor Granodiorite, a coarse-grained biotite-homblende granodiorite and tonalite, which has been dated by K-Ar as 215-220 Ma and by Rb-Sr as 244Ma (more likely). It also outcrops in the core of the Kubor anticline, where it forms a rugged mountain topography, with relief up to 1200m, and summit heights up to 4000m.

The Strickland Granite is thought to be of Upper Permian age, and comprises pink granite with chlorite, calcite, and sericite alteration. It outcrops where the Strickland Gorge cuts through the Muller anticline in the western fold belt; in addition granite has been penetrated in a few of the petroleum exploration wells, and it has been concluded that the platform and fold belt are underlain by remnants of metasediments intruded by granitic rocks (Dow, 1977). 38

2.22 Upper Palaeozoic and Mesozoic Units

2.22.1 Permian and Triassic

Rocks of Permian and Triassic age are found in the Kubor anticline, and in parts of the Mobile Belt. The Permian?-Triassic Kuta Formation, consisting of slightly recrystallised limestone reef remnants, commonly with quartz and feldspar, especially towards the base, lies unconformably on basement in the Kubor anticline. The Middle to Upper Triassic fossiliferous black shales which form the Yuat Formation, and the Upper Triassic Kana Volcanics are also found in this area. The Kana Volcanics, which consist of 600-2500m of massive, dark green, basic to intermediate agglomerate, andesitic lavas and dykes, tuffs, pillow lavas, volcanic breccia, rhyolite, coralline limestone, black greywacke and red calcarenite, are extensive throughout the north and east of the fold belt. They are also found in the Mobile Belt to the north, and 80m of dacite penetrated in the Komewu-1 well may well represent an extension of them to the south of the fold belt (Dow, 1977).

2.222 Wahgi Group

The Lower Jurassic to Upper Cretaceous Wahgi Group includes a diverse collection of stratigraphic units which are found throughout the northern and eastern fold belt. It has been divided into five formations in the Kubor region, though the five formations are never seen in sequence. Elsewhere, not all formations can be identified. The Wahgi Group is regarded as being the partly volcanogenic basinal equivalent to the Kuabgen and Feing Groups and the Om beds. Of the five formations within the Wahgi Group, the most extensive in the fold belt region are:

(i) The Kerabi Formation is found at the base of many thrust sheets in the eastern fold belt, and comprises silty sands and muds of Neocomian to Cenomanian age which were deposited in an open marine environment. It is laterally equivalent to the Kondaku Tuff, a greenish-grey coarse lithic sandstone or greywacke, with tuffaceous sandstone, and dark grey siltstone or shale, which is found further to the north.

(ii) The Upper Cretaceous Chim Formation was deposited in a subsiding open marine basin (Kutubu Trough) and is transitional south-westwards to the Feing Group. It comprises grey, calcareous shale with minor interbeds of flaggy limestone; grey to near black micaceous mudstone and siltstone with rare clean quartz sandstone interbeds. It outcrops extensively to the [ 39

north of the eastern Muller anticline, and also forms the base of many thrust sheets in the fold belt to the southeast of Nipa and Mendi.

Among the other constituent formations of the Wahgi Group are the Upper Jurassic Maril Shale, a dark calcareous siltstone and shale deposited in an outer shelf to basin environment, and the Balimbu Formation, comprising polymictic quartz sandstone, including some altered volcanics and probably resulting from reworking of the underlying Kana Volcanics, both of which are found in the Kubor region.

2 2 2 3 Om Formation

The Om Formation and the Om Metamoiphic Formation outcrop extensively in the Lagaip River region to the north of the Muller anticline, and they are inferred beneath much of the western fold belt. The Om Formation, dated as Middle to Late Jurassic consists of black, carbonaceous siltstone and mudstone with minor quartz sandstone, and black, pyritic chert nodules. Fossils include ammonites, belemnites, bivalves, and wood fragments. Microdiorite dykes are also found. The Om Formation is the basinward equivalent of the Kuabgen Group, and is a potential hydrocarbon source rock.

Om Metamorphic Formation, of Middle to Late Jurassic age, comprises black, graphitic mica schist, sericite schist, phyllite and slate. It also includes a fine quartzite, with black cherty nodules and lenses, and dykes of meta-microdiorite. It is the metamorphic equivalent of the Om Formation.

222.4 Kuabgen Group

The Middle to Late Jurassic Kuabgen Group has its type section in the core of the Muller anticline, in the Blucher Range sheet area. It has been divided into two formations, the Koi-Iange Sandstone and the Imburu Mudstone. Elsewhere, their lateral equivalents, the Bol Arkose and the Atemin Shale have also been identified.

The Middle to Upper Jurassic Bol Arkose is a poorly sorted, indurated, conglomeratic arkose, containing sub-angular clasts of granite, pink orthoclase and adamellite, which fines upwards into a fine to coarse quartz sandstone. The granite clasts have been interpreted as suggesting that the 40

Bol Arkose unconformably overlies the Strickland Granite (Osborne, 1945), and the Bol Arkose has been inferred to be covering the Strickland Granite beneath much of the foreland, having been deposited locally within troughs in a fluviatile or estruarine environment

The Atemin Shale, of Upper Jurassic age, is a hard, micaceous, grey silty shale, sandy and calcareous in parts, with abundant bivalves and belemnites. It lies conformably on the Bol Arkose, and both formations are lateral equivalents of the Upper Jurassic Koi-Iange Formation, a fine to coarse, micaceous quartz sandstone, which is feldspathic at its base, and was deposited in a shallow marine environment. It has hydrocarbon reservoir potential towards the top, where bioturbated sandy siltstones and mudstones are common. A pebble bed and coal are found at the base in the Strickland River valley.

The Upper Jurassic Imburu Shale is a micaceous mudstone and siltstone, partly calcareous with concretions, interbedded with flaggy, micaceous, silty quartz sandstone towards base. There are abundant bivalves in the mudstones, and the Imburu Shale is regarded as a possible petroleum source rock.

2.2J2S Feing Group

The Upper Jurassic to Lower Cretaceous Feing Group, which is less indurated than the Kuabgen Group which it conformably overlies, also has its type section in the Strickland valley. Within these 1000-2000m of marginal marine sediments, several local unconformities have been identified, representing non-deposition both in the early and late Cretaceous. Two formations are recognised:

(i) The Toro Sandstone - these Late Jurassic to early Cretaceous beach and barrier-bar sands comprise well sorted, fine to medium-grained clean quartz sandstones, with siltstone and mudstone interbeds and variable (<10%) glauconite content, and are recognised as being the major potential hydrocarbon reservoir rocks of the area. In the type area, a 45m basal section of cross-bedded medium quartz sandstone is separated from an upper 220m sandstone section by a 170m recessive siltstone member. Elsewhere, the Toro Formation is estimated to be 200-400m in thickness.

(ii) The Upper Cretaceous Ieru Formation comprises fine bioturbated glauconitic quartz sandstone and siltstone interbedded with glauconitic mudstone. It yields ammonites, belemnites 41 and some volcanic (trachytic) clasts. The Ieru Formation was deposited in a fluviatile to marginal marine environment, and is relatively impermeable, providing a potential cap rock for any hydrocarbon accumulations within the Toro Formation.

In some areas, the Feing Group is not subdivided. The lithologies are similar to, but less glauconitic than, those of the Toro and Ieru Formations, but the ages are not always identical - the Undivided Feing Group has been dated as Lower Cretaceous (Valanginian) to Upper Palaeocene/Lower Eocene. It is transitional into the Ieru and Toro formations, and conformably? overlies the Om Formation.

2.23 Cenozoic Units

The Cenozoic stratigraphy of the area has become complicated and confused, with the nomenclature inconsistently used by the numerous authors. This is largely a result of the small size of area which it is possible to map in detail, and lateral facies changes. The following discussion attempts to relate the Cenozoic stratigraphy of the four 1:250000 geological map sheets covering the fold belt area.

2.23.1 Salumei Formation

The Early Cretaceous to Middle or Late Eocene Salumei Formation, or more accurately, synthem, is a collection of trench and arc sediments, over 5000m in thickness, which are thought to have been emplaced during Eocene subduction and Eocene to Oligocene collision (Davies and Hutchison, 1982). They are found in outcrop to the north of the Lagaip River in the northwestern fold belt, and are, in part, lateral equivalents of the Feing Group. In some areas they may be partially equivalent to the Chim Formation; elsewhere the Salumei Formation overlies the Wahgi Group. Five "sub-units" have been identified, though the relationship between the sub-units is unclear, and where field data is poor the Salumei Formation has been mapped as an undifferentiated mixture of submarine volcanics, volcaniclastic sediments, calcareous pelitic sediments, and red radiolarian shale. 42

2.23 2 Moogli and Urubea Formations

The Moogli Mudstone Formation of Late Palaeocene age unconfonnably overlies, and is partly a lithostratigraphic equivalent of the Chim Formation (Wahgi Group). It is a soft, grey, foraminiferal calcareous mudstone, with quartz sandstone at the base, and has been identified in the fold belt to the west of Mendi and south of Kagua.

The Late Palaeocene Urubea Sandstone Formation conformably overlies the Moogli Formation. It comprises medium grained calcareous sandstones and siltstones containing glauconite and fragments of volcanic origin which are thought to have been deposited in a shallow marine (shelfal) environment.

In many areas the Chim, Moogli, and Urubea Formations cannot be identified individually, and their chronostratigraphic equivalents have been mapped as approximately 4000m of Late Cretaceous to Late Palaeocene sediments. In the Tsak valley area (Wabag Sheet), these sediments are thought to be tightly folded and weakly metamorphosed.

2.2.33 Mendi Group

The Early Eocene to Early Oligocene Mendi Group was defined by Buchan and Robinson (1969) from exposures in the Kutubu sheet area. It outcrops on the flanks of the Kubor anticline, and in the thrust sheets of the northern parts of the fold belt, in a belt extending southeastwards from Laiagam in the north to beyond Mt Giluwe. To the south of Mt Giluwe it has been divided into five units based on discontinuous lithology and micropalaeontology:

(i) Early to Middle Eocene basal siltstone and mudstone unit, with silty micrite and siliceous limestone. (ii) Middle Eocene foraminiferal micrite and fine calcarenite. (iii) Middle Eocene fine calcareous and glauconitic quartz sandstone grading into sandy limestone. (iv) Middle Miocene foraminiferal micrite and fine calcarenite with chert nodules. (v) Late Eocene to Early Oligocene upper unit of foraminiferal micrite with chert nodules, glauconite, and iron oxides. 43

To the northwest, the Mendi Group has been divided into two formations:

(i) The Tongul Calcilutite, of Middle Eocene age, consisting of 1000-2000m of hematitic, foraminiferal argillaceous micritic limestone, calcilutite, and calcareous mudstone.

(ii) The Eocene Nebilyer Limestone, which comprises 100-20Qm of partly recrystallised micrite and fine calcarenite, outcrops on the flanks of the Kubor anticline.

2.23.4 Darai Limestone Formation

The Darai Limestone Foimation, dated as Late Oligocene to Middle Miocene, consists of massive to thick bedded limestones, varying from calcilutite to calcirudite, with algal foraminiferal biomicrite and blocks of coralgal biosparite, indicating deposition in a shallow marine environment. Varying in thickness from 1500m in the southern and western palaeoshelf areas, the Darai limestone thins northeastwards and is not present in the northeastern parts of the Papuan fold belt where its lateral equivalents, the Nipa and Aure Groups, are dominant. The Darai formation has been extensively weathered; it outcrops extensively throughout the southern and western areas of the fold belt, where it forms a karst terrain, with tall pinnacles and deep sink holes feeding extensive cavern systems, which is typical of much of the highlands and the uplifted Darai plateau to the southeast.

2.2.3.S Nipa Group

The Nipa Group is an Early Oligocene to Late Miocene sequence of limestone and shale, up to 2000m in thickness, which is the first basinward equivalent of the Darai limestone, having been deposited in an open shelf to slope environment. It outcrops throughout the fold belt and, along with the Darai Limestone, it forms many of the thrust sheets, which are responsible for the steep cliff-bounded valleys and high peaks of the region. It has been sub-divided into five formations:

(i) Darai Limestone - generally where the Darai Limestone is less than 450m in thickness, it is regarded as the lower member of the Nipa Group.

(ii) Nembi Limestone - a pale grey to brown foraminiferal micrite and fine grained calcarenite, argillaceous in places, and with occasional cherty nodules, the Nembi limestone is also a lateral 44 equivalent of the Darai Formation.

(iii) Kera Formation - another Late Oligocene to Early Miocene equivalent to the Darai formation, the Kera Formation comprises of over 1000m of fine micritic limestone and soft grey siltstone, shale and marl. In places it forms the base of the Aure Group, and is in part the basinward equivalent of the Nembi limestone.

(iv) Lai Siltstone - these Middle Miocene calcareous siltstones and mudstones, with micritic limestone, and, in the north, tuffaceous sandstone, indicating the presence of a contemporaneous volcanic source, probably to the north, are also laterally equivalent to the Aure Group.

(v) Mala Limestone - Forming a protective cap to the less resistant Lai siltstone, the Middle to Late Miocene argillaceous micrites, calcarenites and corraline limestones of the Mala Formation are thought to be lateral equivalents of the Warre limestone member of the Pynyang Formation which has been identified in the northwest of the fold belt

(vi) Orubadi Formation - The upper part of the Nipa Group consists of approx. 250m of Late Miocene calcareous mudstone and siltstone, with minor limestone. Some glauconite, carbonaceous material, wood fragments and shell fragments are also found. The Orubadi formation thickens eastwards, and in the southeast of the fold belt it is in excess of 750m thick.

2.23.6 Aure Group

Recognised throughout the north and east of the fold belt, the Middle to Late Miocene Aure Group is a thick (5km) sequence of rapidly deposited Miocene deep water clastic sediments which are the basinward equivalent of the Nipa Group. Generally three formations are recognised: a Middle Miocene sequence of andesitic lavas, agglomerate and tuff, known as the Tarau Volcanics, are overlain by the Laialam Limestone. This thin (~100m) shelly band of calcarenite, thought to be yet another lateral equivalent of the Darai Formation, and is overlain by a dark volcanic sandstone known as the Burgers Formation.

To the east, undifferentiated Aure Beds overlie the Upper Oligocene Omaura Greywacke, and are lateral equivalents of the Movi Beds, a 500-4000m sequence of volcanolithic and calcareous siltstone, shale, and polymictic conglomerate, which thickens to the southeast. 45

223.1 Pliocene Units

Orubadi Formation - The Orubadi formation is composed of Upper Miocene to Pliocene mudstone, siltstone, and sandstone. It unconformably overlies the Darai Limestone in the foreland, and in places in the southern fold belt. Often regarded as the upper formation of the Nipa Group, the Orubadi Formation thickens to over 400m in the southeast; to the west it is regarded as a lateral equivalent of the more calcareous Wai Asi Formation.

Wongop Sandstone - A sequence of 600m of Pliocene fine to coarse sandstones, tuffaceous in part, with minor conglomerate, siltstone, and shale, commonly with carbonaceous plant remains, and thin lignite seams towards the top. The Wongop formation was deposited in a shallow water to estuarine and fluviatile environment, and conformably overlies the Orubadi formation in the foreland. Along with the Orubadi formation it has now been eroded from much of the fold belt It is a lateral equivalent of the lower Era Beds.

Era Beds - Unconformably overlying the Darai Limestone, but conformable on the Orubadi formation, the fine to coarse (partly tuffaceous) sandstones, siltstones, mudstones and conglomerates of the Pliocene to Pleistocene Era Beds are around 400m in thickness, and are parity lateral equivalents of the Wongop Sandstone. They are more limited in northwards extent than the Orubadi formation.

The Liddle and Birim Formations, which consist of Pliocene clastic sediments, and are recognised only in the west of the fold belt, are thought to be equivalent to the Wongop and Era formations. Some authors map all four together as the Strickland Formation.

2.2.4 Quaternary Stratigraphy

The increase in volcanic activity from the Late Pliocene has led to the depostion around volcanic centres throughout the area of Late Pliocene to Pleistocene basaltic and andesitic lavas, with agglomerates and tuffs. Associated with these are large debris flows, chaotic scree deposits, and river gravels. In the higher mountain valleys glacial morraine deposits are also found. 46

23 The Structure of the Papuan Fold Belt

The Central Orogen of Papua New Guinea has formed as a result of the transpressional interaction of the Australian and Pacific Plates. The Papuan fold belt, which can be divided from the rest of the Central Orogenic Belt on the basis of the dominant structure, forms the southern part of the orogenic belt, and comprises disturbed Australian continental margin material. The northern part of the Orogen is dominated by the left-lateral wrench fault zone known as the New Guinea Mobile Belt, comprising mostly terxanes suspect to the Australian craton. Its southern boundary is defined by the Bismarck-Lagaip fault zone. To the south of this fault, fold and thrust structures are dominant. Thrust-faulted anticlines, trending NW-SE, can be traced along strike for 60-100km. Their associated thrust faults dip northeastwards and, as a group, can be traced along strike for around 200km (Durkee et al., 1986). At the southeastern end, the fold belt trends north-south, and broad synclines separate the faulted anticlines whose thrust faults dip both east and west This area is known as the Aure Tectonic Belt, and is often regarded as being separate from the Papuan fold belt, though no distinct boundary has been identified.

The Papuan fold belt has been subdivided on the basis of the dominant structural features into six zones (Fig 2.3), which broadly reflect the changes from northeast to southwest across the fold belt (Jenkins, 1974).

2.3.1 Zone 1 - The Muller-Kutubu Zone

The Muller-Kutubu zone is the southernmost belt, closest to the foreland. It varies from up to 50km across in the vicinity of the Muller and Blucher ranges, to around 20km across in the Lake Kutubu region. It is characterised by large overthrust folds, with northeastwards dipping fault planes, the largest of which is the Muller Anticline, which has a wavelength in excess of 50km and an amplitude of approx. 2.5km, and extends for some 150km along strike from the Bol River at the western end of the zone to the Lavani Valley. The Muller Anticline is flanked by several smaller structures such as the Juha and Cecilia anticlines, which have amplitudes of around 1km, and wavelengths of 10km or so. Similar structures are also found further east, including the Hides, Mananda, Iagifu and Hedinia anticlines, all of which have exhibited hydrocarbon potential to some degree. 47

141 142 143 U4 145

Mobile Belt Ok2 apm in

a T a ia g a m

,Lavanl-1

K andep Mount Hagan Dom a P eak s Cecilia-1 Mt Sisi Mendi Lake\ Nipa Kutubu P orom a iglfu-2* Mt Bosavi e Mt t1 arimui

O M t Murray

Kanau-1

Darai Plateau Foreland Basin

141 142

Fig 2.3 Major structural subdivisions of the Papuan Fold Belt. 48

2 3 2 Zone 2 - The Oksapmin-Tari-Mt Murray Zone

To the north of the Muller Range there is an abrupt change in structural style to a complex zone of Miocene limestone thrust sheets and overthrust folds. The western end of this zone| is approx. 15km wide, and comprises both thrust sheets and fold strutures, such as the Oksapmin Bowl, where the Darai Limestone is repeated four times in a synclinal feature. However, to the east of the volcanic centres of Mt Kerewa and the Doma Peaks, fold structures are rare, and a stack of flat limestone thrust sheets is observed, continuing from the Tari Gap to the Poru Plateau.

2.3.3 Zone 3 - The Porgera-Nipa-Poroma-Karimui Zone

This zone, around 40km wide, extends to the north of the Doma Peaks from Porgera in the west to Nipa and the Nembi River in the centre of the fold belt, and Karimui in the east. It is characterised by sinuous complex folds, occasionally overturned and overthrust, typically 5km in wavelength and <500m in amplitude, with the anticlinal features being tighter. In the west these fold structures can rarely be traced for more than 15km along strike, though further east the folds, such as the Andebare and Wage anticlines, become much more laterally persistent.

2.3.4 Zone 4 - The Laiagam-Giluwe-Kubor Zone

This extensive zone is characterised by the broad folding of the trough sediments of the Papuan basin. Lying beyond the limit of the Miocene shelf limestones which determine and dominate the structure to the southwest, this zone includes the large anticlinal feature known as the Kubor Uplift, a reactivated basement feature which separated the Southern Highlands basin from the northern outer trough of the Papuan Basin during the Mesozoic.

2.3.5 Zone 5 - The Darai Plateau

The Darai Plateau is formed by a single large asymmetric anticlinal structure, over 90km long and 40km wide (Brown and Robinson, 1982), which exposes the Darai Limestone at the surface. The structure rises steeply in the southwest above a northeast-dipping thrust fault, but the northeastern limb dips gently down to the Hegigio River. The surface of the plateau is extensively karstified, and tensional fractures are well developed. 49

2.3.6 Zone 6 - The Aure Tectonic Belt

The Papuan fold belt is separated from the metamorphic belt of the Owen-Stanley Range by an area over 50km across which is characterised structurally by laterally persistant north-south trending fold structures, known as the Aure Tectonic Belt. The anticlinal structures, which have steeper, often overturned, western limbs, and are often cut by east-dipping thrust planes, can be traced along strike for 60-80km. Typical fold wavelengths are 3-5km, and amplitudes are around 500m. Two phases of folding have been identified; the first phase in the Early Miocene also affected the Owen-Stanley Metamorphic Belt but has been strongly overprinted by the second late Middle Miocene phase which produced the fold structures (Dow et al., 1974).

2.3.7 Relationships between Stratigraphy and Structure

The changes in structural style which occur across the fold belt can be directly related to stratigraphic changes at many levels within the sedimentary sequences. For example, where the Miocene Darai Limestone overlies a relatively thin sequence (~2km) of Mesozoic sediments, the large scale faulted anticlines (Zone 1 in Fig 2.3) are the dominant structure.

As the Mesozoic sedimentary sequence thickens to the northeast, and the lithologies change from those of the Kuabgen and Feing Groups to those of the Om Formation and Wahgi Group, so the structural style changes to imbricate stacks of thrust sheets.

Further to the northeast, the change in structural style from thrust sheets to small scale thrust- faulted anticlines, corresponds with another Cretaceous facies change, resulting in the extensive occurrence of the Chim Formation, which thickens rapidly northeastwards. The final change in style (from zone 3 to zone 4) corresponds to the Miocene facies change from the shelf and slope sediments of the Darai Limestone and Nipa Group in the southwest to the trough sediments of the Aure Group in the northeast.

Interpretations illustrated by Brown and Robinson (1982) show the decollement zone beneath the southwestern structural zone (Zone 1 in Fig 2.3) at the base of the Feing Group, in sediments of Jurassic age. To the northeast the decollment zone beneath zones 2 to 4 (Fig 2.3) is shown to be within the Cretaceous Wahgi Group sequences. This difference in detachment level will also have a significant effect of the style of deformation. 50

2.4 Geological History

The history and evolution of the Papuan fold belt can be divided into three phases of sedimentation separated by the two major tectonic events. The first phase, representing the sedimentary history up until the first tectonic episode at the end of the Mesozoic, is identified as the Papuan Basin phase. This is followed by a Foreland Basin phase, which represents the post tectonic development of a basin in which initial thick carbonate sedimentation is inevitably superceded by terrigenous inundation (Pigram et al.,1989), a state which persists to the south of the deformation front today. The final phase, that of the Central Orogen, is characterised by the deformation and molasse type sedimentation associated with the Neogene formation of the fold belt, and is currently restricted to the fold belt area.

2.4.1 The Papuan Basin

Throughout the Mesozoic the subsiding and rifting Australian continental margin was the site of littoral, deltaic, and shelf/slope sedimentation. The main phase of sedimentation commenced in the mid Jurassic, and continued throughout the Cretaceous. Sediment was mainly derived from the Australian craton to the south, though there were occasional uplifted local sources, such as the Lake Murray basement high and the Pasca Ridge. There is also evidence of Upper Triassic volcanism, possibly related to rifting episodes.

Uplift of both the Muller Range and Kubor Range occurred during the Cretaceous; the Cretaceous succession over the Muller Range is much thinned, and the Kubor Range separates the thick Cretaceous clastic sequences of the Southern Highlands Basin to the south from the greywackes, pillow lavas and tuffs to the north.

The whole region was uplifted in the Early Palaeocene as a result of spreading in the Coral Sea, and much of the Upper Cretaceous succession was eroded, especially in the south, where the base-Tertiary sub-crop now comprises sediments of Albian age (Jenkins, 1974). Sedimentation commenced again in the north and east during the Early to Mid Eocene, with the deposition of typical continental margin sediments (Mendi Group), but deposition in these areas ceased during the Early Oligocene. Metamorphism of the sediments of the outer trough occurred in the Mid Oligocene, at which time it is thought the ultramafic rocks of the Mobile Belt (April Ultramafics and Marum Ophiolite) were emplaced. 51

2.42 The Foreland Basin

The emplacement of the ultramafic rocks in the Mid Oligocene, and the subsequent thickening of the continental margin, resulted in a downwaiping of the continental crust, and formation of a narrow foredeep. A peripheral forebulge also developed, thus providing an ideal environment for carbonate deposition on the platform to the south (Darai Fm and Nipa Group). In the deep, narrow basin to the north, which was starved of detritus, pelagic sedimentation and gravity flows were predominant (Aure Group). On the northern flank of the trough, locally derived detritus allowed a limited amount of clastic sedimentation. Carbonate sedimentation continued on the relatively stable shelf to the south throughout the Lower and Middle Miocene, but the locus of deposition was gradually pushed further south. By the Late Miocene, carbonate sedimentation had finally been terminated by the influx of detrital material from the volcanic Maramuni Arc and the onset of collision with the island arc units now forming the New Britain Province (Chapter 1).

2.4.3 The Central Orogen

Though the initial continent-arc convergence and ophiolite emplacement took place in the Mid Oligocene, deformation and uplift of the Papuan fold belt, and the Mobile Belt to the north, did not commence until the Late Miocene. This followed a shortlived, but violent episode of arc volcanism in the Middle Miocene, and resulted in the rapid and continuing uplift of the fold belt, and concomitant erosion, characterised by frequent landslips and molasse-type sedimentation in terrestrial basins (Milsom, 1985). Quaternary processes have been dominated by the eruption of large quantities of potassic lavas from vents throughout the orogenic belt. 52 3. Gravity Data

Much of the gravity data used in this study was provided by BP Australia in the form of a database containing over 22,000 readings from the whole of Papua New Guinea. In addition, a further 19 stations were acquired by Chevron during 1989, and some 300 gravity stations were established during fieldwork in PNG during November and December 1989. These three data sources are discussed in more detail below. Ultimately around 8400 of these gravity readings were used, the study being restricted to the fold belt area bounded by eastings 500000 and 810000, and northings 9100000 and 9430000. The location of these surveys is shown in Figure 3.1.

3.1 The Papua New Guinea Gravity Database

The Papua New Guinea Gravity Database was produced in 1984 by Digimap Pty. under contract to Robertson Research in conjunction with Flower, Doery, Buchan Pty., and included most of the gravity data collected in PNG before 1982. This data was purchased by BP, who have since added their own data, collected in 1986 and 1988.

This data was passed to UCL in the form of a magnetic tape containing the gravity data, and an accompanying detailed report describing the history and location of all the surveys (Thomas and Lawry, 1988). The database was divided on a survey by survey basis, and the information on those surveys used in this study is shown in Table 3.1. Each survey was assigned a code by Robertson Research (1984), using a four digit numerical code in the form year-3* for stand-alone gravity surveys, and year-4* for seismic and gravity surveys (Thomas and Lawry, 1988). For example, the 1974 Lavani Valley seismic and gravity survey has code 74-41. However, the survey codes used in the database are in the BMR format, which is confusingly similar to the Robertson’s format. For clarity, both codes are shown in Table 3.1.

