late tertiary and quaternary geology of the east java basin, indonesia
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University of WollongongResearch Online
University of Wollongong Thesis Collection University of Wollongong Thesis Collections
1995
Late tertiary and quaternary geology of the EastJava Basin, IndonesiaSusilohadiUniversity of Wollongong
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Recommended CitationSusilohadi, Late tertiary and quaternary geology of the East Java Basin, Indonesia, Doctor of Philosophy thesis, School of Geosciences,University of Wollongong, 1995. http://ro.uow.edu.au/theses/1973
LATE TERTIARY AND QUATERNARY GEOLOGY
OF THE EAST JAVA BASIN, INDONESIA
A thesis submitted in (partial) fulfilment of the
requirements for the award of the degree of
DOCTOR OF PHILOSOPHY
from
UNIVERSITY Of WOLLONGOMfi
LIBRARY
THE UNIVERSITY OF WOLLONGONG
NEW SOUTH WALES, AUSTRALIA
by
SUSILOHADI
(Ir. ITB)
School of Geosciences
1995
ABSTRACT
The East Java Basin is classified as a retro-arc basin, situated on the southeastern
margin of the Sunda Shelf, Indonesia. A detailed Late Tertiary and Quaternary sedimentary
evaluation of this basin is based primarily on seismic stratigraphic interpretation, combined
with onshore outcrop studies of equivalent geological units. The primary aim of this study was
to assess the sedimentation patterns and sedimentological history in the East Java Basin.
During the Middle Miocene and earlier, the northern part of the East Java Basin was
strongly controlled by northeast-trending half graben systems which occurred along sutures
between the earlier Late Cretaceous-Early Tertiary subduction systems. Little is known about
the basin configuration in the southern part of the basin prior to the Middle Miocene.
Since the Late Miocene, east-trending anticlinal zones developed and were
superimposed on the previous northeast-trending structures. The anticlinal Rembang Zone,
which extends from the Blora area to Madura Island, became a dominant structure which
divided the East Java Basin into two major basinal (synclinal) areas. In the north, a basinal
area occurred between the Karimunjawa and Bawean Arches and the Rembang Zone. In the
south, a basin developed between the Rembang Zone and the chain of volcanism along the
median line of Java. Peak sedimentation in these east-trending basins occurred during the Late
Miocene and Pliocene. The sedimentation patterns, such as local thickening, in the north
basinal area (southern Java Sea) suggest that the basin development ceased in the Early/Middle
Pleistocene. In contrast, rapid basin subsidence still occurred in the south basinal area at least
until the late Middle Pleistocene, when extensive east-trending folding accompanied by diapiric
movement occurred to form the anticlinal Kendeng Zone.
Middle Miocene and earlier sedimentation have not been comprehensively studied in
the East Java Basin, particularly because of the paucity of outcrop and well data. The present
seismic stratigraphic study of the onshore Rembang Zone revealed that the Middle Miocene
sedimentation (Tawun Formation) patterns were strongly controlled by half graben systems,
the southern extension of the northeast-trending stmctures indicated earlier. These structures
still partly controlled the Late Miocene and Pliocene sedimentation in the Rembang Zone and
the southern Java Sea.
Hemipelagic marl deposition (Wonocolo Formation) occurred in the southern part of
the East Java Basin (onshore area) during a sea level highstand in the early Late Miocene, but
little is known about equivalent deposition in the southern Java Sea at this time. The marl
deposition was contemporaneous with the deposition of volcaniclastic turbidites (Kerek
Formation) derived from Late Miocene volcanism along the median line of Java. This
volcanism appears to have ceased in the late Late Miocene. A shallow shelfal facies (Ledok
Formation) occurred in the southern part of the Rembang Zone (onshore area) which
southward, grades laterally into the marl of the Kalibeng Formation.
During the Pliocene, the sedimentation was confined within the northern and southern
basinal areas. Widespread pelagic marly deposition occurred during a prolonged highstand of
sea level in the Pliocene which reflects a global eustatic sea level rise recognised by Haq et
al. (1988). In the southern basinal area, the deposit is represented by the M u n d u (seismic unit
PL1), Kalibeng and Atasangin Formations respectively; whereas in the northern basinal area
(southern Java Sea) the deposit is represented by seismic units JP1 and JP2. A relative sea
level fall occurred in the Late Pliocene and led to the development of an anoxic lacustrine
deposit (Lidah Formation) in the southern basinal area.
During the Quaternary, the seismic stratigraphic study indicates that sediment stratal
patterns were highly controlled by frequent relative sea level changes, and sedimentation was
still confined within the two major basinal areas indicated earlier. In the southern basinal area,
seismic sections clearly display an intertonguing between volcaniclastics derived from the
Quaternary volcanism along the median line of Java, and shallow marine sediments derived
from the exposed Rembang Zone in the north. Field observations indicate that volcaniclastic
deposits are coarse-grained, whereas on seismic sections from Madura Strait these deposits are
associated with relatively steep depositional slopes (of alluvial fan/fan delta systems). The
seismic stratigraphic study in Madura Strait indicates that the depositional timing of these
deposits was strongly controlled by the relative sea level changes in this area.
In the southern Java Sea (northern basinal area), the basin filled during the Pliocene
as revealed from the seismic stratigraphic study, giving a flat basin topography at the end of
the Pliocene. Nine thin seismic subunits are recognised in this area and are believed to
represent at least nine Quaternary sea level fluctuations. Each subunit is relatively thin, and
tends to be distributed widely because of deposition on a relatively flat lying area. The seismic
characters of these subunits are very similar, with subparallel reflection or almost reflection
free patterns representing marine deposits, topped by extensive fluvial channelling. This
repetitive succession represents highstand and lowstand of sea level respectively.
The integration of seismic sequence stratigraphic analysis and outcrop sedimentological
studies has provided a comprehensive view of the East Java Basin on the timing of deposition
and deformation of the basin fill, mode of deposition and sedimentation patterns within the
sub-basins, and the influence of sea level changes on stratal patterns.
ACKNOWLEDGMENTS
I gratefully acknowledge the Australian Government through the Australian Agency for
International Development (AusALD) for providing a scholarship for this research.
I would like to express my deep appreciation to my supervisors Dr L.E.A. Jones and
Dr C. Murray-Wallace, and also to Associate Professor B.G. Jones for their help,
encouragement and guidance during the study.
I am also indebted to the former Heads of the Geology Department, Wollongong
University, Dr A.C. Cook and Associate Professor A.J. Wright, for their constructive
suggestion to do this research. Special gratitude also goes to all academic, administrative and
technical staff of the department, particularly M s B. McGoldrick, M r M . Perkins, M r D.A.
Carrie, M r A.M. Depers and M s P. Williamson.
Special thanks are forwarded to the former Director of the Marine Geological Institute
of Indonesia, M r Ismail Usna, and the present Director M r Aswan Yasin, who permitted the
author to do this research. Special appreciation is also forwarded to m y colleagues in this
institute: M s K.T. Dewi for micropalaeontological analysis, and Messrs S. Tjokrosapoetro, B.
Dwiyanto, A. Masduki, A. Setiabudhi, P. Astjario, S. Lubis and B. Dharmawan for then-
support and suggestion.
Special thanks are also addressed to the management of Pertamina, PPT Migas and
Lemigas for the permission to use the geological, geophysical and well data from the Java Sea
and Cepu area. The author would like to acknowledge Drs N. Hasjim, A. Muin and S. Soeka;
Messrs T. Ratkolo, A. Sutrisman, N. Suparjono, and S. Musliki for their help in searching for
some references in these institutes.
Many thanks are also forwarded to my post-graduate colleagues, including Messrs
Surono, Y. Kusumahbrata, A. Pujobroto, Moradian, Soltani, A. Mandile, A. Wahib, D.
Dharmayanti, Destini, H. Sobir, Prasetyo, A. Hamedi and others, for lively discussions during
this research. Finally, I am thankful to m y wife and daughter, Kresna and A m m , and the rest
of m y family for their support.
Except where otherwise acknowledged, this thesis represents the author's original research and
the material included has not been submitted for a higher degree to any other institution.
Susilohadi
TABLE OF CONTENTS
ABSTRACT ACKNOWLEDGEMENTS
CHAPTER 1 INTRODUCTION 1 1.1 PREVIOUS INVESTIGATIONS IN THE EAST JAVA BASIN 1 1.2 ATMS OF STUDY 4 1.3 AREA OF STUDY 5 1.4 ORGANISATION OF THESIS 6
CHAPTER 2 GEOLOGICAL SETTING OF THE EAST JAVA BASIN 7 2.1 ESTTRODUCTION 7 2.2 TECTONIC SETTING 7 2.3 REGIONAL GEOLOGICAL STRUCTURES 8 2.4 REGIONAL DEPOSITIONAL EVOLUTION 9 2.5 PHYSIOGRAPHY 11 2.6 SUMMARY 12
CHAPTER 3 STRATIGRAPHY OF THE ONSHORE EAST JAVA BASIN 15 3.1 INTRODUCTION 3.2 LLTHOSTRATIGRAPHY 15
3.2.1 Rembang Zone 17 3.2.2 Kendeng Zone 22
3.3 BIOSTRATIGRAPHY 26 3.3.1 Miocene and Middle Pliocene strata 27 3.3.2 Late Pliocene and Pleistocene strata (Lidah Formation) 29
3.4 CLTMATOSTRATIGRAPHY 29 3.5 MAGNETOSTRATIGRAPHY 30 3.6 SUMMARY 31
CHAPTER 4 SEISMIC STRATIGRAPHY OF THE REMBANG ZONE 33 4.1 INTRODUCTION 33 4.2 SEISMIC SEQUENCE AND FACIES DEFINITION 34 4.3 SEISMIC SEQUENCE STRATIGRAPHY OF THE REMBANG
ZONE 37 4.3.1 Sequence boundaries 39 4.3.2 Seismic fades 41 4.3.3 Relative sea level change 51
4.4 SEDTMENTARY PETROGRAPHY 51 4.4.1 Tawun Formation 52 4.4.2 Bulu Formation 53 4.4.3 Wonocolo Formation 53 4.4.4 Ledok Formation 54 4.4.5 Mundu Formation 55
4.5 DISCUSSION 56 4.6 SUMMARY AND CONCLUSIONS 58
CHAPTER 5 SEDIMENTOLOGY OF THE KENDENG ZONE 59 5.1 TNTRODUCTION
60 60 61
64 65
5.2 KEREK FORMATION 5.2.1 Lithofacies 5.2.2 Petrography 5.2.3 Depositional setting and discussion 61
5.3 KALIBENG FORMATION 6 4
5.3.1 Lithofacies 5.3.2 Petrography 5.3.3 Depositional environment and discussion 66
5.4 ATASANGIN FORMATION 68
5.4.1 Lithofacies 69
5.4.2 Petrography 72 5.4.3 Benthonic fauna 73 5.4.4 Discussion 74
5.5 SONDE FORMATION 75
5.5.1 Lithofacies 75
5.5.2 Discussion 76 5.6 LID A H FORMATION 76
5.6.1 Lithofacies 76 5.6.2 Petrography 78 5.6.3 Discussion 80
5.7 PUCANGAN FORMATION 81 5.7.1 Lithofacies 81 5.7.2 Petrography 92 5.7.3 Discussion 94
5.8 KABUH FORMATION 95 5.8.1 Lithofacies 95 5.8.2 Discussion 97
5.9 S U M M A R Y 97
CHAPTER 6 QUATERNARY SEISMIC STRATIGRAPHY OF MADURA STRAIT 99 6.1 INTRODUCTION 99 6.2 SEISMIC DATA 101 6.3 WELL DATA 102 6.4 PHYSIOGRAPHY OF M A D U R A STRAIT 105 6.5 SEISMIC FACIES ANALYSIS AND PNTERPRETATTON 105
6.5.1 Late Tertiary units 106 6.5.2 UnitPre-A 110 6.5.3 Unit A 110 6.5.4 UnitB 112 6.5.5 UnitC 114 6.5.6 UnitD 118
6.6 DISCUSSION 122 6.6.1 Correlation with onshore data 122 6.6.2 Depositional timing 123 6.6.3 Sedimentary source 124
6.7 S U M M A R Y A N D CONCLUSION 125
CHAPTER 7 LATE TERTIARY AND QUATERNARY SEISMIC SEQUENCE STRATIGRAPHY OF THE SOUTHEAST JAVA SEA 127 7.1 INTRODUCTION 127 7.2 DATA 128 7.3 STRUCTURAL GEOLOGY AND BATHYMETRY 129 7.4 SEISMIC STRATIGRAPHY 131
7.4.1 UnitJM 131 7.4.2 UnitJP 132 7.4.3 UnitJQ 138
7.5 DISCUSSION 143 7.6 SUMMARY 144
CHAPTER 8 SUMMARY AND CONCLUSIONS 147 8.1 INTRODUCTION 147 8.2 BASIN DEVELOPMENT 147 8.3 SEDIMENTATION HISTORY 148
8.3.1 The onshore area during the Middle Miocene and Pliocene 148 8.3.2 The onshore area (eastern Kendeng Zone) and Madura Strait
during the Quaternary 149 8.3.3 Pliocene and Quaternary sedimentation in the southern Java
sea 150 8.4 NOTES ON THE APPLICATION OF SEQUENCE STRATIGRAPHY EST
THE EAST JAVA BASIN 151 8.5 CONCLUSION 153
REFERENCES 155
LIST OF FIGURES LIST OF TABLES APPENDICES LIST OF SEISMIC SECTIONS
FIGURES
FIGURES
FIGURES
FIGURES
FIGURES
FIGURES
FIGURES
FIGURES
FIGURES
TO CHAPTER 1
TO CHAPTER 2
TO CHAPTER 3
TO CHAPTER 4
TO CHAPTER 5
TO CHAPTER 6
TO CHAPTER 7
TO CHAPTER 8
FIGURE 1.1
FIGURE 2.1
FIGURE 3.1
FIGURE 4.1
FIGURE 5.1
FIGURE 6.1
FIGURE 7.1
FIGURE 8.1
FIGURE 2.8
FIGURE 3.10
FIGURE 4.23
FIGURE 5.58
FIGURE 6.32
FIGURE 7.15
FIGURE 8.2
TABLES
TABLES TO CHAPTER 3
TABLES TO CHAPTER 4
TABLE 3.1 - TABLE 3.4
TABLE 4.1
APPENDICES
APPENDIX A. LIST OF AVAILABLE VELOCITY DATA FOR LFNE PWD-23
APPENDIX B. PETROGRAPHIC AND X-RAY DIFFRACTION DATA
APPENDIX C. PLANKTONIC AND BENTHONIC FORAMTNIFERAL DATA
APPENDIX D. LIST OF SAMPLES USED FN THE STUDY
APPENDIX E. GRAPHIC SUMMARY OF PETROLEUM EXPLORATORY WELLS
FROM THE JAVA SEA
APPENDIX F. SOURCE CODES OF COMPUTER PROGRAMS USED TN THE
STUDY
SEISMIC SECTIONS (Volume 2)
SEISMIC SECTIONS TO CHAPTER 4
SEISMIC SECTIONS TO CHAPTER 6
SEISMIC SECTIONS TO CHAPTER 7
SECTION 4.1 - SECTION 4.2
SECTION 6.1 - SECTION 6.12
SECTION 7.1 - SECTION 7.10
1
CHAPTER ONE
INTRODUCTION
The East Java Basin is located on the southeastern margin of the Sunda Shelf, an extensive
shelf which comprises the whole of western Indonesia. The basin is classified as a retro-arc
basin (Dickinson, 1977) because of its position behind a volcanic arc in a convergent plate
tectonic setting. Like other retro-arc basins in Indonesia, this basin is geologically and
economically important. The East Java Basin is an important petroleum province and has been
the subject of geological investigations since the end of the 19th century. Besides petroleum-
related geological investigations in this area, important research has also been undertaken in
the fields of biostratigraphy and Quaternary anthropology. These interests have high-lighted
the necessity to continuously assess the region's geological evolution in the tight of new
geological concepts.
The development of geological concepts and ideas has been extremely rapid during the last
decade, particularly after the introduction of (seismic) sequence stratigraphy (Vail et al, 1977;
Posamentier et al., 1988; Posamentier & Vail, 1988). Application of these ideas is world wide,
but the numerous applications concentrate on shelf margin settings rather than retro-arc basins.
Since the introduction of these concepts, very few attempts have been made to apply them to
the East Java Basin and these few were aimed at specific problems in limited areas of the basin
(e.g. Yulianto, 1993). The present study critically examines the East Java Basin based on
seismic sequence stratigraphic concepts. Within the limits of the available data, this study
investigates the Late Tertiary and Quaternary strata in almost the whole East Java Basin with
a major emphasis on seismic interpretation.
1.1 PREVIOUS INVESTIGATIONS IN THE EAST JAVA BASIN
a. Pre 1940s
Before the 1940s, geological exploration was carried out primarily by Dutch geologists.
In the inland area, two institutes with different interests were involved. In the northern part
of this area (the anticlinal Rembang Zone; see Chapter 2.5 for an explanation), the geological
exploration, which began in 1887, was carried out by the Dortsche Petroleum Maatschappij and
then continued by the Bataafsche Petroleum Maatschappij (BPM). Both of these Dutch oil
2
companies found some oil fields shortly after exploration commenced, but the results of their
study were proprietary and not publicly available (van Bemmelen, 1949).
Geological exploration in the southern part of the inland area (the anticlinal Kendeng Zone;
see Chapter 2.5 for an explanation) was carried out by the Geological Survey of the
Netherlands Indies, and involved geological mapping and stratigraphic evaluation of Tertiary
and Quaternary strata. Some East Java geological mapping reports are considerably detailed
(e.g Duyfjes, 1938a,b,c,d). Most of the work of the Geological Survey of the Netherlands
Indies was summarised by van Bemmelen (1949). Until recent years this summary remained
the main source of information on the geology of East Java.
Quaternary geological studies in East Java began with the finding of mammalian faunas
and hominid fossils by Dutch geologists in the late 19th century, as reported by Duyfjes (1936)
and de Vos and Sondaar (1982). These faunas occur in three main Quaternary stratigraphic
units, the Pucangan, Kabuh and Notopuro Formations. Examination of the relationships
between Quaternary deposits and sea level changes in East Java is timely, as the present study
suggests that the extent of these stratigraphic units is largely determined by Quaternary sea
level changes.
Previous geological investigations in the East Java offshore area were carried out mostly
as part of a regional study. The early work is credited to Earle (1845) who noted that the
Sunda Shelf, on which the present study area is located, has an extensive, flat sea floor. This
idea was further explored by Molengraaff (1921) who suggested that the Sunda Shelf is a
submerged peneplain resulting from eustatic sea level change. In the 1930s the Snellius
Expedition I, a large scale oceanographical exploration, was conducted in Indonesian waters.
The expedition included bathymetric profiling, the results of which prompted Kuenen (1935b;
1950) to suggest the occurrence of several distinguishable drainage patterns across the Sunda
Shelf.
b. Post 1940s
After the 1940s (post-Dutch colonisation), geological investigations in East Java were
carried out by various geological institutes. The Bataafsche Petroleum Maatschappij ( B P M )
concession and its exploration activities were taken over by Lemigas (a national oil institute),
3
and the Geological Survey of Indonesia continued the work of the Geological Survey of the
Netherlands Indies.
In the late 1960s and early 1970s, Lemigas and BEICIP (Bureau d'Etudes Industrielles et
de Cooperation de l'lnstitut Francais du Petiole) conducted a regional review of the geology
of East Java. Unfortunately, most of the results were proprietary and only a few parts were
published as a general review of the geology of East Java (e.g. D e Genevraye & Samuel, 1972;
Soetantri et al, 1973).
One of the significant achievements of geological investigations in the East Java onshore
area after the 1940s was the application of planktonic foraminiferal biostratigraphic zonation
for unravelling a chronology of deposition during the Neogene (e.g. Bolli, 1966;
Pringgoprawiro & Baharuddin, 1979; van Gorsel & Troelstra, 1981; Pringgoprawiro, 1983;
Hasjim, 1987). This application was made possible by the fact that the Neogene strata are rich
in foraminifers and highly diversified faunal assemblages, and more importantly because of the
increasing interest in using planktonic foraminifers in biostratigraphic correlation since the mid
1950s (Bandy, 1964). Although the chronology of the Neogene strata is now better
understood, sedimentation patterns in sub-basins of the East Java onshore and offshore areas
are still poorly understood.
Quaternary geological investigations in East Java have been carried out by at least two
institutes, Institut Teknologi Bandung (affiliated university) and the Geological Research and
Development Centre (a subsidiary of the Geological Survey of Indonesia). Their activities
were directed towards accurate dating of sedimentary horizons which contain hominid remains
(Sartono et al, 1981; Jacob, 1978; Hyodo et al, 1992, 1993), rather than the relationship
between Quaternary sedimentation patterns and sea level changes.
In the late 1960s and early 1970s, regional marine geological and geophysical surveys
were carried out in the Sunda Shelf by the Woods Hole Oceanographic Institution as a
contribution to the United Nations E C A F E - C C O P program. This shelf was studied as part of
reconnaissance mapping of the continental margin structure off eastern Asia (Ben-Avraham &
Emery, 1973). The study included seismic reflection and refraction profiling, gravity,
magnetics and bathymetry. The results of the Java Sea part were published in two papers by
Emery et al (1972) and Ben-Avraham and Emery (1973). Emery et al. (1972) corroborated
4
the results of the previous surveys by Dutch geologists. They agreed that many fluvial cut-
and-fill structures occur in the uppermost 50 m of the Java Sea floor, and have attributed these
to subaerial erosion of the region as a consequence of sea level lowstands during the
Pleistocene glacial maxima. A study by Ben-Avraham and Emery (1973) stressed the regional
structure of the Tertiary basement that underlies the Sunda Shelf rather than the sedimentation
pattern within the basins on this shelf.
Since the mid 1960s some major oil companies have been granted concessions in the Java
Sea and Madura Strait. Although numerous geological and geophysical studies have been
undertaken (Tanner & Kennett, 1972), the results are seldom available. Only a few papers
have been published (e.g. Najoan, 1972; Sudiro et al, 1973; Kenyon, 1977; Bishop, 1980).
The regional stratigraphy of the Java Sea was discussed by Najoan (1972), while Sudiro et al
(1973) concentrated on the Tertiary basement structural configuration of this area. Bishop
(1980) discussed the relationship between structure, stratigraphy and hydrocarbon prospects in
the eastern part of the Java Sea. They agree that northeast-trending structures controlled the
Early and Middle Tertiary sedimentation patterns, but no further evaluation has been made of
the Late Tertiary and Quaternary sedimentation.
Today, offshore geological research is mainly carried out by the Marine Geological
Institute (a subsidiary of the Geological Survey of Indonesia). Its activities are centred on sea
floor mapping, including the sedimentary strata extending a few hundred metres below the
surface (Quaternary and recent sediments). However, until now, no comprehensive study has
been undertaken using these data.
1.2 AIMS OF STUDY
The initial aim was to investigate the distribution of the Late Tertiary and Quaternary
sediments of the offshore East Java Basin (Madura Strait and southern Java Sea), beyond the
level of understanding reached by early workers through inferences based on the geology of
the inland areas and shallow soundings. Preliminary investigations highlighted the necessity
to include the onshore area to provide a comprehensive overview of the East Java Basin.
5
In general, the study concerns three different parts of the East Java Basin: the onshore (the
Rembang and Kendeng Zones), Madura Strait and the southern Java Sea. The main objectives
of the study include:
1. a critical reassessment of the sedimentation history of the onshore East Java Basin during
the Middle to Late Tertiary, focusing on the relationship between the northern (Rembang
Zone) and southern (Kendeng Zone) anticlinal zones;
2. a reinterpretation of the onshore Quaternary succession in the light of extensive new data,
determined from shallow seismic investigations carried out in Madura Strait, which forms
the eastern offshore extension of the East Java onshore area. The seismic interpretation
forms the major part of this study;
3. Seismic stratigraphic interpretation of the onshore portion of the East Java Basin and its
relationship to tectonism and global sea level changes; and
4. Seismic stratigraphic interpretation of the southern Java Sea to assess the sedimentary
sequences extending from the Rembang Zone onto the margin of the stable Sunda Shelf.
1.3 AREA OF STUDY
Geographically the study area is located between longitudes 111° and 114° East
and latitudes 5°55' and 7°40' South, situated on the southeastern part of the Sunda Shelf (Fig.
1.1). It includes the northeastern part of East Java, Madura Island, Madura Strait and the
southeastern part of the Java Sea. For the present study emphasising seismic sequence
stratigraphic interpretation, the geophysical data coverage (particularly offshore seismic data)
is relatively comprehensive in this area and includes almost the whole East Java Basin.
The study area which experiences year round maximum temperatures between 29 and
32°C, is about 600 k m distance from the Indonesian capital city, Jakarta, and can be reached
by plane to Surabaya airport or by public transport. During the offshore surveys, the research
vessel used three main ports in the study area, Surabaya, Probolinggo, and Tuban, to load and
unload research materials.
The onshore part of the study area consists of low hills commonly covered by teak forests,
and plains which are mostly covered by alluvium of two major rivers, the Solo and Brantas
Rivers. These rivers originate in the mountainous area of Quaternary volcanoes along the
longitudinal line of Java, and traverse toward the Java Sea and Madura Strait. Rapid rural
6
development in East Java has made transportation within the study area easy, as almost all
remote areas are linked by public transport. However, as in other tropical areas, geological
outcrops are rare, which has been the major constraint during this study. Geological outcrops
are mostly located in river banks and quarries. Geological field work is impractical during the
rainy season (southern hemisphere summer, November to February), as flooding caused by
both major river systems is the main hindrance.
1.4 ORGANISATION OF THESIS
This thesis comprises eight chapters. The geological discussion begins with a review of
the geology of the study area within a regional geological framework, which is presented in
Chapter Two. Chapter Three is a review of the stratigraphy of the study area, focussed
particularly on the onshore part of the study area. The sedimentology and sequence
stratigraphic discussions are presented in Chapters Four, Five, Six and Seven. These four
chapters are arranged so that each chapter is concerned with a different geographic area. This
arrangement is also necessary because of the different methods used. Multifold seismic
sections combined with some geological observations were used in the study on the northern
part of East Java (Chapter Four), while the study on the southern part of East Java (Chapter
Five) was carried out based on outcrop observations and laboratory work only. Chapters Six
and Seven, which discuss the sequence stratigraphy of Madura Strait and the Java Sea
respectively, rely heavily on shallow seismic profiles. The final chapter, Chapter Eight,
concludes and summarises the discussion presented in the preceding chapters.
7
CHAPTER TWO
GEOLOGICAL SETTING OF THE EAST JAVA BASIN
2.1 INTRODUCTION
This chapter discusses and summarises the geological setting of the East Java Basin. The
discussion concentrates on the tectonic setting in terms of plate tectonic concepts, regional
geological structures and regional depositional history. The aim of the discussion is a regional
review of the geological processes that have controlled the development of the East Java Basin
since the Early Tertiary.