3.2 The Lavani - Juha Regional Profile

In addition to the gravity data contained in the database, 19 stations were acquired by Chevron in 1989 along a line from Lavani in the northeast, through Baia and Juha, towards the foreland 53

9425000

9400000

9375000 Ok Ted!

1 av am Juha Regional

9350000

C e c ilia 9325000

9300000

K.nm Stncklar.d

92^5000 Kutubu Orokana

River-Lake Mura* 9225000

920000C

9150000

9125000

9100000 500000 600000550000 650000 700000 750000 800000

Fig 3.1 Gravity survey location map. 54

oo m

T> O.

CL CL

00 00

On oo o n c o 3 £ ON C/3

ON co co co oo cs CN CS

CS (S m VO cs

rs

CS CS CS roi ro co Ov O n CS CS CO VO

CO 00 CO CO © CS VO 55

•-» 00 tl& <=*^

VO On ro 00 cs »o cs

>> Table Table 3.1 which surveys were Gravity the database from used in this study.

CO

cs ro CS ro CO co■ 00 0000 0000

cs Ti­ cs ro CO ro co CO CO »n 00 00 00 0000 56 to the west of Cecilia-1. This data was supplied to UCL through BP Australia in spring 1990, in Ed con’s data format. This was converted into the database format, and incorporated into the database, and provided a useful link between the foreland data of the Cecilia and Nomad surveys and the highland data of the Lavani Valley and Komo surveys.

3.3 Tiengo Gravity Survey

3.3.1 Introduction

The Tiengo Gravity Survey was carried out during November and December 1989, in the Southern Highlands of Papua New Guinea (Figure 3.2). Approximately 300 gravity stations were established along the road network in PPL 27 and PPL 86, from Komo in the southwest to Mendi in the east, and towards Laiagam in the north, with a station separation of approximately 2km. The aim of the survey was to provide semi-regional gravity coverage of PPL86, connecting into the existing detailed profiles in PPL27 (Komo and Nembi surveys), and filling a significant gap in the existing regional data coverage.

3.3.2 Field Survey

The Southern Highlands of Papua New Guinea rise to in excess of 4000m above sea level. The high peaks are separated by deep, often cliff-bounded valleys, some of which are only 1200m above sea level. The region is covered by thick primary jungle, though some of the valleys have been cleared for cultivation, and some of the high mountain passes are sparsely vegetated because of the low temperatures. The annual temperature range is relatively small, though the diurnal variation is typically greater than 20°C. Typical temperatures at Tiengo were 1-2°C at 6am, and 22-26°C at 2pm.

Papua New Guinea as a whole is sparsely populated, and the Southern Highlands are no exception. Small villages are scattered throughout the region, many in the most inaccessable places, often requiring several days walk, or a helicopter flight to reach. Along the recently built Highlands Highway there are many small settlements, but the only places which could be classified as towns by European standards are Mendi (the provincial capital) and Tari. 9700000 6 57

Fig 3.2 Tiengo survey location map. 58

The local people mostly belong to the Huli tribe from the Tari basin, and they were generally found to be friendly, though care was always needed. In the Kandep area, part of the Enga province, more problems were encountered. One survey had to be abandoned because of a tribal battle, and on another occasion money and equipment! were stolen. However, these were only isolated incidents, and on most occasions a full day’s work, generally from 7:30am to around 4:30pm, was possible.

3.3.2.1 Equipment

A LaCoste and Romberg gravity meter (G-513), on loan from the Open University, was used for the survey. This instrument proved very reliable, and stood-up well to the rigours of transportation along the rough roads. Elevation control was obtained using two Baromec microbarometers (#937 and 1102) and a System Paulin Precision Altimeter (A-2). A wet/diy thermometer set was also carried to determine temperature and humidity.

33 2 2 Accommodation

Hotel accommodation in Port Moresby and Melbourne was provided by BP Australia. Accommodation in the field was provided by Exploration PNG in their camps at Tiengo, Leru and Kopalu. Four nights were spent at Leru while the Tari to Komo section was surveyed, and two nights were spent at Kopalu following the completion of the survey; the remainder of the time was spent at Tiengo.

3.3.23 Logistics

Transport was provided by Exploration PNG of Mount Hagen, and by Pacific Helicopters of Goroka. A Mazda B2600 Utility 4x4 vehicle was used throughout the survey, and a driver was provided by Exploration PNG. The base station ties from Lem to Komo and Hides-1 were carried out using a McDonnell Douglas MD-500ER helicopter (P2-PAH), operated by Pacific Helicopters. A similar machine (P2-PAU) was used for transportation between Tiengo and 59

Kandep High School, where the vehicle had to be left for five days because of the poor road conditions between Kandep and Tiengo.

3,32.4 Road Quality

With the exception of the Highlands Highway, the quality of the roads used was not generally known prior to the survey. In addition, many seemingly reasonable roads can become impassable during or immediately after heavy rain. With this in mind, a road quality map has been produced (Figure 3.3), and the roads used during the survey have been classified into three types:

(i) Good all-weather roads, such as the Highlands Highway, where reasonable progress can be made at all times (shown in red in Figure 3.3).

(ii) Poor all-weather roads, where progress can be made at all times, though it may often be very slow (<15km/h), (shown in green in Figure 3.3)

(iii) Dry weather only roads, along which progress is slow in dry weather conditions, and difficult or impossible in wet conditions | (shown in blue in Figure 3.3).

3.3.2.S Base Stations

New gravity base stations were established at Leru and Tiengo, and a subsidiary base station was established at Kandep High School. The Leru base station was tied to the existing stations at Komo, Hides-1, and Tari, and the Tiengo base station was tied to Tari and Nipa. The Kandep base was tied solely to the newly established Tiengo base, and all survey loops from Kandep also visited the Tiengo base station. Descriptions of these three new base stations are contained in Appendix A. In addition, the base stations at Mendi and Tari (St John, 1967), Komo (Solo, 1982), and Nipa, Hides and Karamera (Edcon, 1988a) were also visited during the course of the survey. The IGSN71 gravity value was used in each case, and where this was not published, it was calculated from the 1930 value using the equation given by Thomas and Lawry (1988):

gobs=978197.946+1.0005118(gobs-978212.67) 60

CO OJ o CM o CD CM CM

o.

05. .05

.CO

.CO

CO.

05, .05 Fig 3.3 survey Tiengo road quality map. CD CO CM CD ■»T O CD CM r-» CM co CO * 61

A mistie error of approx. 2.5mGal between the base stations to the east of Tiengo and those to the west was observed, a similar error to that previously noted by Edcon (1988a). To maintain consistency with existing surveys, stations west of Tari were tied to the Komo and Hides values, and those to the east of Tari were tied to the Nipa/Mendi values. This has the unfortunate effect that the Tari station has two values, one tied to the Leru/Komo/Hides stations, and the other tied to Nipa/Tiengo/Mendi. Ideally all the incorrect data would have been altered to fit one or the other of the Tari values; however, since it was not known which of the existing surveys were tied to which base stations, and whether these were the only base stations affected by this problem, very little could be done about the situation. As the data is presented in the form of contour maps, which are produced by smoothly contouring a grid of data which has been inteipolated from several of the original station values, the mismatch is smoothed over. In any case, in the regional context, a mismatch of 2.5mGal in a region where the gravity gradient is ~2mGal/km is of little significance.

3.3.2.6 Station Positions

The locations of all the gravity stations are shown in Figure 3.4. The position of each station was determined from the 1:50000 topographic maps with an estimated accuracy of +/-50m in most cases. However, in certain cases, where the position of the road had been changed since publication of the maps, the new route of the road had to be estimated from the topography, and the station location could only be determined to within an estimated +/-200m.

332.1 Elevation Control

Elevation measurement was carried out using aneroid barometers, and a precision altimeter. A base barometer was read at 15 minute intervals throughout the duration of each day’s survey by a local employee, and a second barometer and the altimeter were read at each station. The temperature and humidity at each station were also measured, and were taken into consideration when reducing the elevation values; the full method is described in Section 3.3.3.2. Tiengo Gravity Survey - Station Locations

Fig 3.4 Tiengo survey station location map. 63

33 2& Terrain Corrections

For each station, an estimate was made of the difference between the station elevation and that of the terrain within a 50m radius. A standard graticule was used to convert these estimates to terrain correction values, which are contained in Appendix B. The remaining terrain corrections were made as outlined in Section 3.3.3.3.

3 3 2 .9 Field Operations

The first survey loop was carried out from Tiengo to just west of the Tari Gap, and a number of unexpected problems were encountered. Firstly, the gravity meter galvanometer could not be steadied for more than a couple of seconds at many stations in the loop, and secondly, station location proved extremely difficult for some of the stations, including the Tiengo base station. Finally, the base barometer readings, which had been taken by a member of the camp staff (from the local village), showed a surprising number of fluctuations.

The problem with the gravity meter only repeated itself for a few hours about two weeks later, and was attributed to presumed seismic activty. The station location problem was found to be caused by the fact that the new highway did not follow the route of the old road which was marked on the topographic maps - in the Tiengo area, the new route was up to 2km to the south of the old route. Similar problems were faced on other occasions, and where they occurred the station location, and the new route of the road, had to be estimated from the topography. As a result of these two problems, the results of the first day’s survey were not used in the final analysis.

The third problem, that of the irregular base barometer profile, was attributed to operator error. The Baromec aneroid barometers are not the easiest instruments to read, and it is very easy to make mistakes. Since the barometer reader was, by European standards, not well educated, these mistakes generally went unnoticed, especially in the first week or so of the survey. The problem of an unlikely base barometer profile occurred on other occasions, but was totally unpredictable; many of the diurnal profiles were perfectly good, and there seemed to be no pattern to the operator’s errors. The solution of this problem is detailed in Section 3.3.3.2. 64

Having encountered these problems at an early stage, the remainder of the survey went well. Initially only the areas close to the base camps at Tiengo and Leru were covered, allowing morning and afternoon survey loops to be carried out. However, as the survey area necessarily became more and more distant from the base camps, only one loop per day could be carried out For example, it took over 2V2 hours to drive from Tiengo to Mendi, which meant that when the surveys to the north of Mendi were being carried out, over 5 hours per day was spent travelling to and from the survey area, leaving only 4 hours in which to complete the survey in order to arrive back at camp before dusk; travelling by road at night was regarded as too dangerous.

Towards the end of the survey period it became necessary to leave the vehicle in the Kandep High School compound, and to fly by helicopter from Tiengo each morning. This meant that full day surveys were undertaken, the duration of which was restricted by the availability of a helicopter to make the return journey, and by the weather conditions, which could often prevent helicopter operations. During these surveys, a VHF radio for communication with the helicopters, and an HF radio for emergency communication with the base camp, were carried.

The entire Tiengo survey was carried out within the region of high mountains in the Papuan fold belt. The roads used ran both along the valleys, at elevations of 1200-1800m, and over the mountain passes, with elevations of 2800-3000m. In most cases, a large swathe of ground had been levelled during road construction, which meant that station sites could be chosen where local terrain corrections were small (<0.1mGal). In all cases, they were less than 0.5mGal.

The Tiengo gravity survey was completed on 22nd December, 1989, when the loop from Tiengo to the Tari Gap was repeated, producing much more sensible results. The Tiengo base camp was dismantled on 23rd December.

33 3 Data Reduction

3.3.3.1 Reduction to Observed Gravity

Meter readings were converted into mGal using the conversion table provided with the meter (Appendix A), and corrections were made for tidal gravity variations and instrument drift Meter repeatability was generally better than 0.3 mGal over a woridng day. Drift was perhaps slightly 65 higher than would normally be expected, but quite good considering the rough treatment the meter was subjected to when travelling along the poor roads. Observed gravities were calculated from the differences in measured gravity between each station and one of the base stations mentioned above.

33 3 2 Reduction of Elevation Data

Elevation readings were taken using the barometers and the altimeter, and two different methods of reducing these readings were applied. The altimeter readings, which were in feet, were treated in a similar way to the gravity readings, in that a uniform drift was assumed, using the initial and final readings of each survey to determine the drift. Though this may at first seem a rather inaccurate method given that the daily pressure variation at the base station was measured using the barometer and was definitely not linear, the actual daily variations were generally only 2- 3mBar, and at distances of around 20-50km from the base stations at which most of the surveying was carried out, the linear variation cannot in reality be signifcantly less accurate than using the base station readings.

The procedure for reducing the barometer readings was more complicated, and also took into consideration the air temperature and humidity variations. Firstly, the pressure difference AP between the gravity station and the base station was calculated, and if this difference was not zero when both barometers were read at the same time and place, a correction was made. The base station diurnal curves are shown in Figures 3.5 and 3.6. On most occasions a smooth profile was produced, reaching a peak at around 7am, and a low around 3pm, and any obvious glitches could be smoothed over. On a few occasions, most notably the 2nd and 9th December, the base station pressure readings were unreliable. In these cases, the errors were assumed to be operator errors, and the true readings were estimated from the first and last readings of the survey (which were made by the author), and the typical diurnal curve. 775-1 776-1

770- 770

765- 765

760 760

30-11-69 Tiengo 30-11-60 Tlango

775

770- 770'

765 785

760' 760

775-i 775-1

770- 770

765- 765

760 760

09-12-60 Tiengo

775 776

770- 770

765- 765

760 760

14-12-69 Tiengo

775-| 775-1

770- 770-

785- 765-

760 760

16-12-69 Tiengo

775-i 775-i

770- 770

765- 786-

760 780

20-12-69 Kandep 22-12-69 Tiengo

Fig 3.5 Tiengo and Kandep daily pressure readings. 67 635-1 636-1

830- 830-

825- 625-

620 620'

04-12-60 Uru

835-1 635-1

830- 630-

625-

620' 620'

635-1 635-j

830- 630-

625- 825-

1' i I i I i I i 820' 06-12-69 Nipa 10-12-60 Nipa

835-1

B30- 630-

825- 625-

820' 620 11-12-89 Mendi

635-1

830- 630-

825-

820' 620'

13-12-69 Mendi 15-12-60 Nandi

Fig 3.6 Leru, Nipa and Mendi daily pressure readings. DRY BULB The elevation difference Ah was calculated using the formula the using calculated Ahwas difference elevation The Figure 3.7 (Crone, 1948), and the final fully corrected elevations were stored. were elevations corrected fully final the and 1948), (Crone, 3.7 Figure of elevation the from calculated then was station each of elevation corrected temperature The h bs sain. h hmdt creto ws hn are ot y eeec t te rp in graph the to reference by out carried then was correction humidity The stations. base the (K) temperature air 0O dry where the is iue . uiiyCreto oAeod egt Coe 1948). (Crone, Heights Aneroid to Correction Humidity 3.7 Figure g is the acceleration due to gravity (ms*2) gravity to due acceleration the is g r is the gas constant (278.03 J K'1J kg*1) (278.03 constant gas the is r m'1)K (0.0065 gradient temperature the is s $ DIFFERENCE DRY-WET BULB DRY-WET DIFFERENCE AT I OE THOUSAND ONE IN PARTS o l 4 6 s 0 22 20 7s 16 14 Tl To ORCIN FACTOR CORRECTION wt n dy edns ae en ae at taken for been oorncaon have find reodlngs dry stations and wet f I the bulb Increase wet level the seo with above read that considerably Is working reading If bulb dry The r bl rom I dges arnet used level. Fahrenheit sea degrees In the reodmg bulb dry dry and et w f o difference for curve to horizontally ne gah ih r bl temo tr edn; ov m reading; eter om therm bulb dry with graph Enter o o Ices te ifrne wt n dy bulbs dry and wet f o difference the Increosc not Do per correction obtain to vertically descend and bulbs huad eet e fe thousand rp b te ubr tosns fe above feet f o thousands f o number the by graph P„ s UIIY CORRECTION HUMIDITY 18 O NRI HEIGHTS ANEROID TO

20

22

eoch eoch 24 24 6 28 26

n ue h mean. the use and 26

28 3 to enter enter to both both

30 0

68 69

Thus for each station, two elevation values were obtained. Both elevation values were considered, and for each survey loop, the method which gave the more reasonable results was chosen. Criteria used in this selection included internal consistency, elevation of any other base stations in the loop, and comparison with the topographic maps. Given the potentially greater accuracy of the elevation values determined from the microbarometer readings,| these values were used whenever possible. For most of the surveys, the two elevation results were within 5-10m of one another. In a few cases, only one value was available, usually as a result of instrument or operational difficulties.

The accuracy of the levelling using pressure sensitive devices such as barometers and altimeters is unclear, and is very much dependent on the atmospheric conditions. St John (1967) regarded ± 10m as a reasonable level of accuracy, with a maximum error of ±25m. Milsom (1970) claims a maximum error of ±20m, and a standard deviation between the true and measured values of ±5m. This equates to an accuracy of ±lmGal in the Bouguer anomaly values.

3 3 3 3 Reduction to Bouguer Anomaly

Station latitudes were calculated from the map co-ordinates using a UTM projection onto the IUGG 1967 spheroid, and the theoretical gravity at each station was calculated using the 1967 gravity formula. Free air corrections were calculated, and the simple Bouguer correction was determined using a density of 2500kgm*3.

The extended Bouguer anomaly was calculated using the method outlined in chapter 4, using densities of 2670kgm'3 and 2500kgm’3. This proved to be a time consuming procedure for this survey, since such an extensive area was covered. Not only did it require the generation of 69 new topographic grids, but also the extension of the existing detailed DTM to cover much of the Wapenamanda and Mendi sheet areas. 70

33 .4 Data Presentation

The extended Bouguer anomaly data from the Tiengo survey is plotted in Figure 3.8. The raw data are contained in Appendix C. The Tiengo survey data were combined with extended Bouguer anomaly data from other surveys in the region (Komo (8641 and 8132), Nembi (8831), Hegigio (7531) and Kutubu-Orokana (6932)), to produce the extended Bouguer anomaly map of the Southern Highlands region (Figure 3.9). Despite the mistie noted at the Tari base station, the results of the Tiengo survey were found to be compatible with the existing surveys. Though the data shown in Figures 3.8 and 3.9 has been smoothed, contouring of the raw data revealed very few mismatches which were not within the range of error of the elevation data. iienqo Survey - Extended Hnuquer Anomaly (?SOO) 0 Qi O'o' 120 O' 71 1*8 I'B ix 73 CO l-H •£3 ,8 3 £ •a « cn OO § b £ O C o o cd Q S O 2 C/3 o & a O > §> d> O 33 O (D 00 03 C? O G Southern Hiqhlands - I'xtonrlpcl Dntiquor Anomaly 0OEi 100 C 125.C 72 i I W) 1tu 73 CO a a a x> T3 •o 3 T CO a 1 s CO 5 as* cn a o S e CS ci 0 o o s 1 >> o c c o z o 1 0 o 3> <5 (A 0 s a (L> x e o 73 4. Gravity Data Reduction

4.1 Introduction

The process known as gravity reduction attempts to convert the gravity measured at a point on the Earth’s surface to that which would have been measured at the point on a reference surface, usually the geoid, vertically (normal to the geoid) above or below the true point of measurement Removal of the theoretical gravitational attraction at the reference surface (latitude correction), upwardly continued to the point of observation (free-air correction), leaves solely the effect of topographic masses above the reference surface, and any density contrasts below the reference surface. The effect of topographic masses can be removed (topographic correction), and other corrections for known sub-surface density distributions, such as the Isostatic Correction, may also be made.

The resulting anomaly value thus represents the effects of density variations below the reference surface as observed at the true point of measurement

To ensure data compatibility, all the gravity data described in Chapter 3 were processed in the same manner from the observed gravity values. In some cases, the data included terrain corrections and mistie information. However, the methods of calculation of these terrain corrections were not known, and therefore the terrain correction information was not used. The mistie information was used where necessary.

4.2 Production of a Digital Terrain Model

In order to compute the terrain corrections, it was first necessary to convert all the elevation data into digital form. This was achieved by digitising the topographic maps of the region, using the procedure outlined below. 74

4.2.1 Digitising

The digitising was carried out in the Geography Department at UCL on a Summagraphics AO size digitising table linked to a VAX 11/750 computer system, using a relatively simple software package called "digit". This initially requires the user to define four points on the map in terms of their x and y UTM co-ordinates. The x and y co-ordinates of the cursor position could then be output to a specified file when the appropriate mouse button was pressed. The elevation of the point was then input manually using a keypad on the mouse. Points of known height digitised included points lying on contour lines, spot heights, and triangulation pillar elevations. An output file containing x, y, and z values could be produced relatively easily. Generally, approximately 600-700 points per hour could be digitised, for periods of up to 3 hours duratioa A typical 1:50000 sheet would be represented by approximately 8000 data points, about 12 hours work, which would typically be spread over one week.

4.22 Data Quality

After each digitising session, the output file was checked for obvious errors - generally keying errors, such as double clicking a button. However, to ensure good data quality it was necessary to produce check-plots of each map which could be easily compared with the original map to detect any errors.

Unfortunately the VAX computer was not sufficiently powerful to enable such large data sets to be gridded and contoured at a reasonable speed, so the raw data were transported on magnetic tape to the University of London Computer Centre and loaded onto their Amdahl 5890 high speed computer. The data were gridded at an interval of 200m using the RANGRD subroutine of the GINO-F library. This uses a least squares fitting paraboloid to inteipolate the value at a grid point, and considers a specified number (>6) of data points around each grid point The result of this stage was to produce an array of elevation values, which could then be contoured.

Because of restricted hard output facilities at ULCC, the gridded data were then transferred over a datalink to the University College computer system, a small network of GEC4190 mini­ computers. Here the gridded data were contoured, and 1:50000 scale contour maps were 75 produced. These were then compared with the original maps, and the positions of any discrepancies were noted.

It was then necessary to return to the original x,y,z data to attempt to fmd the source of the errors, which proved a very time consuming procedure. Once the errors had been identified and corrected, the gridding and contouring procedure was repeated. Any remaining errors were identified and corrected, and another check plot produced. This procedure was repeated until all the errors had been eliminated. Generally this could be done having produced 3-4 plots, though some required no correction, and others up to 8 attempts.

The ultimate result of this was a topographic database covering much of the Southern Highland area which could be gridded as required for the calculation of terrain corrections. The area covered by the digital terrain model (DTM) is illustrated in Figure 4.1. In order to carry out the terrain corrections, it was also necessary to digitise some of the surrounding sheets, albeit at a reduced density, and these are also shown in Figure 4.1.

4.23 Data Presentation

As each topographic map was digitised, a 1:50000 contour map was produced of the digital data (Fig 4.2). These plots were excellent for tracing errors in the data, but did not produce a particularly clear impression of the true terrain. More recently, with the acquisition of a Sun workstation and the Uniras graphics package by the department, colour images have been produced of all the areas. These take the form of coloured contour maps, colour-shaded maps, and coloured isometric views (Plates 4.3 to 4.6) and clearly illustrate the nature of the terrain in the region.

4.3 Reduction and Terrain Correction of Gravity Data - Theory

43.1 Initial data format

The gravity data were supplied in the form of a database containing over 22,000 land and marine gravity readings covering the whole of Papua New Guinea. This database was compiled by CO 03 0 * -i. CO 0 III - 0 . - - . ,1 i _ ^;} f, **••$*, &***, ■;■ >^ri *.• *"• . * ?->$■ *-'* •'•■ '■ O) 0 t'•»'f< ‘ >*• jfc: ayjfn'i.:r- * • r^ v >«*■*%:• *i. !y' v Q ’ v*-. \ •„. -. --r Fig Fig 4.1 Area covered by the digital terrain model

Lj;nos S88j68q Fig 4.2 1:50000 topographic map of a 5x5km areas of the Doma Peaks (top) and the corresponding computer-contoured check plot produced from the DTM. 78

DTM High density data

A80VE 3300.0 3000.0 3300.0 2700.0 3000.0 2* 00.0 2700.0 2100.0 2 * 00.0 1000.0 2100.0 1900.0 1000.0 1200.0 1300.0 000.0 1200.0

Plate 4.1 Isometric view from the southeast of the high data density region of the DTM. Area covered is 165x165km.

Karoma sheet — viewed from SSW

*000 0 3000.0 *000 0 3000.0 3600.0 3*00 0 3600.0 3200.0 3*00.0 3000.0 3200.0 2000 0 3000.0 2000.0 2000 0 2*00 0 2600.0 : : c c c 2 * 00.0 2000.0 2200 0 1000 0 2000.0 1000 0 1000.0 1* 0 0 .0 < 6000 1200.0 1 * 00.0 1000.0 1200.0 000.0 1000.0 600.0 600.0 * 00.0 600.0

Plate 4.2 Isometric view from the south-southwest of the area covered by the Karoma 1:100000 topographic map. Area covered is 55x55km. Plate 4.3 Isometric view from the south-southeast of the area covered by the Koroba 1:100000 topographic map. This region includes the Karius Range (bottom centre), the Lavani Valley (top left), the Tari Basin (bottom right), and the western flanks of the Doma Peaks (top right). Area covered is 55x55km.

Plate 4.4 Close-up isometric view of the Lavani Valley. View from the southeast. Area covered is 25x30km. 80

Digimap Pty as a part of the 1984 study by Robertson’s Research in conjunction with Flower, Doery, Buchan Pty entitled "Petroleum potential of the Papuan Basin" (Thomas and Lawry, 1988), and included all the gravity data collected before 1982. Two gravity surveys conducted by BP in PPL27 in 1986 and 1988 had been added to this database, and two more surveys (Tiengo and Lavani-Juha) were added during the course of the study.

All the data had been reduced to their observed gravity using the IGSN71 datum, and those earlier surveys which could not be tied to this datum had been omitted. The establishment of the new international gravity datum in 1971 led to a revision in the absolute gravity at Port Moresby of -14.7mgal, and those gravity stations tied to the old datum were corrected by Robertson Research using the formula

got*=978197.946+1.0005118(gobs-978212.67) (4.1)

This correction had been applied to the Digimap data, and hence was also applied to the two recent surveys not part of the original database (Thomas and Lawry, 1988).

The database also included free-air and Bouguer anomaly data, and in some cases terrain corrections and mistie corrections. However, to ensure similar reduction of all the data, all reduction was carried out from the observed gravity values.

4.32 Latitude Correction

The latitude correction, which is the theoretical gravity measured on the reference surface at a point vertically below (or above) the point of observation, was calculated using the alternative representation of the 1967 Gravity Formula proposed by Mittermayer (1969):

gtal=978031.85(l+0.005278895sinJ(iaf)+0-000023462sin4i(to)) (4.2) where lat is the geographic latitude of the point of observation. 81

43 3 Free-Air Correction

The free-air correction compensates for the decrease in gravity with increased distance from the centre of the Earth. For surveys covering a relatively restricted area, a figure of 0.3086mgal/m is commonly used; however, for surveys covering extensive areas of severe terrain, the free-air correction of Vyskocil (1960)

! /ac=0.308772((l-0.001442sin2(kl))/i-0.000000072/!2) (4.3) is often used, where lat is the geographic station latitude, and h the station elevation. This more accurate correction had been applied to the Digimap data, and was therefore also applied in the re-calculation of the free-air correction during this study. However, the true increase in accuracy using this method to data from New Guinea is insignificant; for a station at 8 degrees south, with an elevation of 1000m, the free-air correction would be 308.8mgal using the accurate method, compared with 308.6mgal using the approximation.

4.3.4 Topographic Correction

The purpose of the topographic correction is to correct for topographic masses above the reference surface. Traditionally this has been done by correcting for an infinite horizontal slab, with a thickness equal to the station elevation, whose gravitational effect is given by

g=2npGh (4.4) where p is the slab density, G the universal gravitational constant, and h the station elevation.

This is the simple Bouguer correction, which is subtracted from the free-air anomaly to give the simple Bouguer anomaly. In areas of subdued topography, the simple correction is often sufficient However, in mountainous areas, the topography can differ substantially from that approximated by the slab, and terrain corrections are made for deviations of the true topography from the slab. Calculating these terrain corrections is a time consuming process in which the height difference between the slab and the true topography is estimated (or calculated) for a predefined area, and its effect at the gravity station is determined. Manual methods use a 82 graticule of compartmentalised circular rings to zone an area, and tables have been produced which show the effect of a particular height difference in a particular zone (Hammer, 1939).