2.2 TECTONIC SETTING
The East Java Basin is located on the southeastern margin of the Sunda Shelf. This shelf
constitutes most of the western Indonesian region and includes the Malay Peninsula, Sumatra
Kalimantan, Java and the region covered by sea between these islands (Fig. 2.1). The plate
tectonic setting of this region has been described in great detail by Hamilton (1977, 1979,
1988). H e reported that the western Indonesian region has experienced a complex geological
history resulting from the convergence of the Eurasian plate and the Indian-Australian plate
(Fig. 2.1). Hamilton (1977, 1979) and Baumann (1982) recognised six major morphologic
elements in western Indonesia resulting from this convergence (Fig. 2.1 & 2.2). These are:
the Sunda (Java) trench, accretionary wedge, fore-arc ridge, fore-arc basin, active volcanic arc
and back-arc/retro-arc basin (Fig. 2.2a). These elements can be traced from the southern
Andaman Sea southeastward to the Banda Sea area. To the south of Java the fore-arc ridge
and basin are submerged (Hamilton, 1977; 1979; Fig. 2.2a). The East Java Basin has in fact
developed as a retro-arc basin to the north of the active volcanic arc in central Java (Fig. 2.3).
The earliest evidence for basin development should correlate with the commencement of
the subduction system in western Indonesia. Hamilton (1977) suggested that subduction
commenced as early as Late Oligocene. Eocene and Oligocene sediments near the Eurasian
plate margin are dominated by non-volcanic facies deposited in a stable continental shelf
setting. This opinion was supported by Baumann et al. (1972) and de Genevraye and Samuel
(1972). They found that the oldest volcanic facies in Java are the Late Oligocene to Middle
8
Miocene Jampang Formation in West Java, and the Middle Miocene to Late Miocene Kerek
Formation in Central and East Java.
Dickinson (1977) suggested that in a back-arc/retro-arc basin, evolution of the basin is
controlled by regional downward flexure of the plate margin in response to local tectonic
loading at the plate edges, and sedimentation keeps pace with subsidence. This is apparently
tme for the East Java Basin where thick and relatively continuous sedimentation has occurred.
A report by Lemigas/BEICIP (1969) indicated that more than 5000 m of sediment has been
deposited in this basin since the Early Tertiary and drill holes have never reached basement.
The present study is an attempt to evaluate the depositional styles in the East Java retro-arc
basin, and focuses on the Late Tertiary and Quaternary strata where data are more available.
This study also evaluates the role of tectonics and sea level change on the depositional stratal
pattern in this setting. The wide coverage of the data used allows a comprehensive study from
the retro-arc basin to shelfal area.
2.3 REGIONAL GEOLOGICAL STRUCTURES
Most workers agree that the stress caused by the northward motion of the Indian-
Australian plate has been responsible for most of the structural deformation in eastern Java and
the surroundings (Sudiro etal, 1973; Situmorang etal, 1976; Bishop, 1980; Baumann, 1982).
Faults are the most prominent structural elements on the Sunda Shelf (Ben-Avraham &
Emery, 1973). In the eastern Java Sea (Fig. 2.2a & 2.3) the most prominent faults trend
northeastward. According to Koesoemadinata and Pulunggono (1971) these faults may have
originated as dextral strike-slip faults, but some differential vertical movement must have also
occurred.
A number of northeast-trending structural highs occur in the eastern Java Sea (Figs 2.2a
& 2.3) and are suspected to be the remnants of an older arc system (Ben-Avraham & Emery,
1973; Katili, 1989). These structural highs extend to southeastern Kalimantan where
serpentinised ultrabasic rocks as weU as radiolarian cherts are exposed, indicating a previous
subduction complex. Early Tertiary sedimentation is in fact confined within the depression
between these highs (Situmorang et al, 1976; Bishop, 1980; Baumann, 1982).
9
In eastern Java, the most prominent structure is the east-west half graben along the median
line of Java, parallel to the plate margin as shown on Figure 2.2b (its suture is now occupied
by the Quaternary volcanic deposits). This feature is very well displayed by the regional
Bouguer gravity mesh diagram (Fig. 2.4). The data for this diagram were digitised from the
gravity maps produced from regional gravity surveys conducted by the Bataafsche Petroleum
Maatschappij (pre 1940s) and by the Geological Survey of Indonesia (post 1940s). In this
diagram the Bouguer anomaly shows, at a regional scale, a progressive decrease in a southerly
direction from about +40 mgal to about -50 mgal. This is interpreted as the result of a
progressive southward deepening of basement. According to Ben-Avraham and Emery (1973)
the gravity niinimum is approximately coincident with the thickest sediments in the East Java
Basin.
Superimposed on this southward regional trend are a number of local gravity highs. On
the northern half of the diagram these highs are linearly oriented in a general east-west
direction. It is suggested that these anomalies result from vertical differential movement
along east trending faults in the basement. The seismic sections interpreted during this study
reveal that these positive anomalies correspond to Tertiary topographic highs.
On the southernmost part of the diagram the gravity highs are less oriented. They are
interpreted as the expression of magmatic activity in this region. Interpretation of the free-air
gravity anomaly over the volcanoes in Java by Ben-Avraham and Emery (1973) has indicated
the presence of a mass excess as expressed by an increase in value of up to +150 mgal.
2.4 REGIONAL DEPOSITIONAL EVOLUTION
According to van Bemmelen (1949) the oldest stratigraphic unit on the Sunda Shelf is
Palaeozoic, consisting of crystalline schists of unknown age that are products of sedimentary
deposits altered by regional metamorphism. These rocks have a very restricted distribution,
mainly on Sumatra, Kalimantan and the Malay Peninsula.
In Java, Early Tertiary sedimentary deposits are not widely observed. Outcrops of Early
Tertiary lithologies are found only in West and Central Java (Fig. 2.5), while in East Java early
Tertiary rocks are known only from well data.
10
In West Java, early Tertiary sequences are represented by about 1900 m of strata
dominated by paralic to neritic facies (Lemigas/BEICIP, 1969; Kusumahbrata, 1994). In
Lokulo, Central Java, the Eocene sedimentary sequence unconformably overlies pre-Tertiary
rocks. The sequence consists of polymict conglomerate and sandstone, with carbonaceous
material at the base and sandstone alternating with shale, marl and lenses of foraminiferal
limestone in the upper portion. In Jiwo Hill, Central Java, the Eocene strata unconformably
overlie pre-Tertiary metamorphic rocks. T w o units can be differentiated: conglomerate, marl,
sandstone and limestone of the Gunung Wungkal Formation and foraminiferal limestone and
marl of the Gamping Formation which overlies the Gunung Wungkal Formation (van
Bemmelen, 1949).
Ben-Avraham and Emery (1973) noted that the Tertiary basins on the Sunda Shelf did not
originate simultaneously. Koesoemadinata (1969) argued that the sedimentary sequences
within these basins are similar, an idea that has been supported by Baumann (1982). As in the
case of the study area, where the basin is fragmented into smaller basins, the similarity of
sedimentary sequences appears to have been strongly controlled by relative sea level changes.
Baumann (1982) recognised five sedimentation cycles in Sumatra and Java during the
Cainozoic. Each of these cycles commenced with a transgression, and ended with a phase of
volcanism and tectonism with the duration of the cycles being up to 10 M a . These cycles are:
Middle Eocene to Early Oligocene, Late Oligocene to Early Miocene, late Early Miocene to
Middle Miocene, Middle to Late Miocene, and Pliocene to Recent. Figure 2.6 displays the
tentative palaeogeography within the first four cycles as summarised by Baumann (1982). This
figure illustrates that sedimentation mostly occurred along the rim of the Sunda Shelf where
tectonic disturbance as well as volcanic activity has been dominant. The core of the shelf was
relatively stable and has been a major source of sediment, in addition to the volcanic arc
(Baumann, 1982).
The most widespread lithologic unit on the southern Sunda Shelf is an Early Miocene
reefal limestone. These deposits were followed by a regressive phase during which mudstone
was deposited in the deeper parts of the western Java Sea basins, while mudstone and
carbonate rocks accumulated in the eastern Java Sea. This regressive phase in the eastern Java
Sea was associated with uplift and folding (Ben-Avraham & Emery, 1973).
11
During the Plio-Pleistocene, marine sedimentation continued over the Sunda Shelf and
gradually changed the morphology of the basin to a relatively flat plain (Ben-Avraham &
Emery, 1973). The present seismic study in Madura Strait and the eastern Java Sea by the
author has confirmed this opinion (see Chapters 6 & 7).
Magmatic activity on the Sunda Shelf occurred throughout the Tertiary and Quaternary.
Volcanism occurs along the rim of this shelf and was thought to be associated with major
tectonic events by van Bemmelen (1949), which the m o d e m plate tectonic concepts would
suggest relate to the plate convergence between the Indian-Australian plate and Eurasian plate
(Hamilton, 1977; 1979).
2.5 PHYSIOGRAPHY
The Sunda Shelf has an extensive flat floor, bordered by a volcanic arc along the southern
and western margins. O n this shelf, water depth is generally shallower than 100 m, in contrast
with the rather complex and irregular bathymetry of the surrounding deep-ocean floor. The
flatness of the sea bottom in the Sunda Shelf was first reported by Earle (1845). After the first
Snellius Expedition in the 1930s Kuenen (1950) suggested the presence of several drainage
systems which have been submerged during the present highstand of sea level (Fig. 2.7). This
finding is very well confirmed by the present seismic stratigraphic study in the Java Sea (see
Chapter 7).
Van Bemmelen (1949) described the physiography of Java in detail (see Fig. 2.8). These
physiographic features have mostly resulted from orogenic movements during the Tertiary and
Quaternary. In East Java, where the study area is located, there are two prominent zones of
anticlinoria, the Rembang and Kendeng Zones.
The Rembang Zone has been recognised as an oil-rich zone in East Java. This zone
consists of a number of roughly east-west trending anticlinoria, alternating with alluvial plains.
The Rembang Zone has an average width of 50 k m and a height of less than 500 m above sea
level. From the Tuban-Paciran area, the folds plunge westward to the northwest of Purwodadi
(Fig. 2.8). In some places the Rembang Zone was eroded to expose Miocene strata. In its
north-eastern half, this zone is characterised by the presence of a Plio-Pleistocene reefal
limestone that overlies the Middle Miocene formations.
3 0009 03155596 9
12
The Kendeng Zone is separated from the Rembang Zone by the depression of the
Randublatung Zone (Fig. 2.8). Little is known about the geology of the Randublatung Zone
due to limited outcrop and well data. As in the Rembang Zone, the Kendeng Zone is also
characterised by a number of east-trending anticlinoria formed during the Middle Pleistocene.
The height, width and fold complexity of the Kendeng Zone gradually decrease eastward, and
it disappears under alluvial deposits in the Mojokerto area (Fig. 2.8). The anticlines are
generally eroded down to marly and clayey sequences of Late Miocene age. The southern
limit of the Kendeng Zone is a depression which has been filled with the Quaternary volcanic
products of the Solo Zone.
The character of the sedimentary units in the Kendeng Zone is considerably different from
those in the Rembang Zone. The differences are mostly expressed by their composition, which
indicates that there were two different sources of sedimentary detritus during the Tertiary and
Quaternary. Volcaniclastic and volcanogenic fragments are common in the sequences from
the Kendeng Zone, while continentally-derived sediment is common in the Rembang Zone.
The Bouguer gravity diagram (Fig. 2.4) indicates that the gravity minimum coincides with the
Kendeng Zone. As Ben-Avraham and Emery (1973) pointed out, this gravity minimum
coincides with the thickest sedimentary sequence, and continuous basin subsidence would have
occurred in the Kendeng Zone.
2.6 SUMMARY
The East Java Basin situated on the Sunda Shelf is considered to represent a retro-arc
basin. This type of basin is commonly characterised by thick sediment accumulations
(Dickinson, 1977), especially towards the arc.
The collision between the Eurasian and Indian-Australian plates has been responsible for
most of the structural deformation in and around the East Java Basin. The resulting structures
largely controUed the Tertiary and Quaternary sedimentation pattern in this region. T w o
prominent structures are displayed: the east-west basement depression running along the
median line of Java, and the northeast-southwest graben in the Java Sea between Java and
Kalimantan. In East Java the east-west depression zone contains two anticlinorial zones, the
Rembang and Kendeng Zones.
13
The oldest sedimentary strata (Early Tertiary) in the Java area are underlain by a pre-
Tertiary basement of metamorphic rocks which crop out in several places in Java. In East Java
the oldest recognised Tertiary sequence is Early Miocene. The Tertiary sedimentary sequence
in the East Java Basin is similar to the other parts of the Sunda Shelf, and was deposited in
five periods: Middle Eocene to Early Oligocene, Late Oligocene to Early Miocene, late Early
Miocene to Middle Miocene, Middle Miocene to Late Miocene and Pliocene to Recent
(Baumann, 1982).
14
15
CHAPTER THREE
STRATIGRAPHY OF THE ONSHORE EAST JAVA BASIN
3.1 INTRODUCTION
In this chapter, a review of the stratigraphy of the Middle Tertiary to Quaternary
sedimentary sequences in East Java is presented. The discussion is based on previous
investigations in this area combined with some field observations and laboratory analyses. In
this chapter, the discussion is confined to the lithofacies variation and relationship between
lithofacies. The results of biostratigraphic, climatostratigraphic and magnetostratigraphic
studies by previous investigators are also discussed to evaluate the lithofacies analysis in a
chronological context. The sedimentological aspects of the Middle Tertiary to Quaternary
strata are discussed in Chapters Four and Five. Chapters Three, Four and Five provide a
geological framework for understanding the geology of Madura Strait and the Java Sea which
will be discussed in Chapters Six and Seven respectively.
3.2 LITHOSTRATIGRAPHY
The depositional environments of formations in the Rembang and Kendeng Zones range
widely from open marine to tidal flat, with fluvial environments also developed in the Kendeng
Zone. The Rembang Zone sequence, particularly during the Middle Miocene, was strongly
influenced by continentally derived sediment. In contrast, the Kendeng Zone succession right
up to the Quaternary has been influenced by volcaniclastic detritus. Volcanic activity has been
present in the vicinity of the Kendeng Zone since the Middle Miocene (van Bemmelen, 1949;
de Genevraye & Samuel, 1972). The lithostratigraphic and lateral facies variations of these
formations are described in the following sections.
Detailed geological mapping of East Java was carried out mostly by the Geological Survey
of the Netherlands Indies in the 1930s. In the Rembang Zone where petroleum is abundant,
geological studies were also carried out by the Bataafsche Petroleum Maatschappij (a pre-
1940s Dutch petroleum company). T w o different stratigraphic subdivisions evolved,
particularly for the Rembang Zone. The stratigraphic subdivision by Trooster (1937; Table
3.1) was commonly used by investigators from the petroleum industries, while the subdivision
16
in Table 3.2 by van Bemmelen (1949) was formally used by the Geological Survey of the
Netherlands Indies. As these stratigraphic subdivisions came before formal stratigraphic
nomenclature was established, the terminology used is mixed between lithofacies characters
and geographic terms.
Pringgoprawiro (1983) proposed a revision of the terminology used by previous authors
to follow the formal stratigraphic nomenclature (Table 3.3). Pringgoprawiro's (1983) proposal
was basically a renaming of the earlier terminology. The ages of the lithological formations
were redefined and fine-tuned based on his biostratigraphic studies and those of other workers,
but no significant re-organisation of lithofacies was proposed. However, there is still a
tendency for individual workers to establish their own stratigraphic subdivisions without
supporting evidence (e.g. Ardhana, 1993).
The lithostratigraphic discussion presented in this chapter follows the stratigraphic
subdivision in Table 3.4. This stratigraphic subdivision has been modified from
Pringgoprawiro (1983; see Table 3.3 and Fig. 3.1) with some modifications based on the
biostratigraphic studies by Bolli (1966) and Blow (1969) for the Bojonegoro-1 well, by van
Gorsel and Troelstra (1981) in the Solo River section, and by the author in the area north of
Kabuh. These modifications are particularly concerned with the timing of the Tawun,
Atasangin and Sonde Formations. The validity of some formations will be discussed based
on the results from the present biostratigraphic and sedimentological studies by the author. It
has been a tradition to separate the stratigraphic discussion for the Rembang and Kendeng
Zones. Discussion presented in Chapter T w o has indicated the presence of different tectonic
and basement behaviour for these two zones as well as different sedimentary sources.
During the Middle Miocene to Quaternary seven major lithostratigraphic units developed
in the Rembang Zone: the Tawun, Bulu, Wonocolo, Ledok, Mundu, Paciran and Lidah
Formations. Nine lithostratigraphic units have been recognised in the Kendeng Zone during
the same and younger time intervals, these are: the Kerek, Banyak, Kalibeng, Atasangin,
Sonde, Lidah, Pucangan, Kabuh and Notopuro Formations. The Lidah Formation in the
Kendeng and Rembang Zones is equivalent, since it was deposited as the result of regional
base levelling at the end of the Pliocene.
17
3.2.1 Rembang Zone
a. Tawun Formation
Sediments of the Tawun Formation were deposited during the Middle Miocene, and the
stratotype is in Tawun-5 well (Bolli, 1966; Lemigas/BEICIP, 1969; Pringgoprawiro, 1983).
This formation is widely distributed along the Rembang Zone (Fig. 3.1) and Madura Island,
mosdy cropping out in the cores of anticlines. The formation may reach 1500 m in thickness
and is typified by brownish grey mudstone with interbeds of calcareous sandstone. The upper
part of the formation is organic-rich and often contains thin beds of lignite. This part is also
associated with a mappable quartz sandstone (Ngrayong Member) and relatively thin (up to 1
m ) orbitoidal (larger foraminifers) limestone beds. The Ngrayong Member has been the main
petroleum reservoir rock in the Rembang Zone.
The age of this formation was determined mainly on the basis of species of Orbitoididae
which are abundant particularly in the upper part. These species are: Lepidocyclina
atuberculata, Lepidocyclina ephippioides, Lepidocyclina sumatrensis, Lepidocyclina nipponica,
Miogypsinoides bantamensis and Cycloclypeus spp. (Pringgoprawiro, 1983). Based on the
planktonic foraminiferal study by Bolli (1966) for the Bojonegoro-1 well, Blow (1969)
indicated that the Tawun Formation is underlain by a surface of unconformity (early Middle
Miocene unconformity). The planktonic foraminiferal zone N10, and part of zones N 9 and
Nil, are missing in this well.
In the past the environment of deposition of the Tawun Formation was interpreted as
paralic (van Bemmelen, 1949; Lemigas/BEICIP, 1969; Pringgoprawiro, 1983). A recent study
by Ardhana (1993) indicated that the depositional facies varied laterally, particularly in the
upper part, from paralic to hemipelagic and was strongly controlled by the palaeogeography.
In the areas north of Cepu and Bojonegoro (Fig. 3.1) the facies are mostly paralic.
Hemipelagic facies occur in areas such as south of Cepu and Bojonegoro. The present seismic
stratigraphic study (Chapter 4) indicates that the upper part of the Tawun Formation formed
under regressive conditions.
b. Bulu Formation
Sediments of the Bulu Formation were deposited on the paralic facies of the Tawun
Formation in the late Middle Miocene. It occurs mainly between Cepu and Rembang (Fig.
3.1) and consists predominantly of reefal limestone with some interbeds of quartz sandstone
18
and calcarenitic limestone with abundant fragments of benthonic foraminifers and red algae.
The age determination was based on the presence of the benthonic foraminiferal (Orbitoididae)
species: Lepidocyclina angulosa, Lepidocyclina sumatrensis, Cycloclypeus annulatus,
Cycloclypeus indopacificus, Lepidocyclina spp. which are associated with a middle neritic
environment (Pringgoprawiro, 1983). This formation attains a thickness of 250 m.
The nature of contact of this formation with the underlying Tawun Formation is debatable,
mainly because of lack of palaeontological controls near the boundary of the formations.
Lemigas/BEICIP (1969) reported that local erosion occurred at the top of the Tawun
Formation. In contrast, based on field observations in the Bulu area, Pringgoprawiro (1983)
argued that the Bulu and Tawun Formations are conformable. The present seismic
stratigraphic study (Chapter 4) suggests that this formation was only locally developed (on the
high area between Bulu and Rembang), during a transgression following the Tawun Formation
deposition, and the subsequent highstand of sea level.
c. Wonocolo Formation
The Wonocolo Formation was deposited during late Middle Miocene to early Late
Miocene. The type locality of this formation is in the Wonocolo area, some 15 k m northeast
of Cepu (Fig. 3.1; Pringgoprawiro, 1983). It is characterised primarily by unstratified
Globigerina-tich marls and clayey marls with rare intercalations of calcarenite. The Wonocolo
Formation is characterised by a 20 to 8 0 % planktonic/benthonic foraminiferal ratio
(Pringgoprawiro, 1983), which indicates an inner to outer shelf depositional environment
(Murray, 1991). Most of this formation developed during a highstand of sea level following
sea level fall in the late Middle Miocene (see Chapter 4). The Wonocolo Formation attains
a thickness exceeding 300 m and shows a coarsening-upward sequence into the overlying
Ledok Formation.
In the areas between Cepu and Rembang this formation conformably overlies or is laterally
equivalent to the Bulu Formation. The Wonocolo Formation in northeastern Java (Tuban and
Paciran areas, Fig. 3.1 & 3.2) does not occur and Pliocene reefal limestones (Paciran
Formation) rest directly on the Middle Miocene Tawun Formation (north of Bojonegoro) and
the Early Miocene Kujung, Pmpuh and Tuban Formations (in the Tuban and Paciran areas).
This hiatus indicates the presence of a topographic high during the Late Miocene in these
areas.
19
d. Ledok Formation
The Ledok Formation conformably overlies the Wonocolo Formation. In the area
northeast of Cepu, the transition between the formations is characterised by an increasing
frequency of calcarenite intercalations within the marl. This upward trend is interpreted to
indicate a shallowing environment of deposition, closer to a high energy environment. In the
area between Bojonegoro and Blora, the upper part of the formation developed as a tidalfy-
influenced deposit, as indicated by abundant large-scale cross-beds of probably tidal sand
waves. The planktonic foraminiferal species Globorotalia plesiotumida is common in the
Ledok Formation and indicates Late Miocene age (Pringgoprawiro, 1983).
In the core of an anticline north of Ngimbang (Fig. 3.1) the equivalent Ledok Formation
is characterised by unbedded marl and marly clay. In the Java Sea, the equivalent unit to the
Ledok Formation is characterised by reefal limestone (Najoan, 1972). In contrast, the Ledok
Formation as well as the Wonocolo Formation was not developed in the Tuban and Paciran
areas (Duyfjes, 1938a) suggesting that topography remained high during the deposition of both
formations.
Pringgoprawiro (1983) found that the planktonic/benthonic foraminiferal ratio decreases
upwards through the formation (from 4 7 % to 30%) which also suggests a shallowing
environment of deposition. The lower part of the formation is typified by benthonic
foraminifers: Eponides praecintus, Nodogenerina scalaris, Uvigerina peregrina, Uvigerina
flintii, Bulimina striata, Laticarinina pauverta, Robulus orbicularis, Epistoma bradyi, Dentalina
neugobarina, Nodosaria spp., Elphidium spp., Bulimina marginata. This faunal association
indicates an inner shelf environment of deposition (Murray, 1991).
e. Mundu and Paciran Formations
The type locality of the Mundu Formation is in the Mundu area, 10 k m west of Cepu (Fig.
3.1). It was deposited conformably on the Ledok Formation during the Early to Middle
Pliocene (van Bemmelen, 1949; Lemigas/BEICIP, 1969; Pringgoprawiro, 1983). This
formation is composed of unstratified yellowish-white to greenish-grey marl. In some intervals
the Mundu Formation shows a sandy character due to a high content of foraminiferal tests.
The formation is also typified by an association of outer neritic benthonic foraminifers. The
original thickness of the Mundu Formation is uncertain due to post-Mundu erosion but 10 k m
north of Cepu, it attains a thickness of 340 m (Lemigas/BEICIP, 1969).
20
The Mundu Formation occurs widely in the East Java Basin. Major facies changes to
reefal and calcarenitic limestone occur mainly along topographically high areas such as the
Tuban area and on Madura Island. In the Tuban area (Figs 3.1 & 3.2) the limestone is termed
the Paciran Formation (Pringgoprawiro, 1983) or the Karren Limestone (van Bemmelen, 1949),
and on Madura Island is termed the Madura Formation (Situmorang et al, 1992; Aziz et al,
1993). Along the southern coast of Madura Island the equivalent unit to the M u n d u Formation
ranges from marls to calcarenite with common bioclasts of foraminiferal tests and algal
fragments.
The Mundu Formation in the area between Bojonegoro and Pati exhibits a regressive
sequence upward from outer neritic to inner neritic. The upper part of the M u n d u Formation
is composed of coquina intercalated with marl (Aminuddin et al, 1981). The bioclasts are
commonly micritised smaller foraminifers and algal fragments. The unit is also characterised
by the presence of interstitial glauconite embedded within foraminiferal tests. Pringgoprawiro
(1983) assigned this limestone to the Selorejo Member of the Mundu Formation. The Selorejo
Member thickens gradually from only few metres in the Bojonegoro area, to approximately 300
m in the southern Rembang area.
The Selorejo Member is also characterised by typical inner neritic benthonic foraminifers
including the genera Nonion, Quinqueloculina and Ammonia/Rotalia (Phleger, 1960; Murray,
1991). Aminuddin et al. (1981) concluded that the Selorejo Formation was deposited during
the Late Pliocene based on the presence of the planktonic foraminifers Globorotalia tosaensis,
Globorotalia acostaensis, Globigerinoides trilobus, Globigerinoides sacculifer,
Globigerinoides immaturus, Pulleniatina praecursor, Pulleniatina obliquiloculata and Orbulina
universa.
f. Lidah Formation
The type locality of the Late Pliocene-Pleistocene Lidah Formation is in the Lidah
Anticline, 10 k m southwest of Surabaya (Fig. 3.1; Pringgoprawiro, 1983). In the Kendeng
Zone the Lidah Formation was known as the clay facies of the Pucangan Formation (Duyfjes,
1938a,b,c,d). In the Rembang Zone it was called the Mergel Ton by the Bataafsche Petroleum
Maatschappij (Table 3.1.) or the Blue Clays by van Bemmelen (1949; Table 3.2.).
21
The Lidah Formation is widely developed in the eastern part of the Kendeng Zone and the
southern and western parts of the Rembang Zone (Fig. 3.1). It consists primarily of
unstratified bluish-grey mudstone of lacustrine/tidal flat origin. The formation thickness varies
between 300 and 550 m In the eastern part of Kendeng Zone, this variation is particularly
marked because of a lateral facies change between the upper part of the Lidah Formation and
the volcanic facies of the lower part of the Pucangan Formation.
The field study indicated that the lower contact with the underlying Atasangin Formation,
as observed near Kabuh and Sumberringin north of Jombang (Kendeng Zone), is conformable
as shown by a gradual change from marl to calcareous mudstone and bluish grey mudstone.
In contrast, in the Rembang Zone the Lidah Formation was deposited on the eroded Mundu
Formation (Lemigas/BEICIP, 1969), which indicates a later development of the Lidah
Formation than in the Kendeng Zone. A similar erosional contact is observed along the south
coast of Madura Island, where the mudstone Lidah Formation rests directly on the tidal sand
wave deposit Mundu Formation equivalent (Pasean Formation).
In areas southeast of Cepu and north of Bojonegoro the lower part of the formation is
carbonaceous and pyritic. The middle part consists of glauconitic calcarenite with thin
interbeds of marl which were called the Malo Member by Trooster (1937) and
Lemigas/BEICD? (1969). The upper part is dominated by calcareous mudstone and marl with
interbeds of coquina up to 2 m thick. The Lidah Formation can also be found in some places
along the south coast of Madura Island (east of Kamal and south of Pamekasan, Fig. 3.2)
where it lies on calcareous muddy sandstone of the Mundu Formation equivalent.
Field observations in the Sumberringin (Kedungwam and Pucangan Anticlines) and
Kedamean (Guyangan Anticline) areas reveal two prominent sandy intervals of up to 27 m in
thickness. These intervals resemble a tidally influenced unit and contain marine molluscs
(Gastropods, Echinoides and Arthropods [Balanus]; Duyfjes, 1938b,d), as well as benthonic
foraminifers which are diagnostic of an inner neritic environment (offshore facies). These
indicate that there were two marine incursions during the deposition of the Lidah Formation
in this area.
22
3.2.2 Kendeng Zone
a. Kerek Formation
The Kerek Formation was deposited during the Middle to early Late Miocene. The type
locality is in Kerek Village (8 k m north of Ngawi) along the Solo River (de Genevraye &
Samuel, 1972). This formation crops out only in the cores of anticlines in the Kendeng Zone
(Fig. 3.1) with a thickness attaining 1000 m. The main components are medium-grained
calcareous sandstone alternating with marl. Benthonic foraminifers such as Eggerella bradyi,
Melonis soldanii and Globobulimina pupoides found within this formation (Van Gorsel &
Troelstra, 1981) suggest an outer neritic environment.
Texturally the Kerek Formation shows a turbiditic character and van Bemmelen (1949)
referred to it as a flysch deposit, whereas Muin (1985) regarded the formation as a submarine
fan deposit. The abundance of volcanogenic coarse clastic fragments in this formation suggests
a close association with the Middle Miocene volcanism [Banyak Volcano, van Bemmelen
(1949)], which was present along the place presently occupied by Quaternary volcanoes. A
sedimentological reinterpretation during the present study suggests that the Kerek Formation
may not be associated with a submarine fan system. The palaeogeographic setting indicates
that the formation was probably deposited on the apron of volcanoes. In addition, benthonic
faunal analysis by Lemigas/BEICIP (1974) indicated deposition in an inner-middle neritic
environment which is seldom characteristic of a submarine fan system (Walker, 1992). It is
suspected that the turbiditic character of this formation may have resulted from storm-induced
currents or been directly related to the volcanic activity. Similar settings to the latter have
been reported by Busby-Spera (1988) in Baja, California (Mesozoic deposit) and Hathway
(1994) in southwestern Viti Levu, Fiji (Oligocene-Miocene deposits). Towards the top of the
Kerek Formation, the environment of deposition was more commonly inner neritic, as indicated
by the occurrence of storm deposits.
b. Kalibeng Formation
The Late Miocene to Early Pliocene Kalibeng Formation is characterised by greenish or
greyish white unstratified foraminiferal-rich marl. Bedding is always difficult to observe
except where intercalations of calcarenitic sandstone are present. The type locality of the
Kalibeng Formation is along the Kalibeng River, 15 k m northwest of Jombang, East Java (Fig.
3.1). The Kalibeng Formation was formerly termed the Kalibeng Beds by Duyfjes (1938b)
which were subdivided into the Lower (unstratified marl) and Upper (limestone and marl)
23
Kalibeng Beds. D e Genevraye and Samuel (1972) renamed them the Lower and Upper
Kalibeng Formations (note the use of term "Formation"). Based on the age and lithofacies
differences, Pringgoprawiro (1983) proposed that the term Kalibeng Formation be retained for
the Lower Kalibeng Formation and the new Sonde Formation be used for the Upper Kalibeng
Formation.
Along the Solo river to the north of Ngawi, some 500 m of Kalibeng Formation occurs
which, as in the type locality of this formation, is mainly composed of marl. A few thin
intercalations of calcarenite and calcareous fine sandstone are present. In this section the
Kalibeng Formation is also very rich in well preserved foraminifers. Benthonic forarninifers
are all open marine species, such as Oridorsalis umbonatus, Gyroidina neosoldanii, Planulina
wuellerstorfi, Anomalina globulosa and Uvigerina auberiana (van Gorsel & Troelstra, 1981).
In the Atasangin area (some 40 km east of Ngawi; Fig. 3.1), up to 200 m of volcaniclastic
deposits occur in the upper part of the Kalibeng Formation. Pringgoprawiro (1983) suggested
that this represented a new member of the Kalibeng Formation, and he called it the Atasangin
Member.
The stratigraphic relationship between the Kalibeng Formation and the adjacent formations
is difficult to define due to lack of bedding structures within the Kalibeng Formation. De
Genevraye and Samuel (1972) and Pringgoprawiro (1983) considered that the Kalibeng
Formation has a conformable relationship with both the Kerek and Sonde Formations. Then-
suggestion was based on the similarity in bedding orientation and faunal content in the
Kalibeng, Kerek and Sonde Formations.
c. Sonde and Atasangin Formations
The term Sonde Formation was proposed by Pringgoprawiro (1983) to replace the term
Upper Kalibeng Beds of Duyfjes (1938a,b) and van Bemmelen (1949), and the Upper Kalibeng
Formation of de Genevraye and Samuel (1972). This formation was deposited in the Kendeng
Zone during the Late Pliocene. The type locality of the Sonde Formation is in the Sonde area,
along the Solo River, 15 k m west of Ngawi (Fig. 3.1.). It is characterised mainly by coarse
bioclastic limestone in the lower part and marl in the upper part. The bioclastic limestone has
been called the Klitik Member of the Sonde Formation (Pringgoprawiro, 1983).
24
About 65 m of limestone in the Sonde Formation occurs in the Solo River section north
of Ngawi. It shows a bedded calcarenite at the bottom and reefal limestone at the top. Van
Gorsel and Troelstra (1981) considered the Sonde Formation as a regressive unit. Their
benthonic foraminiferal analysis of the formation revealed an assemblage of middle shelf
faunas at the base, and inner shelf faunas in the middle part of the formation. The upper part
of the formation consists of about 80 m of sandy marl. The Sonde Formation in the Solo
River section is unconformably overlain by 120 m of debris flow deposits of the Pucangan
Formation.
In the eastern part of the Kendeng Zone this formation was developed in a completely
different setting than in the type locality. The present study by the author (Chapter 3.3, and
Chapter 5.4) suggests that in this area the Sonde Formation was developed in a relatively
isolated basin. It can be divided into two main lithofacies, a sandstone facies (about 100 m
thick) in the lower part and a marly facies (about 300 m thick) in the upper part. The
sandstone facies consists primarily of prograding deltaic strata which rest conformably on the
Kalibeng Formation. The marly facies consists of finely laminated diatomaceous marl with
a few thin interbeds (15 cm) of fine sandstone. Based on the occurrence of benthonic
foraminifers and facies characters, the marly facies is thought to have been deposited in an
marginal marine environment (Chapter 5.4). The planktonic foraminiferal evidence suggests
that the depositional tinting of these facies is slightly earlier (Middle Pliocene, Chapter 3.3)
than in the type locality (Late Pliocene).
The stratigraphic position of the sandstone facies is similar to the volcaniclastic unit of the
Kalibeng Formation (Atasangin Member) found by Pringgoprawiro (1983) in the Atasangin
area (some 40 k m east of Ngawi; Fig. 3.1). Petrographical evidence suggests that the
sandstone facies are volcanogenically derived (Chapter 5.4), and probably of similar origin to
the Atasangin Member of the Kalibeng Formation. The marl facies has no equivalent in the
type locality of the Sonde Formation. These results necessitate a reexamination of the validity
of the Sonde Formation in the eastern part of the Kendeng Zone. In the present study the
author proposes the term Atasangin Formation due to the differences in texture and age from
the Sonde Formation in its type locality, and partly due to the facies similarity with the
Atasangin Member of the Kalibeng Formation. The Atasangin Formation includes the
Atasangin Member of Pringgoprawiro (1983).
25
d. Pucangan Formation
The type locality of the Pucangan Formation is on the south flank of the Pucangan
Anticline (18 k m north of Jombang; Fig. 3.1). It was proposed by Duyfjes (1938a,b,c,d) to
consist of two main facies: bluish-grey mudstones of paludal origin in the lower part and
volcanic facies in the upper part. As discussed earlier (Chapter 3.2.1f), the lower part is
considered to be the Lidah Formation and the upper part remains as the Pucangan Formation
(Pringgoprawiro, 1983).
The Pucangan Formation developed mainly in the Kendeng Zone during the Early
Pleistocene. This formation consists of volcanic deposits and volcanic-derived sediment
deposited in two different settings: tidally-influenced and alluvial fan/fluvial environments.
According to Duyfjes (1938b,d) all of these volcanic materials were derived from the Wilis
Volcano which was the only active volcano during the Early Pleistocene. A hominid skull
fossil and a number of vertebrate fossils have been found within this formation, particularly
in the conglomeratic facies related to the alluvial fan and fluvial systems (Duyfjes, 1938d).
e. Kabuh Formation
The term Kabuh Beds was used by Duyfjes (1938b,d) for the tithologic unit that
conformably overlies the Pucangan Formation. Pringgoprawiro (1983) termed this unit the
Kabuh Formation. It is only developed along the Kendeng Zone and was deposited during the
Middle Pleistocene. It comprises two main lithofacies: marine mudstone and sandstone facies
in the lower part and volcanic facies in the upper part.
The base of the widespread mudstone facies marks the boundary between this formation
and the underlying Pucangan Formation (Duyfjes, 1938d). The present study (Chapter 5.8)
indicates that this facies conformably overlies the fluvial facies of the Pucangan Formation and
represents a shelf mud facies which marked a significant relative sea level rise in East Java
during the Middle Pleistocene. The marine facies of the Kabuh Formation is well exposed on
the flanks of anticlines north of Mojokerto with a thickness of about 15 m. A similar facies
is also well developed in an anticline west of Pasuruan (see Fig. 3.1) where the marine facies
consists of green mudstone in the lower part and mollusc-rich muddy sandstone in the upper
part.
26
In some areas (i.e. near Kabuh and Sumberringin, 18 k m north of Kabuh) the marine
mudstone is absent and the volcanic facies of the Kabuh Formation lies directly on the
Pucangan Formation. The volcanic facies was deposited along the southern edge of the
Kendeng Zone. It consists of conglomeratic cross-bedded sandstone in which the fragments
are mostly andesitic. The volcanic facies is also characterised by the presence of fresh-water
molluscs (Melania and Unio) and abundant vertebrate fossils (Duyfjes, 1938b,d).
f. Notopuro Formation
The Notopuro Formation (Pringgoprawiro, 1983; formerly the Jombang Beds of Duyfjes,
1938b) lies unconformably on the Kabuh Formation. It consists mainly of volcanic breccia,
tuff and tuffaceous sandstone and was deposited in the proximity of Late Pleistocene volcanoes
(i.e. Lawu, Wilis and Arjuno volcanoes). The unconformable relationship with the underlying
formation resulted from Late Pleistocene folding which also produced most of the anticlinoria
in the Kendeng Zone (Duyfjes, 1938b; van Bemmelen, 1949).
3.3 BIOSTRATIGRAPHY
Biostratigraphic studies have been undertaken by several investigators on Miocene to
Pleistocene strata in various places in East Java (e.g. Bolli, 1966; Pringgoprawiro &
Baharuddin, 1979; van Gorsel & Troelstra, 1981; Pringgoprawiro, 1983; Hasjim, 1987). Most
of these studies were based on planktonic foraminifers which are abundant in these strata.
Two commonly used planktonic foraminiferal zonations are the one by Bolli (1970) and
Bolli and Premoli-Silva (1973), and the other by Banner and Blow (1967) and Blow (1969).
Both zonations are similar, based on the first and/or last appearances of species. However, the
zonation of Blow (1969) places greater emphasis on changes within evolutionary lineages, and
zonal boundaries are assigned on the critical level within lineages. Bolli (1970) and Bolli and
Premoli-Silva (1973) divided the Neogene into 19 major zones, with further subdivision carried
out in the Early and Middle Pliocene, and Pleistocene (Fig. 3.3a). Banner and Blow (1967)
and Blow (1969) subdivided the Neogene into 20 zones (Fig. 3.3b), and numerically labelled
from N 3 (Late Oligocene) to N23 (Late Pleistocene). The use of this numerical identification,
in some respects, is more convenient than taxonomical identification as practiced by Bolli
(1970) and Bolli and Premoli-Silva (1973) since stratigraphical order is more easily
remembered. This is probably the reason for the wide use of the Blow (1969) zonation in the
27
Indonesian region. Until recently, no significant revisions have been made to these zonations,
but attempts to numerically date the zonal boundaries have been carried out by Berggren
(1972; 1973) and Berggren et al. (1985), particularly using integrated data of deep sea cores.
The Miocene/Pliocene boundary has been variously placed at the base (Berggren et al,
1985), middle (Banner & Blow, 1967) or top (Blow, 1969) of zone N18. The numeric age for
the base of zone N 1 8 is 5.2 M a (Berggren et al, 1985; Harland et al, 1989). The top of zone
N18 is the evolutionary level of Sphaeroidinellopsis subdehiscens to Sphaeroidinella dehiscens
dehiscens, dated 5.0 M a (Berggren et al, 1985). The Pliocene/Pleistocene boundary is
coincident with the boundary of zones N21/N22, which is at the evolutionary level of species
Globorotalia tosaensis to Globorotalia truncatulinoides (Banner & Blow, 1967). Berggren et
al (1985) and Harland et al (1989) provided an age of 1.64 M a for this boundary.
3.3.1 Miocene and Middle Pliocene strata
The Miocene and Pliocene strata in the East Java Basin are rich in foraminifers. The first
biostratigraphical study carried out in the East Java area was by Bolli (1966) for the
Bojonegoro-1 well (Fig. 3.2). It was followed by a study by Pringgoprawiro and Baharuddin
(1979) for the TO-5 and TO-8 wells in the Tobo area 15 k m south of Bojonegoro. Van Gorsel
and Troelstra (1981) carried out detailed biostratigraphic and climatostratigraphic studies in the
Solo River section north of Ngawi (Fig. 3.2). Biostratigraphic studies on some surface sections
in the Rembang Zone have also been carried out by Pringgoprawiro (1983) and Hasjim (1987).
The discussion below summarises the results of their investigations and estimates the timing
of lithostratigraphic development. The discussion below also refers to the planktonic
foraminiferal zonation developed by Blow (1969).
a. Solo River section
Biostratigraphy of the Solo River section, north of Ngawi, has been studied by
Lemigas/BEICIP (1969) and van Gorsel and Troelstra (1981). These studies were particularly
concerned with the Kalibeng Formation and the underlying Kerek Formation. Figure 3.4
summarises the biostratigraphic and climatostratigraphic studies in this section by van Gorsel
and Troelstra (1981).
The zonal boundaries erected in both studies have been defined by the first appearance of
planktonic foraminiferal species and by the first appearance of the succeeding species. Van
28
Gorsel and Troelstra (1981) have not indicated whether these appearances are evolutionary or
migratory. However, they suggested that the species represent a reliable record of the changes
in planktonic faunas at the time of deposition. This suggestion was based on the supposition
that the Kerek and Kalibeng Formations were deposited in an open marine environment with
no indication of transported or reworked faunas.
The Miocene-Pliocene boundary (N18/N19 boundary) in the section is about 150 m above
the base of the Kalibeng Formation (Fig. 3.4). This boundary was determined by van Gorsel
and Troelstra (1981) using the first appearance of Sphaeroidinella dehiscens immatura. The
Pliocene-Pleistocene boundary (N21/N22 boundary) is determined by the first appearance of
Globorotalia truncatulinoides and falls in the lower part of the Sonde Formation.
b. Puncakwangi and Kali Sampurna sections
Figures 3.5 and 3.6 summarise the biostratigraphic studies from the Puncakwangi and Kali
Sampurna sections respectively by Pringgoprawiro (1983). He used the same species as van
Gorsel and Troelstra (1981) to determine the Miocene-Pliocene and Pliocene-Pleistocene
boundaries.
On these sections the Miocene-Pliocene boundary has been defined at the boundary
between the Ledok and Mundu Formations which coincides with the first appearance of the
planktonic foraminiferal species Sphaeroidinella dehiscens immatura. The first appearance of
Globorotalia truncatulinoides, which marks the Pliocene-Pleistocene boundary, occurs at the
boundary between the Mundu and Lidah Formations.
c. North Kabuh section
A biostratigraphical analysis of the north Kabuh section was also undertaken in the present
study (Fig. 3.2). This section consists of 400 m of sediment represented mostly by the
Atasangin Formation. The lower part of the section consists of the deltaic deposits of the
Atasangin Formation underlain by the Kalibeng Formation. Figure 3.7 summarises the
biostratigraphic analysis of this formation based on planktonic foraminifers in the 12 samples
analysed. The results of the analysis should not be regarded as very accurate because of two
major problems. These are: (1) that the analysis was based on a relatively wide sampling
interval, and (2) the lithofacies, which was deposited in a marginal marine environment (see
discussion in Chapter 5.4), is characterised by a low fauna! diversity.
29
The analysis shows that the whole section is Pliocene in age. Deposition of the Atasangin
Formation apparently commenced in the early N 2 0 zone as indicated by foraminiferal species
contained in the uppermost Kalibeng Formation in this area. The N20/N21 boundary in this
section is determined by lack of Pulleniatina primalis and Pulleniatina praecursor in the
samples from the upper part of the formation. This boundary is about 50 m below the
suspected boundary between the Atasangin and Lidah Formations.
3.3.2 Late Pliocene and Pleistocene strata (Lidah Formation)
Since the discovery of vertebrate and horninid fossils in the Pucangan, Kabuh and
Notopuro Formations in Central and East Java by Dubois in 1891 and von Koenigswald in
1936, as reported by Duyfjes (1936), and subsequent workers (Sartono, 1961; Jacob, 1978),
the Quaternary sediments in the Kendeng Zone have been the subject of various stratigraphic
studies.
In the eastern Kendeng Zone Ninkovich and Burckle (1978) analysed and compared the
diatom assemblages studied by Reinhold (1937) with the diatom assemblages from the
palaeomagnetically-dated deep sea core V28-179 located in the central Pacific (Shackleton &
Opdyke, 1977). The Reinhold (1937) samples were taken from the upper part of the Atasangin
Formation, the marine sandstone intercalated within the mudstone of the Lidah Formation and
from the transitional strata between these formations (Fig. 3.8). Ninkovich and Burckle (1978)
have arrived at an age of around 2.1 M a for the base of the Lidah Formation. This result is
in close agreement with the present biostratigraphic study on the north Kabuh section (Fig.
3.7). The latter study found that the boundary between the Atasangin and Lidah Formations
lies about 50 m above the N20/N21 boundary. According to Berggren et al. (1985) the age
of the N20/N21 boundary is about 3 Ma. Sartono et al (1981) also arrived at a similar result.
They carried out micropalaeontological analysis on the Lidah Formation in Perning area (10
k m north of Mojokerto, Fig. 3.1). The zone N21 age for the Lidah Formation was based on
the planktonic foraminiferal species Globigerinoides extremus, Globorotalia cultrata exilis and
Globorotalia tosaensis.
3.4 CLIMATOSTRATIGRAPHY
A climatostratigraphic study was carried out by van Gorsel and Troelstra (1981) on the
Kalibeng Formation in the Solo River section. Their study was based on the morphological
30
characters, coiling behaviour and abundance of some planktonic foraminiferal species which
are sensitive to climatic change.
A morphological character change in response to varying climatic conditions has been
shown by species Orbulina. Based on samples from surficial sediments and Indian Ocean
waters, Be et al. (1973) has demonstrated that the diameter of the globular form of Orbulina
is inversely correlated with latitudinal occurrence. A faunal test diameter between 600 to 800
p m is c o m m o n in the tropical and subtropical areas, and less than 450 u m diameter occurs in
the subtropical-subpolar transition zone. The species Globorotalia (Neogloboquadrina)
pachyderma is known to be a typical species in the polar and subtropical regions (Kennett &
Srinivasan, 1983). This species according to Jenkins (1967), Kennett (1968) and Bandy (1972)
is sinistrally coiled in polar and subpolar regions, and dextrally coiled towards lower latitude.
These indices have been used by van Gorsel and Troelstra (1981) to reconstruct the climatic
events in the Solo River section. Cold events are indicated by a significant drop in Orbulina
diameter and abundance increase of sinistrally coiled Globorotalia pachyderma. Additionally,
van Gorsel and Troelstra (1981) also used the relative abundance of some other species which
are typical in tropical regions (Globorotalia menardii, Globigerinoides trilobus, Globoquadrina
altispira, Pulleniatina spp.) and subpolar regions (Globigerina bulloides) regions.
Van Gorsel and Troelstra (1981) revealed the presence of five climatic zones (Fig. 3.4).
Zones I, m and V are cold periods and have been interpreted to be associated with sea level
lowering during the early Late Miocene, end of Late Miocene and Pliocene. The base of zone
V coincides with the transition from the Kalibeng Formation to the Sonde Formation. Based
on the correlation with planktonic foraminiferal datums determined from several deep sea cores
from the equatorial Pacific, the numerical ages for the tops of zones I and UI are 5.8-6.1 M a
and 5.0-4.8 M a respectively, and for the base of zone V is 2.6-2.8 M a , according to van Gorsel
and Troelstra (1981).
3.5 MAGNETOSTRATIGRAPHY
In the Perning area (10 km north of Mojokerto, Fig. 3.1), palaeomagnetic investigations
were carried out in 1986-1988 by a Japanese and Indonesian joint research team (Hyodo et al,
1992). This investigation focussed on the Late Pliocene and Pleistocene strata, with the
primary aim of dating the hominid fossil bearing strata.
31
Magnetostratigraphy is based on variations in remanent magnetisation within a sequence
of strata, in response to polarity reversals of the geomagnetic field. The present normal
polarity of the Earth's geomagnetic field (Brunhes Epoch) has persisted since 0.78 M a (Spell
& McDougall, 1992), a revision of the value of 0.73 M a of Berggren et al. (1985). A reverse
epoch (Matuyama) occurred between 2.47 and 0.78 M a (Berggren et al, 1985). There are
three short geomagnetic events of normal polarity within this Matuyama Epoch, which are
known as the Reunion (2.04 Ma), Olduvai (1.88-1.72 Ma) and Jaramillo (0.94-0.88 M a ) events
(Berggren et al, 1985).
The investigations by Hyodo et al (1992) in the Perning area, obtained reversed or
intermediate magnetisations in most horizons of the Pucangan and Kabuh Formations. Two
zones dominated by easterly magnetisation directions have been recognised and correlated to
the Olduvai and Jaramillo events for the lower and upper zones respectively. The upper zone
coincides with the sediment layer that contains hominid fossils. The Bmnhes/Matuyama
boundary is situated near the boundary between the Pucangan and Kabuh Formations. The
interpretation of Hyodo et al (1992) was based on the similarity of the geomagnetic
declination swing pattern to that in the previously defined Quaternary succession in Sangiran,
Central Java (some 20 k m north of Solo; Fig. 3.2).
The results of the palaeomagnetic study differ from the results of K-Ar dating by Jacob
and Curtis (1971) in the same area (Perning, Mojokerto). The radiometric dating was carried
out on a tuff which lies just below the hominid layer and gave a date of 1.9 ± 0.4 Ma. Hyodo
et al (1992) found that the tuff deposit near the hominid layer is secondary and this was
confirmed by the section measured during this study. Consequently, the radiometric dating
appears to indicate the age of the primary tuff which is older than the Jaramillo event identified
by palaeomagnetic measurements.
3.6 SUMMARY
Seven main lithostratigraphic units occur in the Rembang Zone and range in age from
Middle Miocene to Late Quaternary. These include: the Tawun, Bulu, Wonocolo, Ledok,
Mundu, Paciran and Lidah Formations. Nine lithostratigraphic units have been recognised in
the Kendeng Zone during the same and younger time intervals, these are: the Kerek, Banyak,
Kalibeng, Atasangin, Sonde, Lidah, Pucangan, Kabuh and Notopuro Formations. The
32
biostratigraphic studies have allowed an accurate lithofacies correlation between the Rembang
and Kendeng Zones. Figures 3.9 and 3.10 provide the north-south lateral correlation between
these zones.
The sedimentary origin of the Rembang and Kendeng Zones is different. This is
particularly obvious in the Middle Miocene deposits of both zones. The paralic and shallow
marine Tawun and Bulu Formations contain continentally derived elastics, such as quartz
sandstone, while the Kerek and Banyak Formations were sourced from volcanism along the
south of the Kendeng Zone. The Late Miocene and Pliocene deposits in both the Rembang
and Kendeng Zones (Wonocolo, Ledok, Mundu, Paciran, Kalibeng, Atasangin and Sonde
Formations) are highly calcareous (marly) which suggests deposition mainly in an open marine
environment during a period of reduced terrigenous supply.
The widespread lacustrine/tidal flat mudstone of the Lidah Formation marked the end of
Pliocene deposition in the Kendeng Zone, but the deposition of this mudstone continued during
the Pleistocene in the Rembang Zone, while coarse clastic deposits (Pucangan, Kabuh and
Notopuro Formations) associated with the Quaternary volcanism were developing in the
Kendeng Zone.
33
CHAPTER FOUR
SEISMIC STRATIGRAPHY OF THE REMBANG ZONE
4.1 INTRODUCTION
The Rembang Zone is known as the petroleum-rich province of East Java. Intensive
geological studies since the end of the 19th century have been undertaken principally by
petroleum companies. Published geological reviews were mostly concerned with the general
geology of the area. Detailed micropalaeontological studies, however, have been carried out
by Bolli (1966), Pringgoprawiro and Baharaddin (1979), Pringgoprawiro (1983) and Hasjim
(1987). These studies were concerned with the constmction of biostratigraphic zonations for
the Middle Miocene to Pliocene strata. Other detailed sedimentological studies which
concentrated on the Middle Miocene strata were carried out by Muin (1985) and Ardhana
(1993).
The objectives of this chapter are: (1) to delineate the chronostratigraphic units on seismic
profiles; (2) to infer the palaeoenvironment and relative sea level changes under which the
units were deposited; and (3) to compare the seismic stratigraphy and facies with the
stratigraphy and facies defined by previous authors. This study is based primarily on the
recognition of sedimentary facies defined on the basis of a seismic stratigraphic study.
Unfortunately, only two regional seismic lines were made available for the study, which
practically eliminates possible inferences about the lateral facies distribution. These two
seismic lines have been provided by PPT Migas (an Indonesian petroleum and gas educational
institute) and Pertamina (the state oil and gas company); both lines are located on East Java
(Fig. 4.1).
The seismic stratigraphic analysis was carried out by applying the seismic stratigraphic
method proposed by Vail et al (1977), Posamentier et al (1988) and Posamentier and Vail
(1988). The interpretation of environmental facies from seismic facies units follows the criteria
of Sangree and Widmier (1977), supported by petrographic analysis on selected rock samples
obtained during field observations, and by the benthonic faunal analysis of Pringgoprawiro
(1983) and Aminuddin et al. (1981).