More recently methods of performing the correction using digital computers have been developed. These involve dividing the terrain into regular-shaped bodies, the attraction of which is relatively simple to calculate. Methods using cones, square prisms, and flat plates have all been documented (eg Kane, 1962; Takin and Talwani, 1966; Zhou et al., 1990), as have methods involving the approximation of the topographic surface with a function f(x,y) which can be integrated to determine the topographic volume (eg Krohn, 1976; Granser, 1987).

In addition to the simple terrain correction, other corrections may be needed. These include a correction for the curvature of the Earth’s surface, and for density variations in the topography, both of which can be accommodated when using computer processing. Several workers have suggested methods of replacing the two-stage approach with a single stage topographic correction calculated by summing the effects of all topographic masses above the reference surface. All the suggested methods involve approximating the topography using a large number of small bodies with a shape whose gravitational attraction is easy to calculate. This method also allows for the easy correction of curvature, and density variations.

The method used in the study is similar to that used by St John and Green (1967), (St John, 1967) in that it uses square prisms to approximate the true topography. A regular grid of elevation values was produced from randomly digitised data as described in Section 4.2, the grid spacing varying with the prism-station distance and severity of the terrain in order to speed up the correction process. The method of calculating the gravitational attraction of each prism was also varied with the prism-station distance; for distant topography, the thin rod approximation was used (Eqn 4.6), but for prisms close to the gravity station, the full equation for the gravitational attraction of a right rectangular prism (Nagy, 1966) was used (Eqn 4.7). The errors involved in using the thin rod approximation have been documented by Bott (1959) and providing the station-prism distance is at least 5 times the prism spacing, the error involved will be less than 2%.

The line mass approximation (St John and Green, 1967), which is used to calculate the vertical gravitational attraction of a prism, given by 83

g ^ C p x A - - - } (45) rt rk where x and y are the prism dimensions, and r, and rb the distances from the station to the top and bottom of the prism, does not take into consideration the fact that the "vertical" at the prism may not be the same as the "vertical" at the gravity station, due to the Earth’s curvature. St John claims that multiplication by a factor cosa where a is the angle between the two "verticals" accounts for this, but it is not the case, and the more detailed equation

~ 1 rr-zshia r-dsina, {A g=Gpxy [ ] (4.6) rcosfl r t r b

should be used (see Fig 4.3 for explanation of variables.) For a=0, this is the same as the line-mass approximation, but the line-mass approximation can be in error by around 2-3% when the more distant prisms are considered, which equates to 5mgal in some cases.

The distances r, and rb can simply be calculated by converting the positions of the station and prism from geographic latitude and longitude (lat,long,h) into a cartesian (x,y,z) system centred on the Earth’s centre, using the relationships (Morgan, 1987):

x=(v+h)cos(lat)cos(long), y=(v+h)cos(lat)sin(long)t z=((vZ72/a2)+/i)sin(/ar) where -e2sin2(Za/))1/2, a= semi-major ellipsoid axis, semi-minor ellipsoid axis.

Since all the post-1965 surveying in Papua New Guinea used projection onto the IUGG’67 ellipsoid (Bomford, 1967), the appropriate values of 0=6378160. and 1#=298.25 were used, wheref-(a-b)/a

For prisms close to the gravity station for which the thin rod approximations was not sufficiently accurate, the full equation for the vertical attraction of a prism (Nagy, 1966) < r >

Fig 4.3 The thin rod approximation.

p S

o

Fig 4.4 Calculation of the angle subtended at the centre of the Earth. 85

H Vi

was used, where x, y, and z define the prism position relative to the gravity station. When expanded, this equation has 24 terms, which can be reduced to 12 if one of the z limits is taken to be zero, as is the case in the terrain correction program.

It was also necessary to determine which topographic masses act upwards on the gravity station, and which act downwards. All masses which lie above an infinite horizontal (tangential) surface through the gravity station will exert an upwards force at the station, whilst those below the surface will act downwards. It was therefore necessary to calculate the distance d (Fig 4.4) from the reference surface to the tangential surface at the gravity station, for all prisms. If the prism height is greater than d, then a proportion of the prism, Pt-d, will act upwards at the point of observation, and the remainder, d, will act downwards.

Distance d can simply be calculated from triangle OSP (Fig 4.4), where d=(OS/costf)-OP0

The angle a can be determined since the position vectors OS and OP0 are known, and using the scalar product relationship, cos#=(OP0.OS)/( | OP01 |OS|)

Having calculated the vertical effects of both parts of the prism, the net attraction was simply taken as the sum of the two parts. Summing the gravitational effects of all the prisms in a given area will result in the total effect of the area, and providing the area is sufficiently large, this will produce the necessary topographic correction for each gravity station.

4.4 Reduction of Gravity Data - Method

4.4.1 Introduction

This section describes the method of reduction of the gravity data to the extended and final Bouguer anomaly stages, and is based on the theory described in the previous section. Generally, the gravity data were reduced on a survey by survey basis, the exact method depending on the location of the survey and the severity of the surrounding terrain. Most survey data were reduced 86 to the extended Bouguer anomaly, but where misties were noted, or had been noted by Robertson Research, a mistie correction was added to the extended Bouguer anomaly, producing the final Bouguer anomaly.

4.42 Reduction to Simple Bouguer Anomaly

A simple computer program was written which read in the gravity data from the database and calculated the free-air and Bouguer anomalies for each station. Firstly the latitude correction was calculated by substituting the station latitude into Equation 4.2. Then the free-air correction was determined by substituting the station latitude and elevation into Equation 4.3. The free-air anomaly was the calculated, using the relationship

§f»» Sob*_§ltt"^§f»c

The Bouguer correction was then calculated, using a uniform crustal density of 2670 kgm'3 for the Bouguer plate. The Bouguer anomaly

8b«—Sft«”8bc was then calculated. This was generally within 0.05mGal of the original database Bouguer anomaly value.

4.43 Topographic Correction

A computer program was also developed to calculate the topographic correction. During the development of the program, both the one-stage and the two-stage methods of determining the effects of topographic masses were used. Both programs were written on the theory detailed in the Section 4.3.4, and it was hoped that the two-stage approach would provide a validity check on the method, since the final results should be approximately the same, though reached from different directions (one from the free-air anomaly, and the other from the Bouguer anomaly). 87

Both programs were written in FORTRAN77, and a commented listing of the one-stage correction program is contained in Appendix D. The programs were run on the Amdahl 5890 mainframe computer at the University of London Computer Centre, which runs the IBM MVS operating system. Compilation by the Fujitsu F77 compiler using maximum optimisation took around 1 second. Using the 1986 Komo (8641) survey data as an example, and topographic blocks containing up to 40000 prisms, run times were of the order of 15 minutes cpu time.

The logic of the one-stage (topographic correction) program is shown in Figure 4.5. Once the projection and spheroid constants had been defined, the gridded topographic data were read in. The data format was as produced by the gridding routine (Appendix D), the first line containing the array dimensions, the second the x and y limits of the data, and the remainder of the file the gridded data, stored in the standard FORTRAN order, starting with 2X1,1), Z(2,l) ... Z(NUMXNUMY). The positions of each prism are then converted from UTM co-ordinates to geographic latitude and longitude in subroutine LATLON. This subroutine is a modified version of a program supplied by the Photogrammetry and Surveying Department at UCL, and of particular help in making the modifications were Morgan (1987) and Snyder (1982).

Following this conversion, data for the first station were read from the gravity datafile and the limits of the area requiring the most topographic detail (which is restricted to the immediate vicinity of the gravity station) were calculated. Then the topographic correction was calculated by considering each topographic prism in turn. The program tested to establish the prism size, and if the prism was from the fine grid, and lay within the fine grid area, the topographic effect was calculated, using either the thin rod approximation (subroutine TCFAR), or the full equation of Nagy (1966) (subroutine NAGY), depending on the station-prism distance. For prisms which were not part of the fine grid and lay outside the fine grid area, the thin rod approximation was used in all cases.

For the calculation of the topographic effect using the thin rod approximation, the next step was to convert the prism and station positions into spherocentric rectangular co-ordinates (subroutine RECT), and calculate the height at the prism position of the tangential surface at the gravity station. The vertical gravitational attraction of the prism at the gravity station could then be calculated using the thin rod approximation. ( START ) 88

SET UP CONSTANTS -DENSITY -PROJECTION

READ TOPOGRAPHIC DATA

CONVERT STA LAT/l.ONG NEXT TO EARTII-CENT. PRISM .RECT. CO ORDS.

ALL PRISMS?

READ GRAVITY STATION DATA

CALCULATE 1.1MITS OF FINE GRID AREA

CONVERT PRISM CO ORDINATES I TO LAT/l.ONG

c a l c u l a t e : NEXT SURFACE DIST. PRISM PRISM STA.(DD)

A

Page 1 of 2

Fig 4.5 Flowchart illustrating the logic of the one stage correction program. 89 a a A

V IS DD

/PRISM \ / PRISM \

/rRlSM \ CALCULATE

CALCULATE CALCULATE CALCULATE TOPO EFFECT TOPO. EFFECT UP EFFECT BY THIN ROD BY 11 UN ROD BY NAGY

CALCULATE DOWN EFFECT BY NAGY

ADD PRISM EFFECT TO TOTAL

CALCULATE EBA SO FAR

WRITE STATION DATA

LAST STA7

Fig 4.5 Page 2 of 2 90

For prisms close to the gravity station, for which the thin rod approximation was not valid, the equation for the attraction of a rectangular prism was used. This equation is contained in subroutine EQN9. Subroutine NAGY was used to set up the variables for substitution into the full equation, and was called twice: firstly with z equal to the station elevation, and secondly with z equal to the difference between the station and prism elevations. The vertical gravitational attraction of the whole prism as observed at the gravity station is simply the difference between these results.

The attraction of the prism was then summed with the attractions of the previous prisms, and once all the prisms had been summed, the result for the gravity station was output. The whole procedure was then repeated for the next gravity station, and so on until the topographic correction due to the particular block of topography had been calculated for all stations. The program then had to be re-run with different topographic data until the whole of the desired area had been covered. For the 1986 Komo survey, 27 topographic blocks were required to cover an area up to 90km from each gravity station.

The program to calculate the terrain correction (ie the two-stage method) was simply a modified version of this program which used the difference between the station and prism elevations rather than purely the prism elevation when calculating the terrain effects.

The results of the two programs are shown as profiles along the survey lines shown in Fig 4.6. Figure 4.7 shows the elevation of each station, and the simple Bouguer, one-stage extended Bouguer, and two-stage extended Bouguer anomalies. For most of the stations, particularly those with lower elevations, the one and two stage profiles are within 2-3mgal of each other, which is as expected. Figure 4.8 shows the difference between the two results for survey lines 1 and 2, plotted as a profiles along the lines. For the stations with higher elevations, the difference between the two profiles increases significantly. This is a result of the increased divergence of the Bouguer plate and the true topography at these altitudes, where the terrain correction is at its largest, and scatter plots of the difference between the one and two stage anomalies plotted against station elevation, terrain correction and topographic correction (Figs 4.9a,b,c) support this suggestion. They reveal that the best direct correlation of the difference between the one and two stage anomalies is with the terrain correction value, and the worst is with the one-stage correction, indicating that the terrain correction is the more likely source of error. Tests were also carried out using flat topography, and similar differences were observed. Though these 91

Line Line 5

Line 2 Line 3

Line 4

Line 6

9315000 685000 735000

Fig 4.6 Location of th e Komo (8641) survey lines. 92

(mgai) -50 | Line 1 Line 2 Extended Bouguer Anomaly (One-Stage)

Extended Bouguer Anomaly •' (Two-Stage) |\/ I: / v A , Simple Bouguer Anomaly r \ A

r C %

‘ S ' ,

%

-160

3000

Station Elevation (m) r'

1000 30 60 Distance along profile (km)

Fig 4.7 Profiles along Komo survey lines 1 and 2 showing both the one and two stage correction results using a standard crustal density o f 2670kgm 3. 93

20

Difference between Line 1 Line 2 one-stage and two stage J extended Bouguer anomalies '

(mgal)

A,

30 60 Distance along profile (km)

Fig 4.8 Difference between the results of the one and two stage corrections.

Difference between one-stage and two-stage extended Bouguer anomalies

(mgal)

1000 3000 Siation elevation (m)

Fig 4.9a Plot of the difference between the one and two stage correction results against station elevation. Difference between one-stage and two-stage extended Bouguer anomalies

(mgal)

0 Terrain Correction (mgal) 30 Fig 4.9b Plot of the difference between the one and two stage correction results against the terrain correction.

Difference between one-stage and two-stage extended Bouguer anomalies

(mgal)

0 270 120 Topographic Correction (mgal)

Fig 4.9c Plot of the difference between the one and two stage correction results against the total terrain correction. 95

results cannot be regarded as conclusive, it was felt that the results from the two methods were sufficiently similar to warrant continuing only with the one-stage approach.

A further development of the one-stage approach was to allow for vertical and lateral density variations in the topography. This was achieved by digitising the geological map of the area, and assigning typical density values to each formation. For formations which were thought to vary in density with depth, such as the Darai Limestone, and volcanic rocks which form large cones on a flat lower surface, a depth was also defined at which the density was thought to change. The results of this approach are shown as profiles along the Komo survey lines in Figure 4.10. It can be seen that although the results are shifted from those of the uniform density topographic correction, there are no new anomalies, and similar results could be achieved by multiplication of the uniform density data by a single factor (Fig 4.11). A much better correlation could be achieved by allowing the assumed density of the Bouguer plate to increase with increasing elevation, as more dense rock will make up the lower parts of the plate where the plate is thicker. The results of this approach are shown in Figure 4.12. However, it was felt that the effects of these variations were of little regional significance, and the use of a topographic correction using a single uniform density would be sufficient.

4.4.4 Application of the Topographic Correction

4.4.4.1 Komo 8641

This survey was used in the development of the topographic correction program, and in the determination of the effects of topographic grid spacings and density variations. It was felt that a 50m grid spacing was the minimum which could reasonably be generated from the digitised data and processed efficiently. Experimentation showed that the same result would be achieved with a 200m grid spacing for prisms beyond 3km from the gravity station as with a 50m grid. Similarly, beyond 15km a 1000m grid interval could be used, and beyond 25km a 5000m grid would be sufficiently accurate.

To correct all the stations in this survey out to 90km, 27 topographic grids were required. Exlcn > > c: E

Fig 4.10 Profiles along Komo survey lines 1 and 2 using variable density and uniform density of 2670kgm \ 96 97

-40 Extended Bouguer Anomaly Variable Density L ine 1 Line 2 ■1 Fixed Density (2500kgin 3)

(mgal) VJl V I / 'N> !/ \ Simple Bouguer Anomaly ) vV W V \

V.

-160 30 60 Distance along profile (km )

Fig 4.11 Profiles along Komo survey lines 1 and 2 using Variable density and uniform density of 2500kgm3.

-40 Extended Bouguer Anomaly

Variable Density i H t. Depth factor |V ^ (mgal)

Simple Bouguer Anomaly

-160 30 60 Distance along profile (km)

Fig 4.12 Profiles along Komo survey lines 1 and 2 using variable density and constant density multiplied by a depth factor such that EBA=EBAx(l-(2.9xlCT5ZSTA)). 98

4.4.4.2 Komo 8132

This survey covers an area to the south of the 1986 Komo survey, and overlaps with both the 1986 survey and the Hegigio 7531 survey. Consequently, those stations in the Hegigio survey which were close to the 1981 survey stations were corrected along with the Komo data. The method used to calculate the topographic correction was the same as used for the 1986 Komo survey. Only 4 new topographic grids needed to be generated, as the existing grids from the 1986 survey correction could be used.

4.4.43 Hegigio 7531

Those stations in the two northwestern lines of the Hegigio survey which had not been corrected with the Komo (8132) survey were corrected using the same method as for the Komo 8641 survey. To correct these data to 90km required the generation of 21 new topographic grids - 10 at 50m prism spacing, 4 at 200m spacing, 4 at 1000m spacing, and 3 at 5000m spacing. The southern 5000m grid was not generated because the topographic data was not available at the time; however, because of the flat, low lying nature of the terrain in this region, it is unlikely that this omission made any significant difference to the results, since the one-stage correction uses the prism height, not the difference in height between the station and the prism, to determine the topographic correction.

4.4.4.4 Lavani 7441

All the Lavani survey data were corrected together, requiring the generation of 12 new topographic grids. Three of the 5000m grids generated for the Komo 8641 survey were also used in correcting this survey. Because of a lack of topographic information north of northing 9350, a few of the stations could not be corrected out to the lull 90km. However, the contribution of the outer areas to the total topographic effect is very small (<3%), so the end result was not thought to be significantly in error.

4.4.43 Kaim-Strickland 7047 and Tomu River 7231

These two surveys, along with a few stations from the Western Papua 3830 survey, were corrected together. Because of the low and gentle nature of the terrain in this area, a slightly 99

different method of calculating the topographic correction was used. The topographic correction program was modified so that the finest grid required had a prism spacing of 1000m. This meant that the full equation for the attraction of a rectangular prism had to be used for prisms within 6km of the gravity station, rather than 300m as with the 50m grids. Only three topographic grids were generated for these surveys: one 1000m grid covering the area containing all the gravity stations, and two 5000m grids. No topographic data was available to the south and west of the survey area, but the area is low (<50m) and flat, so the lack of data was not thought to be significant. The -8mGal mistie correction propoed by Robertson Research (1984) needed to be added to those stations which were part of the Western Papua survey.

4.4.4.6 Cecilia 7043 and Nomad 6841

The Cecilia and Nomad surveys were corrected along with stations from the Western Papua survey which were in the same area. The same method was employed here as with the Kaim- Strickland and Tomu River surveys. Five new topographic grids were generated, one with 1000m prism spacing, covering the area around the gravity stations, and four 5000m grids. Some of the 5000m grids covered areas for which topographic data were not available, and in these cases approximate data were fabricated.

Serious problems were encountered with the Cecilia survey. The station elevations differed from the topographic elevations by up to 300m, resulting in the stations apparently being at the bottom of deep holes when the topographic correction was calculated. The problems were compounded by the fact that there was a 100m "virtual cliff' running east to west through the topographic grid caused by a mismatch of two topographic maps.

The cause of the station-topography difference was eventually traced to the conversion of the elevation readings from feet into metres. The original survey was carried out in feet (CGG, 1970), and it appears that for the elevations in the gravity database, the conversion from feet into metres had been carried out twice, resulting in the station heights being reduced by a factor of approximately 3.25. Little could be done about the mismatch of the topographic maps, which appeared to be an unrelated problem, though the effects of the "cliff' were minimised by assigning the station elevation to all prisms close to gravity stations affected by the mismatch. 100

The most significant effect of the change in station elevation was to remove the need for the 15mgal mistie correction for the Cecilia survey used by Robertson Research (1984). However, the -8mGal mistie correction was required for the stations in the Western Papua 3830 survey. The steep gravity gradients which were present in the original data were also eliminated by the elevation changes.

4.4.4.7 Palmer 7230

The topographic correction method applied to the Palmer survey was similar to that applied to the Kaim-Strickland survey. Three topographic grids were generated: one with a 1000m prism spacing covering the survey area, and two with a 5000m spacing, to the north and east. The effects of the flat, gentle terrain to the south and west were considered insignificant in comparison with the mountains to the north, since the one-stage correction uses only the prism heights, which are very small, to determine the topographic correction. The northernmost stations of this survey could only be corrected to around 65km (minimum) from the station because of the restricted DTM coverage, though the contribution of the outer 25km would only be approximately lmGal because of the large station-prism distances.

4.4.4.8 Western Papua 3830, Aworra-Ok Tedi 7442, Upper Fly River-Lake Murray 7341, Komewu-Darai 5741, and Bamu 7334

These surveys were all carried out in areas of flat and gentle terrain, even less rugged than the Kaim-Strickland area, where terrain corrections were less than lmgal. Therefore no topographic corrections were calculated for these surveys, and only the simple Bouguer anomalies, including any misties (-8mGal for both Western Papua and Aworra-Ok Tedi) were calculated.

4.4.4.9 Kutubu-Orokana 6932, Nembi 8831, and Waro 8131

These surveys, which covered adjoining areas of rugged terrain, required the generation of 44 topographic grids: 32 50m grids, and 4 each of 200m, 1000m and 5000m grids. 4.4.4.10 Kanau 7037

The Kanau survey was corrected along with those stations of the Hegigio (7531) survey which lay adjacent to it. Only four topographic grids were used (1 at 1000m spacing, and 3 at 5000m), because the survey was carried out in an area of gentle terrain, allowing the simpler terrain correction method to be applied. No terrain data were available south of northing 9170000, but the area is very low-lying, and would have little effect on these observations.

4.4.4.11 Lake Kutubu 6738

Although this survey was carried out in an area of rugged and relatively high terrain, the correction method used in areas of gentle terrain had to be used. This was because the differences between the station elevations and the topographic prism heights were unacceptably large. The cause of this was almost certainally the nature of the terrain over this, the Darai Plateau, where karstification is extensive, and elevation changes can be very large over short distances. This has two effects:

(i) The topographic maps often do not depict the terrain very accurately.

(ii) The gridding process may be adversely affected by the apparent randomness of the elevation values; in any case, the gridded data will be considerably smoothed.

4.4.4.12 Lavani-Juha Regional

Though there were only 19 stations in this survey, they covered a large area of rugged terrain, and consequently required the generation of 12 50m grids, 4 200m grids and 4 1000m grids, and 3 5000m grids, a total of 23 new topographic grids. The detailed method of calculating the terrain correction was used.

4.4.4.13 Tiengo

This survey was by far the most time consuming to terrain correct since it covered such a large area of the Highlands. A total of 69 grids of topographic data were required: 49 with a 50m node spacing, 12 with a 200m spacing, and 4 each with 1km and 5km spacings. The northern 102 stations of this survey could not be corrected out to the lull 90km as the DTM did not extend sufficiently far to the north; all stations were corrected to at least 70km.

4.4.5 Overview

Once the terrain corrections had been carried out, the data from the individual surveys were combined to form a database covering much of the Papuan fold belt region and its foreland basin. This database included over 8400 individual gravity readings, and was used as the basis of all the interpretation and further processing which is detailed in the following chapters. 103 5. Qualitative Interpretation and Further Processing

5.1 Extended Bouguer Anomaly Maps.

The extended Bouguer anomaly data produced as a result of this study is presented in the form of a 1:500000 map (enclosed). Annotated versions are shown in Figures 5.1 and 5.2. The maps were produced by interpolating a regular 1km grid of values from the original data set For each grid node, a combination of linear, quadratic and distance weighted methods was used to calculate the value at the node from all the data points within a 20km radius. The resulting grids were then computer-contoured, using a routine which produced smooth contours (at the expense of a certain amount of accuracy), and the contour plots were modified by hand where it was felt the computer contouring algorithm had produced an unlikely contour pattern.

5.1.1 General Subdivision

The extended Bouguer anomaly map can be usefully divided into six areas (Fig 5.1). The southern area (Area 6) shows the more gentle gravity gradients associated with the foreland, and although the dominant trend is the expected WNW-ESE trend running roughly parallel to the fold belt deformation front and the axis of the foreland basin, there is also clear evidence of a secondary, NE-SW trend.

This secondary trend is not at all obvious in areas 2 and 3, where steep gravity gradients strongly dominate the anomaly pattern. To the south of Area 3 there is a line of gravity lows (Area 4), which also follow the ESE-WNW trend. This line lies along the northern edge of the Southern Plains province (identified in Chapter 1).

The general ESE-WNW trend of the Bouguer anomalies in all areas is parallel to the regional structural grain of western PNG, and roughly perpendicular to the direction of convergence of the Australian and Pacific plates. In the northern and eastern parts of the map, the steeper gravity gradients of areas 2, 3 and 5 are characteristic of the Papuan fold belt, where there is both structural thickening of the sedimentary section and depression of the Moho under the mountain belt. 104

M20000 o

1 6 - B3BOOOO -2 0

8370000

8340000

3300000-

8 2 8 0 0 0 0

8 2 6 0 0 0 0

3240000 3230000- 8220000 OS

9170000 9160000

9130000

9110000 9100000- 575000525000 600000 625000500000 700000575000 725000550000 750000 775000 0 0 0 0 0 0

Fig 5.1 Extended Bouguer Anomaly Map showing general subdivisions. CI=5mGal

Subdivisions based on character of the gravity field: 1 NE fold belt, gentle gravity gradients. 2 SW fold belt, steep gravity gradients. 3 Muller Anticline, steep gravity gradients. 4 Gravity lows to the south of die Muller Anticline. 5 Local gravity high over the Darai Plateau. 6 Gentle gravity gradients of the Foreland/Platform. 105

#410000 G # 4 0 0 0 0 0

8 3 8 0 0 0 0 —20 Muller Anticlines 8370000

8 3 6 0 0 0 0

8340000 Hides ** Cecilia

#300000 Mananda

8200000' Hedinia v Mt. Bosavi

8240000

8 2 3 0 0 0 0 -

8 2 2 0 0 0 0 Lake Murray

8200000 9180000 3180000

9160000 9160000 #140000 9130000

9100000- 5 2 5 0 0 0 575000 625000 675000500000 700000 725000750000 775000

Fig 5.2 Extended Bouguer anomaly map showing the location of the major geological features. 106

Area 1

This area contains the lowest gravity values in the fold belt region, and in the extreme NE the gravity gradient is reversed, with gravity increasing to the northeast. However, this area as a whole is characterised by gentle gradients and the lack of distinct features in the anomaly pattern.

Area 2

Area 2 forms a band across the NE comer of the map. It is bounded to the northeast by the -130 mGal contour, and broadens from being 50km across in the north to over 90km across to the northeast of Mt Bosavi (Fig 5.2). The gradients in this zone are of the order of 2mGal/km, which is more than double that in the neighbouring areas (1,4 and 6). The trend of the contours in this zone changes from NW-SE at the eastern end to almost N-S at the northwestern end. This N-S trend must have important implications for the structure of this region but unfortunately there is a significant data gap between the northern end of this area and the Muller Anticline (Area 3) where the contours trend E-W.

Area 3

The wavelength and amplitude of the major positive anomaly in Area 3 shows that the structure in this area is large and must involve the continental basement and possibly deeper levels of the crust. This is the Western Muller Anticline (Hill, 1989a).

Because of the lack of data in the eastern part of Area 3, the gravity field of the Eastern Muller Anticline is barely represented. However, at the extreme eastern end of Area 3 the -35mGal contour does suggest the plunging nose of this feature.

Area 4

To the south of the Muller Anticline (Area 3) there is a major gravity low. This low is restricted to the western part of the map, and apparently does not extend further to the east (ie to the south of Area 2). 107

Area 5

Area 5 lies to the south of Area 2 at the eastern edge of the map, and is characterised by a NW- plunging gravity high. This area corresponds to the Darai Plateau, a major uplifted area which is bounded to the south by the Darai Fault. Seismological evidence suggests that this fault extends deep into the continental basement (Abers and MacCaffrey, 1988).

Area 6

This area is characterised by gentle gravity gradients, with the contours generally following the NW-SE trend which dominates the fold belt area to the northeast (Area 2). The gentle northeastwards decrease in gravity can be attributed to the gentle deflection of the lithosphere in response to the loading to the north, resulting in a deepening of the foreland basin and an increasing thickness of recent sediments.

In the SW comer of Area 6 the gravity contours show a NE-SW trend which is particularly interesting as it is not well represented in either the geological or topographic features of the fold belt region. A similar trend is also shown by a small "nose" in the contours which extends to the northeast from Area 6 into Area 4 to the southwest of the Cecilia anticline (Fig 5.2). It is perhaps of note that if this "nose" in the anomaly field is continued northeastwards, it crosses the Muller Anticline area to the east of the West Muller Anticline, in an area where there is a major change in the direction of the gravity contours.

5.1.2 Qualitative Interpretation

The low gravity values of Area 1 are likely to be a result of both a low density root beneath the highlands, and the greater thickness of sediments thought to lie in this region (cf Jenkins, 1974). The gentle gradients suggest that there are either no structures involving the strata where the major density contrasts occur, or that any such structures are deeply buried.