34
4.2 SEISMIC SEQUENCE AND FACIES DEFINITION
Vail et al. (1977) developed the concept of seismic stratigraphy as a method of analysing
depositional systems which involves a genetic approach for interpreting seismic reflection
profiles. Seismic stratigraphy assumes that seismic reflections derived from sedimentary layers
are essentially isochronous, which permits a grouping of seismic reflections into packages of
sequences based on their stratal pattern. Mitchum et al. (1977) defined a depositional sequence
as "a stratigraphic unit composed of a relatively conformable succession of genetically related
strata bounded by unconformities or their correlative conformities". Vail et al (1984)
recognised two major types of unconformities which they called type I and type n
unconformities. A type I unconformity is formed during subaerial exposure of the shelf
produced by a rate of eustatic fall that exceeds the subsidence rate. In contrast a type U
unconformity is formed when the rate of eustatic fall is less than the rate of subsidence and
no subaerial exposure occurs. Sequences are named according to their type of underlying
unconformity.
Mitchum and Vail (1977) outlined three generalised steps in studying stratigraphy through
seismic data. These are:
(1) recognition, correlation, and age determination of seismic sequences
(2) recognition, mapping, and interpretation of seismic facies, and
(3) regional analysis of relative changes of sea level.
Seismic sequences are defined by recognising surfaces of discontinuity from reflection
terminations. Surfaces of discontinuity are recognised by interpreting systematic patterns of
reflection termination along the surfaces (Fig. 4.2; Vail, 1987). Mitchum et al. (1977, p. 57-
59) have categorised a set of relationships between seismic reflectors (summarised in Figure
4.3) and the following are their explanations for the terms used.
Lapout is the lateral termination of a stratum at its original depositional limit. This is a general term for lateral termination of strata at their depositional pinchout. Baselap is lapout at the lower boundary of a depositional sequence. This term is used when the onlap pattern cannot be distinguished from downlap Onlap is a baselap in which an initially horizontal stratum laps out against an initially inclined surface, or in which an initially inclined stratum laps out updip against a surface of greater initial inclination.
35
Downlap is a baselap in which an initially inclined stratum terminates downdip against an initially horizontal or inclined surface. Toplap is lapout at the upper boundary of a depositional sequence. Erosional truncation is the lateral termination of a stratum by erosion. Structural truncation is the lateral termination of a stratum by structural disruption. Hiatus is the total interval of geologic time that is not represented by strata at a specific position along a stratigraphic surface. Offlap is the progressive offshore shingling of sedimentary units within a conformable sequence, in which each successively younger unit leaves exposed a portion of the older unit on which it lies.
In seismic sequence stratigraphy, onlap, downlap and toplap are evidence of non-
depositional hiatus. However, this hiatus may not necessarily be identical with a tme
sedimentary hiatus because of the limited resolving power of the seismic method. Sheriff
(1977; 1985) demonstrated that this resolving power decreases for beds thinner than about 1/4
wavelength. In the case where 25 H z seismic frequency is used, a bed thinner than 20 m (v
= 2000 m/sec) will not be represented. In multichannel seismic data, deeper sedimentary strata
will suffer more from this limitation, because during processing, frequency filtering parameters
are commonly depth dependent and lower frequencies are passed. These filtering parameters
are chosen on the basis that higher frequencies are attenuated more rapidly with distance (and
time) travelled. Thus the terms: onlap, downlap and toplap may also be interpreted as a
thinning of strata below the seismic resolution.
In recent developments, the seismic stratigraphic method has been adapted to the sequence
stratigraphic concept outlined by Posamentier et al. (1988) and Posamentier and Vail (1988).
This development appeared be an attempt to overcome the criticisms of the proprietary nature
of seismic data on which the seismic stratigraphic concept relies, and to apply the principles
previously used on seismic sections to outcrop and well log sections.
In the new development, Posamentier et al (1988) and Posamentier and Vail (1988) have
introduced some new terms and the concept of a systems tract. T w o commonly used terms
are accommodation space and parasequence. Accommodation space is defined as "the space
made available for potential sediment accumulation" (Jervey, 1988). The term parasequence
was introduced by van Wagoner (1985), and is defined as "a relatively conformable succession
of genetically related beds or bedsets bounded by marine flooding surfaces and their correlative
surfaces". Although it was stressed as a descriptive term (Posamentier & James, 1993), the
use of this term is rather ambiguous. There is a tendency to use the term as a part of the
36
sequence hierarchy (e.g. Swift et al, 1987; Boyd, 1994) similar to the hierarchy of supercycle,
cycle and paracycle of Vail et al. (1977) which reflects different orders of duration of relative
sea level change.
Analysis of seismic facies within a seismic sequence is carried out primarily to define
depositional environments. The descriptions of seismic facies are based on their reflection
patterns (Fig. 4.3), as classified by Mitchum et al (1977), and combined with other reflection
attributes such as continuity, amplitude, frequency and geometry. Table 4.1 from Sangree and
Widmier (1977) summarises the seismic characteristics and environmental facies interpretation
of clastic seismic facies units.
The recognition of coastal onlap and toplap is fundamental to the delineation of seismic
sequences. Vail et al (1977) have proposed that eustacy is the driving mechanism for
sequence evolution and he assumed that coastal onlaps which define sequence boundaries are
globally synchronous. The elementary assumptions given by Vail et al (1977) are that a
relative sea level rise is indicated by a landward shift of coastal onlap, while downlap in the
reverse direction indicates a relative sea level fall, and coastal toplap indicates a relative
stillstand of sea level. M a n y arguments have been put forward regarding these proposals.
Miall (1986) argued that seismic features such as coastal onlap are not necessarily related to
sea level change, but could be due to an autocyclic switching of a distributary channel, as in
the case of a submarine fan system, or indicate flexural subsidence rather than sea level
change. Moreover, Miall (1986) pointed out that quantifying coastal onlap could face a serious
problem when dealing with a divergent continental margin where a gradual thinning out of
strata is more c o m m o n than coastal onlap.
Schlager (1993) investigated the role of sediment supply within a supply-dominated
system. In the case of constant eustatic sea level, a similar stratal pattern to that produced
by sea level change can be generated through the varying of sediment supply. Allen and
Posamentier (1993) have documented that the highstand systems tract on the Gironde (France)
estuary has developed a seaward prograding pattern since about 4000 ka, inconsistent with the
adjacent shoreline which shows sediment starvation. These studies indicate that in the absence
of chronological data, stratigraphic correlation between parts of the basin based on stratal
pattern may face errors. Posamentier and Allen (1993) suggested that relative sea level change
determines the timing of sequence bounding surfaces rather than stratal pattern. Thus, a
37
stratigraphic correlation can only rely on chronostratigraphic surfaces caused by relative sea
level change, such as unconformities and flooding surfaces.
The sequence stratigraphic concept has been applied in various basins, but many of these
examples are from passive continental margins on which the subsidence rates are greater
basinward. This chapter presents two examples of seismic sequence stratigraphy from two sub-
basins in the Rembang Zone which have different tectonic behaviour. The first example (line
PWD-23) is from a basin which resembles a half graben basin in character, and the other (line
DNR-12) is from a basin which has similarities with a passive continental margin setting. A
seismic facies interpretation was carried out along the seismic lines, but, it was not possible
to generate maps showing lateral facies variations because of the limited extent of the available
seismic data.
The inferences of sea level change in the study area were made using the method of Vail
et al. (1977). By referring to the limitations pointed out by Miall (1986), the constmcted curve
has to be regarded strictly as a qualitative indicator of the relative position of assumed coastal
onlap or toplap. The Vail et al (1977) and Vail (1987) method also allows an estimation of
palaeowater depth of the basin floor based on the profile of a prograding clinoform, a feature
which has been used for the present study. However, since this method ignores post-
depositional effects, such as compaction, the results have to be regarded as rninimum depths.
4.3 SEISMIC SEQUENCE STRATIGRAPHY OF THE REMBANG ZONE
Two unmigrated 12-fold seismic lines, DNR-12 and PWD-23, were used for this study
(Fig. 4.1). These lines lie mostly in the Rembang Zone and ran north-south across the major
east-west structural trend. The data for line PWD-23 was acquired in 1974 and reprocessed
in 1988, whereas line DNR-12 was shot and processed in 1979. The source was dynamite, and
24 groups of 18 geophones were used (group interval 75 m for PWD-23 and 60 m for DNR-
12). Processing included relative statics (sea level datum), pre-stack deconvolution, velocity
analysis, residual statics, C D P stack and time variant filtering and scaling.
The interpreted lines for PWD-23 and DNR-12 are provided as Sections 4.1 and 4.2,
respectively. Figures 4.6 and 4.7 show the chronostratigraphic correlation sections for lines
PWD-23 and DNR-12, respectively. Appendix A contains velocity information for some shot
38
points along line PWD-23. Unfortunately, similar velocity information was not provided with
seismic line DNR-12. The left portion of both seismic lines lies in the northern part of the
Kendeng Zone, with extensive deep-seated faults on line DNR-12 (Sect. 4.2) representing the
zonal boundary. In contrast to the Rembang Zone, the reflection quality is extremely poor in
the Kendeng Zone, which definitely prevents a seismic facies study in this region, and
moreover, makes it impossible to correlate units between these zones. The cause of these poor
reflections is not clearly known, but it is probably due to the geological conditions in the
Kendeng Zone where tight folding and faulting occur.
Strata deposited during the Middle Miocene to Pliocene can be sub-divided and assigned
to seven seismic sequences E M I , M M 1 , M M 2 , L M 1 , L M 2 , PL1 and PL2. During the
discussion, systems tracts of these sequences are referred to as the lowstand (LST),
transgressive (TST) and highstand (HST) systems tracts, respectively, to indicate a rock unit
deposited within a certain interval with respect to relative sea level position. Ages of these
sequences were based on correlations with the biostratigraphic studies carried out in East Java
by previous investigators (see Chapter 3). In addition, biostratigraphic studies by Bolli (1966)
and Blow (1969) for the Bojonegoro-1 well (Figs 4.1 & 4.4), and by Pringgoprawiro and
Baharuddin (1979) for the Tobo-08 well (Fig. 4.1) are extremely useful in defining the
sequence ages on line DNR-12. The study by Kadar (1992) for the B N G - 1 well (Banyubang
area, Figs 4.1 & 4.5) is useful for line PWD-23. The reliability of these data relies on the
presence of sandy shelf deposits that are widespread and relatively constant in thickness. This
facies, the Ledok Formation, displays a very strong and laterally continuous reflection pattern,
and is an easily identifiable seismic unit. This facies has been used as a seismic marker on
both seismic lines P W D - 2 3 and DNR-12. A biostratigraphic study by Pringgoprawiro (1983)
of samples from the Ledok Formation from various places in the Rembang Zone suggested
correlation with planktonic foraminiferal zones N 1 7 and N18.
This study found that although the seismic lines PWD-23 and DNR-12 are only 25 km
apart, the sequence characters and boundaries are expressed differently on the two sections.
The two seismic lines may be regarded as representing the two main basinal areas in the
Rembang Zone: the southwest extension of the East Bawean Trough (PWD-23) and the Central
Deep (DNR-12, Fig. 4.10). The two elongate basins are separated by the JS-1 Ridge. The
differences in seismic features may reflect differences in tectonic behaviour in these basins.
39
O n line P W D - 2 3 (Sect. 4.1) which stretches nearly 50 k m from south of Rembang to
south of Cepu, downlap directions and general facies trends suggest that all Middle Miocene
to Pliocene sediments were deposited on a relatively gentle southward slope. The rate of basin
subsidence during the Early and Middle Miocene was high and appears to have been more
pronounced in the landward (margin) area which produced relatively thick sequences (EMI,
M M 1 and M M 2 ) on the northern portion of the seismic line (PWD-23). This basin subsidence
is suspected to be associated with the downward movement of a half graben system on the
northern boundary of the basin, and the area beneath the Kedinding Anticline acted as a pivot
point. During the Late Miocene and possibly Pliocene the rate of subsidence was reduced and
the sequences deposited during this interval (LM1, L M 2 and PL1) are relatively constant in
thickness along the line.
Line DNR-12 (Sect. 4.2) spans about 30 km and was recorded across the Tertiary
topographic high of the Tobo High and the basin to the north of it. The Tobo High stretches
from east of Ngimbang to south of Cepu and is very well expressed by the gravity map (Fig.
2.4). Compared with the previous line (PWD-23), all Middle Miocene to Pliocene strata were
deposited toward the north on relatively steeper topography (?shelf slope). Excessive lateral
thickness anomalies of sequences are not indicated which suggests a low sediment supply and
a relatively uniform subsidence rate throughout the Middle Miocene to Pliocene.
4.3.1 Sequence boundaries
The sequence boundaries were mostly defined by recognising surfaces on which seismic
reflections systematically terminate. Criteria for recognising sequence boundaries were given
by Vail et al (1977) and outlined earlier. The boundaries of the recognised sequences on the
studied seismic lines are characterised by onlaps and truncation of reflections. Ages of these
boundaries were defined through correlation with the biostratigraphic studies carried out by
Bolli (1966), Blow (1969), Pringgoprawiro and Baharuddin (1979) and Pringgoprawiro (1983)
in several places in East Java.
Most of the sequence boundaries on line PWD-23 are type U. On this line (Sect. 4.1) the
base of the sequence E M I cannot be defined due to the poor seismic record. The base of its
regressive unit (HST) was deduced from inconspicuous reflection appearances which are
interpreted as a downlapping reflection pattern. The seismic velocity analysis at some
shotpoints has indicated an abrupt decrease of interval velocity at the suspected positions of
40
this regressive base. The estimated age is Early Miocene, based on the palaeontological data
of the BNG-1 well (Banyubang).
On line DNR-12 (Sect. 4.2) the lower boundary of sequence MM1 is difficult to define
due to the character of the seismic data which are almost free from reflections. The base of
sequence M M 2 on this line is expressed by a very strong continuous reflection which is
interpreted as an unconformity surface (type I unconformity). This boundary has not been
reached by wells drilled in the area of line DNR-12. A rather strong and continuous reflection
occurs on top of sequence M M 2 , which is also interpreted as an unconformity surface. This
surface may represent the late Middle Miocene unconformity at the top of the Tawun
Formation. Biostratigraphic studies on the Bojonegoro-1 well (Figs 4.1 & 4.4) by Bolli (1966)
and Blow (1969) reveal that this hiatus occurs between 855 and 914 m in depth, which is in
foraminiferal zone N14. The other type I sequence boundaries on line DNR-12 occur between
sequences M M 2 and L M 1 , L M 1 and L M 2 , and probably between sequences L M 2 and PL1.
In contrast with line DNR-12, erosional surfaces on line PWD-23 (Sect. 4.1) are mostly
local and occur on the northern end of the section between sequences L M 1 / L M 2 and probably
between sequences LM2/PL1. The boundaries of sequences M M 1 / M M 2 and M M 2 / L M 1 on
this line (Sect. 4.1) are not expressed by pronounced surfaces due to continuous deposition
from sequence E M I to M M 1 and from M M 1 to M M 2 . The prominent surfaces shown on this
line between M M 1 (TST) and M M 1 (HST), and between M M 2 (TST) and M M 2 (HST) are
in fact maximum flooding surfaces (surfaces of non-deposition) which occurred during the
maximum elevation of sea level. On line PWD-23 these surfaces are marked by basinward
terminations of reflectors beneath the surfaces, which are interpreted as apparent truncations,
following a geometrical relationship suggested by Vail (1987) (Fig 4.2). By definition
(Mitchum et al, 1977), these surfaces cannot be regarded as sequence boundaries. The
sequence boundaries, thus, have to be found half way between the H S T unit of the earlier
sequence and the TST unit of the subsequent sequence.
A maximum flooding surface is also recognised within the sequence LM1 and lies on top
of its transgressive unit (TST). Compared with the position of the transgressive units of
sequences M M 1 and M M 2 , the transgressive unit of sequence L M 1 was shifted more
landward, suggesting a relatively more rapid sediment starvation which may have resulted from
an insufficiency of terrigenous clastic materials to keep pace with a rapid sea level rise. Due
41
to this, in the area south of Banyubang Anticline, the regressive unit (HST) of M M 2 is
separated by a surface of non deposition from the regressive unit (HST) of sequence L M 1 .
Correlation between these seismic sections, outcrop and well data (BNG-1, NGBU-1,
Tobo-04 and Tobo-08; Fig. 4.1) confirms that the sequence boundaries L M 1 / L M 2 and
LM2/PL1 correspond to the lithostratigraphic boundaries of the Wonocolo/Ledok and
Ledok/Mundu Formations respectively. The base of the Wonocolo Formation, according to
the data from BNG-1 and NGBU-001 wells, was picked at the boundary between sequences
M M 2 (HST) and L M 1 (HSTa) for the area south of Banyubang Anticline. The
stratigraphically equivalent Bulu Formation (developed on the north flank of the Banyubang
Anticline in the area between Blora and Rembang; Figs 3.1, 4.1 & 4.5) represents the
transgressive unit of sequence L M 1 . Its reefal facies may have developed during the high sea
level stillstand (HSTa). Stratigraphically, the sequence just below the Wonocolo and Bulu
Formations should represent the Tawun Formation. The boundary of the Tawun and
underlying Tuban Formations according to the NGBU-001 well data was picked at the
boundary between the transgressive and regressive units of sequence M M 1 . Ages of the
formation boundaries have been discussed in Chapter 3 and summarised in Table 3.4. Here,
these ages are used to infer the ages of the seismic sequences.
4.3.2 Seismic facies
a. Line P W D - 2 3
Discussion presented in this section refers to the interpreted seismic section, Section 4.1,
and the chronostratigraphic correlation section Figure 4.6. The sequence terminology E M ,
M M , L M and PL correspond to their periods of formation: Early Miocene, Middle Miocene,
Late Miocene and Pliocene respectively.
a.l Sequences EMI (regressive unit), MM1 and MM2
These sequences are composed of nearly symmetrical regressive-transgressive systems
(Sects 4.1 & Fig. 4.6). As discussed earlier the boundaries between these sequences are type
II unconformities. The base of sequence E M I is unidentified due to poor data quality in the
lower part of the seismic section, and only the regressive unit (HST) of sequence E M I is
recognised.
42
All of these sequences thicken landward (northward) suggesting greater subsidence and
sedimentation rates landward. To the south, the deposition of the sequences was limited by
a topographic high beneath the Kedinding Anticline. Early and Middle Miocene sediments
above this high are relatively thin. The absence of erosional features suggests that this high
was syndepositional, and the development ceased from the deposition of sequence L M 1
(Wonocolo Formation) onward. This high is interpreted to be the southwest extension of the
JS-1 ridge in the Java Sea (Fig. 4.10).
The transgressive and regressive units of these sequences are distinctive. The regressive
units are characterised by a subtle sigmoid progradational pattern, and due to thinning
southward the pattern tends to be shingled at the toes. The transgressive units are characterised
by a progressive landward shifting (northward) of the apparent tmncation features on which
the subsequent highstand units downlap in a basinward direction . These regressive and
transgressive patterns suggest that the combination of subsidence rate, sediment supply and sea
level fluctuations might have successfully maintained a constant accommodation space for the
terrigenous clastic to be deposited. The Banyubang well, BNG-1 (Fig. 4.5), indicates that these
sequences are not characterised by pronounced vertical changes in sedimentary facies. They
are represented by thick paralic deposits, consisting mostly of carbonaceous mudstone with
thick interbeds of bioclastic limestone and glauconitic quartz sandstone. The bioclastic
limestone and quartz sandstone dominate the upper part of sequence M M 2 . It is interpreted
that at the end of this sequence the relative sea level slowly increased which is also evident
from the significant landward shift of the transgressive unit (TST) of sequence L M 1 .
The sequence EMI and the lower part of sequence MM1 are equivalent to the Tuban
Formation (Ardhana, 1993). The boundary between this formation and the overlying Tawun
Formation in the B N G - 1 well is in fact difficult to define due to the similarity in sedimentary
facies. However, the Bojonegoro-1 well (Fig. 4.4) has indicated the presence of a hiatus
(early-Middle Miocene hiatus, at depth of 1842 m ) between these formations. O n line P W D -
23 (Sect. 4.1) this boundary is tentatively placed on the prominent surface (maximum flooding
surface) between the transgressive (TST) and the regressive (HST) units of sequence M M 1 .
The regressive unit (HST) of sequence MM2, according to BNG-1 well, corresponds with
the Ngrayong Member of the Tawun Formation. Seismically, this formation is characterised
by a parallel clinoform pattern with variable amplitude and continuity. The upper part of the
43
Tawun Formation (corresponding to sequence M M 2 ) , as observed in the Bulu, Plantungan and
Jepon areas (Fig. 4.1), consists of paralic deposits which transgress into shallow marine facies.
In the Jepon area, the paralic deposits consist of well-bedded, organic-rich mudstone (Fig.
4.11) with some thin interbedded coal seams. A quartz sandstone of littoral facies (Ngrayong
Member) is very well exposed in the Plantungan area. Here, this sandstone is often
intercalated with beds of orbitoidal (larger foraminiferal) limestone of up to 1 m thick (Fig.
4.12). Some of the foraminifer genera have been attributed to the Tawun Formation (van
Bemmelen, 1949; Pringgoprawiro, 1983). The upper part of this formation is typified by the
genera Lepidocyclina, Miogypsinoides and Cycloclypeus. The living genera of larger
forarninifers occur in shallow marine waters of the tropics or subtropics, and are common
constituents in rocks containing coral reef deposits and calcareous algae (Moore et al, 1952;
Serra-Kiel & Reguant, 1984).
As indicated previously, the basin in which sequences EMI, MM1 and MM2 were
deposited is characterised by a greater landward subsidence rate. This style is commonly
associated with a tilt-block of a half graben basin (Leeder and Gawthorpe, 1987; Schlische,
1991), and the sediments are derived from the uplifted foot wall. It should be noted that such
basins commonly occur in an extensional tectonic system. Hamilton (1979) suggested that
the present northward movement of the Indian-Australian Plate is about 10 cm/year. A similar
northward compressional stress would occur if the tectonic setting during the Middle Miocene
were similar to the present day, as indicated by Baumann (1982). Dewey (1980) indicated
that an extensional tectonic system could occur in a retro-arc setting along an old line of
weakness obliquely oriented to the subducting plate. This has been proven to occur along the
Sumatran Fault system (Katili, 1970), and may also occur in the study area where lines of
weakness (Fig. 4.10) resulting from a previous pre-Tertiary subduction complex are suspected
to be present.
A greater landward subsidence rate has allowed a landward thickening of the Middle
Miocene deposits. This style of sedimentation, in some cases, resembles foreland basin
sedimentation where the greater basin subsidence is adjacent to the uplifted area where the
sediments are sourced. Examples of the application of sequence stratigraphy to foreland
basins are few. Swift et al. (1987) studied the Late Cretaceous Mesaverde Group, a foreland
basin deposit in the Book Cliffs (Utah). They found that the anatomy of sedimentary
sequences differs significantiy from those found on passive margins. In the Mesaverde Group,
44
the transgressive deposits are offshore sand bodies derived from the high area in front of the
basin, intertongued with the deltaic regressive deposits derived from the uplifted mountainous
region. Figure 4.8 illustrates the sequence architecture models from passive margin and
foreland basin settings. In the foreland basin setting, a sea level fall-induced sequence
boundary has never been well represented. The sea level fall has never effectively exposed
the landward margin because of more rapid subsidence than the seaward margin. This
situation appears to be similar to the study area (tine PWD-23), but both the transgressive and
regressive deposits were derived from the uplifted area in the north. However, the influence
of eustatic sea level change on the Middle Miocene sediment deposition is strongly displayed
in the study area, as indicated by the almost sinusoidal advance and retreat of these deposits
(Fig. 4.6).
a.2 Sequence LM1
Three units (TST, H S T a and HSTb) can be recognised within this sequence. The
transgressive unit (TST) was deposited during a relatively rapid sea level rise. The deposition
was restricted to landward (north) of Banyubang Anticline, while a surface of non deposition
occurred in the basinal area (south of the Banyubang Anticline) between this sequence and the
underlying sequence. The unit HSTa is the early regressive unit, and occurred during the
highstand of sea level.
Unit TST is characterised by a medium to high amplitude and high continuity parallel
reflection pattern. The spacing of reflections is somewhat irregular, probably indicating an
irregular spacing of sedimentary interbeds. Data from the B N G - 1 well confirms that this
transgressive unit corresponds with the Bulu Formation which consists of calcarenite
interbedded with quartz sandstone and marly mudstone. The total thickness of the formation
in this well is about 80 m. A relatively thick (>10 m ) reefal limestone consisting of larger
foraminifers (Orbitoididae), bryozoans, echinoids and coral fragments occurs in the area near
Bulu (Fig. 4.1). In this area the thickness of the formation reaches 240 m.
Unit HSTa is typified by a low to medium amplitude subparallel reflection pattern. Most
of this unit was deposited in a low energy depositional setting. The m a x i m u m thickness is
about 270 msec T W T (about 270 m ) decreasing landward to about 150 msec T W T (about 150
m for a velocity of 2000 m/sec). Its stratigraphic position, which lies between the
transgressive (TST, Bulu Formation) and prograding (HSTb) units implies deposition during
45
a highstand of sea level. With regard to its seismic character, most of unit H S T a may be a
hemipelagic deposit. Well data (NGBU-001) confirm that it consists of foraminiferal-rich marl
and corresponds with the lower part of the Wonocolo Formation. The planktonic/benthonic
ratio in this formation ranges from 20 to 5 0 % in the Bulu area and 60 to 8 0 % in the
Kawengan area (Pringgoprawiro, 1983), which indicate a facies variation from inner to outer
shelf (Murray, 1991). Toward the north (Bulu area), the facies might have changed into reefal
which is regarded as the upper part of the Bulu Formation.
The late regressive unit (HSTb) has sigmoidal prograding character which changes into a
shingled reflection pattern due to basinward thinning. It is characterised by medium to high
amplitude and good continuity. It thickens landward from about 70 to 110 msec T W T (about
70 to 110 m ) . In the area between the Banyubang and Plantungan Anticlines the upper part
is locally truncated, indicating the presence of sedimentary bypassing. However, the presence
of basinward thinning combined with a shingled progradation pattern instead of outbuilding
clinoforms suggests that sediment supply was relatively low. This unit corresponds with the
upper part of the Wonocolo Formation.
a.3 Sequence LM2
The sequence L M 2 was deposited at the end of the Late Miocene. Recognition of
reflection terminations and system tracts was difficult as the apparent depositional slope is
relatively flat on section PWD-23. Seismically, this sequence is characterised by high
amplitude continuous parallel reflection patterns. A relatively constant lateral thickness
suggests a sheet-like or tabular external form. These features indicate a uniform subsidence
rate. A weakening of reflection amplitude occurs in the area south of the Kedinding Anticline.
This is interpreted as a facies change into marl-prone facies which possibly represents a deeper
marine environment.
Sequence LM2 is correlated with the Ledok Formation which, elsewhere in the Rembang
Zone, stratigraphically overlies the Wonocolo Formation. Field study in the areas near Ledok
and Wonocolo found that the Ledok Formation consists of interbedded calcarenite and
marl/marly mudstone. In the field, the transition from the Wonocolo to Ledok Formations is
recognised by an increase in the number of calcarenite interbeds (Fig. 4.13). The upper part
of the Ledok Formation is very sandy and glauconitic with an average grain size of about 1
m m . The benthonic foraminifers Uvigerina, Bulimina, Laticarinina and Nodosaria are
46
common in the Ledok Formation (Pringgoprawiro, 1983), suggesting a middle to outer neritic
environment. But, the abundance of large crossbeds (> 1 m ) in this formation suggests the
deposition was influenced by tidal currents, which implies a shallower environment. Samuel
and Johannes (1986) measured the orientation of these structures in an area north of Cepu and
obtained a general southward trend of the dominant tidal currents. To the south, the Ledok
Formation is correlated with the pelagic marl of the Kalibeng Formation in the Kendeng Zone.
These sedimentary features suggest that the Ledok Formation was deposited in a tide-
dominated shelf environment. Dalrymple (1992) indicated that in this setting, tidal currents
commonly range from 0.5 to 1.5 m/sec, which are strong enough to generate tidal sand waves.
A m o d e m example of this environment has been reported by Berne et al (1991) to occur in
the Bay of Bourgneuf (France). Here, sand waves range in amplitude from 0.7 to 9.4 m, and
exhibit marked asymmetry, which has been attributed to the combined effect of wave and
tidal currents. Dip measurements by B e m e et al (1991) on the internal laminae of sand waves
suggest the occurrence of sandflow similar to those found in eolian dunes.
a.4 Sequence PL1 and PL2
Sequence PL1 was deposited during the Pliocene. Most parts of the sequence observed on
the seismic sections show transgressive features, since to the south of the Kedinding Anticline
it progressively onlaps landward (northward) on top of sequence L M 2 . The seismic character
is a low amplitude and poor continuity subparallel reflection pattern, indicating a low energy
depositional environment. The lithofacies interpretation was based particularly on correlation
with outcrops along the seismic line, and suggests a correspondence with the marl forming the
Mundu Formation. This marl is an outer neritic hemipelagic deposit characterised by abundant
planktonic foraminifers.
A study by Aminuddin et al. (1981) in the area between Cepu and Rembang indicated that
the Mundu Formation shoaled upward into an alternating coquina (Fig. 4.14), sandy limestone
and marl beds [Selorejo Member, Fig. 4.14, Pringgoprawiro (1983)]. This unit thickens
northward from few metres in the area south of Cepu to more than 200 m in the area south
of Rembang. Aminuddin et al (1981) found that many forarninifers contained in this unit
represent reworked faunas. This was confirmed by the petrographic study (chapter 4.4.5) that
showed most foraminifers in this unit are micritised. The micritisation may have occurred
during a reworking of older units or during diagenesis. The Selorejo Member may be regarded
47
as the regressive unit of sequence PL1 where the sea level was nearly or at the high stillstand
position.
Sequence PL2 is the uppermost sequence identified. South of the Kedinding Anticline
(Sect. 4.1) the base of this sequence is characterised by an erosional surface. The internal
character shows a high amplitude wavy seismic reflection pattern. It appears to lie horizontally
which implies that deposition occurred after or contemporaneous with the Kedinding Anticline
formation. This sequence is correlated with the tidal flat deposits of the Lidah Formation
which is widely exposed in the Rembang Zone and the eastern part of the Kendeng Zone.
Some authors agree that the development of this formation has led to the Late Pliocene base
levelling through most of East Java (van Bemmelen, 1949; Lemigas/BEICIP, 1969). The real
thickness is not known, but the seismic sections suggest up to 400 msec T W T (about 400 m ) .
b. Line DNR-12
Since Tertiary sedimentary outcrops along and near this seismic line are absent, the
lithofacies interpretation of this seismic line was carried out through correlation with well data
from Tobo-04, Tobo-08 and Bojonegoro-1 wells and additional reference to the criteria given
by Sangree and Widmier (1977). The Tobo-04 and Tobo-08 wells are shallow wells, with total
depths of 410 m and 432 m, respectively. They are only effective for the correlation of the
upper sequences: L M 2 , PL1 and PL2. The Bojonegoro-1 well has a total depth of 2034 m.
However, due to its position which is about 17 k m away from line DNR-12, a good correlation
with this line has not been achieved. Section 4.1 displays the interpreted section for line D N R -
12 and Figure 4.7 summarises the chronostratigraphic correlation of this line. The sequence
terminology used in these figures is similar to line PWD-23, and implies chronostratigraphic
equivalence.
b.l Sequences EMI, MM1 and MM2
A detailed discussion on sequences E M I and M M 1 is hampered by their seismic character
which is nearly reflection free. The sequences probably consist of mudstone facies with some
sandstone lenses as indicated by the presence of a few discontinuous and variable amplitude
reflections. The recognition of systems tracts, on the basis of this seismic character, would be
speculative and is not attempted.
48
Sequence M M 2 on line DNR-12 is a type I sequence. Three systems tracts can be
recognised in this segment of the line: the lowstand, transgressive and highstand systems tracts.
The lowstand systems tract is represented by a relatively thin [< 110 msec T W T (110 m)]
shingled prograding unit. The internal seismic character of this unit shows a high amplitude
and continuous parallel reflection pattern. The lowstand unit is thought to have been deposited
in a shallow marine environment and consists of terrigenous clastic strata. The water depth
of basin floor was probably equal to the maximum thickness of the unit, 110 msec T W T (about
110 m).
The subsequent sea level rise stopped deposition of the lowstand unit and the transgressive
systems tract developed on the topographic high. The reflections along the upper boundary
of this tract backstep landward (southward), indicating a further retreat of sedimentation (note
the difference in sedimentation direction compared with line PWD-23). The seismic reflection
pattern which shows a mounded form and lack of internal continuity, suggests that it is
composed of carbonate buildups.
The highstand systems tract is characterised by an oblique progradational pattern with
reflectors toplapping against the upper sequence boundary. The internal reflection pattern
shows a high continuity of medium amplitude parallel reflectors. Mitchum et al. (1977)
attributed this type of seismic pattern to rapid basin infill and sedimentary bypassing resulting
from a combination of high sediment supply and sea level stillstand. O n the seismic line
studied, a rapid basinward shift of the toplapping pattern, particularly above the northern flank
of the topographic high, may have resulted from a relative sea level fall. In the distal part
(basinward), the presence of northward toplapping suggests that this unit merged with a
northerly derived depositional system.
b.2 Sequence LM1
This a type I sequence in which two main systems tracts can be differentiated from the
seismic section: the lowstand and highstand systems tracts.
The lowstand systems tract consists of three distinct features which are interpreted to
represent the sheet-drape, basin-floor fan and lowstand wedge. This interpretation refers to the
criteria outlined by Mitchum (1985), Posamentier and Erskine (1991) and Posamentier et al
(1991). The basin-floor fan is characterised by a mounded external form, and the internal
49
reflectors are somewhat variable in both amplitude and continuity indicating a variable current
during deposition. Basinward (northward), it grades into a sheet-drape of probable hemipelagic
deposits which is characterised by a medium amplitude parallel reflection pattern. Shelfward
(southward), the basin-floor fan grades into a thin oblique progradational clinoform of a
lowstand wedge.
Deposition of the highstand systems tract occurred as a result of the subsequent relative
sea level rise. It can be separated into three main parts: proximal, middle and distal. The
proximal part downlaps on the lower sequence boundary. The middle part consists of an
oblique prograding wedge which grades into a tangential-oblique reflection pattern in the distal
part. Toplapping occurring on top of the tract suggests sedimentary bypassing combined with
a very high sediment supply. Similar to the regressive unit of sequence M M 2 , the most distal
(prodelta) facies may have merged with a northerly derived depositional system as indicated
by the presence of southward prograding clinoforms. The Tobo-04 well has penetrated about
40 m of the highstand deposit and indicated a lithofacies of greenish grey, glauconitic sandy
marl. The well report assigned this facies to the Wonocolo Formation.
The highstand systems tract may have been deposited as a highstand prograding delta
complex. However, a further detailed seismic study would be necessary to support the
interpretation, with the specific aims of defining the lateral distribution, external geometry and
associated seismic features (such as fluvial channelling). Highstand prograding clinoforms
associated with a prograding delta complex have been reported by some authors. Suter and
Berryhill (1985) recognised the occurrence of five Late Quaternary deltas on the northwest
shelf margin of the Gulf of Mexico. They used only geophysical data to identify these delta
systems, but the extent and coverage of the data allowed a recognition of mappable
geomorphic patterns which are associated with high-angle clinoform reflection patterns of delta
lobes and buried fluvial systems. A similar method combined with core drilling was used by
Sydow and Roberts (1994) on the recognition of a shelf edge delta in the northeast Gulf of
Mexico.
b.3 Sequence LM2
Sequence L M 2 on line D N R - 1 2 is a type I sequence. The seismic character of this
sequence is very similar to that on line PWD-23. Its depositional setting, which was on a
50
relatively steeper slope, has allowed recognition of two major systems tracts, the lowstand
(LST) and highstand (HST) tracts.
The lowstand systems tract is characterised by relatively low amplitude and high continuity
tangential-oblique reflections which may be associated with a lowstand wedge mudstone facies.
The highstand systems tract was deposited following this facies. The landward onlapping of
the lower part of the highstand systems tract (HST) probably represents transgressive
conditions and may be regarded as the transgressive systems tract. The m a x i m u m relative sea
level rise is difficult to estimate because of poorly defined onlapping features in the landward
direction.
The highstand deposits pinch out basinward, and are characterised by a high amplitude
reflection pattern with frequent internal discontinuity suggesting a fairly shallow marine
environment. These deposits downlap on top of sequence L M 1 and the lowstand deposits of
L M 2 . T w o shallow wells, Tobo-04 and Tobo-08, penetrated the highstand deposits and
indicate similar facies to the sequence L M 2 on line PWD-23, which was assigned to the
Ledok Formation in the earlier discussion. In these two wells, this formation is also
characterised by interbedded fossiliferous, glauconitic sandstone and marl; the main terrigenous
fragments are fine to medium-grained quartz.
b.4 Sequences PL1 and PL2
Sequence PL1 is identified by its progressive landward onlap on top of sequence L M 2 .
This part may represent a transgressive systems tract. Seismically, this unit is characterised
by a low amplitude and high continuity parallel reflection pattern. Data from the Tobo-04 and
Tobo-08 wells indicate that the sequence consists mostly of foraminiferal-rich marl similar to
the M u n d u Formation recognised on line PWD-23. In the Tobo-08 well this formation is 70
m thick, thinning southward to about 30 m thickness in the Tobo-04 well. Thinning of the
formation appears to be related to the depositional topography which was higher toward the
south. The coquina and sandy limestone of the Selorejo Member, which characterise the upper
part of the Mundu Formation, are still recognisable in the Tobo-08 well as a 30 m thick unit.
Sequence PL2 lies horizontally and onlaps on top of sequence PL1. Seismically, it is
characterised by a relatively stronger amplitude reflection pattern than PL1, suggesting a more
pronounced alternation of fine and coarse clastic strata. This uppermost sequence is widely
51
exposed along the seismic line and has a similar appearance to the tidal deposits of the Lidah
Formation in Cepu area.
4.3.3 Relative sea level change
The inferred relative sea level changes for the Rembang Zone are presented in Figure 4.9.
The curves were based only on the Middle Miocene to Pliocene coastal aggradations in line
DNR-12 (Sect. 4.2 & Fig. 4.7) and are not normalised to remove local tectonic effects.
Recognition of coastal onlaps on line PWD-23 is difficult, due to limited extent and poor data
quality in the landward areas, and the interpretation of its relative sea level change is not
attempted.
The inferred Pliocene sea level magnitude in Figure 4.9 was not based entirely on the
seismic features on line DNR-12 due to the difficulty in defining the topmost position of onlap.
The seismic study in Chapter Seven suggests that the relative sea level rise in the Early
Pliocene was about 115 m. The sea level magnitude was also based on the foraminiferal
examination by Pringgoprawiro and Baharuddin (1979) in the Tobo-08 well. The Pliocene
sequence, PL1, consists of hemipelagic marl and was deposited in an outer neritic environment.
Figure 4.9 indicates that the Middle Miocene and Pliocene maximum sea level magnitudes
were relatively higher than in the Late Miocene. These are consistent with the Late Tertiary
global relative change of coastal onlap proposed by Haq et al. (1988). However, the number
of Middle Miocene and Pliocene sequences in the study area is fewer than the number of
sequences recognised by Haq et al. (1988) for the same period, and appears to be a
generalisation of Haq's curve. Posamentier and James (1993) indicated that the number of
sequences observed in a certain area will depend on many local factors. The four major
variables include tectonic subsidence, eustatic change of sea level, sediment supply and climate
(Vail, 1987). Because of these, each area may have its own relative sea level curve (Suter et
al, 1987).
4.4 SEDIMENTARY PETROGRAPHY
This section briefly discusses the petrographic character of the Middle Miocene to Pliocene
lithofacies in the area between Cepu and Rembang. All rock samples for the petrographic
analysis came from outcrop. The analysis mostly used a transmitted light microscope. The
52
main objective of the petrographic study was to document the overall character of the
lithofacies, particularly composition and provenance. Appendix B.l summarises the results of
the petrographic analyses.
4.4.1 Tawun Formation
Four samples from the Tawun Formation were analysed for this study. All of these
samples were from exposures in the Bulu, Plantungan and Jepon (Banyubang) areas (Fig. 4.1)
which represent the upper part of the Tawun Formation. T w o main facies are developed:
organic-rich mudstone with thin beds of fine quartz sandstone, and quartz sandstone (the
Ngrayong Member). The mudstone is commonly parallel laminated. In the samples studied,
the amount of fine-grained (average grain size of 0.2 m m ) quartz is up to 2 0 % (Fig. 4.15).
Quartz sandstones in the Tawun Formation are mostly free from mud, and commonly
contain organic spots of possibly dried oil in their pore spaces. Quartz (up to 86%) is
commonly characterised by undulose extinction and cracking (Fig. 4.16). Some
polycrystalline quartz grains are sutured and sheared (Fig. 4.16). The grain size of quartz
varies from 0.1 to 1.5 m m with mostly subangular shapes.
Folk (1974) used the quartz undulatory extinction and polycrystallinity as parameters to
indicate its origin. H e proposed that highly undulose quartz is diagnostic of metamorphic
origin. A study of the significance of quartz undulatory extinction and polycrystallinity by
Basu et al. (1975) concluded that these features can be used to distinguish plutonic quartz from
low-rank metamorphic quartz in first-cycle sandstone. In the Tawun Formation, in all four
sandstone samples studied, undulose extinction and polycrystalline quartz are present. Their
metamorphic origin is also suggested by the presence of metamorphic rock fragments such as
slate.
The seismic interpretation has indicated that all of the Tawun Formation in the area
between Cepu and Rembang was derived from a northerly source. This suggests the
importance of the Bawean Arch as the source of sediment during the Middle Miocene (Fig.
4.10). The paralic environment which characterised the Tawun Formation would have
developed along the margin of the subaerially exposed Bawean Arch.
53
4.4.2 Bulu Formation
A study of an exposure about 50 m long in an area near Bulu (Fig. 4.1) indicates the
formation is characterised by a transgressive sequence. The sequence begins with an organic-
rich dark brown mudstone associated with a channel facies containing quartz-rich sandstone
and conglomeratic-lag deposits. Above these deposits is about 30 m of interbedded
fossiliferous mudstone and fine-grained quartz sandstone which are topped by a thick orbitoidal
bioclastic limestone (Fig. 4.18).
The petrographic study was carried out on four limestone samples and one sandstone
sample from the above section. In character, the channel sandstone resembles the quartz
sandstone of the Ngrayong Member, but most of the pores contains organic streaks (?dried oil).
The quartz grains (Fig. 4.17) have a similar appearance to those in the Ngrayong Member and
are suspected to be derived from a similar source. The bioclastic limestone is bedded and
typified by an abundance of larger foraminifers (Orbitoididae) and algal fragments (up to 6
m m ) . It also contains quartz grains, commonly less than 16%, with the grain size varying from
0.1 to 2 m m (Fig. 4.18). In most of the limestone samples studied, finely crystalline sparry
calcite cement is well developed (as in Fig. 4.18).
4.4.3 Wonocolo Formation
Samples for the petrographic study of this formation were from the Kawengan area (some
20 k m northeast of Cepu) and the upper part of the formation. Here, this formation is
characterised mainly by foraminiferal-rich sandy marl. Figure 4.19 shows a typical specimen
with carbonate m u d constituting about 5 0 % , mainly as matrix, pore fillings and micritisation
of bioclasts. The bioclasts are mainly planktonic foraminiferal tests (< 1 m m size), mollusc
and algal fragments (up to 6 m m size) and make up 4 0 % . The remainder consists of
terrigenous clastic detritus, composed mainly of angular to subangular monocrystalline quartz
with an average grain size of 0.15 m m . Sparry calcite cement appears to be poorly developed,
and is found only as a pore filling within some foraminiferal tests (e.g. Fig. 4.19).
X-ray diffraction analysis has been carried out on a sample from this formation, and the
results are shown in Figure 4.23. The analysis has confirmed the microscopical examination
and shows that the main constituents of this formation are carbonates. The most abundant
carbonate is calcite which displays a very pronounced response at 29.43° 26, along with most
of its subordinate peaks. Aragonite is also present and is suspected to represent calcareous
54
faunal fragments. The analysis also indicates the presence of very small amounts of smectite
and kaolinite.
The presence of substantial carbonate mud is commonly interpreted to indicate deposition
under fairly low energy conditions (Matthews, 1966). But the notable presence of quartz and
algal fragments suggest that periodic bottom currents were active. Quartz, in this formation,
is interpreted to have been derived from the Tawun Formation since in the Late Miocene, the
topographic high (JS-1 Ridge) was still developing near Tuban (van Bemmelen, 1949;
Lemigas/BEICIP, 1969) and exposing the Middle Miocene strata subaerially.
4.4.4 Ledok Formation
Four samples for petrographic analysis were obtained from the middle and top of the
formation in the Ledok area (some 15 k m north of Cepu) and the Wonocolo area (some 20 k m
northeast of Cepu). In the samples from the middle part of the formation, carbonate mud,
bioclasts of foraminiferal tests and algal fragments constitute up to 69-88%. The terrigenous
fragments, mostly monocrystalline quartz, form less than 12-37%. These fragments display
subangular form with an average grain size of 0.3 m m .
In the upper part of the formation, the average grain size (0.6 mm) is significantly coarser
than in the middle part. This is associated with shoaling of the sequence. The upper part is
characterised by the appearance of up to 4 7 % glauconite (Fig. 4.20). Almost all of the
glauconite occurs as rounded peleoid grains, accompanied by some interstitial filling within
foraminiferal tests. The rim of some glauconite grains appears to have been altered into iron
oxides, probably during the present weathering cycle. Carbonate content is up to 69%,
consisting mainly of foraminiferal tests and algal fragments. Undulose monocrystalline quartz
constitutes up to 1 7 % , with the grain size varying between 0.1 and 0.5 m m . Compared with
the samples from the middle part of the formation, carbonate mud is essentially minor
indicating deposition under a strong bottom current. This interpretation is supported by the
fact that large crossbeds are common in the upper part.
The occurrence of glauconite in the upper Ledok Formation has never been discussed in
detail. The mode of deposition, which was under strong bottom currents led Samuel and
Johannes (1986) to the opinion that the glauconite formation was not in situ. The processes
which lead to the formation of glauconite are still uncertain. Burst (1958) and Hower (1961)
55
believed that glauconite pellets are derived from a three-layer silicate lattice (clay minerals)
through a transformation process involving supplies of potassium and iron, and a suitable
oxidation potential. Odin and Matter (1981) argued that glauconite formation is independent
of parent materials. They suggested that glauconite authigenicaUy grows in substrate pores
favoured by the microenvironment in the substrate, and is accompanied by progressive
alteration and replacement of the substrate. The glauconite occurrence in the Ledok Formation
appears to favour the Odin and Matter (1981) argument. There are three modes of glauconite
precipitation in the Ledok Formation, as rounded peleoids, interstitial fillings and casts of
bioclasts. A m o n g these, the rounded peleoid type is the most common. The usual microscopic
appearance is that it often contains well preserved very fine-grained quartz fragments, which
imply that the glauconite may have been altered from quartz-bearing pellets.
There is a common agreement that glauconite formation is associated with a low
sedimentation rate (Hower, 1961; Odin & Matter, 1981; van Houten & Purucker, 1984), and
most commonly occurs in the open marine environment beyond the zone of fluvial influence
(Odin 8c Matter, 1981). In contrast, abundance glauconite in the Ledok Formation occurs at
the upper part of the formation, and was associated with strong tidal bottom currents. These
suggest that glauconitisation might not have occurred simultaneously with substrate deposition,
and possibly occurred later during the subsequent transgression.
4.4.5 Mundu Formation
The M u n d u Formation consists of marly facies and limestone facies (Selorejo Member).
The marly facies consists of hemipelagic marl and is very rich in planktonic and benthonic
foraminiferal tests (up to 4 0 % of the total carbonate; Fig. 4.21). Terrigenous clastic detritus
is minor and mostly consists of very fine-grained quartz and feldspars. In the area near Cepu,
the limestone facies of the M u n d u Formation is known as the Selorejo Member
(Pringgoprawiro, 1983). A sample from this unit shows that subangular plagioclase grains with
an average size of 0.05 m m make up about 1 % and glauconite amounts to less than 1%. The
field study indicates that this member is typified by a high content of bioclastic material,
including foraminiferal tests and molluscan shell fragments. All these elastics are cemented
by sparry calcite, resulting in a total carbonate content of more than 9 0 % . Microscopic study
found that these bioclasts have been intensely micritised (Fig. 4.22).
56
4.5 D I S C U S S I O N
a. Middle Miocene strata
The Middle Miocene strata economically are the most important strata in the Rembang
Zone. The primary petroleum reservoir rock lies in the regressive quartz sandstone of the
Middle Miocene Tawun Formation (Ngrayong Member). Little is known about organic source
rocks, but the source potential of organic-rich mudstones in this formation has been suggested.
Several suggestions have been put forward regarding the setting and mode of deposition
of the Middle Miocene strata. O n a regional scale, Van Bemmelen (1949) postulated that the
Rembang Zone was the marginal part of Sundaland (Sunda Shelf). The Early and Middle
Miocene basin subsidence and sedimentation that occurred in this zone were results of
diastrophism in southeast Sundaland. Regional studies by Lemigas/BEICIP (1969), Sudiro et
al. (1973) and Ardhana (1993) indicated that during these periods the sedimentary basin was
fragmented by northeast trending ridges which can be traced into the Java Sea (Fig. 4.10). As
shown by the seismic interpretation, sedimentation occurred essentially within locally
developed sub-basins which have different rates of subsidence and sediment supply.
Particularly high subsidence and sedimentation rates occurred in the sub-basin forming the
southwestward extension of the East Bawean Trough [Fig. 4.10 & Sect. 4.1 (seismic line
PWD-23)], and most petroleum accumulations in the Rembang Zone occur in this sub-basin.
Some inferences about the depositional environment and palaeogeography for Middle
Miocene sediments in this sub-basin have been made by various workers. Ardhana (1993) re
examined the depositional setting of the Tawun Formation based on outcrops and well data
mostly from the area between Surabaya and Rembang (Fig. 4.1). His study relied heavily on
the benthonic faunal association in the sediments and the turbidite model developed by Mutti
and Ricci-Lucchi (1972). He assumed that most of the Tawun Formation was derived from
the Bawean Arch (Fig. 4.10). In the south, Ardhana (1993) interpreted that the Tawun
Formation developed as a slope facies and bathyal submarine fan system which are
characterised by deep marine faunas such as Globocassidulina subglobosa, Hoeglundina
elegans and Uvigerina crassicostata. In the area near the depositional high (Bawean Arch)
Ardhana concluded that the formation developed as a tidally influenced deposit. This confirms
a study by Muin (1985) in the area between Blora and Rembang (Fig. 4.1) in which he
concluded that most of the facies developed as coastal sands and tidally influenced deposits.
57
Posamentier et al (1988) and Posamentier and Erskine (1991) pointed out that a submarine
fan system is most likely to occur during the development of a type-I sequence, which is
characterised by the occurrence of a lowstand systems tract. The seismic study on line P W D -
23 (Sect. 4.1) indicates that the Middle Miocene sequences are all type-U, and are not
associated with lowstand systems tracts in which submarine fans commonly occur.
Furthermore, the basin in which the Tawun Formation was deposited never reached a bathyal
environment. A balance between high subsidence and sedimentation rates provided a constant
amount of accommodation space throughout the Middle Miocene, and the sedimentation simply
advanced and retreated depending on the eustatic sea level change. Additionally, faunas such
as Globocassidulina, Hoeglundina and Uvigerina are proven to be poor palaeowater depth
indicators since Phleger (1960) and Murray (1991) documented their occurrence ranging from
middle neritic to bathyal environments.
Less attention has been given to the Middle Miocene strata developed in the basin between
the JS-1 Ridge and the Tobo High (Fig. 4.10). In the past, the clastic sediments were assumed
to be derived from a northerly source. In contrast to the neighbouring basin, the seismic
interpretation (DNR-12, Sect. 4.2) has indicated that the sediment supply was relatively low
and appeared to be derived from bordering highs both to the north and south. However,
whether a similar pattern applies for the whole basin (Central Deep, Fig. 4.10) still has to be
proved through a comprehensive seismic interpretation.
b. Late Miocene and Pliocene strata
In the Rembang Zone the Late Miocene and Pliocene are known as periods of widespread
development of hemipelagic marl (Lemigas/BEICIP, 1969). Terrigenous deposition was
significantly reduced, apart from some rather widespread deposits in the Ledok Formation.
This marly depositional period was contemporaneous with the deposition of similar facies in
the Kendeng Zone - the Kalibeng Formation. The underlying formation, the Kerek Formation,
which is stratigraphically equivalent to the Wonocolo Formation, is characterised by
volcanogenic detritus, indicating a close association with Late Miocene volcanic activity.
The mechanism which led to the widespread deposition of these marly facies is not fully
understood. A n explanation was given by Van Bemmelen (1949) whereby he associated them
with the Middle Miocene "Banyak volcanism", which occurred along the area where the
Quaternary volcanoes lie. A n uplift which accompanied this Miocene volcanism caused some
58
subsidence on the adjacent Kendeng Zone and further tilting on the unstable area in the
Rembang Zone. The present study indicates that the marly deposition was closely related to
a prolonged high stillstand of sea level. A similar relationship is also observed in the Java Sea
(Chapter 7), where the marl deposition during the Pliocene occurred following a relative high
stillstand or slow fall of sea level.
4.6 SUMMARY AND CONCLUSIONS
The Rembang Zone consists of two main sub-basins which were the southwestward
extensions of the East Bawean Trough and the Central Deep. During the Middle Miocene to
Pliocene the number of sequences developed in these sub-basins is the same, but they have
different seismic character as the result of different rates of subsidence and deposition in the
sub-basins. Relatively low sedimentation and subsidence rates, particularly during the Middle
Miocene, occurred in the southwestern part of the Central Deep.
The relative sea level fluctuations and timing during the Middle Miocene to Pliocene in
general appear to be in agreement with the global relative coastal onlap charts established by
Haq et al. (1988). However, the study failed to recognise some of the smaller fluctuations
indicated by Haq et al. (1988).
The seismic and petrographic studies have indicated that sediments deposited in the sub-
basins were derived from the adjacent highs. Thick Middle Miocene paralic deposits in the
basin north of Cepu were derived from the Bawean Arch, whereas in the basin south of Cepu
they were derived from the JS-1 Ridge and Tobo High. Most of these sediments are organic-
rich mudstones and quartz sandstones which are good source and reservoir rocks in these areas.