The boundary between Areas 1 and 2 has been identified as corresponding to many significant and interrelated changes in the geology, these being (Harrison, 1990): 108

- a change in facies of the Mesozoic and Early Cenozoic sediments from predominantly shelf sediments to the southwest to predominantly slope and trough sediments to the northeast.

- a presumed change on basement lithology from the Strickland Granite to the southwest to the Kubor Granodiorite to the northeast (Brown and Robinson, 1982).

- a change in structural style from large thrust sheets and broad, faulted anticlines to the southwest to thrusts and tight, short wavelength anticlines and synclines in the northeast

Though these changes cannot directly account for the gravity anomaly, they are likely to result from fundamental changes in the geology, such as a rapid thickening of the Mesozoic sedimentary succession (cf Jenkins, 1974; Fig 2), which will affect the gravity field.

However, no such simple correlation can be found for the southwestern boundary of Area 2: there are no sediment facies or basement lithology changes, and the boundary does not convincingly correspond with the southwestern margin of the fold belt. It is, however, likely that the SW margin of the fold belt does have a significant effect on the change in character of the gravity field, but it is modified by "local" features:

- the -50mGal "flat" in the centre of Area 2 is a feature which in the context of the general NE gravity gradient represents a local gravity high. This major topographic and geological feature, which almost certainly involves basement, is the Mananda Anticline. There is also a large gap in the data coverage around the Mt. Bosavi volcanic centre, another feature which must have a significant effect on the gravity field.

- to the southeast the Darai Plateau, another significant topographic feature involving basement, also gives rise to a local gravity high.

Broadly speaking, Area 2 corresponds to that part of the fold belt comprising shelf sediments deformed into large thrust sheet and broad thrust-faulted anticlines, features such as the Hides, Mananda and Hedinia Anticlines.

Areas 3 and 4 both show the effects of the western Muller Anticline. The high gravity values of Area 3 are the result of the involvement of crystalline basement and possibly deeper levels 109 of the crust in the wetem Muller Anticline, and the low gravity anomaly values of Area 4 are interpreted as representing the effects of a deepening of the foreland basin, possibly combined with a downwaiping of the lithosphere under the loading of the fold belt to the north.

Area 6 is the only area to show the SW-NE trend in the gravity field. It is likely that this trend reflects features of the region which pre-date the formation of the fold belt. They could well represent areas of thick and thin sediments, such as palaeo-valleys, or extensional features such as horsts and grabens, dating from the Late Mesozoic to Early Tertiary, which may well have been reactivated in the later compressional phase of deformation associated with fold belt formation. Another possible explanation of this trend is that it represents the antithetic set of conjugate strike-slip faults which are associated with many large scale wrench systems and their en-echelon fold belts (Wilcox et al., 1973).

The northern part of Area 6 represents a part of the foreland which is affected by the increase in crustal loading of the fold belt to the north, the northwards decrease in gravity being due to an increase in sediment thickness and a slight downwards deflection of the lithosphere. The isolated gravity high in the southwestern part of this area has been interpreted as representing a basement high, known as the Lake Murray basement high, which was a major feature at least until the Late Jurassic (Durkee et al., 1986).

5.2 Enhancement of Terrain Corrected Gravity Data.

In order to enhance specific features in the terrain corrected data, and to aid its interpretation, a number of digital filters were applied to the gridded data set. A suite of public domain digital filtering programs installed on a Sun 3/60C workstation was used. These programs provided a means of using two-dimensional Fourier analysis to perform a number of filtering operations on a gridded input data set.

5.2.1 The Fourier Transform.

The concept of Fourier analysis is that any waveform, be it periodic or transient, can be represented as the sum of a number of simple sine or cosine waves of differing amplitude and 110 phase, with frequencies or wave numbers that are integral multiples (harmonics) of a basic repetition frequency, the fundamental. Any wave can thus be expressed in terms of the period, waveform and amplitude (time domain) (Fig 5.3), or frequency, amplitude and phase, (frequency domain) (Fig 5.4) of its constituent sine/cosine waves.

The Fourier transform is the mathematical means of changing the representation of a waveform from the time domain, g(r), to the frequency domain and back again. In the frequency domain the wave can be represented mathematically in terms of the amplitude and phase spectra, A if) and (/), or in terms of the complex function Gif) where

Gif) = A (/)e*w

The terms g(t) and G(f) are known as a Fourier pair, and are the transforms of one another.

The reason for the use of the Fourier transform in the application of digital filters is that to apply a digital filter in the time or space domain requires the convolution of the input data with the filter response, a process which even with today’s computers is very slow. However, by transforming both the input data and the filter response into the frequency domain, multiplying them together, and taking the inverse transform, the same result can be achieved in a reasonable time. Using the Cooley-Tukey (1965) method for calculating the fast Fourier transform (FFT), this took about 1 hour for most of the filters applied to the PNG data.

5.2.2 The Digital Filtering Program (DigFil)

The digital filtering program uses two-dimensional Fourier analysis to perform the following operations:

- calculation of the first vertical derivative - calculation of the second vertical derivative - upward continuation of the gravity field - strike/trend analysis - wavelength filtering Ill

(a) (b)

Fig 5.3 Complex waveforms resulting from the summation of two sine wave components of frequencies / and 2/ (from Kearey and Brooks, 1984).

amplitude amplitude

_ 2 - a) 2 (b:

± 1 2f frequency f 21 frequency phase phase rr/ 2 _ n/2 _

f 2f frequency 1 2f frequency

-tt / 2 ■TT/2

Fig 5.4 Frequency domain representation of the waveforms shown in fig 5.3 (from Kearey and Brooks, 1984). 112

The program works by first transforming the space (x,y) domain input data into the wave number (k) domain by means of the fast Fourier transform. The Fourier coefficients are then multiplied by the wave number response of the appropriate filter, and the resulting coefficients are the transformed back into the space domain (Kelvin, 1989). The program also allowed access to the Fourier coefficients, which could then be used to produce the amplitude and power spectra of the data.

The following discussion illustrates the problems which can be overcome by the use of the DigFil program.

5.23 Source depth constraints

A commonly applied method of potential field data interpretation is to separate the "regional" scale anomalies from the "local" anomalies. Quite what "regional" and "local" are depends largely on the data in question, but the method basically involves the separation of the long and short wavelength anomalies, or those which have deep and shallow sources. Gupta and Ramani (1980) identified three methods used to derive the regional and residual values: spectral factorisation, upward continuation, and graphical smoothing. Because of the large amount of data involved in this study, the graphical smoothing method, which involves plotting profiles and estimating the regional field by eye, was not used. However, both the upward continuation and spectral factorisation methods were used.

5.2.3.1 Upward Continuation

The upward continued field is calculated by applying the upward continuation filter H(f)=e2%M/, whereh is the continuation distance, to the Fourier coefficients (Dean, 1958). When transformed back into the space domain, the upward continued result shows the gravity field which would be observed at a distance h above the original surface of observation, and consequently attenuates anomalies due to shallow or narrow sources more than anomalies due to deep or broad sources (Blakely and Jachens, 1990). It should be noted that upward continuation is carried out over a distance ( h) from the original surface of observation, rather than to a height ( h) above a reference surface. Consequently the effects of the surface topography are not eliminated from the upward continued gravity field, though they are reduced along with all the other high wave 113 number components. Ideally, to produce an anomaly map which totally eliminates the effects of surface topography the extended Bouguer gravity value at each station or grid node should be downward continued by the corresponding elevation value. Equally, to produce an upward continued field map which eliminates these effects, continuation should be carried out to a height h (above the reference surface) which is above the highest level of observation. These are, however, daunting tasks in terms of time and effort, and would require considerable modification to the digital filtering programs.

The residual field was produced simply by subtracing the regional (upward continued) field from the extended Bouguer anomaly field (cf Blackwell et al., 1990). There is a certain amount of discussion as to whether the upward continued field should be downward continued to the original level before subtraction so that although the short wavelength information is lost, the magnitudes of the anomalies are preserved (cf Steenland, 1987), though Jacobson (1987) argues that this is not the case. In any event, the downward continuation of the upward continued field was not possible in the DigFil suite, and the residual map was not used for quantitative inteipretation, so this problem was largely academic.

Though this method involves much computation, its success is largely dependent on the choice of the appropriate continuation distance, and therefore the results cannot be regarded as being totally objective.

5.23 2 Spectral Factorisation

This method involves the splitting of the gravity effects arising from deep effects such as crustal thickness variations from those of shallower sources on the basis of the wavelength of their anomalies. This is carried out by first analysing the power spectrum of the extended Bouguer gravity field and then using the results to define band pass filters to separate the long and short wavelength anomalies.

Regarding the gravity field as the effect of an ensemble of blocks of varying size, depth and density (Spector and Grant, 1970), power spectrum analysis involves the statistical partition of the gravity field into components with similar properties using the relationship (Granser et al., 1989) 114 ln(E(k» = ln(Ai/k)f - 2/*/z, where k is the wave number vector, E(k) the radially averaged power spectrum, z0 the maximum source depth, and AJk) the "white spectrum", ie. the spectrum continued down to z0.

Calculation of the power spectrum values requires access to the Fourier coefficients, which are stored in the complex form in the DigFil program. A routine was written to access these variables, and to calculate the power spectrum, which is simply the sum of the squares of the real and imaginary parts of the complex variable, each representing one wave number harmonic.

Plotting a graph of In (E(k)) against wave number k produces a curve similar to that in Fig 5.5, which can hopefully be divided into straight line segments, each representing sources with similar properties. The gradient of this line can be used to give the maximum source depth z0. If the spectrum is completely random, with no grouping of anomaly sources, no straight line segments should be apparent.

The low frequency end of the spectrum can be regarded as representing the regional component of the gravity field, and the high frequency components the residual field. This enables a cut off frequency to be defined from the power spectrum below which regional features are dominant, and above which the residual features are more significant. A wavelength filter can then be defined to separate the regional from the residual using the wavelength filter available in the DigFil suite. The power spectrum also contains a significant amount of noise. This is represented by the very high frequency end of the spectrum, where the spectrum is asymptotic to thex axis.

Again, however, this method cannot be regarded as being totally objective in the separation of the two fields, as the choice of the wavelength is subjective within the constraints of the power spectrum.

Of the upward continuation and spectral factorisation/wavelength filtering methods, Gupta and Ramani (1980) suggest that the spectral factorisation provides the better results. In(Energy) 4. 0 .0 -4 12.00 u.oo U.OO 4.00 0.00 O'.OO iue . Tpcl oaihi pwr pcrm lt fo Grne e a, 1989) al, et ranser G (from plot spectrum power logarithmic Typical 5.5 Figure 56km 0.08 = r e b m u n e v a W 0.16 2TTI wave l h t g n le e v a w - 9km - Z X>< 0.40 115 116

5.2.4 Trend Analysis

Trend analysis is a method by which specific directional trends within the data can be enhanced or rejected. It should be noted that when trends are enhanced, even a completely random data set will show the required trend, so extreme caution must be used when applying such filters.

In the PNG case, this filter was used in an attempt to remove the dominant effects of the deformation trend resulting from the collision of the Australian and Pacific plates, with the intention of looking for any pre-collision features. Obviously any pre-collision features with the same strike as the defoimation would also be removed by such a filter.

5.3 Results of Data Enhancement

The results of application of the digital filters described in the previous section to the Papuan fold belt data are described below. It should, however, be pointed out that the application of the filters produces results which are affected by the data distribution and the mathematics of the Fourier transform. Therefore any features of the filtered maps which cannot be identified in the original data (ie Fig 5.1) should be regarded with caution.

5.3.1 Upward Continuation

The extended Bouguer gravity field was upward continued by 10km (10 grid units) as used by Gupta and Ramani (1980). The result of this operation is shown in Figure 5.6. The significant regional features illustrated by this map are the gradients associated with the Muller Anticline in the northwest (a), and the fold belt in the northeast, with the low in the far northeast comer also being shown as a regional feature. The gradients in the southern part of the map are more gentle, as would be expected in a foreland region, but the Lake Murray high (b) does show up as a regional high (albeit of low magnitude). The most significant feature of this map is the marked change in trend of the contours from the northwest to the northeast comer. Despite the poor data coverage between the two areas, it is clear that on a regional scale the structures of the Western Muller Anticline (a) and the Lavani region (c) are not at all similar. - 1 6

9400000

9375000 -26

9350000

9325000

9300000

9275000

9225000

9200000

9175000

9150000

9100000 550000500000 600000 650000 700000 750000 800000

Fig 5.6 Extended Bouguer gravity field upward continued by 10km. CI=5mGal.

425000

375000

350000

325000'

300000

275000

25000C Jo.

225000

200000

175000

150000

125000'

100000' 500000 550000 600000 700000 750000650000 800000

Fig 5.7 Residual gravity map having removed the upward continued field (fig 5.6) from the extended Bouguer anomaly field (fig 5.1). CI=5mGal. 118

The result of subtracting the upward continued gravity field from the extended Bouguer gravity field is shown in Figure 5.7. This residual gravity map shows a line of gravity highs extending to the southeast from the northwest comer of the map, and does not show the marked change in trend illustrated by the regional (upward continued) gravity map. This would appear to indicate that the sinuosity of the contours in this region on the extended Bouguer anomaly map is a result of large scale, and probably deep, structures. Although the residual map does show positive residuals over the regional highs, and negative residuals over the lows, which may suggest that separation of the regional and residual fields has not been fully achieved, it is probable that there is a coincidence of similar regional and residual features. This is supported by the surface geology which shows crystalline basement close to the surface in the Muller Anticline (a), and thick sediments in the northeastern comer of the study area. It can therefore be concluded that the near surface structures are of little significance in producing the sinuosity of the gravity field.

Also shown on the residual gravity map are the Lake Murray high (b) (which is more extensive on this map than the regional map), the "nose" in the Cecilia area (c), and the line of features extending northeastwards from south of Lake Murray to the southwestern edge of the fold belt (d). These are not represented on the regional map, and therefore must be the result of near surface density variations. The high associated with the Darai Plateau (e) occurs north of the corresponding high on the extended Bouguer anomaly map; this is because the southern edge of the plateau is more prominent on the regional map, probably as a result of the deep structure associated with the Darai Fault which forms the southern boundary of the plateau.

53 2 Power Spectrum Analysis

Initially just two power spectra were plotted, one for the whole data set (Fig 5.8), and the other for the foreland area bounded by northings 9170000 and 9300000, and eastings 500000 and 660000 (Fig 5.9). | Though both logarithmic power spectrum plots are very close to being "white", producing a near perfect exponential decay, several straight line fits have been identified. i

The two lines shown in Fig 5.8 give maximum source depths of 11km and 28km, and those identified in Figure 5.9 give maximum source depths of 5km and 16km. In the case of the foreland region, the depths fit closely with the known depths of "basement" (from seismic data O.jTS

ln(Energy)

z=28km

z=l1km

Wave number

Fig 5.8 Logarithmic power spectrum of the entire region.

ln(Energy)

z=16km

z=5km

Wave number

Fig 5.9 Logarithmic power spectrum of the foreland region. 120 and well control) and the lower crust (using standard continental crustal thicknesses). Therefore it would seem reasonable that for the whole data set, the values obtained represent the same interfaces, ie. that the maximum depth of "basement” is 11km and that of the base of the upper crust is 28km. (Note that "basement" in this case is used to describe those rocks which have a density indistinguishable from the standard crustal density of 2670kgm*3.) The fact that the foreland region foimed part of both power spectra was not felt to be important since the maximum depths for the whole data set would come from the fold belt region, and the inclusion of the foreland data in the whole data set allowed better control of the longer wavelength parts of the spectrum by increasing the size of the area.

These power spectra were also used to define the wavelength cut-offs for the bandpass filtering. For the foreland data, the change in slope (from representing a source depth of 16km to 5km) occurs at a wavenumber value corresponding to a wavelength of ~50km; the similar point on the spectrum representing the whole data set corresponds to a wavelength of ~ 100km.

Similar power spectrum plots were produced for the Eastern and Western Muller Anticlines (Hill, 1989a) (Figures 5.10 and 5.11). The Eastern Muller Anticline area was bounded by northings 9320000 and 9380000 and eastings 650000 and 730000, and the Western Muller Anticline area was bounded by northings 9340000 and 9420000 and eastings 540000 and 600000. Both these power spectra have been interpreted as showing the maximum depth to the lower crust -1 8km for the Western Muller Anticline, and 29km for the Eastern Muller Anticline, though these areas are geographically too small for reliable information on the longer wavelength parts of the sprectrum to be produced. The spectra for the Mananda and Hedinia areas were also produced, but these proved to be completely white, and no depths could be determined. The Muller Anticline spectra do show that major differences exist between the Eastern and Western anticlines; based on the maximal source depths given for the lower crust, it appears that the western anticline affects deeper crustal levels than the eastern anticline, and that the eastern anticline is more typical of the other fold belt structures found further to the east.

However, it should also be noted that these depths are only a reflection of the structure in the geographic areas chosen, and that since the "whole data” spectrum also covers the Eastern Muller Anticline area, it may be that this feature is the sole cause of the inferred deep base of the upper crust 121

ln(Energy)

z=29km

Wave number

Fig 5.10 Logarithmic power spectrum for the Eastern Muller Anticline area.

ln(Energy)

z=18km

Wave number

Fig 5.11 Logarithmic power spectrum for the Western Muller Anticline area. 122

The three spectra from the geographically smaller areas (the foreland, and the eastern and western Muller Anticline) all show an alignment of points at the low-wavenumber end of the spectrum which give maximum source depths in the 75 to 150km range. Though these points lie in straight lines, the areas covered by these spectra are too small to control these long- wavelength parts of the spectrum, and these results were not used.

5 3 3 Wavelength (bandpass) filtering

Three band pass filters were applied to the fold belt data using the cut off values determined from the power spectrum analysis. Firstly a low pass filter, which passed only the field components with a wavelength greater than 100km was applied; the result is shown in Figure 5.12. This shows that a significant proportion of the field of the true fold belt region (Area 2 in Fig 5.1) has a long wavelength, and is therefore likely to have a deep source. Close to the eastern and western edges of the map, the contours are shown trending N-S. This is a result of the proximity of the edge of the data set, and is a common occurrence in digitally filtered data. The northern and southern edges pose slightly more of a problem: there is little data close to the southern edge, and certainly in the SE comer there does appear to be an unusually steep gradient, but this is not seen in the NW comer where the non-filtered gravity field also trends E-W (parallel to the edge of the data). This is the region of the western Muller Anticline, a large feature which would be expected to extend to a significant depth, and be characterised by long- wavelength anomalies. Yet there are no signs of its effect in Figure 5.12.1 believe that this too is a result of edge effects obscuring the features.

A major source of concern with this wavelength filtered map is the apparent regularity of the gravity highs. The high formed by the 40mGal closure close to the western edge of the map (a) is exactly 100km to the south of another local high (b), which in turn is exactly 100km to the west of another local high (c), and separated from it by a low formed by the 25mGal closure (d) which lies exactly half way between the two highs. This indicates that the application of the filter operation on these regions may have introduced errors into the data which are a probable result of the cut off wavelength approaching the dimensions of the area. These errors are more likely in areas of poor data coverage and close to the edge of the map where the grid has to be extrapolated before application of the filter program. Consequently, the only areas of filtered gravity field shown in Fig 5.12 which can be regarded with much certainty are the northeastern 700000 750000 600000

Fig 5.12 Wavelength filtered extended Bouguer anomaly field - 100km+ passed. CI=5mGal.

9 400000'

9375000

9350000

9325000

9300 0 0 0'

9275000 '

9250000

9225000

9200000

9175000

9150000

9125000

9 100000' 500000 550000 600000 650000700000 750000 800000

Fig 5.13 Wavelength filtered extended Bouguer anomaly field - 50km+ passed. CI=5mGal. 124 parts of the foreland and the central part of the fold belt (ie areas away from the edges of the map which also have good data coverage).

The second wavelength-filtered map, again using a high-cut filter, this time illustrating data with a wavelength greater than 50km, is shown in Figure 5.13. In an attempt to reduce the edge effects, the unfiltered giid wts txtcnced by 20kin in £ll diim lors tefore the FFT aid filtering were applied; this does appear to have been successful, and there are few edge-parallel contours close to the eastern and western edges of the map. Again, however, the cases of the northern and southern edges are less clear. Figure 5.13 does show a gravity high associated with the Muller Anticline in the NW (a), and another in the Juha region - the centre of the northern part of the map (b). Gravity lows are shown in the NE of the map, indicating that deeper elements reach their maximum depth in this region. To the west, the lowest values are found to the south of the Muller Anticline.

Other features shown in Figure 5.13 are more difficult to explain: these include the low values at the western end of the Darai Plateau (c), and the steep gradients in the SE comer of the map. The Lake Murray area (d) shows up as a high on this map, unlike Figure 5.8 where the high is to the north of Lake Murray, and was almost certainly affected by the edge of the data. Also of note is the apparent flat in the foreland, running E-W close to northing 9250000.

Since the spectral analysis indicated that the 50km+ filter would not be expected to give complete separation of regional and residual effects in the northern areas (fold belt), the anomalies shown in the fold belt region will also result from the effects of changes on sediment thickness. Thus the similarity in the contour patterns of the upward continued and band pass filltered maps (Figs 5.6 and 5.13) implies that separation of regional and residual was not fully achieved by either method. However, the upward continued map does not show some of the apparently short wavelength features shown in the 50km+ band pass filtered map, such as the high over the Darai Plateau (e), and the low at its western end, and the steep gradients in the southeast comer of the map. This is strong evidence that the upward continuation filtering has provided better separation in the fold belt region than the band pass filtering.

Figure 5.14 shows the fold belt gravity field having been passed through a low-cut filter, with only the anomalies with wavelengths less than 100km being included. Geariy this data is much 125

0400000

B375000'

8225000

8200000

8150000'

9100000 900000 500000 600000 650000 700000 750000

Fig 5 14 Wavelength filtered extended Bouguer anomaly field - 0-100km passed. Cl=5mGal.

Plate 5.1 Wavelength filtered extended Bouguer anomaly map - 0-100km passed. CI=5mGal. 126 more irregular, with near-surface effects having greater significance. Edge effects are apparent along both the eastern and western edges.

The northern part of the map shows a line of gravity highs which decrease in amplitude from the Muller Anticline in the NW (a), through the Lavani (b) and Karius (c) Anticlines (the 15mGal closure in the centre of this area), to the Mananda Anticline (d) (the lOmGal closure) and beyond to the SE, including the Hedinia and Iehi (e) structures, and the NE part of the Darai Plateau (f), close to the Kikori River. These features are more clearly shown in Plate 5.1. Rather surprisingly the Darai Plateau itself does not appear as a major high on this map, though its SW margin is characterised by a steep gradient.

The N-S trending high at the eastern edge of the map is clearly affected by the edge of the data, and does in fact appear to follow the distribution of the stations in this regioa

The southern parts of the map, covering the foreland region, show highs in the Elevala (g) and Lake Murray (h) areas, and suggest the eastward continuation of the Lake Murray high. Significant gravity lows are shown to the southwest of the Muller Anticline, to the southwest of the Elevala high, and in the region of the Bamu survey (i).

Together these three maps should enable the relative depths and sizes of the various anomalies to be determined using the not unreasonable assumption that the longer the wavelength, the greater the depth of the causative structure. It should, however, be noted that this assumption is not always true, as the geometry of bodies will also affect the wavelength of the anomaly. With reference to the fold belt area, it can be seen that the Muller and Juha features have deeper causes than the other fold belt structures, as these are the only ones which can be seen in Figure 5.13. The structures of the Karius, Mananda and Hedinia regions are shown to be larger and to extend to greater depth than the fold and thrust structures which occur further to the NE and which are not apparent on any of the filtered maps.

In the foreland region, the gravity lows immediately to the south of the fold belt deformation front are only significant at intermediate depths close to the Western Muller Anticline and the western Darai Plateau, possibly indicating less crustal deflection in other regions. 127

5 3 .4 Trend Analysis

Only one trend-filtered map was produced; this is shown in Figure 5.15 and Plate 5.2. This shows the study area gravity field filtered to show only anomalies with strikes between 015° and 075°, in an attempt to remove the effects of the dominant NW-SE trend of the fold belt structures. In this case in particular it is very difficult to determine which elements of the filtered data are edge effects, and which are genuine; however I believe the map does confirm the presence of NE-SW trending features in the region. The most striking feature is the line of gravity lows which extends from the southwest comer of the map to the Nipa (a) and Tari (b) areas in the northeast. However, many of the other features shown simply represent the NE-SW component of known structures such as the Muller Anticline (c) and the Lavani/Juha structure

(d).

In the foreland there is a clear gravity gradient, with highs to the northwest and lows to the southeast, extending northeastwards in a broad zone from Lake Murray (e) to the Cecilia area (f), and beyond. The gradients are comparatively gentle (>0.5mGal/km), and are therefore unlikely to be the result of any major structures. More likely causes are either a series of minor structures, such as the antithetic wrench faults, or variations in sediment thickness. The Mesozoic isopach map and the base-Tertiary sub-crop map of Jenkins (1974) do not lend support to the latter postulate; any thickness changes must therefore be in the Tertiary sedimentary sequences.

It should be noted that the magnitude of many of the anomalies shown in Fig. 5.15 is small when compared with the overall extended Bouguer anomaly field, and therefore that the anomalies shown in the trend filtered data are considerably less significant than those of the NW-SE trend.

5.4 General Comments on Enhanced Results

The results described in the previous sections have demonstrated that the digital filtering of the gravity data can enable information to be extracted from the data which would otherwise have gone unnoticed. However, the limitations of the filtering methods, particularly where edge-effects I are concerned, have also been demonstrated. The problem faced in the interpretation of such data is the separation of the genuine effects from the filtering effects. It should also be borne in mind, 9425000 -20

9400000'

8375000

9350000

8325000

8300000

9275000

9250000

8225000

9200000

9175000

9150000

9125000

9100000 800000 500000 550000 600000 650000 700000 750000 Fig 5.15 Trend filtered extended Bouguer anomaly field - trends from 15° to 75° passed Cl-5mGal.

Plate 5.2 Trend filtered extended Bouguer anomaly map - trends from 15° to 75° passed. CI=5mGal. 129 especially in thie case of the Papuan fold belt, that the gravity field is also affected by other factors, such as station elevation; this has the effect of making anomalies appear more significant in some areas than others because of a decrease in source-station separation. In addition, some of the results of the enhanced data, particularly the wavelength filtering, may be susceptible to errors where the wavelength of the topography is relatively constant: in most cases topography can be regarded as having a totally white spectrum, though it is possible that this is not the case in the fold belt region.

Fortunately these are not the only methods of interpretation available, and the technique known as forward modelling is also commonly used; this method does take the station elevations into account, and should therefore allow the positions of the sources of many of the anomalies described in this chapter to be confirmed.

v 130 6. Forward Modelling

The gravity anomaly maps show clearly that the dominant trends in the gravity field are in a NW-SE direction, and the geological maps reveal that the geological structures are elongated in a similar direction. Therefore two-dimensional forms of analysis can be used to solve problems connected with these structures (Talwani et al., 1959). The principal advantages of such a situation are that the concepts behind these two-dimensional methods are relatively simple, and models can easily be displayed in the form of polygonal two-dimensional bodies which are easily input into computers (Cady, 1980). A fundamental assumption of this method is that the bodies effectively have infinite strike lengths (i.e. y extends from to +«>). The two and a half dimensional (2.5D) method, which allows the strike length of the model to be constrained, was considered, but it was felt, given the linear nature of the fold belt structures, that it would not provide any greater accuracy.

6.1 Forward Modelling Theory

The technique known as forward modelling, as applied to gravity data, involves the calculation of the gravity field due to a model of the subsurface geology. Modem applications involve the creation of such a model interactively on a computer, with density values being assigned to polygonal bodies representing geological units on the basis of the contrast between the expected (theoretical) and true rock density; the resulting anomaly can be calculated and displayed graphically. The calculated gravity profile is compared with the observed anomaly profile, and the model refined until the two profiles are similar.