The Late Miocene and Pliocene periods were dominated by hemipelagic marly facies
deposition (Wonocolo and Mundu Formations), interrupted in the late Late Miocene by a
period of relatively coarse clastic deposition forming the Ledok Formation. Relatively thick
marls of the Wonocolo Formation and probably the Mundu Formation were deposited during
a prolonged highstand of sea level during which terrigenous clastic input and deposition was
reduced. Near the topographic highs the Wonocolo marl changed into deltaic (on the Tobo
High) and into limestone dominated sequences (Bulu Formation in the area between Rembang
and Blora).
59
CHAPTER FIVE
SEDIMENTOLOGY OF THE KENDENG ZONE
5.1 INTRODUCTION
This chapter examines the Late Miocene to Pleistocene sedimentology of the Kendeng
Zone. The main objective of this chapter is to provide a parallel discussion with Chapter 4
which provided details about the Rembang Zone. Further, this chapter aims to review the
palaeoenvironments and depositional settings under which the units were deposited. Due to
the paucity of geophysical data the whole discussion presented in this chapter will be based
on the lithofacies analysis of a number of measured vertical sections. Petrographic and
benthonic microfaunal analyses are used to support the interpretation, particularly when the
lithofacies analysis is ineffective.
The sedimentological discussion includes seven lithofacies units, these are the Kerek,
Kalibeng, Sonde, Atasangin, Lidah, Pucangan and Kabuh Formations. The discussion
emphasises the Atasangin, Lidah and Pucangan Formations, since the sedimentological settings
of these units were not discussed in detail by previous authors (e.g. de Genevraye & Samuel,
1972; Pringgoprawiro, 1983). A detailed discussion on the Kerek, Kalibeng and Sonde
Formations will not be given in this chapter. These formations have been widely studied by
previous authors (Lemigas/BEICIP, 1969; van Gorsel & Troelstra, 1981; Muin, 1985).
However, the previously defined depositional environments are discussed in the light of new
data and environmental models. The upper part of the Kabuh Formation and the Notopuro
Formation mainly consist of volcanic deposits (Duyfjes, 1938b,d), and due to the paucity of
the data, will not be discussed.
The facies analysis is based on 14 measured vertical sections, with the eleven most
complete sections presented in the discussion. The localities of these sections are given in
Figure 5.1. This limited number of measured sections is primarily due to the scarcity of good
outcrops which further limits the discussion about lateral facies relationships of some
lithostratigraphic units. The analytical methods for the petrographic study are similar to those
used in Chapter Four (i.e. transmitted light microscope and X-ray diffraction methods), and
provide documentation about the general composition and provenance of the succession. The
60
benthonic faunal analysis aims to identify faunal assemblages which are characteristic of
certain environments. This study focuses on the benthonic foraminifers and follows the criteria
given by Phleger (1960) and Murray (1973; 1991).
5.2 KEREK FORMATION
5.2.1 Lithofacies
The lithofacies analysis of the Kerek Formation is based on a vertical section measured
by Muin (1985) in the Solo River section north of Ngawi (Fig. 5.1). The generalised section
is presented in Figure 5.2. The uppermost 26 m was measured during the study (Fig. 5.3), and
was particularly aimed at examining the upper boundary of the formation.
The exposed Kerek Formation in the Solo River section attains a thickness of about 800
m (Fig. 5.2). The section consists of a monotonous repetition of sandstone and mudstone in
which Muin (1985) found various sedimentary structures typical of Bouma-type sequences.
Complete Bouma sequences (Ta-e) are rare, but some complete Bouma sequences are
characterised by a 10 to 75 cm conglomeratic sandy unit with erosional base (Bouma facies
Ta), grading into parallel-laminated sand (facies Tb; 30 to 110 cm thick), ripple-laminated and
convolute-bedded sandstone of about 20 cm thick (facies Tc), parallel-laminated fine-grained
sandstone of about 5 cm thick (facies Td), and pelagic marl (facies Te). Most of the coarse
clastic units are characterised by abundant plagioclase, lithic fragments and amphibole,
suggesting a very close association with volcanism. Two thick tuff beds (1 and 8 m) occur,
indicating deposition during an active period of volcanism.
The uppermost 26 m of the Kerek Formation in the Solo River section is characterised by
interbedded parallel-laminated calcareous mudstone and cross-bedded, ripple and parallel
bedded medium- to coarse-grained sandstone (10 m thick; Figs 5.3 & 5.4). Bed thickness
varies from 5 to 20 cm. The mudstone and sandstone are calcareous due to the abundant
occurrence of planktonic foraminifers. These underlie a 10 m thick calcareous mudstone which
is topped by about 6 m of calcirudite. The calcirudite is composed of angular to subrounded
clasts of medium- to coarse-grained limestone with dioritic and andesitic rock fragments.
Larger benthonic foraminifers are abundant in the calcirudite, but many of these specimens
show indications of mechanical destruction and are thought to have been reworked from
previous limestone.
61
5.2.2 Petrography
The petrographic study was carried out on four samples from the uppermost units in the
formation (Appendix B.l); the stratigraphic position of the samples is indicated in the lower
part of Figure 5.3. T w o samples (KE1A & KE3) are characterised by medium-grained
sandstone and sandy mudstone. They are composed of a mixture of planktonic foraminifers
and terrigenous fragments which are often differentiated into contrasting laminae (1 to 3 m m
thick; Fig. 5.6). The terrigenous elastics make up 6 0 % of the sequence, with an average grain
size of about 0.6 m m . They consist mainly of angular to subrounded plagioclase, lithic
fragments (volcanic) and mafic minerals.
The texture of the calcirudite in the uppermost part of the formation indicates that it was
deposited in a high energy environment; mud or carbonate mud is nearly absent. It is
characterised by an abundance of coral fragments and larger benthonic foraminifers; genus
Lepidocyclina is the most common (Fig. 5.7). Algal fragments are minor and planktonic
foraniinifers are absent. This unit is also characterised by the presence of about 3 5 %
terrigenous elastics. These are rounded, coarse-grained to pebble size diorite and andesite,
plagioclase and very minor amounts of mafic minerals.
5.2.3 Depositional setting and discussion
Turbiditic deposits equated with the Kerek Formation are widely developed in the Kendeng
Zone (van Bemmelen, 1949; Pringgoprawiro, 1983; de Genevraye & Samuel, 1972; Muin,
1985) where it attains a thickness of more than 1000 m. The main terrigenous clastic
components in this deposit have been reworked from the Middle to Late Miocene volcaniclastic
strata referred to as the Banyak Volcanism (van Bemmelen, 1949). The volcanism occurred
along the area which is presently occupied by the Quaternary volcanic centres. Figure 5.10
shows a tentative palaeogeographic map of East Java in the Late Miocene.
The uppermost part of the formation has attracted some attention (e.g. van Bemmelen,
1949; van Gorsel & Troelstra, 1981; Pringgoprawiro, 1983). Based on sedimentary structures
and vertical facies sequence, this part appears to be regressive, and was capped by the
deposition of a relatively thick (6 m ) calcirudite, which van Gorsel and Troelstra (1981)
suggested was deposited in an outer neritic environment as indicated by a consistent
appearance of benthonic forarninifers Eggerella bradyi, Melonis soldanii and Globobulimina
pupoides. They suggested that the calcirudite represents an olistostromal deposit. This implies
62
that the calcirudite has a local extent. In contrast, a regional study by van Bemmelen (1949)
found that the calcirudite is stratigraphically equivalent with a conglomeratic unit containing
andesitic boulders deposited in some areas near Miyana (some 30 k m east of Ngawi; Fig. 5.1).
H e suggested that these conglomeratic units represent a basal conglomerate which marks an
unconformable relationship with the overlying Kalibeng Formation. The present study in the
Solo River section indicates that the lower part of the Kalibeng Formation, just above the
calcirudite, is characterised by a repetitive graded sequence typical of storm deposits (see
Chapter 5.3). This implies that the depositional environment of the unit was shallow marine,
which further confirms the van Bemmelen (1949) suggestion that calcirudite m a y represent a
wave-dominated environment.
In the past, the Kerek Formation has been regarded as a submarine fan deposit. It has
been referred to by van Bemmelen (1949) as a flysch deposit. This opinion was confirmed
by van Gorsel and Troelstra (1981) who studied the benthonic faunal content, and by Muin
(1985) who studied the lithofacies of this formation in the Solo River section. But the
palaeogeography, depositional setting and timing of deposition suggest that the Kerek
Formation should not be regarded as a normal submarine fan deposit.
Posamentier et al (1991) indicated that the presence of sufficient slope length is essential
to allow significant gravity flows and turbidity currents to develop. In support, Walker (1992)
pointed out that in order to be preserved as turbidite, the deposits must occur below storm
wave base, effectively at a minimum depth of 250-300 m. The extensive Plio-Pleistocene
Mississippi fan system, for example, occurred in a bathyal setting in the Gulf of Mexico
(Weimer, 1991). Textures and sediment composition of the Kerek Formation suggest that the
deposition was near or probably on the apron of the Banyak volcanoes. A benthonic
foraminiferal study in the Solo River section by Lemigas/BEICIP (1974) indicated that most
of the formation in this area was deposited in an inner to middle neritic setting, a region which
is effectively influenced by tidal currents (Dalrymple, 1992), fair-weather currents and more
importantly storm-induced currents (Johnson & Baldwin, 1986; Walker & Plint, 1992).
The biostratigraphic studies by Lemigas/BEICIP (1974) and Pringgoprawiro (1983)
indicated that most of the Kerek Formation in the Solo River section interfingered northward
with the hemipelagic marly deposit of the Wonocolo Formation (Fig. 5.10). The seismic
stratigraphic study in the Rembang Zone (Chapter 4) reveals that the latter was mostly
63
deposited during a sea level highstand. The Kerek Formation submarine fan deposition, if it
occurred, contradicts the sequence stratigraphic concepts which suggest fans are most
commonly deposited during a sea level lowstand (Posamentier et al, 1988; 1991). However,
further validation m a y be required as the sequence stratigraphic concepts were tested mainly
on passive margins, a different tectonic setting from the study area.
The present study suggests that at least two possible depositional modes occurred in the
Kerek Formation. These are a stoim-dorninated mode and turbiditic current deposition related
to the collapse of volcanic edifices. Because of the tectonic setting of the study area, both may
occur simultaneously. This suggestion relies on the fact that (1) most of the formation is
dominated by Bouma-type sequences, (2) deposition occurred in a shallow marine setting
(Lemigas/BEKTP, 1974), (3) sediment composition and the presence of tuff layers suggest a
close association with active volcanism, (4) high depositional slope was probably present, and
(5) the regional tectonic setting suggests that seismicity was high (Hamilton, 1979).
Bouma-type sequences regarded as graded storm deposits have been recognised by Nelson
(1982) who studied the occurrence of graded sand beds on the Bering Shelf. H e estimated that
during the major storms a coastal set up may occur and promote liquefaction of the upper 2
to 3 m of sediment. Storm-associated bottom currents can transport liquefied sediments more
than 100 k m offshore. Johnson and Baldwin (1986) documented that the thickness of a single
graded storm unit can be up to 0.5 m , and is characterised by an erosive base, a basal lag
deposit, horizontal low angle lamination which is possibly related to hummocky cross-bedding,
wave ripple cross-lamination and a burrowed upper interval. With regard to the tectonic
setting of the study area, similar deposition may have occurred through seismically-induced
waves, such as tsunamis, which often characterise shelves near seismically active region
(Coleman, 1969, cited in Johnson & Baldwin, 1986).
Turbiditic sequences related to volcanic activity have been reported by some authors.
Busby-Spera (1988) documented three lithofacies developed during the deposition of the
Mesozoic volcaniclastics in Baja, California; these are (i) tuff, (ii) lapilli-tuff and tuff breccia,
and (iii) primary volcanic lithofacies. Based on the fact that lateral thickness variation is high,
the tuff lithofacies was interpreted to have been deposited from subaqueous turbidity flows
which followed a palaeo-low. The lapilli-tuff and tuff breccia lithofacies are thick bedded, are
inversely graded in the proximal area (debris flow) and channelised, graded and stratified in
64
the distal parts, and resulted from high-density turbidity currents. The primary lithofacies
consist of dacitic pyroclastic rocks and pillow lavas.
Similar facies and modes of deposition as the lapilli-tuff and tuff breccia lithofacies of
Busby-Spera (1988) have been reported by Hathway (1994) in the Oligo-Miocene
volcaniclastic arc assemblage in the intra-oceanic/fore-arc area of southwestern Viti Levu, Fiji.
He recognised that in the proximal area beds are typically massive and may show inverse
grading. In the distal part graded sandstone and mudstone units occur and show incomplete
Bouma sequences. These units are characterised by sharp-bases, and topped by lime mudstone
which may be bioturbated. He interpreted that the deposition was primarily from high-density
mass flows which were probably initiated by the upslope collapse of unstable volcaniclastic
debris. Based on the sedimentological character, most of the Kerek Formation may be
categorised into the distal facies of the lapilli-tuff and tuff breccia lithofacies of Busby-Spera
(1988) or the graded sandstone and mudstone of Hathway (1994). Figure 5.11 shows the
inferred depositional setting of the Kerek Formation, based on Hathway's (1994) model.
5.3 KALIBENG FORMATION
5.3.1 Lithofacies
A typical succession through the Kalibeng Formation is represented by the Solo River
section (Fig. 5.3) where the total thickness is about 425 m. The section was measured on the
southern flank of an anticline north of Ngawi (Fig. 5.1). Texturatfy, this formation is nearly
homogeneous and massive, and consists mostly of forarniniferal-rich pelagic/hemipelagic marl.
Biostratigraphic studies by Lemigas/BEICIP (1974) and van Gorsel and Troelstra (1981)
indicated that deposition occurred from the Late Miocene to Pliocene (planktonic foraminiferal
zones N17 to N21; Chapter 3.2.2 8c 3.3.3). Pringgoprawiro (1983) suggested the formation
is contemporaneous with the Ledok and Mundu Formations of the Rembang Zone.
The base of the formation is characterised by about 10 m of bioturbated (Cruziana
ichnofacies), carbonaceous sandy marl which lies directly on the calcirudite forming the top
of the Kerek Formation. Some thin (15-20 cm) interbeds of sandy calcarenite contain coarse
grained to pebble size mudclasts, and show planar crossbedding and ripple laminations. Some
of these beds are lensoidal and show an erosional base. Upward, the number and thickness
of the sandstone beds diminish and the unit is dominated by unstratified, foraminiferal-rich
marl.
65
The deposition of the middle and upper parts of the formation appeared to have been
influenced periodically by strong bottom currents which are evidenced by the presence of some
relatively thick crossbedded sandy units. The 7 m thick unit in the middle part of the
formation is dominated by calcareous fine-grained sandstone and is characterised by westward-
trending trough crossbedding (Fig. 5.5). The uppermost part of the formation is characterised
by planar crossbedded glauconitic calcarenite. Here, the sandy character is due to an abundant
occurrence of planktonic forarniniferal tests.
5.3.2 Petrography
Six rock samples have been studied microscopically, all of them from calcarenite and
calcareous sandstone interbeds within the marl. The stratigraphic positions of the samples are
indicated by prefix K B in Figure 5.3, and Appendix B.l summarises the results of the analysis.
Calcarenite samples from the lower part are characterised by bioclasts, mostly consisting of
foraminiferal tests which range in abundance from 20 to 7 5 % , surrounded by carbonate m u d
matrix, giving a total carbonate content of up to 95%. Fine-grained terrigenous clastic grains
are minor, less than 6%.
Calcareous sandstone interbeds occur in the middle of the formation. They are
characterised by fine- to medium-grained terrigenous fragments (plagioclase and hornblende)
which form about 3 2 % of the rock (Fig. 5.8). Rounded limestone fragments (about two thirds
of total carbonate) are mostly muddy limestone (reworked marl). Bioclasts (mostly
forarniniferal tests) are minor (< 1 0 % of total carbonate).
In the upper part of the formation an abundance of foraminiferal tests characterises the
calcarenite. They reach 7 0 % of the total carbonate and are cemented by micrite. Glauconite
reaches 2 1 % (sample no. K B 15), and mostly occurs interstitially within forarniniferal tests (Fig.
5.9). Terrigenous clastic grains, which are commonly plagioclase fragments, are minor (less
than 2 % ) .
An X-ray diffraction analysis was carried out on two marl samples, primarily to define the
type of mineral in the formation. All diffractograms (Fig. 5.12) reveal that calcite is the major
component. Clay minerals (kaolinite and smectite), quartz, plagioclase and mafic minerals are
minor, indicating a low terrestrial influence and that most of the material has been reworked
from calcareous faunas.
66
5.3.3 Depositional environment and discussion
The lithofacies and petrographic studies indicate that the Kalibeng Formation shows tittle
facies variation and only minor terrigenous clastic input. This implies that the formation is
composed primarily of marine derived materials which have been deposited by suspension
sedimentation throughout the accumulation of the formation. The dominant composition,
which is mainly an association of carbonate mud and foraminiferal tests, is typical of a pelagic
deposit. This deposit accumulated during a period without major bottom current activity by
particle by particle settling from the overlying water (cf. Reineck & Singh, 1980).
Because of these sedimentary features, in the past, interpretation of the depositional
environment of this formation appeared to rely only on the results of benthonic faunal analysis.
Lemigas/BEICIP (1974) estimated that the environment of deposition was outer neritic to upper
bathyal, based on the association of benthonic foraminifers Bathysiphon, Dentalina, Planulina,
Stilostomella, Nodosaria, Lenticulina, Gyroidina, Pyrgo, Oridorsalis and Bulimina. A n
extreme result has been obtained by van Gorsel and Troelstra (1981) using a similar analysis.
They estimated a palaeowater depth of about 1000 m, based on the benthonic foraminiferal
assemblage of Oridorsalis umbonatus, Gyroidina neosoldanii, Planulina wuellerstorfi,
Anomalina globulosa and Uvigerina auberiana, whereas a slightly shallower condition (about
400 m ) occurred near the upper boundary. Their conclusion, in fact, appears to disagree with
Murray (1991) who documented that these species are commonly present in a wide range of
environments, from shelf to bathyal.
The present study indicates that beside benthonic faunal evidence, there are at least three
parameters which have to be considered before a reliable depositional environment can be
determined for the Kalibeng Formation. These are: (1) detailed sedimentary features, (2)
palaeogeography during deposition, and (3) eustatic sea level changes during the Mio-Ptiocene.
A climatostratigraphic study in the Solo River section by van Gorsel and Troelstra (1981;
Chapter 3.4) recognised the presence of three cold periods during the deposition of the
Kalibeng Formation, which occurred at the beginning (cold period I), middle (cold period m )
and end (cold period IV) of the deposition (Fig. 3.4). Biostratigraphic studies by van Gorsel
and Troelstra (1981) and Pringgoprawiro (1983) indicated that the segment of the Kalibeng
Formation between the cold periods I and m is laterally equivalent with the Ledok Formation
in the Rembang Zone; the segment between the cold periods m and V corresponds with the
67
Mundu Formation (Chapter 3.3 & 3.6, Table 3.4). The cold periods I and HI probably
correlate with the sea level falls recognised on seismic line DNR-12 (Chapter 4.3.2), which
occurred before and after the Ledok Formation deposition (seismic unit L M 2 ) .
Discussion presented earlier indicated that the sandy facies in the Ledok Formation, on
both seismic lines P W D - 2 3 and DNR-12, was deposited on a shallow shelf influenced by tidal
currents (seismic unit L M 2 ) . Thus, from the facies difference, the lower part of the Kalibeng
Formation probably represents a deeper facies of the Ledok Formation. Although van Gorsel
and Troelstra (1981) did not provide any sedimentological description for the Kalibeng
Formation related to their defined cooling periods, a field study carried out by the author
recognised the occurrence of some sedimentary features which possibly relate to these periods.
The lowermost part of the Kalibeng Formation in the Solo River section, where the cold
period I occurs, is characterised by 6 m thick calcirudite, which from the texture and
composition was probably deposited in a high energy environment (see discussion in chapter
5.2). This, in m m , is overlain by about 10 m of bioturbated (Cruziana ichnofacies),
carbonaceous sandy marl with some thin interbeds of sandy calcarenite, containing pebble-size
mudclasts, and showing planar crossbedding and ripple laminations. Some of these beds are
lensoidal and show an erosional base. These features suggest that the unit was probably
deposited during alternation between storm and fair-weather bottom currents. The Cruziana
burrows in the marl suggest that the unit was deposited in a sublittoral environment which is
commonly typified by moderate to low energy levels (Frey & Pemberton, 1984), and was
beyond the reach of normal wave base (Walker & Plint, 1992). The presence of some
lensoidal sandy calcarenites with an erosional base suggests that the unit may have been
subjected to storm-generated waves (cf. Reinson, 1984; Johnson & Baldwin, 1986). During
a storm event, a sandy unit was deposited by a storm-induced current, followed by a
bioturbated marl unit during the next fair-weather period.
The middle part of the Kalibeng Formation in the Solo River section, suspected to be the
stratigraphic position of cold period DJ, is characterised by trough cross-bedded sandstone.
This 7 m thick unit is indicated in Figures 5.3 (position 230 m ) and 5.5, and petrographically,
is characterised by up to 5 0 % terrigenous clastic detritus. This unit appears to have been
deposited as sand waves, which are common in an environment dominated by tidal currents
(Dalrymple, 1992). This implies that during the cold period the depositional environment
68
became a shallower tide-dominated area. Haq et al (1988) indicated that the magnitude of
eustatic sea level fluctuations during the Late Miocene was about 95 m.
With regard to these environments, the bathyal depositional environment for the Kalibeng
Formation, as suggested by van Gorsel and Troelstra (1981), appears to be incorrect, although
there was a combined significant sea level rise and basin subsidence.
The upper segment of the Kalibeng Formation between cold periods in and IV has been
correlated with the Pliocene marl Mundu Formation in the Rembang Zone. The deposition of
this marl was widespread (Fig. 5.13), and appears to be in response to the significant eustatic
sea level rise in the Early Pliocene (cf. Haq et al, 1988). The seismic sequence stratigraphic
study of the Java Sea (Chapter 7) indicates a relative sea level rise of about 115 m. It must
have occurred relatively quickly, as evidenced by the lack of recognisable transgressive
systems tracts (note that the seismic system used has about 2 m resolution), and was followed
by a slow relative sea level fall. Due to these circumstances, in the Java Sea, most Pliocene
deposits appear to have been deposited under regressive conditions, and as sea level slowly
fell, most of the deposits would be in middle and outer neritic environments. This type of
deposition is believed to have occurred in the Kendeng Zone, where due to its tectonic
position, the subsidence rate was probably higher than in the Java Sea. Most of the marl in
the Kalibeng Formation may be the result of prolonged regressive conditions because of the
slow relative sea level fall, following a rapid sea level rise in the Early Pliocene. The low
content of siticiclastic detritus in this formation, as evidenced by the petrographic data, appears
to support this mode of deposition as such detritus would be deposited closer to the basin
margin.
5.4 ATASANGIN FORMATION
The Atasangin Formation was developed in the eastern part of the Kendeng Zone, and was
deposited in the Middle Pliocene (Chapter 3.2.2 & 3.3). The succession was first mapped and
described by Duyfjes (1938a,b,d) and subsequently by de Genevraye and Samuel (1972), but
interpretation of its depositional environment and setting has never been carried out. T w o
facies typically occur: a deltaic sandstone facies in the lower part and a finely-laminated
siliceous marl facies in the upper part. The total thickness in the eastern Kendeng Zone is
about 400 m.
69
Because of the scarcity of good outcrops, study of this formation could only be carried out
on the north Kabuh section, which is located some 1.5 k m north of Kabuh (Fig. 5.1). The
Atasangin Formation in this section is underlain by the unstratified marly (prodelta) Kalibeng
Formation, which here, was probably deposited in a shelf environment, based on the content
of benthonic foraminifers Bolivina, Bulimina, Cibicides, Laticarinina, Planulina and Siphonina
(Appendix C.2, sample no. NKabuh3; cf. Murcay, 1973; 1991). Representative sections of the
Atasangin Formation are given in Figures 5.14 and 5.15.
5.4.1 Lithofacies
a. Sandstone facies
The sandstone facies occupies the lowermost 90 m of the Atasangin Formation; Figure
5.15 shows the representative vertical section of this facies. Overall facies show a regressive
upward sequence which based on sedimentary textures and structures, can be subdivided into
two main subfacies: coarsening-upward sandstone subfacies, and fine- and coarse-grained
sandstone facies.
a.l Coarsening-upward sandstone subfacies
Facies:
This facies occupies the lowermost 35 m of the north Kabuh section, and is directly
underlain by the Kalibeng Formation (Fig. 5.15).
The lowermost 17 m is characterised by alternating foraminiferal-rich mudstone and
calcareous sandstone on a centimetre to decimetre scale, which is overlain by massive and
parallel laminated sandstone consisting of fine-grained bioclastic (foraminiferal tests),
plagioclase and lithic fragments (andesite).
The uppermost 18 m consists of a coarsening-upward sequence of medium- to coarse
grained sandstone. Thickness of successive sandstone units progressively increases from 20
c m to 3 m , and they are separated by 5 to 20 c m sandy mudstone beds which commonly
contain burrow stmctures. Ripple lamination is the most common stmcture in the sandstone
units. Mudclasts, 2 to 5 c m diameter, are common in the upper parts.
70
Interpretation:
A coarsening upward sequence is commonly associated with a prograding wave- or storm-
dominated shoreface (Reinson, 1984; Bhattacharya 8c Walker, 1992) but it m a y also represent
a delta front facies in a wave-dominated environment (Bhattacharya & Walker, 1992). The
differentiation between these two environments appears to rely on the availability of three-
dimensional control. As such control is absent in this study, a reliable interpretation m a y be
difficult to reach. However, overall sequence tendency, which regresses toward an association
of possible channel levee complexes in a delta plain environment (sub-section a.2) suggests
that the coarsening-upward sequence may have been deposited in a delta front environment.
In this environment, a coarsening-upward sequence occurs due to the amalgamation of a series
of coalesced distributary mouth bar crests and their associated distal bars (Cotter, 1975).
The lower part, which is characterised by interbedded fine-grained sandstone and
mudstone, is interpreted to represent a transition from the underlying prodelta facies (the marly
Kalibeng Formation). The abundant forarniniferal tests in the observed facies indicate a
relatively strong open marine influence.
a.2 Fine- and coarse-grained sandstone subfacies
Facies:
This subfacies overlies the upward-coarsening sandstone subfacies and attains about 54 m
in thickness.
The fine-grained sandstone subfacies is characterised by interbedded sandy marl/mudstone
and fine- to coarse-grained sandstone. In the lower facies (interval between 36 and 50 m; Fig.
5.15), the sandstone beds gradually thicken upward from 10 to 25 cm, are separated by 5 to
20 c m thick mudstone beds, and are characterised by common parallel and ripple laminations.
Deformation stmctures (such as convolute bedding, Fig. 5.16) and mudclasts are also common.
Some of these sandstone horizons are normally graded into parallel laminated mudstone (Fig.
5.16). In the upper facies (interval between 60 and 68 m, Fig. 5.15), the facies is characterised
by interbedded parallel laminated mudstone and parallel bedded, fine- to coarse-grained
sandstone. Some of the sandstone units are considerably thicker, up to 1.5 m, and contain very
small amounts of planktonic foraminifers.