Calculation of the gravity anomaly due to a model makes use of the fact that because of the infinite strike length of the polygons, the volume integral (Equation 4.7) can be expressed as a line integral around each polygon by expressing z as a function of x for each side in turn. Cady (1980) derives the expression

y ( t A -l i TJln r i2-Siln r i2+2KJtan -2fC,tan_11 1 k 1 i J- 131 where

2 2 2 si=xi+xoi r^xj.j+x, Oi r i =xi + zJ

2 2 2 r l+ l~ x l + l+ z l+l K*mTT 'OJ '-'J Cj=N

(see Figure 6.1 for explanation of variables).

However, the forward modelling technique does not produce a unique interpretation of a given set of data; any profile can be produced by an infinite number of combinations of density contrasts and polygon shapes. The ambiguity of the forward modelling method can be reduced by the use of other constraints on the nature and form of the anomalous bodies (Kearey and Brooks, 1984). The constraints used in this study are described in Section 6.3.

Another significant and often overlooked point concerning forward modelling is the effect of the station elevations on the observed gravity profile. Although the process of gravity data reduction removes the effects of the theoretical gravity field at the height of the observation, and the effect of rock between the station elevation and the chosen datum (usually sea level), it does not alter the fact that the observations were not made on the datum surface, and the additional distance of the station above the datum must always be taken into consideration when modelling, especially in areas such as the Papuan fold belt, where relief is very significant.

6.2 Forward Modelling of the Papuan Fold Belt Data

A total of seven profiles were selected in the study area, and their locations are shown in Figure 6.2. In the case of the five SW-NE profiles, every effort was made to select profile lines perpendicular to the strike of the contours, and to include at least one well on each profile to improve the geological control.

Because of the very patchy distribution of gravity stations within the study area, the regional gravity profiles were produced from the gridded data set, rather than from the individual survey lines. This has the effect of losing some of the detailed information from surveys with close 132

+z

Fig 6.1 Calculation of the gravitational attraction of an infinite body. 133

8425000

9400000'

8375000 Andebare-l Lavaru-1

9350000 \ > Hides-1

9375000 Elevala-1 70 Nembi-1 9300000

8275000

8250000*

8225000

Lake Murray-1 820C000

8175000* Bamu-2

8150000

8125000

8100000 500000 550000 600000 650000 700000 750000 600000

Fig 6.2 Extended Bouguer anomaly map showing the location of the gravity profiles and models. Axes in UTM metres. 134 station spacings, and also introducing the possibility, in areas of sparse data coverage, of including features on the profiles which are products of the gridding operation. However, as interpretation was carried out on a broad scale, this problem was not felt to be significant. A computer program was written which produced x,z values at a specified regular interval along a line between two specified points within the gridded area. For any point on the profile the z value was calculated from the values of the four nearest grid nodes using an inverse square average. This method was applied to both the gravity and elevation data.

Initially it was hoped to use the DTM, created for the determination of the terrain corrections, in the production of the elevation profiles, but modelling with this data proved impossible in areas of sparse data coverage. In such areas the profile topography from the DTM showed large changes in elevation which were significant in determining the calculated gravity but were not reflected in the observed gravity profile which was extraploated from more distant points. Consequently, a grid of the station elevation values had to be used as the basis of the elevation profiles.

The modelling program used was a commercially available package called GM-SYS which ran on IBM compatible PCs; the PS/2 Model 70 was the system used in this case. This package allowed the profile data to be read directly from files and could cope with up to 200 points per profile, which meant that individual profiles up to 200km in length could be modelled with the lkm interval used. Several of the profiles were more than 200km in length; in these cases the profile was split into two or more shorter profiles. It was also necessary to change the direction of some profiles in order to keep them perpendicular to the contour trend; in these cases the profiles were split into individual straight line sections.

Models could be created and modified either by using a mouse to identify the relevant vertex positions on the screen, or by using the keyboard to type the co-ordinates of the points. The latter feature was particularly useful for ensuring that horizontal lines were truly horizontal, which was often difficult to determine visually. The modelling package also allowed the position and stratigraphy of wells to be defined and displayed on the model, which allowed consistency with the well stratigraphy to be maintained relatively easily.

The modelling program could be used for either 2D or 2.5D modelling. In this case the modelling primarily concerned the regional structure, and for clarity had to be kept reasonably 135 simple, so only the 2D method was used. Values of -30000 and +30000km (default values) were used as the y axis limits of the models, and the same values were also used for the limits of the x axis, though variations in the models ceased to have a significant effect on the gravity profile about 10km beyond the end of the profile line. It was also assumed that bodies deeper than 50km had no effect on the observed gravity profiles.

The models which were produced are described in Section 6.5, and their overall significance is discussed in Section 6.6.

6.3 Constraints on the models.

The gravity anomaly is produced by the distribution of rock of various densities within the subsurface, and any anomaly profile can result from an infinite number of combinations of density contrasts and distributions of bodies within the subsurface. To make the task of modelling manageable, certain assumptions have to be made about the likely densities and positions of the subsurface bodies, generally using information other than the gravity field. This information can include such details as surface geology, well information, seismic data and the geomorphology of the region. The principal constraints used in the modelling of the gravity field of the Papuan fold belt are discussed below.

6.3.1 Surface Geology

The surface geology is perhaps the single most important information available in most regions because it enables both rock density and distribution to be constrained. Rock density can be measured from specimens collected in the field, and even if no samples have been collected, a knowledge of the lithology will place certain limits on the possible densities. Table 6.1 shows typical density values for lithologies which are common in the Papuan fold belt. The geology of the fold belt has been discussed extensively in Chapter 2, so reference to the formation and lithology descriptions, and to Table 6.1 will give some idea of the densities involved. However, these values are only approximate, and often the range of possibilities given is too wide to be really useful. 136

Lithology Density Range (kgm'3)

Wet Alluvium 1960-2000

Sandstone 2050-2550

Quartzite 2600-2700

Limestone 2600-2800

Granite 2520-2750

Granodiorite 2670-3200

Basalt 2700-3200

Table 6.1. Typical rock density values. Data from Kearey and Brooks (1984).

Formation Density (kgm'3)

Darai Limestone 2700*

Feing Group 2550*

Chim Formation 2620

Wahgi Group 2620*

Kana Volcanics 2690

Strickland Granite 2600-2670*

Kubor Granodiorite 2740

Omung Metamorphics 2830

Table 6.2. Density values of the rock units of the Papuan fold belt. Data from Zadoroznyj and Coutts (1973) and St John (1967)*. 137

Measurement of rock densities from the Papuan fold belt has been carried out by several authors, notably St. John (1967), and Zadoroznyj and Coutts (1973), and the results of their experiments are summarised in Table 6.2. However, it should be noted that these density values are affected by factors other than lithology. Weathering affects surface samples, reducing the measured density, and the bulk density of formations is also affected by phenomena such as karstification. This is particularly important in the case of the Darai Formation, which contains many large caverns.

6.3.2 Formal Gravity Interpretation

A formal interpretation can be loosely defined as one which depicts all aspects of the processing and interpretation of gravity observations. Such an interpretation has been carried out over the fault-bounded Darai escarpment by Anfiloff and Flavelle (1982). They combined the processing and interpretation of the gravity data into a fully computerised procedure. The topographic anomaly was calculated by modelling the plateau surface, and the escarpment density was deduced from Nettleton profiles (Nettleton, 1939) to be about 2500kgm'3 (Fig 6.3). Two dimensional modelling was used to allow for the upward continuation effects produced by the rapid increase in station elevations over the scarp slope, giving topographic density values in the region of 2400 to 2450kgm'3.

Anfiloff and Flavelle (1982) also modelled the sub-topographic density distribution, as the boundary fault, which coincides with the base of the escarpment, results in deeper, more dense, strata being at a higher level on the northern (upthrown) side. Using forward modelling methods they determined a density value at the base of the escaipment of 2200±50kgm'3.

Fig 6.3 shows the actual Bouguer anomaly profiles for a range of topographic densities, illustrating the effects of the choice of density. The synthetic Bouguer anomalies, which are symmetrical about the profile corresponding to the topographic density, show the case when the topography has a uniform density, and comparison of the synthetic and actual Bouguer profiles shows those anomalies which result from density contrasts close to the topography.

Though Anfiloff and Flavelle (1982) do not attempt any geological interpretation of their results, it seems likely that topographic density values to the south of the plateau are ~2200kgm‘3 138

COMPUTED

ANOMALY

SYNTHETIC B. A 30mGal

ACTUAL B. A

2.5

2.2

30 3.0 0 10km

Fig 6.3 Model of the Darai Escarpment showing synthetic and actual Bouguer anomaly values for a range of topographic densities (1.5 to 3.5 g/cc) to illustrate the effects of changes in Bouguer density. The computed anomaly includes an upward continuation correction which is a maximum where there are rapid changes in station elevation. (From Anfiloff and Flavelle, 1982) 139

(Pliocene to recent elastics), and those of the plateau itself (Darai Limestone) are in the region of 2400 to 2500kgm'3.

6.33 Density determinations from Borehole Gravimetry.

Additional rock density information for both the terrain (topographic) correction of the gravity data and modelling of the gravity anomalies comes from the results of the borehole gravimeter (BHGM) surveys of the Hides-1 and Angore-IA exploration wells in the Southern Highlands.

The principle behind the BHGM method is that the difference between the observed gravities g! and %2 at two levels within the borehole is given by

gr g2 = (3.086-4?ipG)Ah where Ah is the vertical distance between the two levels, and p is the rock density. This equation is simply a result of applying the free-air and simple Bouguer corrections to the two readings. The BHGM results are also affected by the terrain, and terrain corrections out to 40km had been carried out for the Hides survey by the contractors using a method involving the fitting of conic surfaces to a terrain model (Edcon, 1988b). However, the terrain correction of BHGM surveys is complicated by the fact that a density has to be assigned to the topography in order to evaluate the terrain correction, and this density has a significant effect on the resulting value for the formation density. For the terrain corrections applied to the Hides survey, a method was used in which the terrain density pA is assumed to be the same as the true rock density pT at the well, given by

Pt = Pb/O "^) where pB is the measured BHGM density, and k is a geometric factor dependent upon the shape of the terrain and position of the gravity stations which is determined by a digital terrain fitting exercise (Edcon, 1988b).

The results of the Hides-1 BHGM survey are shown in Fig 6.4. They show a near surface apparent density of 2720kgm'3, rapidly decreasing with depth to 2200-2300kgm'3 below 50m, 140

Depth (m)

LL

500

1000

1500

LL 2000

2500

3000

2000 2500 3000 Apparent density from BHGM (kgm*3)

Fig 6.4 Results of borehole gravimeter survey of Hides-1 141

and then slowly increasing to around 2700kgm'3 at the base of the Darai Formation at a depth of 1100m. The underlying Ieru Formation sediments are of relatively constant apparent density in the 2450-2600kgm'3 bracket. The mean apparent density of the top 500m of the Darai succession is 2370kgm'3, and the mean for the bottom 500m is 2570kgm'3. The mean for the whole Darai succession is 2480kgm*3.

The sub-Darai succession provides fairly unsurprising results, but the apparent density variations within the Darai Limestone are of considerable interest. The initial large drop in apparent density could be partially due to the terrain correction method. The lack of detailed topographic maps, and the bulldozing of the Hides site meant that detailed terrain information within about 200m of the well site was not available (Edcon, 1988b). This means that the terrain corrections for the upper stations in the well were not as accurate as those below (Edcon, 1988b), and because the effect of terrain on the apparent formation densities is a maximum at the surface (Hearst et al., 1980) this lack of information is of greater importance for the apparent densities near the surface. The effect of the terrain correction on the apparent density is illustrated in Table 6.3.

The remaining characteristics of the apparent density profile of the Darai Formation, namely the relatively low density close to the surface, increasing with depth, are of more significance. Two possible explanations have been put forward by Dr P Morris of BP Exploration (pers . comm., 1989), given that the near surface apparent density is far below that which would be expected from a formation comprising rock which has a laboratory measured density of 2700kgm'3.

Depth interval (m) Apparent density Apparent density (kgm'3) (kgm'3) from terrain from non-terrain corrected BHGM data corrected BHGM data.

0-50 2723 1325

210-240 2386 1707

1527-1557 2574 2274

Table 6.3 Effect of terrain correction on BHGM apparent densities. Data from Edcon (1988b). 142

(i) Karstification of the limestone, which will be more prevalent close to the surface, will reduce the apparent density. Given that any such cavities in this region are likely to| be air-filled rather than water-filled, it has been shown that a cavern close to the well could produce the observed effects (Fig 6.5).

(ii) A combination of increasing density with depth (possibly as a result of greater compaction) and dipping strata will also produce the observed effects (Fig 6.6). The Hides well is sited on the crest of a large thrust-faulted anticline which has a sharp crest and limbs dipping at 55° to the NE and -20° to the SW (Hill, 1989b), so such a situation is quite likely. Corrections can be applied for such effects, eg Brown and Lautzenhiser (1982), but these require an independent means of accurately measuring formation density.

Given that both karstification and dipping beds are known to occur, a combination of the two effects is the most likely.

For modelling purposes a density of 2500kgm*3 was used for the entire Cretaceous to Miocene sedimentary sequence. This choice was in part backed up by the results of the BHGM survey in Angore-IA early in 1990, which showed the density of the Darai Formation ranging from 2500kgm'3 at the surface to 2600kgm'3 at 2500m, and the Ieru sequence immediately below averaging 2550kgm*3 (BP Australia, per s. comm., 1990). The values for the upper parts of the Darai Formation are generally higher than those obtained in the Hides survey, possibly because of the less severe terrain and lower elevation at Angore, resulting in less karstification, or the more gentle dips in the strata.

6.3.4 Density values used for modelling the fold belt rocks

Having estimated the densities of the various rock units, it was necessary to group together formations with similar density properties into larger units for modelling purposes. The following bodies were used:

Pliocene to Recent - density 2200kgm'3 - based on the lithologies and the results of Anfiloff and Flavelle (1982), and shown as unit 1 in the regional models. 143

o o CL O o o £ E £ 8

> s to to T n l ^ f A © c o IT) £ n / ° n CVJ 3 ■85 (O a> o O to tj X q . as J£o> CD o tn o © o ■u CM i-E(O X c CD © ■o O *—* O c o CD E CO i - o tc o to Q. in ITa> Q. ■o < CL c o 8> a) to Q. O Q. m (0 CM ■O k_a> E 3 O) t/> (0 o a> o 2 in o CM CM to v-to c o o o o CL o o o >» a> o o in O o in "toc o

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Densities (kgm3) Depth (m) Depth (m) o o to o O w o O to CM o o o o o o o o 10 in in in o o C\J o M O CM o o o CO o o o CM o o o to to CM o o CM O to CM o o

Q 0) O) E

Fig 6.6 Effect of dipping beds on BHGM survey 144 145

Cretaceous to Miocene - density 2500kgm'3 - based on lithology, Nettleton profiles of Anfiloff and Flavelle (1982), and the borehole gravimetry results. Also consistent with the value used for the terrain corrections. Shown as unit 2 in the regional models.

"Basement" (Jurassic and older) and upper crust - density 2670kgm'3 - the standard crustal value was used, based on the lithology and age of the sediments, and the lithology of the upper crustal rocks. Shown as unit 3 in the regional models.

This rather simple density structure was chosen because it made the models much more simple, and hence clearer and easier to understand, with little loss of detail as precise controls on the density of each individual formation were not available. Additionally, the models were only concerned with the broader scale structures, so any small inter- and intra-formational density changes would produce trivial effects, and indeed some formations have larger internal density variations than contrasts with the underlying and overlying formations.

6.4 Overall density structure of the models

As mentioned in Section 6.1, the forward modelling technique requires the density contrast between the actual and expected (theoretical) rock density of each body in the model to be determined. The actual density values have been discussed above, but the source of the theoretical density values may be less clear.

The theoretical density values come from standard models which have been produced for both oceanic and continental crustal structures, and are largely based on deep seismic reflection work and seismic tomography. Though there is general agreement on the depths of the major physical changes within the crust, the method of determining the density from p-wave velocities is not precise, and the results are subject to errors of ±100kgm'3 (Kearey and Brooks, 1984). Dziewonski et al., (1975) give an theoretical density value of 2720kgm'3 for the upper 20km of continental crust, 2920kgm*3 for the lower crust (20-35km), and 3320kgm'3 for the upper part of the mantle (35-70km), which is shown as unit 5 in the regional models. However, St. John (1967) used an expected density value of 2670kgm'3 for the upper crust, increasing to 2800kgm3 at 35km, and Sibuet et al., (1990), studying the UK continental margin, used values of 2730kgm‘ 3 and 2830kgm'3 for the upper and lower crust respectively. 146

For this study, the standard crustal value of 2670kgm'3 was used for the upper crust (unit 3 in the regional models), as this is the generally accepted value for sialic continental crust. Using the depth constraints on the near surface units provided by the well information, and on the deeper units by the foreland power spectrum, a density value for the lower crust (unit 4 in the regional models) was chosen which provided the best fits with the observed gravity profiles in the foreland region. The value obtained by this method was 2870kgm'3, somewhere between the extremes of St John and Dziewonski et al.

The density contrast values used in the modelling process are summarised in Table 6.4.

Depth (km) Above 0-20 20-35 >35 datum Rock Unit Unit number

Pliocene - 1 -300 -470 Recent

Cretaceous - 2 0 -170 Miocene

Basement and 3 170 0 -200 Upper Crust

Lower Crust 4 - 200 0 -450

Mantle 5 -- 450 0

Table 6.4 Density contrasts (in kgm'3) used for modelling.

6.5 Description of the profiles and models

The location of the profiles described in the following section is shown in Fig 6.2. In each model the geological units are numbered 1 to 5, and correspond to the geological units as shown in Table 6.4. 147

Profile 1.1

Profile 1.1 (Fig 6.7) extends across the foreland from Lake Murray in the southwest to Elevala in the northeast, a distance of about 120km. The model shows the thickening of the Pliocene to Recent sediments northeastwards from less than 100m at the southwestern end to ~ 1100m at Elevala. The relatively steep gravity gradient at the southwestern end of the profile has been explained in terms of a steep northern face to the Lake Murray basement high. The base of the upper crust deepens from 17km in the southwest to 18.5km in the northeast. The mantle shows a corresponding change from 33.5km in the southwest to 35km in the northeast

Profile 1.2

Profile 1.2 (Figs 6.8 and 6.9) is the northeastwards extension of Profile 1.1, and continues to the ENE from Elevala-1 through Juha-1 and Lavani-1 to Andebare-1, a total distance of 168km. There is a change in the direction of this profile at Lavani-1.

The western end of the model shows a deepening of the post-Jurassic sediments close to Elevala- 1. Interpretations with a normal fault to the east or the west of the well proved equally acceptable in terms of the gravity profile, although those solutions with the fault to the west of the well fit the well stratigraphy more closely.

The region from 50 to approx. 130km along the profile was modelled using the geological interpretations proposed by Hill (1989b) to constrain the configuration of the bodies representing the Cretaceous to Miocene and Pliocene to Recent sediments. Profile 1.2a (Fig 6.8) models Hill’s "Inteipretation A", which shows thrust planes dipping to the southwest in the Lavani and Juha structures. Hill’s "Interpretation B", which shows the more common geological interpretation with NE dipping thrust faults, is used to constrain the near-surface geology of Profile 1.2b (Fig 6.9).

For the Lavani and Juha structures, both solutions proved equally acceptable in terms of the gravity profiles, though in order to maintain consistency with the observed gravity values the crest of the Lavani anticline in both cases was placed about 2km to the southwest of the Lavani- 1 well position. This difference could be a result of the gridding of the original raw data. Profile 1.1 5211 Aai o - j B iA a A (5120111) o o 00 ID . a a ID a o LD t H • > ID O o CD H w H t S rH Q) > (0 C0 Ps 2 2 (0 I H 1 CO CO \ CN ; a a ID o cn (ui) m DLD ID a o o a t a a d a n H O o o n r 0 0 O LD O n v »r> o ID □ CD TP O O 00 t IT) ID 0^ o o □ O 00 LD o w ° N ^ ID H I •H -p <*/ t O O tn CO PI u 0 0 II H

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151

To the northeast of the Lavani anticline a second peak (A in Fig 6.8) is observed in the gravity profile which does not correspond to any structures shown in Hill’s cross sections. There are two possible explanations of this. Firstly there could be an anticlinal structure (either a pure fold or a hanging wall anticline) beneath the near surface stack of Miocene limestone thrust sheets, and secondly the observed gravity peak could result from the fact that there are very few data gravity stations to the east of Lavani, and in the Lavani-Andebare region the profile does not run precisely perpendicular to either the gravity contours or the dominant geological trend. As a result the gravity profile information has been interpolated from data points up to 20km away, and may therefore be subject to significant errors. The oblique profile will also mean that peaks in the profile will be drawn out along the line of section, and the dips of the units within the models will appear more gentle than in reality. A second profile (Profile 1.3) was produced in an attempt to reduce these errors.

Further to the northeast of Lavani, the gravity profile shows two positive anomalies (B and C in Fig 6.8) on the steep gradient before flattening out in the Andebare region. These correspond well with the two thrust faults affecting the Mesozoic section shown in Hill’s cross sections, though the density contrasts are not sufficient to have the necessary effect on the gravity profile from such depth. It is perhaps more likely that the small peaks in the gravity profile result from small near surface structures and/or density changes associated with the deeper structures.

The deeper crustal structure shown along Profile 1.2 shows the lower crust rising to a depth of 16- 17km below the Juha structure, then descending to a depth of 23-24km to the northeast of Lavani. The base of the lower crust is shown at a relatively constant 34-35km to the southwest of Juha, descending gently to a depth of 39-40km to the northeast of Lavani.

Profile 13

This profile was produced in the Lavani-Juha region to supplement Profile 1.2 and was confined to areas where the original gravity data coverage was good.

Again models were produced representing the two geological interpretations (A and B) produced by Hill (1989b) (Figs 6.10 to 6.13). Interpretation A (Model b) shows southwest dipping thrust Profile 1.3 Model a n) ndsa d tn cnx) 152 Profile 1.3 Model b C ioh sibo o m ) oo CO ' M n) ndsa d tn cnx) © O CN ID ' CN 153 (N (N o f) O J Profile 1.3 Model a Profile 1.3 Model b C*T® 0 «) a 1 * 1*019 a a •4 1 H I (N (*i) (wj) ^a»a m n *

Density contrasts (kgm3): 1 -470 (-300 above sea level); 2 -170 (0 above sea level); 3 170 above sea level, -200 below 20km; 154 4 4 200 above 20km, -450 below 35km; 5 450 above 35km 155 faults at Lavani and Juha, and Inteipretation B (model a) shows northeast dipping thrust faults. Though the magnitude of the peak in the observed gravity profile to the northeast of Lavani (A in Figs 6.10 and 6.11) was reduced along this new profile line, it was still necessary to invoke a significant basement structure to account for observed Bouguer anomaly values.

In the Juha region the base of the upper crust was maintained at a constant depth of 18-19km. This had the effect that neither of the geological models exactly fitted the observed gravity profile, and only "Interpretation B" could be realistically modified to fit by thickening the pre- Cretaceous section involved in the structure relative to that shown in the models along Profile 1.2. Changing the level of the base of the upper crust enables a better correlation of observed and calculated Bouguer anomaly values in the Juha region, but because of the longer wavelength of the gravity anomaly resulting from the deeper bodies geologically unreasonable near surface structures are required to maintain the good correlation of observed and calculated gravity in the Lavani region.

Profile 2.1

This profile (Fig 6.14) extends from the southwest comer of the study area to the Cecilia-1 well, a distance of 198km. A long wavelength lower crustal bulge is used to account for the regional component of the observed gravity profile, with a thickening of the Pliocene-Recent sediments and deepening of the Cretaceous-Miocene units being used to account for the local low in the middle of the profile. Although such a deepening of the pre-Mesozoic surface has been noted before, the location shown here is significantly to the south of the position indicated on the map of Bain et al., (1972), but significantly to the north of the Fly fault zone which they also show producing an increase in depth of this surface.

The northeastern end of the model shows the downwards displacement of both lower crust and mantle as the fold belt is approached. The gravity profile also shows the effects of the Cecilia anticline, and it would appear that the anticline does not extend to basement level; this supports many cross sections of this structure (eg Davies and Norvick, 1974) and to a certain extent Hill (1989b), which show the decollment surface associated with this structure flattening out in the Cretaceous section. Profile 2.1 C sioqti ) o n o a i □ a in i C/5 £ fN o a in a ( i u m ) ni nin in a D in o O in O o o o r-i an> O ca o q^daa + o o a o 4J X > w o o in c © O © fl a §> 5 1 N ^ » ' « > e i :*? 5 1 L o i^ ...... " £ * t"® tn r-~ o ■3 £ CN 1 * f9 S i ^ 3 so 8 o !* . S 5? 156 J8

157

Profile 2.2

This profile (Fig 6.15) extends northeastwards from Cecilia-1 through Hides-1 to Andebare-1, a distance of 94km, and is the [northeastwards continuation of Profile 2.1. The model also suggests that there is no basement involvement in the Cecilia structure. The Wai-Asi anticline, which lies immediately to the northeast of Cecilia, is not reflected in the gravity profile. This could either be because there are no gravity stations over this anticline, the profile data being inteipolated from other nearby points which are off-structure, or because at the position of the profile the Wai-Asi anticline is not a significant structure. The former is perhaps more likely, as the geological maps show the structure persisting in this region.

Moving along the profile to the northeast, the next major structures shown in the model correspond with the positions of the Karius and Hides anticlines (A and B in Fig 6.15); they are represented by the first two of four successive small positive anomalies superimposed on the regional Bouguer anomaly profile, and the morphology and location of the corresponding structures shown in the model are consistent with Hill’s geological interpretation. However, to the east of the Hides anticline only two more structures are apparent from the gravity profile (C and D in Fig 6.15), while Hill (1989b) shows three Mesozoic thrust sheets in this region. The gravity profile also suggests that the Cretaceous-Miocene sediments extend to a greater depth than shown on the geological cross section, though it should be pointed out that the lines of the gravity profile and Hill’s geological section are not identical.

At the eastern end of the gravity profile, which is beyond the end of the geological section, a large structure affecting the Mesozoic sequences is inferred to the WSW of Andebare-1, presumably somewhere beneath the Doma Peaks Pliocene volcanic centre. Though it is possible that the volcanic rocks of this region, and possible sills and dykes associated with the volcanic centre, may be the cause of the gravity anomaly, the thrust faulted anticline proposed in the model is more consistent with the other models from neighbouring profiles.

At deeper levels the model shows a relatively uniform depth for the upper surfaces of both lower crust and mantle beneath the deformation front in the region of the Cecilia anticline. The major downward displacement of both these horizons occurs 20-40km further along the profile, beneath the Karius and Hides structures, which are the first major (>2500m) topographic peaks Profile 2.2 C sip - d *) ^aTABa-o (UI) *nd»a 159 encountered along the profile. At the eastern end of the profile the lower crust and mantle are both 4-5km below their standard depths.

Profile 3

The gravity profile along this line (Fig 6.16) is dominated by the high associated with the Muller Anticline. Such a high could not be reproduced by only involving the basement (i.e. the Pre- Cretaceous, which does outcrop in the core of the anticline) in the model of the structure, and it was necessary to involve both the lower crust and mantle in this structure to reproduce the amplitude of the observed Bouguer anomaly.

This model also includes two areas of thickened Pliocene to Recent sediments, one to the south of the Muller Anticline, and the other to the south of the intersection with Profile 4. The Cretaceous-Miocene unit appears to thicken only slightly in a northerly direction.

To the north of the Muller Anticline the near-surface structures shown in the model are based on the geological interpretations proposed by Hill (1989b), and appear to be consistent with the gravity profile. The lower crustal and mantle structures shown on this model are similar to those of the rest of the fold belt, modified by the effects of the Muller Anticline.