71
The coarse-grained sandstone subfacies consists commonly of fining-upward units with
basal scours and lag deposits. The sandstone is moderately sorted, without muddy matrix and
consists mainly of rounded andesite fragments. Crude bedding and parallel bedding are
common in the coarser-grained sandstone. In the lower part (interval between 50 and 55 m ;
Fig. 5.15), the facies occurs as small scale channel-fills, which have an average width and
depth of about 1.5 m and 60 c m respectively (Fig. 5.17). In the uppermost part, (interval
between 83 and 85 m ; Fig. 5.15), the facies is characterised by coarse-grained sandstone and
breccia, which tend to be cmdely bedded, and by consolidated blocks of mud (up to 10 cm).
The main fragment composition is andesitic rock and mud blocks.
Interpretation
The fine-grained sandstone subfacies in the lower part (interval 36-50 m; Fig. 5.15) is
characterised by some thin fining-upward beds and deformation structures. A fining-upward
sequence commonly develops in fluvial deposits and is related to lateral migration of point bars
(Allen, 1970). However, the absence of basal scours between this facies and the underlying
facies (upward-coarsening sandstone facies) suggests that these fining-upward beds are not
related to a point bar system. Moreover, Coleman and Prior (1982) indicated that fining-
upward units could occur in sandy crevasse splays of a delta system.
The occurrence of deformed structures in this fine-grained sandstone facies (interval 36-50
m; Figure 5.15) indicates that the sediment was partially liquefied during deposition of the
sandy facies (cf. Allen, 1977; Reineck 8c Singh, 1980). These sedimentary structures are not
unique to a single environment and could develop in at least three different environments:
channel bar deposits (Coleman, 1969), subaqueous levees of a delta front environment
(Reineck & Singh, 1980), and turbidite systems (Middleton & Hampton, 1976). The close
association of this fine-grained sandstone facies with fining-upward beds and distinct channel
structures (interval 50-55 m ; Fig. 5.15) suggests an occurrence in a channel-levee complex of
a delta plain environment. The occurrence of foraminifera-bearing sandstone beds (interval
55-68 m; Fig. 5.15) indicates a rather strong marine influence (?an interdistributary bay).
The coarse-grained sandstone facies in the upper part (interval 68-90 m: Fig. 5.15) was
probably deposited as channel fill deposits, as evidenced by the occurrence of basal scours and
lag deposits. The fining-upward sequences in this facies may be the result of lateral migration
of point bars within a meandering river or distributary system, as characterised by Allen (1970)
72
and Walker and Cant (1984). The occurrence of coarse-grained sandstone, breccia and mud
blocks in the uppermost part suggests possible temporary flood events. The coarse-grained
sandstone facies, thus, may represent a mid to upper delta plain environment on which riverine
processes are dominant (Wright, 1982).
b. Diatomaceous marly facies
The diatomaceous marly facies developed during a transgression following the deposition
of the previous deltaic sandstone facies. The facies consists mostly of light grey, finely-
laminated marls with some thin interbeds (15-20 cm) of very fine-grained calcareous sandstone.
Diatom frustules, as well as radiolaria tests and spongidiscs, are abundant, indicating the study
area was sufficiently nutrient-rich to allow a high productivity of diatoms. This diatomaceous
marly facies always shows a consistent parallel-bedding enhanced by the diatom-rich beds.
Massive beds occur locally particularly in the middle part of the succession (Fig. 5.14).
The depositional environment as defined from the benthonic faunal analysis (foraminifers)
during this study (Chapter 5.4.3) is inconclusive, particularly because of the occurrence of
mixtures of inner (brackish/hypersaline) and middle/outer neritic assemblages. Duyfjes (1938b)
reported the occurrence of some fresh water molluscs (Unio) in the upper part of the siliceous
marl in the Atasangin Formation on an area some 6 k m northeast Kabuh (Fig. 5.30). This
suggests that the diatomaceous marly facies was probably deposited in a marginal marine
environment.
5.4.2 Petrography
The petrographic study was based on three representative samples; one from the deltaic
sandstone facies was studied by using a transmitted light microscope whereas the other two
(from the siliceous marly facies) were analysed by X-ray diffraction.
The sandstone sample consists mainly of andesitic (and some basalt) terrigenous clastic
detritus (Appendix B.l). Rock fragments (30%) are the most abundant clastic component
along with plagioclase (17%) and mafic minerals (13%) (Fig. 5.18). Fragments are cemented
by sparry calcite (34%). The rock fragments are commonly oxidised and glauconitised with
the grain size varying from 0.05 to 0.7 m m . They show a rounded form.
73
The X-ray diffractograms of the siliceous marly facies (Fig. 5.19) show minor occurrence
of two clay minerals (smectite and kaolinite) which probably reflects a minor terrigenous
influence or minimal weathering in the source area (Hardy & Tucker, 1988). The high
siliceous content in the Atasangin Formation is reflected by the high content of quartz in the
samples analysed. A high calcite content is displayed by sample no. NKabuh 13. This sample
is from the middle part of the succession which is characterised by abundant calcareous fauna.
5.4.3 Benthonic fauna
Benthonic foraminiferal analysis was carried out on eight marly samples, mostly from the
diatomaceous marly facies. The results of the analysis are summarised in Appendix C.2, and
the stratigraphic positions of the samples are indicated on the vertical section Figure 5.14.
All samples analysed are characterised by low species diversity, averaged at 9 species per
sample. This causes the maximum species diversity index of Murray (1973) to fall below 3.
Murray (1991) indicated that a diversity index of less than 5 is typical of brackish or
hypersaline marginal marine environments. A rather high species diversity (21 species) occurs
in sample no. NKabuhl6, which is stratigraphically located in the middle part of the succession
(Fig. 5.14). This part corresponds with the occurrence of marl beds with massive structure,
and is interpreted to represent normal salinity marine water.
The faunal association in most of the samples studied appears to be a mixture of inner and
middle/outer neritic faunas (Appendix C.2). This is apparent in samples NKabuhl2,
NKabuhl6 and JKB10, which represent the lower, middle and upper parts of the section (Fig.
5.14), respectively. The genera Ammonia and Elphidium, which commonly characterise
brackish and hypersaline lagoons, occur associated with middle/outer neritic genera Planulina,
Laticarinina, Lenticulina, Bolivinita and Gyroidina (cf. Phleger, 1960; 1965; Boucot, 1981;
Murray, 1973; 1991).
Such unusual assemblages have been reported by some other authors, and could result by
two possible means, upwelling or tidal currents. Sen Gupta et al. (1981) documented the
unusually high abundance of Bolivina subaenariensis on the continental slope off Daytona
Beach (Florida), whereas this species is generally very rare along the western North Atlantic
margin. They suggested that because of the varying influence of the Gulf Stream and seasonal
upwelling, the environment undergoes abrupt physical and chemical changes which favour the
74
species Bolivina subaenariensis. Sloan (1992) reported the occurrence of open marine-type
ostracods in the Late Pleistocene Yerba Buena mud, San Francisco Bay (California), which
elsewhere is commonly characterised by estuarine taxa. The Yerba Buena m u d represents the
deposits of a single transgressive event. The deposition commenced during the last glacial
maximum with brackish and low salinity environments continuing to the present day bay (25-
70 m below mean sea level). H e interpreted that the open marine ostracods were introduced
into the bay by tidal currents which became more prominent as sea level increased.
5.4.4 Discussion
The sedimentary texture and composition of the deltaic sandstone facies of this formation
suggest that volcanic detritus was the main source. Lemigas/BEICIP (1974), de Genevraye and
Samuel (1972) and Pringgoprawiro (1983) reported the occurrence of volcanic breccia at the
same stratigraphic position as the deltaic sandstone facies in areas near the Pandan Volcano
(Fig. 5.1). The absence of similar facies in some areas, such as Ngawi and Ngimbang (Fig.
5.1), indicate that the breccia and deltaic deposits are probably very localised around aprons
of some volcanoes. Figure 5.20 displays the inferred palaeogeography during the deposition
of the deltaic sandstone facies in the Middle Pliocene.
The facies difference between the deltaic sandstone facies and the underlying marly
Kalibeng Formation suggests that the deltaic deposition accompanied a relative sea level fall.
The occurrence of deltaic progradation during such times has been indicated experimentally
by Koss et al. (1994). Biostratigraphic studies by Ninkovich and Burckle (1978) and the
author during the present study (Chapter 3.3.1c and 3.3.2) indicated that the Atasangin
Formation was deposited during the Middle Pliocene (planktonic foraminiferal zones N19/N20.
Tentative correlation with the Haq et al. (1988) eustatic sea level curve suggests that the
deltaic deposition corresponds with the eustatic sea level fall 3.8 M a ago.
The siliceous marly facies was probably deposited during the transgression which followed
deltaic deposition. Abundant diatoms, as weU as other siliceous taxa, suggest that the
environment of deposition was rich in nutrients. Jenkyns (1986) indicated that nutrient-rich
water is introduced when deep-water inflow and shallow-water outflow prevails. Calvert
(1966) documented the occurrence of large accumulations of diatomaceous siliceous sediments
in the Gulf of California, and suggested that circulation of upwelling currents supplied
dissolved silica from the Pacific Ocean and was responsible for the formation of these deposits.
75
A biostratigraphic study in the Solo River section (Figs 3.4 & 5.1) by van Gorsel and
Troelstra (1981) indicated that during the Middle Pliocene, deposition of the marly Kalibeng
Formation occurred in a middle/outer neritic environment (chapter 5.3). In the eastern
Kendeng Zone, which is represented by the north Kabuh section, a shallower marginal marine
environment was probably present as indicated by the low species diversity
(?lagoonal/estuarine; Fig. 5.21). The unusual faunal assemblage in the siliceous marly facies
in this section may reflect the occurrence of water circulation between the shallow marine
environment in the eastern Kendeng Zone and the deeper marine facies in the western Kendeng
Zone (represented by the Solo River section), which also led to the introduction of nutrient-rich
water to allow the development of siliceous taxa.
5.5 SONDE FORMATION
5.5.1 Lithofacies
A vertical section of Sonde Formation was measured on an exposure along the east bank
of the Solo River, north of Ngawi (Fig. 5.22). The formation was deposited during Middle to
Late Pliocene mostly in the central and west Kendeng Zone (Chapter 3.2.2). It is conformable
with the Late Miocene-Early Pliocene Kalibeng Formation.
The Sonde Formation consists of two main facies, the limestone and mudstone facies. The
limestone facies is approximately 65 m thick, and consists of bedded calcarenite at the base
and a coarse clastic limestone unit containing coral fragments and algae in the upper part (Fig.
5.23 and Fig. 5.24).
Van Gorsel and Troelstra (1981) identified the presence of benthonic foraminifers
Rectobolivina, Bolivina, Uvigerina, Trifarina, Globocassidulina and Hoeglundina near the base,
and species Amphistegina papulosa and Planorbulinella elatensis in the middle part of the
limestone facies, from which they suggested a progressive shoaling of the sequence from an
outer to inner neritic environment of deposition.
The calcareous mudstone facies of the Sonde Formation is about 80 m thick, and
conformably overlies the limestone facies. This mudstone facies is poorly exposed but, from
a few outcrops, it consists of massive calcareous sandy mudstone with abundant pebble-size
mudclasts (intraclasts). This facies also contains planktonic and benthonic foraminifers, but
according to van Gorsel and Troelstra (1981) most foraminifers were reworked from the Early
76
and Middle Pliocene sediments. These features suggest that the mudstone facies was deposited
in a wave-dominated region.
5.5.2 Discussion
The vertical facies changes from the marl of the Kalibeng Formation and within the Sonde
Formation clearly indicate an upward regressive sequence. In the central Kendeng Zone
(including the Ngawi area), the stratigraphic position of the Sonde Formation, which is overlain
directly by the debris flow deposit of the Pucangan Formation, suggests that the marine stage
has ended in this area. This is in marked contrast to the eastern Kendeng Zone (including the
Jombang and Mojokerto areas, and Madura Strait) where a high rate of basin subsidence
resulted in a thick succession of siliceous marly facies of the Atasangin Formation and
lacustrine mudstone of the Lidah Formation.
5.6 LIDAH FORMATION
5.6.1 Lithofacies
a. Kabuh and Mojokerto areas
The Lidah Formation in these areas was deposited in the Late Pliocene (Chapters 3.2.2
& 3.3.2), and conformably overlies the Atasangin Formation. The first detailed mapping in
these areas was carried out by the Geological Survey of the Netherlands Indies in the 1930s
(Duyfjes, 1938a,b,c,d). Despite a few subsequent studies (e.g. de Genevraye & Samuel, 1972;
Sartono et al, 1981; Pringgoprawiro, 1983), the sedimentological characteristics of the Lidah
Formation have never been examined in detail. Previous studies have mainly examined the
stratigraphic context of the formation.
In the Kabuh and Mojokerto areas the Lidah Formation attains a thickness of 400 m. A
representative vertical section in these areas is provided in Figure 5.25. The section was
measured in the Sumberringin area, some 3 km northeast of Kabuh (Fig. 5.1). Here, the
formation is characterised mainly by unstratified, homogenous dark blue mudstone.
Sedimentary structures are absent and the bedding orientation can be measured only when a
thin interbed of sandstone is present. This facies is also characterised by streaks of organic
material and some fine-grained pyrite crystals indicating anoxic conditions at the time of
deposition.
77
In these areas, there are at least two sandstone horizons which indicate temporary exposure
to a marine environment. The lower sandstone horizon is about 10 m thick. A detailed study
was not carried out on the lower sandstone facies due to poor exposure. A faunal study on
rock samples from this unit indicates the presence of benthonic foraminifers including the
species Uvigerina peregrina, Bulimina marginata and Bolivina (Appendix C.2) which suggests
an inner to middle neritic environment (Phleger, 1960; Murray, 1991).
More detailed observations were made on the upper sandstone facies of the Lidah
Formation. Figure 5.26 shows the generalised section of this facies, measured in the
Sumberringin area, some 3 k m northeast of Kabuh (Fig. 5.1). This facies is characterised by
a regressive sequence, with a total thickness of 27 m, and can be subdivided into lower,
middle and upper parts, which represent three different environments of deposition. The lower
part (between 2.5 and 6.5 m in Fig. 5.26), consists of massive, bioturbated, muddy sandstone
intercalated with ripple laminated, moderately sorted, medium-grained sandstone. Inner neritic
benthonic forarninifers such as those found in the lower sandstone facies are abundant. This
lower part is interpreted as a subtidal deposit, but the absence of slack-water mud drapes
[common in this environment, Terwindt (1988)] and abundant bioturbation suggest an
environment closer to offshore facies. The middle part (between 6.5 and 21 m in Fig. 5.26)
is characterised by muddy sandstone which fines upward into massively bedded carbonaceous
mudstone. Clay and organic pellets are abundant. This part probably represents a tidal flat
environment. The upper part (between 21 and 29 m in Fig. 5.26) is dominated by fining
upward sandstone facies. The base of the facies is erosional and contains coarse-grained to
pebble-size, rounded fragments of andesite as a lag deposit. Planar crossbedding is common
and ripple stratification is found in some thin intercalations of fine sandstone. This part is
interpreted to represent a fluvial or tidal channel facies.
b. Nglampin area
The Lidah Formation is widely developed in the Rembang Zone. The Lidah Formation
in the Nglampin area (some 25 k m southeast of Cepu, Fig. 5.1) has been studied by Nahrowi
et al (1981). Here, only the uppermost 400 m is exposed. Their biostratigraphic study
suggested that the sequence was developed in the Pleistocene (zones N 2 2 and N23) based on
the present of planktonic foraminifers Globorotalia hirsuta and Globigerina calida calida, and
laterally corresponds with the Pucangan Formation in the Kendeng Zone. Figure 5.27
78
represents the vertical section of the Lidah Formation in the Nglampin area, measured by
Nahrowi et al. (1981).
The exposed section can be divided into three main facies: a lower mudstone, an
intermediate limestone and an upper mudstone facies. The lower mudstone facies is
characterised by massive carbonaceous bluish-grey mudstone, and contains streaks of fine
grained quartz sandstone. Fine-grained pyrite crystals are occasionally found. S o m e thin
interbeds (< 10 c m ) of fine-grained calcareous sandstone are present and commonly show
parallel and ripple lamination. This lower mudstone facies may have been deposited mostly
in anoxic conditions, as were similar facies in the Kabuh and Mojokerto areas.
The limestone facies is characterised by sandy mudstone interbedded with calcarenite. The
sandy mudstone interval varies from 40 c m in the lower part to 7 m in the upper part. The
calcarenite is glauconitic and very rich in mollusc fragments which suggests the influence of
strong waves during the deposition. The thickness of calcarenite varies from 15 c m to 1.5 m
and shows a tendency of thickening upward. This facies may have been deposited in a fairly
shallow marine (?shoreface/foreshore environment).
The upper mudstone facies consists of light grey calcareous mudstone and in some places
is carbonaceous and bioturbated, suggesting a tidal flat environment of deposition. This facies
is also characterised by thin interbeds (< 60 cm) of calcarenite, particularly in the upper part,
which may represent a temporary occurrence of strong waves. Faunal evidence in this facies
supports a tidal flat environment of deposition, as indicated by the occurrence of Ammonia
(Nahrowi et al, 1981).
5.6.2 Petrography
a. Microscope analysis
A petrographic study was carried out on three sandstone samples from the upper sandstone
facies of the Lidah Formation in the Sumberringin area (the results are summarised in
Appendix B.l). The subtidal sandstone is characterised by the presence of fine-grained quartz
fragments (5%), and mollusc fragments and foraminiferal tests (16%). These fragments are
supported by a calcareous clay matrix. Feldspars are minor (< 4%) indicating a low volcanic
influence. The quartz is mainly monocrystalline and shows undulose extinction indicating
strain during its crystallisation. It was possibly reworked from the Middle Miocene sediments
79
(Tawun Formation) which may have been exposed along the core of the Rembang Zone
anticlinorium in the Tuban area.
The sandstone in the tidal channel deposit is characterised by a different mineralogical
content to the subtidal deposit, and suggests a volcanic origin. This sandstone consists of
highly altered, rounded to subrounded fragments of basaltic and andesitic rocks representing
up to 7 2 % of the total sediment. Plagioclase and mafic minerals represent up to 3 0 % and 1 0 %
respectively of the total sediment. At the base of the channel the sandstone is glauconitic and
cemented by carbonate mud.
b. X-ray diffraction analysis
X-ray diffraction analysis was carried out on three mudstone samples from the Kedamean
and Nglampin areas (Fig. 5.1; Appendix B.2).
The diffractograms of the samples from the Nglampin area (Fig. 5.28) show a wide
background spectrum which is known to be typical of organic-rich samples (Mandile & Hutton,
1994). This is in agreement with the general appearance of the mudstone facies which is
carbonaceous. The diffractograms also indicate the presence of the evaporite mineral gypsum.
The precipitation of evaporite minerals from brines is controlled by their relative solubility
(Kendall, 1992). Thus the occurrence of gypsum indicates that the environment was saltier
than normal sea water. With the assumption that sea water was the main source of brine
(Kendall, 1992), during the deposition of the mudstone facies the area may have been
subaerially exposed and only temporarily been flooded with sea water. This suggestion is also
supported by the occurrence of siderite in some of the samples (Nglampin 1, Nglampin7 &
Kedamean4). Siderite is commonly associated with pedogenic concretions (Collinson, 1986).
The diffractogram of the sample from the Kedamean area indicates that smectite and
kaolinite are the most c o m m o n clay minerals in the Lidah Formation in this area. Smectite
is often related to the alteration product of volcanic ash and kaolinite is related to the
weathering of Al-rich silicates (Deer et al, 1992; Hardy & Tucker, 1988). With regard to the
greater X-ray response of kaolinite than smectite (2.5 times; Weir et al, 1975), the
diffractograms suggest that smectite is more abundant than kaolinite, indicating the source was
mostly basic or intermediate volcanism where Al-rich silicate minerals are less common.
80
5.6.3 Discussion
The Lidah Formation is up to 500 m thick in the Kedamean area (north of Mojokerto).
It is also extensively developed in East Java. Field observations during the present study along
the south coast of Madura and Gili Raja Islands, and the seismic study on Madura Strait
(Chapter 6.5.2) confirm the presence of similar facies in these areas. O n the south coast of
Madura and Gili Raja Islands, as well as in the Rembang Zone (Lemigas/BEICIP, 1969), this
formation is underlain unconformably by Pliocene deposits (Mundu Formation and its
equivalent). The lateral extension of the Lidah Formation, thus, appears to be confined within
an east-west elongated basin, bordered in the north by the Rembang Zone anticlines which
extend from the Blora area to Madura Island, and bordered in the south by the Plio-Pleistocene
volcanic centres (Fig. 5.29). Petrographic evidence suggests that the Rembang Zone was
exposed subaerially and was the source of sediments in the lower part of the formation, as
well as the Plio-Pleistocene volcanism in the south which became dominant in the uppermost
tidal channel facies.
The evidence presented here suggests that a large lacustrine basin was the site of
deposition of the Lidah Formation in the Plio-Pleistocene, which was characterised by organic-
rich and stagnant water (anoxic conditions as indicated by pyrite content). This stagnant water
condition was probably responsible for the minor facies variations within this formation. The
thickness, which attains 500 m in the Kendeng Zone during a relatively short time (Late
Pliocene-Earle Pleistocene), suggests that the subsidence and sedimentation rates were high.
The occurrence of siderite and gypsum minerals, and interbeds of marine-influenced deposits
in the Lidah Formation, further suggest that the lacustrine environment had been subjected to
drying and flooding events which are suspected to be related to eustatic sea level changes. The
Haq et al (1988) eustatic sea level curve indicated that there are at least four recognisable
fluctuations during the Late Pliocene-Pleistocene. A more detailed study in the Java Sea by
the author (Chapter 7) indicates that during just the Late Pliocene-Early Pleistocene there are
five sea level fluctuations (represented by seismic units JP1, JQla,b and JQ2a,b; Chapters 7.4.2
& 7.4.3).
The development of this stagnant lacustrine basin ceased when there was a significant
relative sea level rise, as indicated by greater marine influence on the upper mudstone facies
in the Nglampin section (Fig. 5.27), and the development of the Pucangan Formation in the
Kabuh and Mojokerto areas, which is characterised by intercalations between tidally influenced
81
and volcaniclastic deposits. The transition from the Lidah to Pucangan Formations in Madura
Strait is further discussed in Chapter Six.
5.7 PUCANGAN FORMATION
The Late Pliocene/Early Pleistocene Pucangan Formation (Chapters 3.2.2 & 3.5) of the
Kendeng Zone conformably overlies the Lidah Formation. Its sedimentation mostly relates to
the Quaternary volcanism along the south of this zone. In the eastern Kendeng Zone, where
the present study is concerned, it is related to the Wilis Volcanism (Duyfjes, 1938b; van
Bemmelen, 1949).
The first study of the Pucangan Formation was carried out by Duyfjes (1938b) and was
followed by de Genevraye and Samuel (1972). Duyfjes's (1938a,b,c,d) studies were very
detailed, and his comprehensive geological maps are still very useful for contemporary
researchers. However Duyfjes (1938b), as well as de Genevraye and Samuel (1972), did not
examine the sedimentological aspects of the Pucangan Formation. The present study re
evaluates the environments of the depositional units of the Pucangan Formation, particularly
in the tight of contemporary facies models (e.g. Reineck & Singh, 1980; Walker, 1984, 1992;
Reading, 1986). Poor outcrop is the main constraint to the study, and discussion presented in
this section is based on widely spaced vertical sections.
5.7.1 Lithofacies
Examination of the Pucangan Formation was carried out at seven locations, that were
considered to provide representative exposures of the general character of this formation: the
Solo River, Kabuh, Kemlagi, Perning, Banyuurip, Kedamean and Raci areas (Figs 5.1. & 5.30).
Because detailed mapping was not carried out during the study, lithostratigraphic correlation
between the sections was facilitated by a detailed geologic map produced by Duyfjes (1938b,d)
(Fig. 5.30).
The study reveals that the facies associations developed in the Pucangan Formation mostly
result from the introduction of volcaniclastics into shallow marine settings, as shown by
intercalations between marine-influenced and coarse volcaniclastic deposits. Based on textures,
sedimentary structures, faunal content and facies association, most marine-influenced
sedimentary units in the Pucangan Formation can be explained by using the tidal flat facies
82
model of Reineck and Singh (1980), Terwindt (1988) and Dalrymple (1992) (Fig. 5.31); and
the barrier island facies model of Reinson (1984) (Fig. 5.32).
Textures, sedimentary structures and faunal content within units in the Pucangan Formation
suggest that this formation can be divided into three major regressive sequences, where each
sequence commences with a marine-influenced deposit and terminates with a vertebrate-bearing
conglomerate (volcaniclastics). The marine-influenced deposits in all three sequences are
always associated with abundant marine molluscs (Volva, Erosaria, Erronea and Palmadusta;
Duyfjes, 1938b,d). Duyfjes (1938b,d) has used these three intervals of marine molluscs for
biostratigraphic correlation and named them Molluscan Horizons 1,2 and 3. The present study
found that conglomeratic facies (volcanic facies) also represent useful lithostratigraphic
markers. This latter facies occurs widely in the study area and is easily identified.
a. Lower sequence
Three measured vertical sections provide the basis of the discussion; these are located in
the Solo River (north of Ngawi, Fig. 5.22), Kabuh (Fig. 5.33) and Kemlagi (Fig. 5.34) areas.
Stratigraphically, this sequence directly overlies the Lidah Formation (in the eastern Kendeng
Zone) or the Sonde Formation (in the central Kendeng Zone). Relatively good exposures occur
in the Solo River section and on a road cut north of Kabuh village, but the sequence is poorly
exposed in the Kemlagi area. The sequence subdivision refers the Kabuh section in which six
units can be identified, referred to as units LI to L6.
a.l Muddy and matrix-supported conglomerate
Facies:
In the Solo River section this deposit attains 120 m thickness and comprises at least six
sequences dominated by massive, unsorted, matrix-supported conglomerate (Figs 5.22, 5.35 &
5.36). Fragments vary from pebble-sized to about 1.5 m and include igneous lithologies and
older pyroclastic fragments. Indications of inverse and normal grading are not obvious which
suggests that the sediment gravity flows were highly viscous. In the uppermost part,
conglomerate clasts are imbricated and planar cross-bedded, and the unit is overlain by
interbedded planar cross-bedded sandstone and non-fossiliferous mudstone.
he Kabuh section this facies occurs in two main horizons, denoted by L2 (about 32 m )
(about 20 m, Fig. 5.33). Texrurally these units are similar, characterised by matrix-
83
supported conglomerate which normally grades into coarse-grained sandstone. The lower
portion of unit L 2 displays inverse grading. The conglomerate fragments include pebble-sized
andesite and boulder-sized tuffaceous mudstone clasts. A microscopical examination on the
sandstone matrix of unit L 4 revealed the presence of foraminiferal tests, suggesting subaqueous
deposition.
Interpretation:
This facies consists mostly of debris flow deposits. This type of deposit is commonly
characterised by muddy-matrix supported clasts, is unstratified and poorly sorted and shows
no preferred orientation of clasts (Reineck & Singh, 1980; Walker, 1984). However, the
strength of the muddy matrix may force larger clasts to concentrate above the base of the bed
which results in a basal inverse grading or preferred clast alignment (Walker, 1984) as shown
in unit L2. Reineck and Singh (1980) indicated that in more fluid debris-flow deposits, graded
bedding is c o m m o n and flat pebbles are oriented more horizontally. Bull (1977) and Walker
(1984) suggested that debris flows are commonly associated with a steep depositional slope,
and enhanced by lack of vegetation and short periods abundant water supply (Bull, 1977). The
abundance of volcanic materials in this deposit suggests an association with volcanism.