Profile 4

Profile 4 (Fig 6.17) runs roughly parallel to the trend of the gravity contours in the foreland region, to the south of Elevala-1 and Cecilia-1, and lies approximately 30km to the southwest of the mountain front. Because of the proximity of major structures along strike from the profile line the 2D modelling method is less justifiable in this case; however in the interests of clarity and simplicity the 2D method was used. It has been assumed that there is no mantle topography affecting the gravity profiles in this region, and consequently the model does not extend below 25km in depth. The magnitudes of the gravity anomalies along this profile are much less than along the perpendicular profiles; here the amplitude of the largest anomaly is only ~10mGal. Profile 3 o SBt) iUTAPJO (STBOtt) o I I o o ) X H ( Wdaa

Density contrasts (kgm3): 1 -470 (-300 above sea level); 2 -170 (0 above sea level); 3 170 above sea level, -200 below 20km; 160 4 200 above 20km, -450 below 35km; 5 450 above 35km Profile 4

Density contrasts (kgm 3): 1 -470 (-300 above sea level); 2 -170 (0 above sea level); 3 170 above sea level, -200 below 20km; 161 4 200 above 20km, -450 below 35km; 5 450 above 35km 162

With the information from this profile alone it would have proved very difficult to produce a single model as the number of possible combinations of the Pliocene-Recent and Cretaceous- Miocene rock units is limitless, and information from the models constructed along the intersecting profiles (1.1,2.1 and 3) had to be used to place some constraints on this model. This profile also shows a Bouguer anomaly high in the Ok Tedi region (western end of the profile), and two small highs between the intersections with profiles 1.2 and 2.1. These have been attribute to changes in the thickness of both the Cretaceous-Miocene and Pliocen to Recent successions (Fig 6.17), though there is very little to constrain the relative effects of these units, and no firni conclusions could be attached to this information.

Profile 5

This profile (Fig 6.18) extends from Bamu-2 in the foreland, over the Darai Plateau and through Kanau-1, to Nembi-1 and beyond, a total distance of ~200km. There is a change in direction of this profile at the Kanau-1 well.

The near-surface structure can be divided into three distinct sections, roughly separated along the profile by the Kanau-1 and Nembi-1 wells. From Bamu-2 to just south of Kanau-1 the profile runs through the foreland region, and the model shows a gradual thickening of the post- Jurassic succession northwards. The sharp increase in thickness approximately half way between the two wells is representative of the effects of the Komewu fault system (Bain et al., 1972).

In the region between Kanau-1 and Nembi-1 there are five small positive anomalies superimposed on the northwards-dipping regional gravity gradient. The first peak, over the crest on which the Kanau-1 well was drilled, corresponds to the Darai Plateau. This particular structure proved difficult to model in a way which was consistent with the geology, since although the model does show the Darai Fault dipping at -40° and extending to great depth (>20km), as suggested by Abers and MacCaffrey (1988), the gravity profile does not show the large positive Bouguer anomaly which would be expected from the known position of basement. There is no doubt that the highest observed Bouguer anomaly values have been reduced by the smoothing of the original data when producing the grid from which the profile data was sampled, but this is not sufficient to account for the inconsistency between the gravity field and the surface geology. The effects of the increase in station elevations over the plateau must also 163

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Csteoti ) (Ul) in da a 164

be considered, especially since the elevation values used in the modelling were produced using the same method as the gravity values (i.e. sampling gridded data), producing reduced peak elevation values. This would have the effect of increasing the computed anomaly values. Thus the combination of the lower observed Bouguer anomaly values and the lower station elevations giving rise to higher computed Bouguer anomlay values could well account for the differences between the model and the surface geology.

The second local gravity high on this part of the profile has been modelled as another thrust- faulted anticline, lying to the north of the Kikori River. This is the Hedinia anticline. There then follows a sequence of three basement involved thrusts in close succession. Hill (1989b) suggests there are more, but smaller thrusts, and also that this duplex in the Mesozoic succession extends at depth to the northeast, beyond the Nembi and Mendi Rivers. The gravity profile, however, suggests a thick (7-8km) sequence of relatively low density sediments to the northeast of Nembi - 1.

At deeper levels a lower crustal high is modelled beneath the Hedinia anticline, possibly thrust faulted; whether this represents the roots of the Darai Plateau or the Hedinia anticline is unclear. The Moho remains very close to its standard crustal depth of 35km for a considerable distance beneath the fold belt, with a significant downward displacement only occurring approximately half way between the Kanau and Nembi wells. In this model the downward displacement of the upper/lower crust interface is about 4km, and of the Moho about 7km, slightly more than is shown beneath the northeastern fold belt in other models.

Profile 6.1

This profile (Fig 6.19) runs in a WNW-ESE direction, parallel to the deformation front, through the Lake Murray region. All the gravity variations shown on the profile have been modelled as resulting from the density variations within the upper 5km of crust, and (as with Profile 4) the magnitudes of the anomalies shown on the gravity profile are small when compared with those on profiles running perpendicular to the mountain front.

20 bv 2k, 40 eo 3k; 40 bv 35km above 450 5 35km; below -450 20km, above 200 4

est cnrss km ) 1 40 -0 aoe e lvl; -7 ( aoe e lvl; 10 bv sa ee, 20 eo 20km; below -200 level, sea above 170 3 level); sea above (0 -170 2 level); sea above (-300 -470 1 3): (kgm contrasts Density

(nx) qadsa CsiBO®) DO ‘ OE 166

Of particular note in this model is the relatively steep eastern flank of the Lake Murray high, and the thin Pliocene-Recent cover which is neeeded to model the basement unit at the depth at which it was reached in the Lake Murray well.

Profile 6.2

This profile (Fig 6.20) runs in a WNW-ESE direction, to the south of Bamu-2. The amplitude of the anomalies in this region is less than 5mGal, suggesting very little deviation from the theoretical density distribution in this region. The overall decrease in gravity values to the east can be accounted for in terms of downwards displacement of either the base of the upper crust or the base of the Cretaceous sequences. The latter solution is shown in the model. The smaller scale fluctuations on this overall gradient can easily be attributed to variations in sediment thickness. These can either be in the Cretaceous-Miocene unit (as shown), or in the Pliocene- Recent unit, which is shown as being uniformly thin in this region. Very minor variations in the thickness of the Pliocene to Recent unit will have equivalent effects to those shown as deriving from the Cretaceous-Miocene units.

Profile 7

Profile 7 (Fig 6.21) extends from the northeastern foreland to the extreme northeast comer of the study area, crossing the Mananda anticline and the fold belt to the northeast. As with many of these profiles, the near-surface structures were initially based on the geological interpretation of Hill (1989b).

The sharp peak in the gravity profile close to its southern end (A in Fig 6.21) has been attributed to a quirk of the gridding algorithm. There are no gravity stations in this region, and the rather unlikely contour pattern shown on the gravity anomaly map (Fig 6.2) has been interpolated from data points up to 20km from the profile. A more realistic observed gravity profile which smoothed out the peak was estimated, and the model was created such that the calculated Bouguer anomaly matched the estimated values. Profile 6.2 SUW X.1TABJ0(SIUOW) ®) taa ttia T C®X)

4 200 above 20km, -450 below 35km; 5 450 above 35km 167 168

(SIEOTI) itQTABJO CTIX) T nd9(I 169

The southernmost major structure shown on the model is the Mananda anticline. Here the model bodies do not correspond exactly with the configuration proposed in the geological cross section (Hill, 1989b), and the deeper units appear more involved in the smaller structures (B in Fig 6.21) in front of the main Mananda anticline, such as the Libano anticline. Whether this indicates a decollment deeper than the upper parts of the Ieru Fm., or simply folding of the deeper units beneath these small anticlinal structures, is unclear. The geological cross section shows these structures as hanging wall anticlines, which is more compatible with a deeper decollment. However, the gentle folding of the units beneath the foreland shown in the gravity model implies that folding may well pre-date the faulting, so the structures may well be thrust faulted anticlines, in which case the latter solution is more sensible.

The configuration of the Mananda anticline on the model is very similar to the geological cross section, and shows the Jurassic succession rising to within a few hundred metres of sea level. To the northeast of Mananda the gravity profile shows three small peaks (Cl to C3) superimposed on the steep northeasterly regional gravity gradient. The first of these (Cl) has been modelled as a smaller anticline on the northeastern flank of the Mananda structure, and lies in a similar position to the Maro anticline (cf Bain et al., 1972). The second peak proved very difficult to reproduce in the calculated gravity profile as it is relatively small in magnitude, and while it was thought to be a result of another small thrust or fold, it could not be reproduced by the model. The third peak on this part of the profile, which lies to the southeast of the Doma Peaks, has been interpreted as representing a larger thrust/fold combination. Here the structure is quite deep (>5km below sea level), representing a significant thickening (structural and possibly stratigraphic) of the Cretaceous-Miocene unit.

Further to the northeast the gravity profile flattens out, and values even start to increase. The significant gravity high (D) in this region has been interpreted as resulting from a major structure affecting both the pre-Cretaceous and Cretaceous-Miocene units, which raises the Jurassic succession from ~8km to ~2km below sea level. This occurs in the Andebare River region, and is in common with similar structures modelled on other profiles.

At the northeastern end of this profile, where a significant increase in observed gravity is shown (E), the pre-Cretaceous unit is modelled as coming even closer to the surface, and this is consistent with the outcrops of the Chim Formation in this regioa The occurrence of high density rocks of oceanic origin further to the northeast, coupled with the increase in elevation 170

of the crystalline continental basement shown on the geological section, will have the observed effect on the gravity profile.

At deeper levels the downward displacement of the Moho to the northeast occurs initially to the southwest of the Mananda anticline, and at the northeastern end of the profile the base of the upper crust is -4km below its standard level, and the Moho ~7km below the standard level for continental crust. The geological cross section shows fault structures beneath the northeastern end of the profile extending to depths of about 25km. A model of this type was produced and found to be consistent with the observed gravity profile, though no better or worse in terms of the difference between calculated and observed Bouguer anomaly than the model shown in Fig

6 .21.

6.6 Overview of Regional Profiles

6.6.1 Foreland region

Three of the profiles described in the previous section cross the foreland region. These are Profiles 1.1, 2.1 and 5. All show a northwards thickening of the Pliocene to Recent unit which becomes more significant as the southern margin of the fold belt is approached. This thickening is much greater to the west of Mt Bosavi than to the east, in front of the Darai Plateau. There is a corresponding increase in the depth of the "basement" unit so the thickness of the Cretaceous to Miocene sedimentary sequence remains more or less constant.

All models show evidence of an east-west trending bulge in the lower crust, with the boundary with the upper crust rising to about 16km below sea level. The extent of the bulge is shown in Fig 6.22. In the west the bulge is up to 130km across, whereas to the east, close to the Darai Plateau, where it is much closer to the fold belt, the bulge less than 50km across. All profiles show the mantle beneath the foreland at a constant depth of 33-35km.

6.6.2 Fold belt region

All five SW-NE profiles (3,1.2/1.3,2.2,7 and 5) cross this region. Each profile shows one, and only one, large anticline which, with the exception of the Western Muller Anticline, is shown 171

Nembi-1

LakeMarriy-i]

Bi78000

Bi00000 700000 Fig 6.22 Extent of the lower crustal bulge (shaded), superimposed on the extended Bouguer anomaly map (Cl lOmGal). Profile locations are also shown 172 as being thrust faulted. To the northeast of each anticline is a sequence of two or three large thrust sheets affecting the pre-Cretaceous section. These become buried beneath stacked thrust sheets of the Cretaceous to Miocene succession in the northeast, the top of the pre-Cretaceous unit ultimately reaching a depth of about 10km. The four western profiles all show a major thrust fault at the northeastern end of the models. However, because the thrust is close to the end of the gravity profiles, the constraints on it are considerably less than for the other fold belt structures. The thrust fault is shown much further to the south on profile 7, and is not shown at all on profile 5.

Three of the profiles show structures to the southwest of the major anticlines. Profiles 1.2 and 2.2 both cross the Cecilia anticline. This anticline has very little effect on the deeper (pre- Cretaceous) units, being associated with thrusting from decollments within the Cretaceous succession. The gravity anomaly over the anticline results almost entirely from the folding and uplift of the Darai Limestone and the erosion of the Pliocene to Recent sediments from the crest. Profile 5 crosses the Darai Plateau, which lies above the hanging wall of a thrust fault which affect all levels of the sedimentary sequence and the crystalline basement, and possibly extends deep into the crust.

Profile: 3 1.2/1.3 2.2 7 5 Distances in km Mantle at max depth 70 85 45 85 85 to deformation front Mantle at max depth 50 55 25 85 65 to major anticline Mantle at 35km to 50 30 30 -15 45 deformation front Mantle at 35km to 30 0 10 -15 25 major anticline

Table 6.5 Distances from fold belt structures to deep crustal features.

At deeper levels, all the profiles show the eventual increase in depth of the lower crust and mantle beneath the fold belt. The position of the point at which this occurs relative to the major structures of the fold belt is shown in Table 6.5. In relation to the deformation front the distance at which the Moho dips below 35km decreases from ~50km in the west to ~30km in the Lavani/Hides region, and 15km near Mananda. There is then an increase to ~50km in the Darai 173

Plateau region, though if the Darai Plateau is not considered as a part of the fold belt, this is reduced to 25km . This overall decrease is due to the effects of the deep structure of the Western Muller Anticline, below which both the lower crust and mantle are tectonically elevated relative to neighbouring areas. The increase observed at the eastern end of the fold belt may also result from deeper structures; in this case the deep structure of the Darai Plateau involves elevation of the lower crust.

6.7 Detailed Modelling

The Bouguer anomaly data used in the regional profiles were sampled from the 1km grid of extended Bouguer anomaly values. However, in some areas, where gravity surveys had been carried out over structures of interest to the hydrocarbon exploration industry, much more detailed gravity information was available. The lines of such surveys often have a station spacing of only 100-200m, and are generally located perpendicular to the structural axes, thus providing profiles ideal for 2D modelling.

A program was written to project the station positions for a given survey onto a straight line from the first to the last point in the survey line, and this information was then used in the modelling program. Three detailed profiles were created: two over the Hides anticline, and one over the SE Hedinia anticline which extends northeastwards to the Nembi area. In all cases these profiles were close to regional gravity profiles, so the existing models could be used to constrain the deeper structure.

Because of the greater resolution available, the Cretaceous to Miocene unit used in the regional modelling was divided into two, essentially separating the Darai and Ieru Formations. Densities of 2500 and 2600kgm'3 respectively were used.

The location of the Hides profiles are shown in Fig 6.23. Again the geological cross sections of Hill (1989b) were used as the basis of the models. The results of modelling the gravity anomalies observed along Profile 1 are shown in Figs 6.24 and 6.25. The major point illustrated by this profile is the lack of effect of the density contrasts within the sedimentary sequences. The overall gradient of the observed gravity profile results almost totally from the deeper crustal structure, with the structural thickening of the sediments having some effect at the northeastern 174

i

9365000

• 140

Line 1 .me

Line 2 -130

Line 3 -120

-110 -90

-100 Line 4 \ -80

Line 6

-70

9315000 735000685000

Fig 6.23 Hand contoured extended Bouguer anomaly map of the Hides survey showing the location of the gravity profiles. toptk (*■) Onntr (All) #«Ftl (Bi) Onnlj (i OS DC S -O a o o t - -I.OCL > DC. 0 1 • • 0 0OD. 0 . . 00 E = Ieru Formation (Cretaceous), 2600kgm'3 (Cretaceous), Formation Ieru = E F = Pre-Cretaceous sediments and basement, 2670kgm' basement, and sediments Pre-Cretaceous = F B = Heavily karsitified Darai Fm, 2200kgm'3 Fm, Darai karsitified Heavily = B D = Darai Limestone Formation, 2500kgm'3 Formation, Limestone Darai = D 2200kgm*3 sediments, Recent to Pliocene = C 2200kgm'3 volcanics, Quaternary and Pliocene = A . o i.o . ▼ 1.00.X. • a Model 1ProfileHides ie Poie Mdl b Model 1 Profile Hides SW SW • • i .4 ie gaiypoie ,mdl a. model 1, profile gravity Hides 6.24 Fig 1 Fig 6.25 Hides gravity profile 1, model b. model 1, profile gravity Hides 6.25 Fig . 00 OlttMW (D) 10 . oo . NE NE 176 end of the profile. The majority of the remaining differences between the observed and calculated gravity values have very short wavelengths (~10km), and the amplitudes observed can only result from the effects of small, near surface bodies with high density contrasts, such as the recent sediments and volcanics of the Tari Basin (A in Fig 6.24) which are shown between 9 and 15km along Model a and were modelled using a density of 2200kgm‘3.

In the second model (Model b) along Hides Profile 1 a further body, representing karstified limestone with a density of 2200kgm'3, has been added (B in Fig 6.25). This drastically improves the overall match between observed and calculated gravity values: in some areas the correlation is near perfect, suggesting that such a near surface, low density body is highly likely. However this could not be used to explain all the observed Bouguer gravity anomalies shown along the profile, most notably the peak at the southwestern end of the observed gravity profile. It is certainly significant that this peak corresponds to a local topographic low, so even a flat body within the model would give the correct overall shape to the calculated gravity profile. However, to produce the observed anomaly amplitude the density contrast would need to be 400kgm'3, an unlikely value given the geology of the area.

A possible explanation of this is that the karstification of the Darai Limestone is much less severe in the valley between the Hides and Karius Ranges, and more severe on the exposed summits. Thus the bulk density of the Darai Formation would be higher (than the 2500kgm'3 used in the models) in the valley area, and lower in the exposed summit areas.

A second profile was produced in the Hides region - along lines 4 and 2 of the 1986 Komo survey. This profile (Profile 2) is shown in Figs 6.26 and 6.27. Again it was necessary to use small near surface bodies to reduce the difference between the observed and calculated gravity values; in this case a small body of density 2700kgm*3 representing non-karstified Darai Limestone was used to enhance the gravity peak at the southwestern end of the profile (Fig 6.27).

It is also clear that Profile 2 crosses the Hides anticline (h) at a point where it is greatly subdued both in terms of its topography and gravity anomaly. The structure immediately to the southwest of Hides, forming the Karius Range (k), is more dominant on Profile 2 than Profile 1. Whether the Karius anticline is part of the overall Hides structure (Model a, Fig 6.26), or a separate structure (Model b, Fig 6.27) is not clear from this profile alone, but considering the relationship D«ptk (la) OraTiti (aO«l*) D«ftk (la) Oranty (adali) OQ . 0 0 1 - OQ . 0 0 1 - 00, 0 .0 0 6 - ID. oo. -o. oa -s.oa 10. oa| -B.oa CO. 6 6 oo-6 f 00 0 -fi ie Poie Mdl b Model 2 Profile Hides E ▼ ie Poie Mdl a Model 2 Profile Hides __ SW SW

F = Pre-Cretaceous sediments and basement, 2670kgm'3 basement, and sediments Pre-Cretaceous = F 260Qkgm'3 (Cretaceous), Formation Ieru = E D = Darai Limestone Formation, 250Gkgm*3 Formation, Limestone Darai = D loeet eet eiet, 220Qkgm'3 sediments, Recent to Pliocene = C 2200kgm'3 Fm, Darai karsitified Heavily = B 2700kgm*3 Formation, Darai Non-karstified = A 00 1 Fig 6.27 Hides gravity profile 2, model b. model 2, profile gravity Hides 6.27 Fig i .6 ie gaiypoie2 oe a. model 2, profile gravity Hides 6.26 Fig 00 5 1 5 1

00 taaoa (la) a l ( a o a a it t D Ki) (K i e a a t i l S 35 00 NE 177 178 between the Hides and Karius anticlines shown by Profile 1, and also the geological interpretation of Hill (1989b), it is more likely that they are both part of the same large thrust- faulted anticline.

A gravity profile was also produced across the region of the SE Hedinia structure, continuing northward towards the Nembi region. This used data from the Kutubu-Orokana and Nembi gravity surveys (Fig 6.28). On this gravity profile (Fig 6.29) there are very large changes in the Bouguer anomaly values over short distances. Local gradients can be as great as 50mGal/km. These must result from karstification and terrain effects which are too localised to be corrected for in the gravity reduction process. Unfortunately these anomalies make modelling along this profile very difficult indeed as any anomalies resulting from the fold belt structures would be completely obscured by these effects. In this case better results were obtained from the smoothed data used for regional profile 5.

In summary, the detailed modelling over the Hides and other structures demonstrated the following points:

(i) in many cases the overall gravity gradient can be attributed exclusively to changes in the depth of the lower crust and mantle.

(ii) near surface bodies, be they the result of karstification (or lack of it) or recent sediments, have a considerable effect on the gravity profiles which can obscure effects due to the fold belt structures.

(iii) the effects on the calculated gravity profiles of structures in the sedimentary sequences are not always sufficient to separate them from the effects mentioned in (i) and (ii) above. Of the three profiles discussed in the preceding section, only one (Hides profile 2) provided any significant information on the structures affecting the Darai, Ieru and older sedimentary units. 179

9315000

Nembi-1

*o

■55

Kutubu-1 /

SE Hedinia field

9265000 760000 790000

Fig 6.28 Kutubu and Nembi gravity surveys - station locations and location of the Kutubu-Nembi gravity profile. Axes in UTM metres. Gravity C mGa1s 3 - - - -60. 120 100 - -40. 00 0 2 . . . . 50 50 1.0 50 35.00 25.00 15.00 5.00 -5.00 00 ' V TO^ “ , OQ oa DO DQ) ' .

□ = Observed, Observed, = □ ~ " i - , ■ . — — ■1 — i — i . , ■ ■ , ,- ■ - i i i"" ~ " r S NNE SSW Fig 6.29 Kutubu-Nembi gravity profile. gravity Kutubu-Nembi 6.29 Fig 0

Jl mJ *« itne (In) Distance '***" ^ 180 181 7. Geological Implications and Principal Conclusions

New Guinea is the active northern margin of Australia, and has a history of rifting and terrane disaggregation and, more recently of terrane accretion. As part of this latter series of events, a fold and thrust belt has developed, a foreland basin has formed, and volcanic activity has been widespread.

There have been a number of studies of this area which have aimed at an understanding of the fundamental processes involved. St John (1967) was the first to assemble the data which allowed some estimate of the nature of the isostatic compensation of the topographic mass to be made. Abers and McCaffrey (1988) considered the information on deep structure provided by natural seismic activity, and Abers and Lyon-Caen (1990) have modelled the isostatic response along two selected profiles, demonstrating the existence of fundamental differences between the eastern and western ends of the fold belt. Many authors have analysed the surface geology, and most recently Hill (1989a, 1989b, 1991) has used this information to produce balanced cross sections.

This study has had the benefit of additional gravity data and more detailed processing and has focused both on the surface structures, including the accumulation of hydrocarbons, the transition between thin and thick skinned tectonics and the structure of the foreland, and the different types of isostatic response to the topographic loading in order to constrain the timing of events in the evolution of the fold belt.

7.1 Structures affecting the sedimentary cover in the fold belt

The forward modelling has demonstrated that the cross-sections presented by Hill (1989b, 1991) are largely compatible with the measured gravity field, although some additional complications and concealed structures are suggested. The detailed gravity profiles presented in Chapter 6, in conjunction with the regional models, show the position and extent of both the large fold belt anticlines and many of the smaller structures. Of the larger structures, those which also involve crystalline basement and the deeper levels of the crust, namely the Muller Anticline and the Darai Plateau, are discussed separately in Sections 7.1.2 and 7.1.3; the remainder are discussed in Section 7.1.1 below. 182

7.1.1 Shallow structures

The residual and short wavelength filtered maps both show gravity highs along the southwestern margin of the fold belt (structural Zone 1, Fig 2.3) in which the large basement anticlines are known to lie. Though there is little evidence from the gravity contours of further involvement of the deeper strata in any of the other fold belt structures, the NE-SW gravity profiles do show subtle peaks superimposed on the regional gradient in a number of places, generally immediately to the northeast of the major anticlines. These have been inteipreted as being the effects of thrust faults involving the pre-Cretaceous succession, many of which had been predicted from the geology. In terms of the structural zones identified in Chapter 2, these features lie at the northeastern edge of Zone 1 (the Muller-Kutubu zone), which contains the major surface anticlines, and beneath the imbricate stacks of Miocene limestones of the Oksapmin-Tari-Mt Murray zone (Zone 2).

The original models derived from the cross sections of Hill (1989b) along gravity profiles 1.2/1.3 required alteration, and a large thrust fault and corresponding hanging-wall anticline to the northeast of the Lavani anticline is proposed in the final models presented in Chapter 6. No such feature is shown on Hill’s (1989b) geological cross sections, but because it is buried beneath a stack of thrust sheets of Miocene limestones it will not be obvious from the surface geology.

The gravity profiles over the Juha anticline suggests that the structural inteipretation with only northeast-dipping thrust faults (Hill (1989b), Interpretation B) is more likely than that involving southwest-dipping backthrusts on the southwestern flank of the Lavani anticline (Interpretation A). Hill (1991) also opts in favour of Interpretation B. Further to the northeast, beneath the sinuous complex folds of the Porgera-Nipa-Poroma structural zone (Zone 3, Section 2.3.3) the existence of a large structure involving the pre-Cretaceous section has also been demonstrated. This thrust-faulted anticline is shown at the northeastern end of many of the profiles and extends to the northeast of the Doma Peaks, forming the deep core of the Andebare and Wage anticlines. It affects the basement unit of the models, and almost certainly involves crystalline basement as well. The involvement of the deeper levels in structures this far to the northeast of the deformation front is supported by the seismological evidence presented by Abers and McCaffrey (1988) (their events 2 and 3, Fig 12). These thrust faults also run parallel to the axis of the Mesozoic Papuan Basin (Jenkins, 1974), and may well have been formed by the reactivation of Mesozoic rift related faults during the Pliocene. 183

7.1.2 The Muller Anticline

Of all the surface anticlines of the fold belt only the Muller Anticline has its roots deep in the lower levels of the crust. From the profiles alone it is difficult to determine the extent of the anticline at depth. However, the gravity anomaly maps which have been/filtered to show the longer wavelengths of the anomaly field indicate that although the deep anticlinal structure is closer to the surface beneath the western part of the anticline, it can be traced plunging to the southeast beneath the Juha region. Doubtless the slight elevation of the lower crust shown on Profile 1.2/1.3 close to Juha also represents this deep anticline. This is in contrast to the axial trace of the surface anticline which is shown on the extended Bouguer anomaly map, the short wavelength filtered map and the surface geological map, as extending southeastwards to the Lavani region, over 25km to the northeast of the deeper structure (Fig 7.1).

Hill (1989a) noted many differences in the near surface structure of the eastern and western parts of the Muller anticline, and proposed the names Eastern Muller Anticline (EMA) and Western Muller Anticline (WMA) to distinguish between the two. It is very likely that these differences in the surface geology reflect deeper crustal differences. For example, Hill (1989a) notes that the dips in the northeastern limb of the EMA are much more gentle than those in the corresponding limb of WMA. This was used as evidence to support the idea that the EMA and WMA are separate structures separated by a major tear fault, and while this is not in conflict with the ideas presented here, the more gentle dips can also be explained because the EMA lies some distance to the northeast of the deep crustal anticline, away from the steeper dips associated with the crest which are seen in the WMA. So while there are clear differences between the eastern and western parts of the anticline as a whole, the terms EMA and WMA imply that they are different parts of the same anticline, which is not supported by the evidence from the gravity field of the deeper structure.

On the 1972 1:1000000 Geological Map of Papua New Guinea (Bain et al., 1972) the term Muller Anticline is restricted to the surface anticline extending east from the Strickland River to the Lavani Valley (ie Hill’s EMA), which is perhaps a more sensible nomenclature. The WMA would then require a different name. The name Blucher Anticline is here proposed to apply to the surface anticline west of the Strickland River, where it coincides with the deep 184

142 E 5 S 143 E

Oksapmin

Quaternary alluvial clastn

I Quaternary Volcanics j Fold axial trace J Pliocene classics Cecilia ] Miocene to Eocene limestone — Thrust Fault

□ Cretaceous mudstone and sandstone * Town Jurassic elastics *

Fig 7.1 Simplified geological map of the Muller Anticline region. 185 crustal high. Therefore it would be logical to name the entire deep structure, both beneath the Blucher Anticline and further to the southeast beneath the Juha and Wai-Asi anticlines, the Blucher crustal high.

7.13 The Darai Plateau

The Darai Plateau also has its roots deep in the crust, though this fact is much less obvious on the gravity maps because the deep structure lies to the northeast of the Plateau. However, the existence of the deeper roots beneath the Hedinia region is shown by the modelling along Profile 5, in which the southern boundary fault of the Darai Plateau extends to a depth in excess of 20km. Although this is not the only possible model for this structure, it does support the idea proposed by Abers and McCaffrey (1988) on the basis of seismological evidence that the Darai Plateau is one of a sequence of deep-rooted features, similar to the Blucher crustal high, which form the southern edge of the fold belt, and which Abers and McCaffrey (1988) demonstrated to exist further to the north beneath the so-called "thin-skinned" zone where thrust faulting of the Miocene and more recent sequences is dominant

Another important feature of the Darai Plateau is that it is not in isostatic equilibrium. Therefore either the strength of the crust is sufficient to support the extra topographic load or the plateau was formed so recently that compensation has not yet occurred. Since there is evidence of a break in the crust resulting from the compressional forces forming the fold belt, with the Darai Fault extending to depths in excess of 20km, it is unlikely that the crust has enough strength to support the topographic load. In addition, the table of distances between the point at which the mantle is at a depth of 35km and the deformation front (Table 6.5) shows Profile 5 to be anomalous in that there is a general overall reduction in this distance from west to east along the fold belt on all other profiles. If, however, the Darai Plateau is ignored and the Hedinia anticline is taken as the deformation front, Profile 5 fits the overall trend much more closely. Furthermore, the topographic elevations immediately to the northeast of the Darai Plateau, including the Hedinia Anticline, are relatively low. Consequently the topographic loading of the crust is also low when compared with adjacent profiles, and compensation of this part of the fold belt would not be expected. This suggests that at its eastern end, the entire fold belt (excluding the Darai Plateau), shows an isostatic response to the topographic loading (see also Fig 7.2), and therefore that the Darai Plateau is a recent structure with compensation yet to occur. 186

There is also geological evidence for the uplift of the Darai Plateau having occurred relatively recently. At its western end is the Mt Bosavi volcanic centre, and to the north, above the point at which the Darai Fault would reach the mantle, are a number of volcanoes including Mt Murray and Mt Duau, which are all areas of Quaternary volcanic activity. Whether these volcanoes formed as a result of the tectonic activity which uplifted the Darai Plateau, or pre-date the Darai uplift and simply provided a line of weakness along which displacement occurred, is unclear. However, both possibilities suggest that the Darai Plateau is a feature which foimed within the Quaternary period (ie. the last 2 million years). Given this time frame, the plateau must have been uplifted in response to the collision of the New Britain Arc with the continental margin to the north. This collision has provided the additional convergent stress required to reactivate an existing normal fault which lay to the southwest of the Pliocene deformation front as a reverse fault (the Darai Fault) which now forms the southwestern boundary of the plateau.

7.1.4 Significance of the Hydrocarbon Accumulations

Estimation of the timing of the maturation, migration and accumulation of the hydrocarbon reserves of the fold belt can be used as evidence to support the timing and sequence of the formation of the fold belt structures. The interpreted sequence and timing of the fold belt deformation events (Section 7.3) supports the theory that many of the fold belt structures formed during the Pliocene, and that the Oligocene-Miocene accretion of terranes which now lie within the Mobile Belt and beneath the Sepik plains affected only the northern parts of the continental margin, producing the lithospheric deflection which foimed a foreland basin. This depression would have been greater in the western fold belt region, where continental margin was narrower, and the lithosphere stronger (see Section 7.3). The Late Palaeogene foreland basin became a region of extensive sedimentation and up to 1500m of Late Oligocene to Miocene sediments (mostly carbonates), and in some areas up to 1000m of Pliocene clastic sediment, were accumulated in what is now the eastern fold belt region. Thus, before much of the deformation in the fold belt region had occurred, the source of the hydrocarbon accumulations (the Upper Jurassic Imburu Fm) could have been at a depth of 4-5km in the depressed regions, and with the increased heat flows associated with the orogeny to the north, maturation and migration of fluid hydrocarbons would have commenced before many of the fold belt structures had formed. This being the case, the existence of significant hydrocarbon accumulations within the fold belt requires further explanation. 187

The large anticlines forming the southwestern edge of the fold belt formed very early during the deformation phase (Hill, 1989b), and therefore may have existed when much of the generation of fluid hydrocarbons took place. A further explanation of the fold belt hydrocarbon accumulations is that the trapped hydrocarbons represent the later parts of the maturation and migration; this is supported by the predominance of gas accumulations, particularly in the central fold belt (eg Juha, Hides) region, which would have been generated after most of the liquid fractions.

However, a more likely explanation is that these large anticlines were formed as hanging wall anticlines when existing extensional faults were reactivated as thrust faults, as suggested by Hill (1989a) for the Blucher/Juha anticlines. These may have formed initially during the Mesozoic rifting of the Australian continental margin, or perhaps more recently during the development of the foreland basin, and therefore traps could have existed when generation and migration of the hydrocarbons occurred. During the subsequent Pliocene deformation of the fold belt, when the sense of motion along these fault planes was reversed and the large hanging-wall anticlines were formed, remigration of the hydrocarbons accumulations took place. Providing all the migration paths out of the trap remained sealed during the fault reactivation process the hydrocarbon accumulations would remain within the reservoir as deformation occurred.

7.13 Thin v thick skinned tectonics

The existence of large anticlines beneath the complex duplexes (Zones 2 and 3) casts doubt on the applicability of the thin skinned tectonic model to the Papuan fold belt (cf. Hobson, 1986). While there is no doubt that the deformation of the Miocene limestones of the Darai duplex (Hill, 1991) is consistent with thin skinned deformation, the interpretation presented above (and in Section 7.3), in conjunction with the seismological evidence (Abers and McCaffrey, 1988) and geological interpretations (Hill, 1991), confirms the existence of large, basement-cored anticlines, both at the surface (Zones 1 and 5) and beneath the duplex of Zones 2 and 3, which have their roots deep in the crust. %

Evidence presented above and in Section 7.3 confirms that these thrust faults and associated anticlines have formed as a result of the Neogene reactivation of normal faults which formed initially during either the Mesozoic rifting or the Palaeogene loading of the continental margin. So although many of the surface structures have been formed by thin skinned processes, they 188 conceal larger structures resulting from thick skinned inversion tectonism. The application of the thin skinned model to the whole of the Papuan fold belt cannot be justified.

7.2 The Foreland Basin

Both the gravity anomaly maps and the models along the gravity profiles have revealed very interesting features in the foreland region. The models show very few changes in the depth of the base of the crust, though a significant elevation of the mid-crustal density contrast (i.e. the base of the upper crust) was described in the previous chapter. This elevation was necessary to account for the Bouguer anomaly profile which shows a long wavelength high in this region. The observed gravity high could not be reproduced by varying the mantle topography as this interface is too deep to only affect the calculated gravity values of the foreland region and not the fold belt as well. The elevation of the mid-crustal density contrast has the effect of producing a thinned upper crustal section, which appears to be even more thinned in places where there is also localised thickening of the Pliocene to Recent succession.

The apparent upbowing of the lower crust could be a result of the loading of the continental margin to the north. As has been discussed in Section 2.4, the initial accretion of terranes and consequent loading of the Australian continental margin resulted in the development of a foreland basin and an associated forebulge. It is possible that parts of the upper crust were eroded as a result of this upbowing, and that the bulge in the lower crust is a deep remnant of this forebulge.

It is, however, the near surface structure of the foreland which is more interesting. Both the NW- SE and SW-NE profiles show significant variations in the depth of the basement and the depth and thickness of the Cretaceous to Miocene and Pliocene to Recent units. Profile 4 (Fig 6.19) shows very sudden changes in the level of the basement units which are the result of either strike-slip or normal faults running approximately in a NE-SW direction. The existence of such a structural trend is supported by the trend filtered gravity map (Fig 5.15) which reduces the effects of the NW-SE trending features. This shows a significant NW-SE gravity gradient (i.e. NE-SW trending contours) bounding a line of gravity lows extending northeastwards from the south of Lake Murray to the Tari and Nipa areas of the fold belt. This trend is coincident with the line of volcanic centres including the Doma Peaks, Mts Kerewa and Sisa, Mt Bosavi and the 189 numerous associated volcanic centres further to the southwest. There are two possible explanations of this relationship: firstly the relatively low density recent volcanigenic sediments and lavas associated with these centres could produce the observed gravity anomalies, or secondly this could be further evidence of the existence of a major NE-SW fault extending deep into the crust, which crosses the western end of the Darai Plateau and extends into the foreland to the southwest, producing a trough in which low density molasse sediments derived from the highlands have been deposited.

There are two areas in which relatively thick sequences of Pliocene to Recent sediments have accumulated in the foreland. The most northerly of the two lies to the north of the bulge in the lower crust (Fig 6.22), and immediately to the south of the deformation front. The thickness of sediments in this area increases westwards along the southern boundary of the fold belt. The second accumulation lies over the crest of the bulge in the lower crust to the north and northeast of the Lake Murray basement high.

The occurrence of thick Pliocene to Recent sediments close to the deformation front is not surprising, as large quantities of sediment will have been and still are being derived from the highlands to the north and deposited in a molasse basin which has developed as a result of the loading of the crust by the topographic mass. It is less easy to explain the accumulations of Pliocene to Recent sediments further to the south. One possibility is that a small basin could have formed a result of extensional movements in response to the additional crustal loading to the north. Another explanation is that as the forebulge developed in response to the terrane accretion events, the crest of the bulge was eroded, and with further recent accretionary events and the southwards movement of the deformation front, the modem forebulge now lies much further to the south and the old eroded crest now lies in a subsiding region which has filled with Recent sediments.

7.3 Regional Tectonics and Isostasy

Abers and Lyon-Caen (1990) presented a picture of the isostatic compensation of the New Guinea Highlands based on the two end members of a complicated transition between two types of isostatic response. Within the transition region there is a divergence of surface and subsurface trends, and mechanisms for the support of crustal loads vary considerably over short distances. 190

Though both the gravity maps and the surface geological map show sinuous trends within the fold belt, the areas in which these occur are different. Both show approximately E-W trends in the west of the fold belt, and WNW-ESE trends at the eastern end, separated by a region in which the dominant trends are in a NNW-SSE direction. On the surface geological map this latter region lies to the east of the Doma Peaks volcanic centre, whereas the gravity map shows the same trend occurring to the west of the Doma Peaks, close to the Lavani area. The implication of this is that there is a significant difference in this central fold belt region between the surface geological structures shown by the geological map, and the deeper crustal structure illustrated by the extended Bouguer anomaly and longer wavelength gravity anomaly maps.

The eastern end of the Blucher crustal high strongly affects the gravity field of the central fold belt region, and appears to affect the surface structure further to the northeast, with the surface structures apparently piled up against the deeper crustal high; this indicates that the Blucher crustal high pre-dates much of the fold belt deformation. However it is the isostatic compensation of the highlands which has the greater effect on the gravity field of the central fold belt region, and it is in this central region that the greatest reduction occurs in the distance of the point at which the mantle reaches a depth of 35km from the deformation front

The free air gravity map (Fig 7.2a) was produced by interpolating a 1km grid of values from the raw data, and contouring the grid with a high degree of smoothing. Figure 7.2b, which is an interpretation of the free-air gravity map, shows those areas in which there is no isostatic response to the crustal load. Even with point for point isostatic compensation the southern edge of the fold belt would be shown as a free air gravity high because the effects of the deeper compensating root have a substantially longer wavelength than the topographic anomaly. However, the lack of a consequent and significant free air gravity low to the southwest of the mountain front, where the gravity effect of the root would exceed that of the topography, and also the results of the gravity modeling presented in Chapter 6 confirm the elastic (i.e. non­ isostatic) response of the crust beneath the southern margin of the fold belt (Fig 7.2b).

The free air anomaly map apparently shows that neither of the areas exhibiting some form of isostatic response - the northeastern comer and the southwestern part of the map - appear to be in total isostatic equilibrium; both show positive free air gravity values. However, a positive free air anomaly of about 20mGal would be expected because of the relatively large (+75m) difference between the geoid and the ellipsoid in this region (Marsh et al., 1988). Thus the free Fig 7.2a Free air anomaly map.

to mouths of the 9400000 Ramu & Sepik Rivers

Kandej

tevaJa-1 Nemtx 9300000 ■

9200000-

Region of no isostatic response to crustal loading O Volcanic Centre 0 km 100 — — • Major anticline 9100000 i i ■ ■ v * 500000 600000 700000 800000 Fig 7.2b Regions showing no isostatic response to crustal loading - information from free-air anomaly map (Fig 7.2a) and gravity models (Chapter 6). 192 air anomaly of the southwestern part of the map is consistent with isostatic equilibrium. The negative free air values to the south of the Blucher Anticline can be explained in terms of the deflection of the lithosphere as a result of the crustal loading to the north.

Fig 7.3 shows the regional lithospheric structure based on the gravity models presented in Chapter 6. The infinite slab approximation can be used to calculate the gravity anomaly due to each body in the model and to determine the calculated free-air gravity and degree of isostatic response to the crustal loading. For the northeastern end of the model (Fig 7.3), the calculated free air anomaly is +3mGal, a result which is consistent with complete isostatic equilibrium (Airy-type compensation).

The most striking feature of the uncompensated zone (Fig 7.2b) is the offset between the Mananda anticline and the Darai Plateau. If the trend of the offset is continued to the north, it crosses the north coast of New Guinea close to the mouths of the Ramu and Sepik rivers, at the boundary of the North New Guinea and New Britain provinces identified in Chapter 1. A possible implication of this is that the feature identified affecting the western end of the Darai Plateau, and delineated in this region by the volcanic centres of Mt Bosavi and the Doma Peaks, is related to the accretion of part of the New Britain Arc to the northern continental margin. This collision event, which only occurred to the east of the Sepik, and which is the most recent accretionary event in New Guinea, has therefore also produced the recent uplift of the Darai Plateau (Section 7.1.3) and the intense deformation observed in the eastern fold belt where the amount of crustal shortening is much greater than in the west (Abers and McCaffrey, 1988).

This is consistent with the conclusion from the central fold belt region where the difference in trends of the gravity and surface geological features is interpreted as evidence that the deeper structures pre-date the surface structures, and gives added weight to the idea that the large surface anticlines of Zone 1 were formed by reactivation of older, extensional features at an early stage in the evolution of the fold belt.

Additionally, the eastward thinning of Zone 1 indicates that the large anticlines within it were produced by an earlier event which had greater effect in the west - the accretion of the North New Guinea Arc. Both Hill (1989b) and Pigram (in press) date the North New Guinea arc accretion as Late Miocene (6-1 IMa), and Hill suggests that because of its frontal position, the "Muller Anticline" (ie the proposed Blucher Anticline) was the last-formed structure of the 193

SW NE

---- 2 Topography (2500) --- 0 Sedimants (2500)

U Crust (2670) ------8

- -20 U Crust (2670) - -25 L Crust (2870)

— 35 L Crust (2870) --41 Mantle (3320)

Fig 7.3 Typical regional density structure of the eastern Papuan fold belt (densities in kgm'3, depths in km)

142 143 144 145

Turbidites Oksapmin and Volcanics

Ja ri Mount Hagen Marine daslics

Cecilia-1 Mendi

■15CD0

Fluviatile, deltaic anc marginal marine sediments

142 15000

km 100

Isopach interval 2500tt.

Fig 7.4 Mesozoic isopach map of the Papuan fold belt (not showing recent erosion effects), from Jenkins (1974). 194 western fold belt (at 4±0.5Ma from fission track data). This appears to conflict with the suggestion that the large anticlines have formed a barrier to the more recent deformation in the region to the northeast (Zones 2 to 4), but since these zones are not present in the western fold belt, this conflict does not arise. It follows that in the eastern fold belt the large surface anticlines of Zone 1, including the Mananda and Hedinia anticlines, were also formed by the accretion of the North New Guinea Arc. If, as argued in the previous sections, much of the fold belt deformation occurred during the Pliocene, the formation of the structures in the northeast of the fold belt (Zones 2 to 4), and the Darai Plateau to the south, must have been due to the subsequent accretion of the New Britain Arc.

The degree of isostatic compensation can be used along with the depth of the foreland basin to constrain the strength of the lithosphere beneath the fold belt. Abers and Lyon-Caen (1990) suggested that the lack of a compensating root beneath the western fold belt indicates that the lithosphere is much stronger than in the east where the highlands are in isostatic equilibrium. Gravity modelling and drilling have also shown that the Pliocene to Recent succession is much thicker in the west than in the east, supporting the idea that the Australian plate is much weaker beneath the eastern part of the fold belt. Abers and Lyon-Caen (1990) presented evidence that there is a dramatic change from west to east along the fold belt. They suggested that beneath the eastern fold belt the Australian plate must be weak everywhere and therefore unable to support the topographic load (they determined a flexural rigidity IKlO^Nm), or else it extends only part way beneath the highlands before becoming weak or broken. In contrast a strong elastic plate (D=1023-1025Nm) is inferred beneath the western fold belt. However, this ignores the substantially greater width of the fold belt in the east, and the present more detailed analysis suggests that there is little variation in the width of the elastically supported zone along the fold belt, and only the northeastern part of the fold belt, which has no direct equivalent in the west, is supported isostatically.

Abers and Lyon-Caen (1990) also noted that their computed deflections for the foreland basin suggest that the load in the west is less than the topographic weight of the mountains. However they assumed a density for the topography of 2700kgm'\ whereas 2500kgm‘3 is a more realistic value (perhaps 2600kgm*3 in the far west). They also used the fact that there are no obvious large structural boundaries between 141° and 143°E beneath the foreland to indicate that a west to east change in strength along the fold belt is unlikely. Although the filtering of the gravity data (Chapter 5) and the alignment of the volcanic centres (Section 7.1.3) suggests the existence 195 of a crustal weakness in this region, it has been clearly demonstrated above (Fig 7.2b) that the transition between the two types of isostatic response does not occur from east to west along the fold belt, but from WSW to ENE across it, and therefore the identification of this weakness is of little significance in terms of this transition.

The reason for the observed change in the mechanisms of support of the topographic load is less clear. Fig 7.2b, along with the models presented in Chapter 6, show the region in which the highlands become compensated, and this would be the likely location of any break in the lithosphere which might be propsed to account for the change in isostatic response. Southwest to northeast lithospheric weakening must also occur in the same region. The gravity models show that the topographic load is in isostatic equilibrium within 20km (horizontal distance) of the point at which the base of the crust reaches a depth of 35km (Table 6.5) and this is compatible with either a break or a rapid weakening of the lithosphere. The collision of the New Britain Arc had a major effect on the eastern fold belt, and this may have produced the additional loading necessary to cause a break in the lithosphere, but the accretion of the North New Guinea Arc should have had a similar effect on the western fold belt. It may be, however, that the units of the New Britain Arc lie on top of units of the North New Guinea arc in the east, thus constituting an additional extra load on the continental margin.

It is possible that a weakness existed in the lithosphere beneath the eastern fold belt before the collision of either arc, and that this has been reactivated following the loading of the continental margin. The rifting of the northern Australian continental margin during the Mesozoic would have produced weaker areas of lithosphere which have subsided and been filled with thick sequences of Mesozoic sediments. The Mesozoic isopach map (Fig 7.4) shows a thick (22500ft, 6500m) accumulation of Mesozoic open marine fine clastic sediment in a NNW-SSE trending basin (the Southern Highlands basin) to the northeast of Mendi, bounded by the platform to the southwest and by the Kubor high to the northeast, both areas with Mesozoic sequences less than 7500ft (2500m) in thickness. At the western end of the fold belt no such basin exists because there is no northern boundary (Fig 7.5); only turbidites and volcanics are found to the north. This suggests that in the west the rifting which produced the Southern Highlands basin resulted in the detachment of the northern part of the Australian continental margin, which was then carried away as a terrane, presumably into the eastern Tethyan region (cf Pigram, in press). In contrast, in the east the rifting ceased before the Kubor region was detached from the continental margin, leaving the Kubor High separated from the continental shelf by a deep basin (Fig 7.5). 196

Thus when the continental margin was loaded during the Pliocene there was already a substantial area of weak lithosphere in the east, whereas in the west much of this weak area had been detached. The Kubor High was also uplifted during the Pliocene convergent tectonic phase, and now forming the Kubor anticline, which is an additional crustal load in the eastern fold belt.

Fig 7.5 shows the Mesozoic crustal structure of the Mesozoic Papuan Basin (Jenkins, 1974), with the corresponding structural zones of the Papuan fold belt. The major anticlines (Zone 1, Fig 2.3) were formed immediately to the southwest of the southern edge of the Mesozoic rift zone, and the Cenozoic compression of the region has led to the reactivation of the Mesozoic normal faults as reverse faults and the thrusting of the basement in a southwesterly direction over the more stable continental shelf. The zones of more intense deformation (Zones 2 to 4) lie in the region which is shown by the gravity models and the isopach map (Fig 7.4) to have the greatest thickening of the Mesozoic succession. This is shown by a series of normal faults in Fig 7.5, and it is proposed that these structural weaknesses have had a major impact on the distribution of the Pliocene to Recent structures in this part of the fold belt. The existence of deep faults has also been proposed using evidence from natural seismic activity by Abers and McCaffrey (1988).

This deep faulting can also be used to account for the formation of the Quaternary volcanic centres of the northeastern part of the fold belt (Mts Giluwe, Ialibu etc, Fig 7.6) which lie along the axis of the Mesozoic depocentre; as the eastern fold belt settled into isostatic equilibrium, the weak, rifted crust sank deeper, was heated and eventually melted, rising to the surface along the pre-existing lines of weakness - the normal faults formed during the rifting process. A similar explanation can also be applied to the Doma-Sisa-Bosavi volcanic trend (Fig 7.6), since the magma created as a result of the crustal subsidence would also be able to rise to the surface along the NE-SW trending crustal weakness proposed earlier in this section, and in Sections 7.1.3 and 7.2.

In summary, the distribution of surface structures within the fold belt, deeper structures beneath it, and the isostatic response to the additional crustal loading which they have created, is strongly influenced by the distribution of Mesozoic margin parallel rift related faulting, and to a lesser extent by NE-SW cross features. Certain lineaments have been reactivated by the collisions of the North New Guinea and New Britain Arcs, and control the relative significance of thick and thin skinned tectonism and the distribution of Quaternary volcanism within the fold belt. 197

SW Papuan Basin------NE

Fluviatile, deltaic and marginal manna sadimenis

Turbidiles and volcanics

Continental Margin

T " T fo ic B e " 7 ' " ' " Foreland Basin 1 Blucher Crustal High 1 i Mobile Belt i ^ i (Zone 2| j Compensated [ Not compensated J Compensated?

Western Fold Belt

SW Papuan Basin NE Fluviatile, deltaic and marginal manne sediments

Marine elastics

Turtndites and volcanics

Southern Highlands Basin Oon6 nemai Margin

— r- " — r ------1— ------1------i Mapr Anticlines i Fold Belt Quaternary i i Foreland Basin Kubor Anticline ' Mobile Belt [ (zone 1) | (Zones 2 to 4) Volcanic Centres , I " 1 i 1 Compensated i Not C om pensated i Compensated Partially com pensated i Compensated?

Eastern Fold Belt

Fig 7.5 Diagram showing the Mesozoic crustal structure of the Papuan Basin and its relationship to the Pliocene to Recent fold belt deformation. Not to scale. 198

143 144 ■ Towns O Volcanic centres . ^ Volcanic trend Mobile Belt km 100

145

O Hagen Tari |D°m§ Peaks /G ilCtwj, Kubor Anticline*

■ Nipa ■ lalibu

Suaru

P b te r

Darai Plateau Duau O r

Foreland Basin

143 144 145

Fig 7.6 Major volcanic trends of the Papuan fold belt. 199 2 0 0

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Appendix A

- Gravity meter G-513 calibration table - Gravity base station descriptions 208

Appendix A

MILLIGAL VALUES FOR LAGOS^E & ROMRERC, INC. MODEL G GRAVITY METER ''G- 513

COUNTER VALUE IN FACTOR FOR COUNTER . VALUE IN FACTOR FOR READING* MILLICALS INTERVAL READING* MILLIGALS INTERVAL

000 000.00 1.02118 3600 3679.68 1.02422 100 102.12 1.02106 3700 3782.10 1.02435 200 204.22 1.02099 3800 3384.54 1.02446 3 00 306.32 1.02096 3900 3986.98 1.02458 400 403.42 1.02095 4000 4089.44 1.02470 500 510.51 1.02097 4100 4191.91 1.02481 600 612.61 1.02100 4200 4294.39 1.02492 700 ' 714.71 1.02104 4300 4396.88 1.02502 800 816.82 1.02109 4400 4499.38 1.02512 900 918.92 1.02116 4500 4601.90 1.02522 1000 1021.04 1.02124 4600 4704.4 2 1.02531 1100 1123.16 1.02131 4 700 4806.95 1.02540 1200 1225.30 1.02140 4800 • 4909.49 1.02-54 5 1300 1327.44 1.02]48 4900 5012.03 1 .02 5.52 14 00 14 29.53 1.02159 5000 5114.59 1.02557 1300 1331.74 1.02]67 5100 5717 .14 1.02 561 1600 1633.01 1.02177 5200 5319.70 1.02564 1700 1736,09 • 1.02186 5300 54 22.27 1.02566 1800 1839.27 1.02197 5400 5524.83 1.02565 1900 1940.47 1.02208 5500 5627.40 1.02563 2000 2042.68 1.02219 5600 57 29.96 1.02561 2100 214 4.90 1.02230 5700 5832.52 1.02558 2200 2247.13 1.02242 5800 5935.08 1.02554 2300 ’ 2349.37 1.02254 5900 6037.64 : 1.02549 2400 2451.62 1.02266 6000 6140.18 1.02543 2500 2553.89 1.02279 6100 6242.73 1.02535 2600 2656.17 •1.02292 6200. 6345.26 1.02525 2700 2758.46 1.02305 6300 6447.79 .1.02512 2800 2360.76 1.02318 6400 6550.30 1.02499 2900 2963.08 1.02333 6500 6652.80 1.02434 3000 3065.41 1.02346 6600 6755.28 1.02468 3100 3167.76 1.02359 6700 6357.75 1.02450 3200 3270.12 1.02371 6800 6960.20 1.02430 3300 3372.49 1.02334 6900 7062.63 1.02404 3400 3474.87 1.02395 7000 7165.03 3500 3577.27 1.02409

* Note: Ri^ht-hnnd vhcel on counter indicates approximately 0.1 roilligal. 209 COUNTRY NEAREST CITY Papua New Guinea Margarima GRAVITY STATION DESCRIPTION STATE Southern STATION NAME STATION NO PROVINCE Tiengo COUNTY Highlands TNGO LATITUDE LONGITUDE ELEVATION 5° 57’ 10.05" S 143° 15’ 32.76" E 2420m

POSITION CONTROL GRAVITY VALUE Reference to local topographic maps 977459.15mgal

ELEVATION CONTROL BOUGUER ANOMALY Barometric levelling

DESCRIPTION

At W end of lay-by on S side of track, between the centre two of four large polythene huts.

DESCRIBED BY: DATE David Harrison 30-11-89

DIAGRAM

DIAGRAM BY: DATE David Harrison 30-11-89

STATION HISTORY

Established November 1989

SOURCE ORGANISATION Dept Geological Sciences, University College London 210

COUNTRY NEAREST CITY Papua New Guinea Tari GRAVITY STATION DESCRIPTION

s t a t e Southern s t a t io n n a m e s t a t io n n o . p r o v i n c e YY , . . Leru c o u n t y Highlands LERU LATITUDF LONGITUDE ELEVATION 5° 53’ 14.71" S 142° 50’ 49.74" E 1731m POSITION CONTROL GRAVITY VALUE Reference to local topographic maps 977632.17mgal

ELEVATION CONTROL BOUGUER ANOMALY Barometric levelling

DESCRIPTION

At SE comer of concrete foundation block adjacent to the northern wall of the mess hut at Leru camp. The camp is located approximately 100m from the main Tari to Komo road, along a rough track.

DESCRIBED BY: DATE David Harrison 03-12-89

DIAGRAM

C o siu zere AbiwD4T<0

DIAGRAM BY: DATE David Harrison 03-12-89

STATION HISTORY

Established December 1989

SOURCE ORGANISATION Dept Geological Sciences, University College London 211

COUNTRY Papua New Guinea NE1RKandqp GRAVITY STATION DESCRIPTION STATE STATION NAME STATION NO. p r o v i n c e Enga COUNTY Kandep High School KNDP(0211) LATITUDE LONGITUDE 5° 49’ 26.05" S 143° 29’ 30.77" E ELEVATI0N 2368m POSITION CONTROL GRAVITY VALUE Reference to local topographic maps 977461.27mgal

ELEVATION CONTROL BOUGUER ANOMALY Barometric levelling

DESCRIPTION

On flat ground approximately 10m west of the northern water collector at the western end of the woodwork classroom at Kandep High School. The school is approx 4km north of Kandep, on the road to Laiagam.

DESCRIBED BY: DATE David Harrison 17-12-89

DIAGRAM

glassy . CAinIwAtsQ H£c-tPAi> Goc»_£cTtxiS

o e

y N

DIAGRAM BY: DATE David Harrison 17-12-89 STATION HISTORY Established December 1989

Dept Geological Sciences, University College London 212

Appendix B

- Tiengo gravity survey local terrain corrections Appendix B - Local terrain corrections. 213

Station Local Station Local Station Local Number TC (mGal) Number TC (mGal) Number TC (mGal)

0004 0.00 TARI 0.00 0050 0.00 0005 0.03 0061 0.00 0114 0.00 0007 0.00 0062 0.01 0115 0.06 0008 0.00 LERU 0.00 0116 0.00 0009 0.00 LERU 0.00 0117 0.04 0010 0.00 0063 0.10 0118 0.06 0011 0.00 0064 0.00 0119 0.09 0013 0.12 0065 0.08 0120 0.01 0014 0.02 0066 0.02 0121 0.05 0015 0.03 0067 0.03 0122 0.00 0016 0.00 0068 0.04 0123 0.00 0017 0.05 0069 0.02 0124 0.10 0007 0.17 0070 0.44 0125 0.02 0018 0.02 0071 0.17 0126 0.00 0019 0.06 0072 0.05 0127 0.02 0020 0.01 0073 0.04 0128 0.02 0021 0.00 0075 0.00 0129 0.01 0022 0.00 0066 0.02 0130 0.00 0023 0.20 0076 0.00 0131 0.02 0024 0.04 0077 0.00 0132 0.06 0025 0.00 0078 0.08 0126 0.00 TARI 0.00 0079 0.00 0133 0.02 0027 0.00 0080 0.10 0134 0.02 0028 0.00 0081 0.06 0135 0.00 0007 0.17 0082 0.01 0136 0.00 0029 0.00 0083 0.00 0137 0.00 0030 0.12 LERU 0.00 0138 0.03 0031 0.37 0084 0.07 0139 0.04 0032 0.20 0081 0.06 0140 0.01 0033 0.00 0085 0.01 0141 0.03 0034 0.30 0086 0.00 0142 0.00 0033 0.00 0087 0.00 0143 0.00 0035 0.15 0088 0.00 0144 0.01 0036 0.03 0089 0.02 0145 0.00 0037 0.19 0090 0.00 0146 0.02 0038 0.06 0091 0.01 0147 0.02 0039 0.05 0092 0.00 0148 0.00 0040 0.06 0093 0.15 0149 0.05 0041 0.10 0094 0.21 0150 0.00 0042 0.00 0095 0.03 0151 0.00 NIPA 0.00 0096 0.00 0152 0.01 0044 0.09 0097 0.00 0153 0.00 0045 0.01 0098 0.02 0154 0.08 0046 0.05 KOMO 0.00 0155 0.03 0047 0.13 0081 0.06 0156 0.00 0048 0.09 0100 0.00 MNDI 0.00 0049 0.11 0101 0.00 0157 0.00 0050 0.00 0102 0.05 0140 0.01 0051 0.18 0103 0.00 0158 0.01 0033 0.00 0104 0.03 0159 0.00 LERU 0.00 0105 0.04 0160 0.00 0052 0.09 0106 0.01 0161 0.00 0053 0.00 0107 0.01 0162 0.00 0054 0.00 0108 0.00 0163 0.00 0055 0.00 0109 0.00 0164 0.00 0056 0.04 0110 0.01 0165 0.03 0057 0.00 0111 0.07 0166 0.00 0058 0.01 0112 0.03 0167 0.00 0059 0.00 0113 0.00 0168 0.00 0060 0.00 KMRA 0.00 0169 0.00 214

Station Local Station Local Station Local Number TC (mGal) Number TC (mGal) Number TC (mGal)

0.00 0170 0.02 0229 0.02 0286 0171 0.01 0224 0.01 0271 0.00 0172 0.01 0131 0.02 0287 0.00 0.00 0173 0.09 0230 0.00 0288 0.00 0174 0.00 0231 0.00 0289 0.00 0175 0.07 0232 0.00 0290 0.01 0176 0.07 0233 0.00 0291 0.00 0177 0.03 0234 0.02 0292 0.00 0178 0.02 0235 0.00 0293 0.00 0179 0.05 0236 0.04 0294 0237 0.00 0180 0.29 0.00 0295 0.00 0181 0.08 0238 0.00 0296 0182 0.12 0239 0.00 0183 0.00 0240 0.00 0184 0.00 0241 0.00 0185 0.00 0242 0.00 0186 0.04 0243 0.00 0187 0.09 0244 0.00 0188 0.00 0245 0.00 0189 0.03 0246 0.08 0190 0.06 0247 0.00 0191 0.04 0239 0.00 0192 0 .00 0248 0.01 0193 0.01 0249 0.26 0194 0.00 0250 0.00 0195 0.03 0251 0.02 0196 0.02 0252 0.03 0197 0.04 0253 0.16 0198 0.43 0254 0.05 0199 0.01 0255 0.01 0200 0.02 0256 0.41 0201 0.00 0257 0.04 0202 0.09 0258 0.00 0203 0.01 0259 0.00 0204 0.00 0260 0.01 0205 0.00 0261 0.00 0206 0.00 0262 0.00 0207 0.00 0263 0.00 0208 0.00 0264 0.00 0209 0.00 0265 0.00 0210 0.00 0266 0.00 0171 0.00 0267 0.00 0211 0.00 0268 0.00 0212 0.00 0269 0.00 0213 0.00 0270 0.00 0214 0.00 0271 0.00 0215 0.00 0272 0.00 0216 0.00 0273 0.05 0217 0.04 0274 0.00 0218 0.34 0275 0.00 0219 0.04 0276 0.00 0220 0.00 0277 0.00 0221 0.00 0278 0.00 0222 0.00 0279 0.00 0223 0.02 0280 0.00 0224 0.01 0281 0.00 0225 0.00 0282 0.00 0226 0.00 0283 0.05 0227 0.06 0284 0.00 0228 0.04 0285 0.01 Appendix C

- Tiengo gravity survey extended Bouguer anomaly data P — a s O rH aoM,or~CMCM,=J,CMQoor)r~r-LOVDQooocM'x>oP O s o a s rHs r~ vo ^ a CP £ p a s r~ a s •H o inin in 10 CQ rH rH rH I I I I I I I I I I I I I I I I I I I I 1 1 1

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- FORTRAN program listings 225 //UCFBEDHG JOB 'GRID', TIME=(15,0) / /*PASSWORD QANTAS2 //STEP EXEC PHOENIX3 //A EXEC PHOENIX,REGION=4096K SETUPLOCAL (FT04F001=.KO5018/NEW/FB ) FILE .D5000S+UCFBDG1.KUTUBU+UCFBDG1.EDCON+.D5000E2+.KOHTS +++ TO &GRIDAT/NEW/FB/MB=5,1 FVSCLG PROGRAM=%RESTI DD= ' FT05F001=&GRIDAT/ FB' LIBRARY=SYS2.GINOFVS.LIB C C This program interpolates a regular grid of z values C from random x, y, and z data. Before use the grid limits, C grid separation, approximate number of random data points, C and the number of random points to be used for each C interpolation (arg of CIRPTS) should be defined, C DIMENSION ZARR(200, 200),XX(97000) , YY(97000) DIMENSION WW(500000), ZZ(97000) C Grid separation (m) GRISEP=50. C Coordinates of gridded area XMIN=775025. XMAX=784975. YMIN=9295025. YMAX=9304 975. C Number of data points NP=95000 NUMX=((XMAX-XMIN)/GRISEP)+1 NUMY=((YMAX-YMIN)/GRISEP)+1 NW=(2 *NUMX*NUMY) + (4 *NP)+10 0 0 DO 20 1=1,NP 20 READ(5,*,END=500)XX(I),YY(I) , ZZ(I) 500 CONTINUE CALL GINO CALL ERRSWI(0) CALL ERRMAX(-1) CALL CIRPTS(10) CALL RANGRD(NP,XX,YY,ZZ,NUMX,XMIN,XMAX,NUMY, +YMIN,YMAX,ZARR,NW,WW) CALL GINEND 30 CONTINUE 2000 WRITE(4,'(213)')NUMX,NUMY WRITE(4, ' (2F7.0 ,2F8.0)r) XMIN,XMAX,YMIN,YMAX WRITE(4 ,'(10F7.2)') ZARR STOP END 226 / /UCFBEDHT JOB 'TOTAL', TIME '2 SECS' //^PASSWORD QANTAS2 //STEP EXEC PHOENIX3 F77C PROGRAM=%RESTI OPTIMIZE=3 OBJECT=UCFBEDH.TTCOBJ41 0* ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ * ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ * ★ ★ ★ c* C* This program carries out a Total Terrain Correction on C* gravity data using the line mass approximation for distant C* prisms, and Nagy's Equation for those close to the gravity C* station. C* Input data required are gravity data (stream 9) and gridded C* elevation data (stream 3 - max 200x200). C* C* Before use, the spheroid constants and the UTM projection C* constants must be defined. C* 0* ★ ★ ★ ★ -k -k ★ ★ ★ ★ ★ * ★ ★ * * ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★ ★

DOUBLE PRECISION SMAJ,ECC2,FF,SF,DLONO,STALAT,STALON + ,PDLAT,PDLONG,PDX,PDY,F, SMIN INTEGER STANUM DIMENSION ZARR(200,200),PLATS(200,200),PLONGS(200,200), +PXS(200,200),PYS(200,200) COMMON/STA/XSTA,YSTA,ZSTA COMMON/ALL/G,RHO,GRISEP COMMON/CEQN9/X1,X2,Yl,Y2,Z,TC COMMON/CPRIS/PLAT,P LONG, P Z COMMON/CFAR/SCRX,SCRY,SCRZ COMMON/LATCON/F,SMAJ,SMIN,ECC2,SF,DLONO,RLAT,RLON G=6.67E-11 C Density (kg/m3) RHO=2 670. C Spheroid Constants - FF is l . / f SMAJ=6378160.0D0 FF=298.25D0 C UTM projection constants C Central meridian DLONO=141.DO C Scale factor SF=0. 9996D0 F-l.D0/FF ECC2=(2.D0*F)- (F*F) TA s h atato o oe prism one of attraction the TCAC is Cnet tto psto t shrcnrc etnua coords rectangular spherocentric to position station Convert C here Data C Read Gravity Cluae iis f ie 5m gi area (50m) grid fine of limits Calculate C Cluae a ad og f ah prism each of long and lat Calculate C prisms) square (assumes dimension prism GRISEPC the is Eeain data Elevation C area gridded of UTMC limits array(grid) of DimensionsC Ra eeain data C Read elevation 2900 READ( END=990 , 0 0 0 3 , FORMAT(IX,F8.3 9 ,6F10.2) 3000 ,IX,2F11.1 SURVEY,STANUM,STALAT,STALON,XSTA,YSTA, 0) ,IX,F12.7 ,4X,A4, F12.7 nnonn CONTINUE2100 CONTINUE2200 C i te oa atato o te oo lc (mgal) (ms-2) block topo block the topo of this attraction of total attraction the TCT total is running the TCL is 2 FORMAT(2F7.0,2F8.0)12 FORMAT(213)10 + Z S T A , G O B S ,FA A , T E R C O R , Z D I F F , E B A B = . tee r st o eo ny o te first the for only zero to set are these - EBA=0. P L O N G = P L O N G S ( I ,PZ=ZARR(I,J) J ) PY=PYS(I,J) PLAT=PLATS(I,J) PX=PXS(I,J) TCA=0. DO NUMY J=1, 4000 DO 1=1,NUMX 4100 FYMAX=( (INT( (YSTA+3000.+100.)/200.))*200.)-25. FXMAX=( (INT( (XSTA+3000.+100.)/200.))*200.)-25. C A LRE L C T ( S T A L A T , S T A L O N , Z S T A , SFYMIN=FYMAX-(2.*(3000.-25.)) C R X , S C RFXMIN=FXMAX-(2.*(3000.-25.)) Y , S C R Z ) P L O N G S =PLATS(I,J)=PDLAT ( I,J) P D L O N G PDY=PYS(I,J) PYS(I,J)=YMIN+(J—1 (YMAX-YMIN)/PDX=PXS(I,J) * ) (NUMY-1) R E A D M (3,12)X I N , X M A X , Y M I N , Y M A X TCL=0. TCT=0. C A LLA L T L O N ( P D X , P D Y , P D L A T , PPXS(I,J)=XMIN+( D (XMAX-XMIN)/ L * ) O 1 - 1 N GDO ) NUMY J=1, 2200 (NUMX-1)DO 1=1,NUMX 2100 (ZARR(I,J),1=1,NUMX), ( ) READ( J=1,NUMY) ,'(10F7.2)' 3 R E A D (3,10)N U M X , N U M Y TERCOR=0. G R I S E P = ( X M A X - X M I N U ) /(N M X - 1 ) STALON=STALON*3.141592654D0/180.DOSTALAT=STALAT*3.141592654D0/180.DO D F=. u wt ay rvt data gravity any with run ZDIFF=0.

227

228

C Determine surface distance from prism to station DX=PX-XSTA DY=PY-YSTA DX=SQRT(DX*DX) DY=SQRT(DY*DY) DD=SQRT(DX*DX+DY*DY) C Initially only terrain correct up to 90km from station IF(DD.GT.90000.)GOTO 4000 C Fine (50m) grid, or other? IF(GRISEP.GT.100.)THEN C Prism outside fine grid area? IF(DX.GT.3 0 0 0..OR.DY.GT.30 0 0.)THEN CALL TCFAR(TCA) ELSE CONTINUE END IF ELSE C Prism within fine grid area IF(PX.LT.FXMIN.OR.PX.GT.FXMAX)GOTO 3100 IF(PY.LT.FYMIN.OR.PY.GT.FYMAX)GOTO 3100 C Prism close to gravity station - requires full equation IF(DX.LT.350.AND.DY.LT.350.)THEN C Prism at gravity station IF(DX.LT.25.AND.DY.LT.25.)THEN ZDIFF=ZSTA-PZ END IF Z=SQRT((PZ-ZSTA)**2) CALL NAGY(PX,PY,DX,DY,TCA) TC1=TCA Z=ZSTA CALL NAGY(PX,PY,DX,DY,TCA) TCA=TCA-TC1 ELSE CALL TCFAR(TCA) END IF 3100 CONTINUE 3300 END IF TCL=TCL+TCA 4000 CONTINUE 4100 CONTINUE C Convert ms-2 into m illigals TCT=TCL*100000. TERCOR=TERCOR+TCT EBA=FAA-TERCOR STALAT=STALAT*180.DO/3.141592654D0 STALON=STALON*180.DO/3.141592654D0 C Output gravity data with running total terrain correction TERCOR WRITE(4,3000)SURVEY,STANUM,STALAT,STALON,XSTA,YSTA, +ZSTA,GOBS,FAA,TERCOR,ZDIFF,EBA GOTO 2900 9900 CONTINUE STOP END 229

C C TCFAR uses the thin rod approximation to calculate the C gravitational attraction of a prism. C

SUBROUTINE TCFAR(TCA) COMMON/CPRIS/PLAT,PLONG,PZ COMMON/CFAR/SCRX,SCRY,SCRZ COMMON/STA/XSTA,YSTA,ZSTA COMMON/ALL/G,RHO, GRISEP C Covnvert top and bottom of prism to spherocentric rect. co-ords. CALL RECT(PLAT,PLONG,PZrRPX,RPY,RPZ) CALL RECT(PLAT,PLONG,0 .,RPXO,RPYO,RPZO) C Distance from origin to gravity station DIST=SQRT(SCRX**2+SCRY**2+SCRZ**2) C Distance from origin to top of prism DISTP=SQRT(RPX*RPX+RPY*RPY+RPZ*RPZ) C Distanse from origin to base of prism DISTPO=SQRT(RPX0*RPX0+RPY0*RPY0+RPZ0*RPZ0) C A is the angle at the origin between the station and prism COSA=(SCRX*RPX+SCRY*RPY+SCRZ*RPZ)/ (DIST*DISTP) DATUM=DIST/COSA C DATHT is the elevation of the tangent at the gravity station at the prism DATHT=DATUM-DISTPO C Distances from station to top and bottom of prism RTOP=SQRT((SCRX-RPX)**2+(SCRY-RPY)**2+(SCRZ-RPZ)**2) RBOT=SQRT((SCRX-RPXO)**2+(SCRY-RPYO)**2+(SCRZ-RPZO)**2) C All "INT" variables relate to the DATHT prism position CALL RECT (PLAT,PLONG,DATHT,XINT,YINT,ZINT) RINT=SQRT((SCRX-XINT)**2+(SCRY-YINT)**2+(SCRZ-ZINT)**2) SINA=SQRT( 1 (COSA*COSA)) IF(SINA.NE.0)THEN XMZS=RINT-(DATHT-PZ)*SINA XCOSA=RINT*COSA XMDS=RINT-DATHT*SINA ELSE XCOSA=l. XMZS=1. XMDS=1. END IF C TCA is the attraction of the prism TCA=G*RHO*GRISEP*GRISEP/XCOSA*(XMZS/RTOP-XMDS/RBOT) RETURN END I pim rse oii, l fu qarns ut e evaluated be must quadrants four all origin, crosses prism If C I ti ruie te rgn s ae t b te rvt station gravity the be boundarycoordinates to prism taken the is are S and P,Q,R origin the C routine, this InC Coss -xs to avs vlae separately evaluated halves two Y-axis, C Crosses separately o o o o n n A Y s o pim fr hc te ul euto fr the for equation ll fu the which for prisms NAGY for is RR=SQRT(QQ=SQRT((XSTA-(X+25. (YSTA-(Y-25. )**2) ) )**2) ) R=MIN(RR,SS) CALL EQN9 X2=P X1=0 . PP=SQRT( (XSTA-(X-25.))**2) Q = M A XP ( P = M P , I Q N ( Q P P ) , Q Q ) Y2=R Y1=0. CALL EQN9 X2=Q CALL EQN9 X2=P X2=P X2=Q C O M M OC N O / M A L M L O / G N , / R C H E O Q , N G 9 R / I X S 1 E , P , X 2 , Y l, Y 2 , Z , T C TC=0 . Y2=S X1=0 . ELSE IF(DX.LT.25.)THEN X2=Q CALL EQN9 SS=SQRT( (YSTA-(Y+25. )**2) ) COMMON/STA/XSTA,YSTA,ZSTA S = M A X ( R R , S S ) Y1=R CALL EQN9 CALL EQN9 S U B R O U T I NNA E G Y ( X , Y , D X , D Y , T O N ) IF(DX.LT.25..AND.DY.LT.2 .)THEN5 Y2=S f rcaglr rs i necessary. is prism rectangular a of

230

o o o o o o quadrant. one within entirely C Prism separately evaluated halves two X-axis, C Crosses Q 9 s ays qain (9) equation Nagy'sEQN9 is END IF ELSE T7=0 . T7=LOG( (X2+SQRT(ST2))T2=LOG((Yl+SQRT(ST2) (X2+SQRT(ST2+ST5))) / (Yl+SQRT(ST2+ST5))) / ) END IF ELSE COMMON/ALL/G,RHO,GRISEP END IF T2=0. T5=LOG( (X2+SQRT(ST1))Tl=LOG( (X2+SQRT(ST1+ST5))) (Y2+SQRT(ST1)) / (Y2+SQRT(ST1+ST5))) / END RETURN T5=0 . Tl-0 . C O M M O NS /C U E B Q R N O U ,X2 ,Y1 ,Y 9/XI T I , N T 2,Z EQ C E N 9 TCN=TC CALL EQN9 X2=Q Xl-P ELSE X2=Q Y2=S Yl-R Xl-P ELSE IF(DY.LT.25.)THEN CALL EQN9 CALL EQN9 IF(ST2.EQ.0 .)THEN ST5=Z**2 ST4=X1**2+Y1**2ST3=X1**2+Y2**2ST2=X2**2+Y1**2ST1=X2**2+Y2**2 Y2=S Y2=R Y1=0 FSlE.) THENIF(STl.EQ.O)

232 IF(ST3.EQ.0)THEN T3=0 . T6=0 . ELSE T3=L0G((Y2+SQRT(ST3)) / (Y2+SQRT(ST3+ST5))) T6=L0G((Xl+SQRT(ST3)) / (Xl+SQRT(ST3+ST5)) ) END IF IF (ST4.EQ.0 .)THEN T4=0 . T8=0. ELSE T4=L0G((Yl+SQRT(ST4) ) / (Yl+SQRT(ST4+ST5))) T8=L0G((Xl+SQRT(ST4) ) / (Xl+SQRT(ST4+ST5))) END IF IF(Y2.EQ.0.AND.ST5.EQ.0 .)THEN T9=0 . T10=0. ELSE T9=ASIN((Y2**2+ST5+Y2*SQRT(ST1+ST5)) /( ( Y2+SQRT(ST1+ST5))* +SQRT(Y2**2+ST5))) T10=ASIN( (Y2**2+ST5+Y2*SQRT(ST3+ST5)) / ( (Y2+SQRT(ST3+ST5))* +SQRT(Y2**2+ST5))) END IF IF(Y1.EQ.0.AND.ST5.EQ.0 .)THEN T11=0. T12=0. ELSE T11=ASIN((Y1**2+ST5+Y1*SQRT(ST2+ST5)) / ( (Yl+SQRT(ST2+ST5))* +SQRT(Y1**2+ST5))) T12=ASIN( (Y1**2+ST5+Y1*SQRT(ST4+ST5)) / ( (Yl+SQRT(ST4+ST5))* +SQRT(Y1**2+ST5))) END IF TC=TC+(G*RHO*(X2*(T1-T2)-X1*(T3-T4)+Y2*(T5-T6) +-Y1*(T7-T8)+Z*(T9-T10-T11+T12))) RETURN END

C* C* RECT converts geographic lat and long into C* rectangular (X,Y,Z) coordinates with origin C* at the centre of the spheroid. C*

SUBROUTINE RECT(LAT,LONG,HT,X,Y,Z) DOUBLE PRECISION SMAJ,ECC2,F,SMIN,LAT,LONG COMMON/LATCON/F,SMAJ,SMIN,ECC2,SF,DLONO,RLAT,RLON REAL NU,HT S2LAT-SIN(LAT)*SIN(LAT) NU-SMAJ/SQRT(1.-(ECC2*S2LAT)) X=(NU+HT)*COS(LAT)*COS(LONG) Y=(NU+HT)*COS(LAT)*SIN(LONG) Z=(((1 -(ECC2))*NU)+HT)*SIN(LAT) END onnonn A L N ovrs T codnts no egahc latitude LATLON geographic UTM into converts coordinates +D2*DSIN(8.D0*W) + +D2*DSIN(8.D0*W) + 2 »1.5D0*Y*(1.DO-9.D0*Y2/16.DO)A2 END RETURNDLON=LON*180.D0/3.141592654D0DLAT=LAT*180.DO/3 .141592654-DO =NN/R1W NN =NN+(Q-V)*R =EE*(1.D0-F2*E2*E2)E3 =EE*EEE2 EE =EE/R =NN/R1W NN (N-NO) = /SF+M1 N0=10000000.DO =DATAN2(V .DO 2 DO+T*T*DCOS(LON)) . 1 *T*S*S, =E3*(1.DOV +E32*(1.DO/6.DO +E32*(1.DO/12.D1 +E32/504.Dl))) CALL SUB2(El,FI,F2,F3,J,LAT,L3,R) .1D0*Y2*Y2 2 D2 = 2 =7.D0*Y2*.25D0*(B2 .75D0-Y2) (E-EO)EE = /SF RLON=DLONO*3.141592654D0/18,D1 L O N = L O N / J + R LLAT=W+A2*DSIN( O N .D0*W)+B2*DSIN( 2 .D0*W)+C2*DSIN 4 D0*W) . 6 ( RLAT=0.DO n longitude. and =DTAN(LAT)T E32=E3*E3 LAT=L3+Q =E2*E2*(F3*E2-F1)Q LAT=W+A2 *DSIN( .D0*W)+B2*DSIN( 2 .D0*W)+C2*DSIN( 4 .D0*W) 6 C2 =6.3D0*Y2*Y*. 25D0 CALL SUB1(Al,Bl,Cl,El,M1,Rl,Y,Y2) E0=500000.D0 C O M M O N / L A T C O N / F , S M A J , S M I N , E C C 2 , S F , D L O N O , R L A T , R L O N =DSIN(LON*.5D0)S L O N = D A T A NDC 2 ( V O , S ( L A T ) ) S M I N = S M A J - ( S M A J * F ) S U B R O U T I NLA E T L O N ( E ,N,L A T ,L O N ) IMPLICIT DOUBLE PRECISION (A-Z)

233 234

SUBROUTINE SUB1(Al,B1,C1,E1,M1,Rl,Y,Y2) IMPLICIT DOUBLE PRECISION (A-Z) COMMON /LATCON/F,SMAJ* SMIN,ECC2, SF,DLONO,RLAT,RLON Z =RLAT Y - ( SMAJ-SMIN)/ ( SMAJ+SMIN) Y2 =Y*Y Rl * (SMAJ+SMIN)/(2 .DO-. 5D0*Y2) Al -1.5D0*Y*(1.DO-3.D0*Y2*. 125D0) B1 -15.D0*Y2/16.D0 Cl -35.D0*Y2*Y/48.DO Ml -Rl*(Z-A1*DSIN(2.D0*Z) +B1*DSIN(4.D0*Z) -C1*DSIN (6.D0*Z)) El -ECC2/(1.D0-ECC2) END SUBROUTINE SUB2(El,F1,F2,F3,J,LAT,L3,R) IMPLICIT DOUBLE PRECISION (A-Z) COMMON /LATCON/F,SMAJ,SMIN,ECC2,SF,DLONO,RLAT,RLON S -DSIN(LAT) S2 =S*S C =DCOS(LAT) C2 =C*C T -S/C J =DSQRT(1,D0+E1*C2*C2) L3 =DATAN2(S,DSQRT(J*J-S2)) V =1.D0/(1.D0-ECC2*S2) R =SMIN*V FI -E1*S*C*DSQRT((1.D0-ECC2)*V)/6.D0 F2 =El*C2/3.Dl F3 —FI*(3.D0*T*T-2.D0)/3.D1 RETURN END PNG Gravity Study - Extended Bouquer Anomaly Map - 1:500000 141° E 142* E 143* E