Duyfjes (1938b) and van Bemmelen (1949) related these deposits to the Wilis Volcanism.
Sediments in the upper portion of the Pucangan Formation in the Solo River section,
display traction current structures associated with mudstone. These deposits can be interpreted
as the result of reworking by fluvial processes. Van Bemmelen (1949), who studied this
formation regionally, indicated that these deposits are widespread in the western part of the
Kendeng Zone and are dominated by black mudstone. H e reasoned that the deposits represent
a lacustrine environment, as the result of damming drainage systems by debris flow deposits.
A recent study by Stollhofen and Stanistreet (1994) on the interaction between volcanism and
fluvial sedimentation in the Permo-Carboniferous Saar-Nahe Basin (southwest Germany) found
a similar case. They demonstrated that sedimentary sections show a transformation from a pre
emptive meandering fluvial system to a lacustrine system following lava eruption.
a.2 Fossiliferous muddy sandstone
Facies:
The units included in this facies are units LI, L3a and L5 of the Kabuh section (Fig. 5.33)
and almost the whole lower sequence of the Kemlagi section (Fig. 5.34), except the
84
conglomeratic sandstone unit L6. This facies is essentially highly bioturbated muddy sandstone
and sandy mudstone, but parallel bedding is often observed in units in the Kemlagi section,
and cross-bedding occurs in unit LI of the Kabuh section. A notable feature of these units is
the c o m m o n occurrence of planktonic and benthonic foraminifers. Benthonic foraminiferal
analysis in unit L3a (Appendix C.2, sample no. JKB7) indicates the presence of an inner neritic
faunal association such as the foraminiferal genera Ammonia, Cibicides, Eponides,
Pseudorotalia and Quinqueloculina (Munay, 1973; 1991).
Interpretation:
The inferences of depositional environment based on the sedimentary structures are
difficult due to strong bioturbation. But faunal evidence suggests that this facies may have
been deposited in a shallow marine offshore facies. Unit LI, which is characterised by cross
bedding, probably was deposited in an environment closer to subtidal facies. This facies is
commonly characterised by sandbars (Reineck & Singh, 1980). For the Kemlagi section, the
interpretation is also supported by the fact that in this section, the conglomeratic units L2 and
L4, and tidal flat unit L3b of the Kabuh section are missing. This is interpreted as the
wedging out of these facies into deeper marine facies.
a.3 Cross- and ripple-bedded sandstone
Facies:
This facies includes units L3b and L5 of the Kabuh section, with a thickness of 40 m and
2 m respectively. Various types of large scale cross-beds bundled by m u d and shell fragments
are common in unit L3b, and bioturbation is common in unit L5. Cross-bedding cosets in unit
L3b vary in thickness from 20 c m to 1.5 m, and often feature plant remains and mudclasts.
Scouring and reactivation surfaces on cross-stratification are also observed. Interbeds of fine
to medium sandstone with parallel to ripple laminations, combined with m u d draping (mud
flasers), are often developed between the cross-bedded cosets with the interbed thickness
ranging from 5 to 75 c m (Figs 5.38 & 5.39).
Interpretation:
The depositional environments were defined particularly based on the sedimentary
stmctures. The determination based on the vertical sequence of facies is practically impossible
due to the occurrence of conglomeratic facies (debris flow) which replaced the normal trend
of vertical facies change.
85
The associated sedimentary stmctures suggest that units L3b and L5 were deposited in a
tidal flat environment. This environment occurs on a flat to gently sloping area and
geographically extends from the supratidal zone through the intertidal zone, into the upper
portion of the subtidal zone (Dalrymple, 1992). The facies distribution and sedimentary
stmctures in these environments largely reflect the strength of tidal currents generated during
flood and ebb relative to the geographic position. Tidal channels commonly occur on the outer
edge of a tidal flat environment. A relatively strong tidal current occurs in such channels and
produces a sandy character adjacent to the channel which gradually passes into mud near the
high tide line (Dalrymple, 1992). As a common practice, the tidal flat has been divided into
sand flat, mixed flat and m u d flat subenvironments (cf. Reineck & Singh, 1980; Dalrymple,
1992; Fig. 5.31).
Unit L3b is characterised by various types of large scale cross-bedding bundled by mud
and shell fragments. Terwindt (1988) and Dalrymple (1992) pointed out that bundle mudstone
occurs during slack water periods, and reactivation surfaces represent erosion by subordinate
currents of to the preceding stmctures (dunes) formed during flow of the dominant current.
The dominance of cross-bedded sandy units further suggests a subtidal and outer portion of
intertidal (sand flat) environment. The tidal current speeds during both flood and ebb tides are
high in these environments (Dalrymple, 1992). Unit L5 is bioturbated, and may represent a
mixed flat environment. Bioturbation is commonly weakest in the sand flat and becomes
stronger toward the mudflat environment (Reineck & Singh, 1980).
a.4 Normally graded conglomeratic unit
Facies:
This facies includes unit L6, and is observed in the Kabuh and Kemlagi sections (Figs 5.33
& 5.34). It is the most widespread unit and is an important stratigraphic marker. The facies
is about 50 m thick in the Kabuh section, thinning to about 25 m in the Kemlagi section, and
completely wedges out in the western part of the Guyangan Anticline (Fig. 5.30). In the
Perning section (Fig. 5.37), the facies is not entirely exposed but forms a broad plain of
conglomeratic sandstone in the core of the Kedungwam Anticline.
The main character of this facies is conglomeratic, but the overall facies is dominated by
medium- to coarse-grained sandstone. The conglomeratic facies, as observed in the Kabuh and
Kemlagi sections, commonly represents the base of the facies. It is composed mostly of
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massive to cmdely bedded cobble-size fragments of rounded andesite, supported by a matrix
of muddy coarse-grained sandstone. In the Kabuh section, the thickness of the conglomeratic
unit attains 3 m. The conglomeratic facies tends normally to grade upward into medium- to
coarse-grained sandstone. Here, the sandstone is commonly associated with planar and trough
cross-beds. But it also includes pebble-size mudclasts and marine molluscs in the foreset
larninae which indicate the presence of a marine reworking process.
Interpretation:
This facies probably represents an alluvial fan deposit as indicated by the dominance of
coarse elastics. However, the occurrence of marine molluscs associated with crossbedding
suggests a strong influence of tidal cunents. These currents might also be responsible for its
widespread occurrence.
b. Middle sequence
The middle sequence is a regressive sequence, about 50 m thick which lies on the
conglomeratic unit L6 at the top of the lower sequence. The sequence is very well exposed
in the Perning section (Fig. 5.30 & 5.37), but poorly exposed in the Kemlagi section (Fig.
5.34).
b.l Fossiliferous muddy sandstone
Facies:
This facies includes unit M l , and lies on the conglomeratic sandstone of unit L6 (lower
sequence). In the Perning section it is characterised by glauconitic sandstone with interbeds
of carbonaceous mudstone. Comminuted mollusc shells as well as benthonic forarninifers
(Operculina) are abundant, which is why Duyfjes (1938a,b,c,d) called this facies Molluscan
Horizon 2.
Interpretation:
This facies is interpreted to have been deposited in a subtidal environment. Operculina
commonly occurs in a slightly hypersaline inner shelf environment (Murray, 1991). The
common occurrence of comminuted shells in this unit suggests a strong wave influence.
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b.2 Interbedded mudstone and parallel, ripple-laminated sandstone
Facies:
This facies includes unit M 2 in the Perning section (Fig. 5.37) and the unit above unit L6
in the Kemlagi section (Fig. 5.34). Its exposure in the Perning section is relatively good, and
three main subfacies can be recognised: sandstone; mixed mudstone and sandstone; and
conglomeratic subfacies.
The sandstone subfacies is relatively thick (2 to 6 m). There are three prominent
sandstone beds. The lowermost sandstone is characterised by 2 m high epsilon cross-beds
with well developed m u d flasers between the foreset laminae (Fig. 5.40). The middle and
upper sandstone horizons are similar, characterised by clean, planar and trough cross-bedded
sandstone.
The mixed mudstone and sandstone subfacies is characterised by alternating parallel and
ripple laminated sandstone and mudstone beds (Fig. 5.41). In the lowermost 12 m, mudstone
is dominant with a few interbeds (5 to 10 c m thick) of parallel and ripple laminated fine
grained sandstone. This part features an abundance of plant remains and leaf fossils. In the
middle and upper parts, mudstone and sandstone finely alternate (Fig. 5.43). Some of the
sandstone beds reach 10 c m and are bundled by thin wavy and flaser bedded mudstone (Fig.
5.42).
The conglomeratic subfacies occurs as units less than 40 cm thick. Some of the beds in
the upper part (just below unit M 3 ) are normally graded and accompanied by basal scouring
(Fig. 5.42). In the middle part, some beds (about 30 c m thick) show inverse grading from
coarse sand to pebble size.
Interpretation:
The inferred depositional environment was based on sedimentary stmctures and facies
association. This facies is interpreted to have been deposited in a tidal flat environment. The
diagnostic criteria for tidal deposits have been discussed by Ginsburg (1975), Reineck and
Singh (1980), Terwindt (1988) and Dalrymple (1992).
The most diagnostic features in this facies are the occurrence of mud flasers, epsilon cross
bedding and plant remains. M u d flasers occur due to small scale alternations of slack and
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strong water current (Reineck and Wunderlich, 1968). Epsilon cross-bedding, which occurs
in the lower sandstone subfacies, is commonly associated with a lateral migration of tidal
channels, and the thickness of the crossbeds represents the channel height (Reineck & Singh,
1980); in this subfacies the channel height was about 2 m. The clean, planar and trough
cross-bedded sandstone subfacies are believed to result from winnowing sand which should
occur in a relatively strong tidal current region such as the sand flat and upper subtidal
environments (Dalrymple, 1992).
Temporary strong wave (storm) conditions probably occurred in this tidal flat environment,
and resulted in some conglomeratic subfacies. Basal scouring in these subfacies m a y represent
storm erosion (Johnson & Baldwin, 1986), and inverse grading may represent temporary
turbiditic currents which occurred during the storm.
b.3 Normally graded conglomerate
Facies:
This facies is well exposed in the Perning section, where it overlies the tidally-influenced
deposit (Fig. 5.37). In the Kemlagi and Banyuurip sections this facies is absent, suggesting
a limited lateral extent.
In the Perning section the facies is labelled M3, and consists of at least seven fining
upward units. Each unit, which is separated by sandy mudstone, begins with a crudely bedded
conglomerate fining upward into planar cross-bedded, medium- to coarse-grained sandstone.
The evidence for subaerial deposition in this facies is the presence of some terrestrial
vertebrate and hominid fossils [Pithecanthropus modjokertensis, Duyfjes (1938d), Fig. 5.45].
Interpretation:
Although the geomorphic evidence (associated fan-shaped geometry) is not obvious, this
facies is interpreted as an alluvial fan deposit as indicated by the occurrence of conglomerate
as the main facies (Rust & Koster, 1984) and vertebrate/hominid fossils. Most alluvial fans
are characterised by water-laid deposits, dominated by horizontally stratified gravel (Rust &
Koster, 1984). The occurrence of fining upward sequences in each unit, which shows a
transition from coarse gravel to sand, is an indication of a water-laid deposit which reflects
a gradual decrease in the particle/water depth ratio due to decrease of stream competence
(Reineck & Singh, 1980; Rust & Koster, 1984).
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The westward (Kemlagi section) and northeastward (Banyuurip section; Fig. 5.30) wedging
out of this facies, with regard to the geographic position of the Wilis Volcano, suggest that
this alluvial fan system was not directly related to the Wilis volcanism. It may have been
developed locally due to a local topographical anomaly.
c. Upper sequence
The upper sequence is observed in the Perning, Banyuurip, Kedamean and Raci sections.
The localities of these sections are shown in Figures 5.1 and 5.30, and the generalised sections
are shown in Figures 5.37, 5.44, 5.48 and 5.50. Figure 5.49 shows the typical exposure in the
Kedamean section. The thickness varies locally due to fluvial erosion in the uppermost
sequence; in the Banyuurip section the sequence attains 50 m.
cl Interbedded mudstone and parallel-ripple laminated sandstone
Facies:
This facies includes unit UI, and stratigraphically lies directly on the normally graded
conglomerate facies of the alluvial fan system (unit M 3 ) in the Perning section. The lower
part, as observed in the Banyuurip section, is characterised by intercalations of parallel-
laminated, tuffaceous sandy mudstone and parallel to ripple laminated, fine- to medium-grained
sandstone. The bed thickness varies from 2 to 10 cm. The facies is also characterised by the
abundant occurrence of plant remains, including leaf imprints.
The middle part of the facies is characterised by poorly-sorted tuffaceous sandy mudstone.
Bioturbation is strong which almost destroyed the parallel laminated bedding. Large
Gastrochaenolites burrow stmctures (3 to 5 c m diameter; Glossifungites ichnofacies) are
abundant (Fig. 5.46) in both the Perning and Banyuurip sections. This burrow shows the
protrusive appearance of a tear-like boring structure. In the Perning section this part is also
characterised by the abundance of marine mollusc fragments which Duyfjes (1938d) termed
the Molluscan Horizon 3. Because the latter is widespread, he used it as a stratigraphic
marker.
The upper part is characterised by finely laminated tuffaceous mudstone (Fig. 5.47), with
interbeds of parallel and ripple laminated, medium-grained sandstone up to 70 c m thick. This
subfacies is only observed in the Kedamean section.
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Interpretation:
The abundance of plant remains in the lower part of the facies suggests that the
environment was relatively close to or on a vegetated area. But textures and sedimentary
stmctures suggest that the facies may not represent a supratidal environment, as commonly
characterised by muddy facies (Terwindt, 1988). This lower part of the facies m a y represent
a foreshore or shoreface region, where wave energy is present. In the upper part of the facies,
the sedimentary texture and stmctures are similar to those in the lower part, and suggest a
similar depositional environment.
In the middle part, bioturbation is strong which suggests a deeper setting than the lower
and upper parts of the facies. Bioturbation has been regarded as a sensitive indicator of the
depositional energy, and commonly diminishes toward the shore as wave energy increases
(Reinson, 1984; Elliott, 1986). In addition, the Gastrochaenolites burrows, according to Frey
and Pemberton (1984) and Pemberton et al. (1992), are included in the Glossifungites
ichnofacies which commonly develops in firm unlithified substrates (dewatered muddy
substrates) of a littoral or sublittoral environment.
The overall facies (unit UI), thus, appears to represent a transgressive-regressive sequence,
as suggested by the transition of depositional environment from foreshore/shoreface to lower
shoreface and back to foreshore/shoreface.
c.2 Cross-bedded sandstone
Facies:
This facies is very well exposed in the Banyuurip, Kedamean (Figs 5.44, 5.48 & 5.49) and
Raci (Figs 5.50, 5.51 & 5.52) sections. There are two units included in this facies, units U2a
and U2c, which are separated by interbedded sandstone and mudstone unit U2b.
The thickness of the cross-bedded sandstone facies varies from 1.5 to 7 m. They consist
of coarse- and medium-grained sandstone, associated with erosional base and conglomeratic
lag deposits. Planar and trough crossbedding are the c o m m o n sedimentary stmctures. In the
Kedamean section palaeocurrent direction, as indicated by foreset orientation, is toward the
southwest.
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Interpretation:
This facies is interpreted as fluvial channel facies based on the occurrence of basal erosion
associated with channel lag deposits. Planar and trough cross-beds may represent linguoid or
transverse bars and dunes in the channel (Miall, 1992). The patchy nature of the outcrops,
unfortunately, does not allowed a further examination of the fluvial type.
This fluvial facies is well recognised on the seismic sections from Madura Strait and
assigned to seismic unit Cla (Chapter 6.5.5). Its widespread occurrence is interpreted to
correspond with a major sea level fall in the study area (particularly in the Kendeng Zone and
Madura Strait).
c.3 Interbedded mudstone, sandstone and conglomeratic sandstone
Facies:
This facies includes subunits U2b and U2d, which are characterised by interbedded
mudstone and sandstone; and conglomeratic sandstone. The interbedded mudstone and
sandstone is typified by parallel and wavy laminated sedimentary stmctures. Subunit U2b in
the Kedamean section, contains plant remains and leaf imprints, and subunit U2d in the
Banyuurip section is bioturbated and rich in marine mollusc fragments and benthonic
foraminifers (Cibicides & Elphidium). The conglomeratic sandstone is typified by massive
structure, and consists dominantly of medium- to coarse-grained sandstone with abundant
andesitic pebbles and mudclasts. The conglomeratic sandstone occurs in the Raci and
Kedamean sections where the thickness varies from 0.5 to 1.5 m.
Interpretation:
The interbedded mudstone and sandstone facies is interpreted as a tidally-influenced
deposit, based on the sedimentary textures and stmctures. The occurrence of benthonic
foraminifers and bioturbation in subunit U2d suggests a lagoonal environment. The occurrence
of plant remains in subunit U2b suggest a close relationship with a supratidal or estuarine
environment. The conglomeratic sandstone in subunit U2b is massive and is composed of
abundant intraclasts, suggesting a very quick depositional mode which may have occurred
during a storm event.
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5.7.2 Petrography
Fifteen samples from sandy units were examined, and most of these samples are poorly
cemented and show varying degrees of weathering. The study found that there is no c o m m o n
classification of rocks in the Pucangan Formation due to the large variety of lithologies.
However, the study has documented the occurrence of terrigenous clastic deposits of two
different compositions which are strongly suspected to represent two different sources. The
first clastic source is dominated by quartz, and the other contains volcanic-derived fragments
such as andesite, plagioclase, pyroxene and hornblende. The occurrence of quartz is always
affiliated with the marine facies at the beginning of sequences, while the second clastic is more
widespread. Appendix B.l summarises the results of the petrographic examination.
a. Lower sequence
The tidal sandstone unit (unit L3b; Fig. 5.53) is characterised by angular to subangular
grains of plagioclase (21%), pyroxene (17%) and andesitic rock fragments (17%); the grain
size varies from 0.4 to 0.75 m m . The sandstone is very porous, uncemented, with practically
no muddy matrix, indicating a high energy depositional current.
The sandstone of the debris flow unit is characterised by angular to subangular grains of
plagioclase (up to 6 0 % ) , rock fragments (up to 22%) and pyroxene (up to 1 6 % ) ; the grain size
varies from 0.3 to 1 m m . This sandstone is also characterised by the presence of foraminiferal
tests (less than 2%) with no indication of reworking. This indicates that the debris flow unit
was deposited in a marine environment.
Five petrographic samples have been studied from the subtidal facies in the lower sequence
(from the Kemlagi section). The lower part of this offshore facies is characterised by muddy
sandstone which is almost free of volcanogenic fragments (< 1%). The main components are
angular to subangular quartz grains (up to 34%), and bioclasts of foraminiferal tests, algae and
mollusc fragments with the average grain size 0.2 m m . All these fragments are cemented by
carbonate m u d which makes the total carbonate component up to 5 9 % of the rock. A similar
texture occurs in the middle sequence (sample no. K G 7 ) , but around 3 0 % of the terrigenous
clastic material is volcanogenic fragments (plagioclase and andesitic rock fragments). This part
also contains streaks of organic materials. The uppermost part, which can be included in unit
L6 (sample no K G 1 ) , consists of subangular to subrounded andesitic rock fragments (17%),
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angular plagioclase (29%) and pyroxene (< 4 % ) , with the grain size varying from 0.2 to 1.5
m m (Fig. 5.54).
The upward decreasing bioclastic content in the samples from the Kemlagi section suggests
that the lower sequence is regressive. A fully marine environment may have occurred in the
lower part which gradually became shallower upward where the deposition became influenced
by volcanic rock fragments and plant debris. The quartz in the lower part has never been
associated with Quaternary andesitic volcanism. The reason for the quartz abundance is
probably the same as for the sandstone facies in the Lidah Formation in the Sumberringin
section (some 3 k m east of Kabuh). The quartz was derived from the exposed quartz-bearing
formations in the Rembang Zone and transported southward to the Kendeng Zone through
marine processes. The same case occurs in the marine fossiliferous sandstone facies in the
middle (unit M l ) and upper (unit UI) sequences.
b. Middle and upper sequences
The middle and upper sequences are represented by seven petrographic samples. Units
M l and U I are quite similar, and are characterised mainly by carbonate mud matrix. As
indicated earlier, these marine derived facies contain angular to subangular fine-grained quartz
which is interpreted to represent a northerly derived terrigenous clastic component (Fig. 5.55).
In unit M l , quartz is a minor constituent, but in unit U I it is up to 7%. It is associated with
volcanogenic fragments such as andesite (some are glauconitised) and plagioclase (about 14%),
and bioclastic foraminifers (Operculina) and marine molluscs (Lamellibranch).
Three sandy samples from the channel facies (unit U2c) contain relatively little mud
matrix; most of the fragments are lithic (up to 56%). These fragments (andesite) are rounded,
highly altered, with an average grain size of 0.4 m m . Plagioclase and pyroxene attain 1 7 %
and 2 % , respectively, and commonly are less altered and show angular form.
Unit U2b is characterised by the presence of mudstone beds which consist mostly of
carbonate mud. Fine-grained (less than 0.15 m m ) plagioclase and quartz are up to 1 2 % and
7 % of the rock respectively. It also contains less than 1 % bioclasts. The sandstone beds in
this deposit are moderately sorted, consisting of altered volcanic rock fragments (up to 5 6 % ) ,
angular plagioclase (up to 17%) and pyroxene (< 1%).
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5.7.3 Discussion
The Pucangan Formation in the eastern Kendeng Zone developed on the Lidah Formation
which was, in rum, developed as an extensive lacustrine deposit. Thus the Pucangan
Formation was deposited on a relatively flat-lying surface predefined during the deposition of
the Lidah Formation. The lithofacies study indicates that the flat depositional slope persisted
during the deposition of the Pucangan Formation in the eastern Kendeng Zone, as evidenced
by the c o m m o n occurrence of tidal flat deposits in this formation. There are three major
conglomeratic and sandy facies (units L6, M 3 and U2) in the Pucangan Formation which have
been extensively developed across this relatively flat-lying area. They are believed to be the
product of the Wilis Volcano through an alluviation process. Thus, the gross depositional
setting of this formation can be regarded as a fan delta, similar to the ones described by Nemec
and Steel (1988).
The first depositional model for the Pucangan Formation was proposed by Duyfjes
(1938b). This model regarded the vertical development of the old Wilis Volcano in the
surrounding marine facies. He described the volcanic facies laterally wedging out and
changing to the muddy Lidah Formation (Fig. 5.56). The lateral development of the Pucangan
Formation was assumed to correspond with the vertical development of the Wilis Volcano (Fig.
5.57). This model appears to have satisfied many authors working in the eastern Kendeng
Zone (e.g. van Bemmelen, 1949; de Genevraye & Samuel, 1972; Pringgoprawiro, 1983) as
there has been no attempt at further examination by the later investigators.
The development of coarse clastic deposits associated with alluviation in the Pucangan
Formation appears to be intermittent. The subdivision into three major sequences in fact
represents this. Marine processes occurred during the time intervals in between the deposition
of these coarse clastic sequences.
The lower sequence is characterised by debris flow deposits. They may characterise a
steep depositional slope and rapid deposition (Walker, 1975). Observations of the debris flow
deposits in the Solo River section reveals a substantial amount of fine debris, which is
apparently strong enough to support clasts of more than boulder size. The occurrence of these
deposits is probably closely associated with the Wilis volcanism.
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The conglomeratic facies in unit L6 (lower sequence) and unit M 3 (middle sequence) lack
a debris flow character. A close association with vertebrate fossils, particularly in unit M 3 ,
is rather suggestive of a period of volcaniclastic alluviation. Figure 5.58 displays the
depositional model for unit M 3 .
5.8 KABUH FORMATION
5.8.1 Lithofacies
The Kabuh Formation formed during the Middle Pleistocene (Chapter 3.3.2 and 3.5). A
sedimentological discussion of this formation can only be carried out on its lowermost part
where it is well exposed in the Banyuurip, Kedamean and Raci areas (Figs 5.30 and Fig. 5.1).
The vertical sections are shown in Figures 5.44, 5.48 and 5.50. The upper part of the Kabuh
Formation consists of volcanic and fluvial facies, and presently lies in the synclinal area. In
the easternmost Kendeng Zone it is rarely exposed and is often covered by thick recent
alluvium.
In the Kedamean and Banyuurip sections, the lower part of the Kabuh Formation
developed as a regressive barrier island setting (Fig. 5.32). Detailed observations in these
sections (Figs 5.44; 5.48 & 5.49) identified four different facies which are commonly present
in this setting: a shelf mud, lower shoreface, upper shoreface and foreshore facies. In addition,
a non-marine facies overlies the foreshore facies. As shown in Figures 5.48 and 5.49, these
marine facies are relatively thin (about 20 m ) . Duyfjes (1938b,d) reported that these facies
gradually wedge out westward. In the synclinal area north of Perning (Fig. 5.30) the marine
facies have been replaced by a non-marine facies characterised by the presence of non-marine
molluscs and vertebrate fossils.
a. Green mudstone
This facies is interpreted to represent a shelf m u d environment. It consists of massive,
greenish grey, sandy mudstone, and was deposited on the fossiliferous, molluscan-rich tidal
facies unit U2d (Pucangan Formation; Fig. 5.49). The present faunal study of rock samples
from this facies in these sections indicates that benthonic and planktonic foraminifers are rarely
present. Duyfjes (1938d) reported the occurrence of some forarniniferal species (e.g. Ammonia
and Globigerina) in this facies from the Raci area.
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The change from fluvial facies (unit U2c), passing through tidal facies (unit U2d) to the
shelf mud facies appears to represent a very rapid transgression which only deposited thin
units. The shelf m u d facies was widely developed, probably due to a very low depositional
gradient whereby a few metres increase of sea level will lead to the flooding of an extensive
area. A similar facies can be recognised in the Raci area (Fig. 5.50). Here, the facies is
unconformably overlain by the fluvial channel deposit forming the upper part of the Kabuh
Formation.
b. Bioturbated muddy sandstone and parallel laminated mudstone
The first facies represents a lower shoreface environment; it consists of poorly sorted,
muddy sandstone. In the Kedamean section this unit was channelised in the upper part (Figs
5.48 & 5.49) and filled with muddy sandstone which is rich in marine molluscs and benthonic
foraminifers. Genus Ammonia is the most common fauna and typifies an inner neritic
environment (Murray, 1973; 1991). Sedimentary stmctures are absent, possibly due to
intensive bioturbation. Reinson (1984) suggested that the lower shoreface region occurs
seaward of the break in the shoreface slope, at the toe of the barrier-island sedimentary prism
(Fig. 5.32). The normal condition of the lower shoreface region is relatively low energy.
Upward, the facies changes to parallel-laminated mudstone with thin interbeds (< 10 cm)
of ripple laminated fine-grained sandstone. Bioturbation is much less than in the lower
shoreface deposit. This part is interpreted to represent a low energy upper shoreface
environment.
d. Cross-bedded sandstone
Facies:
This facies may represent foreshore and non-marine environments. The lower portion
represents the foreshore region, and is characterised by a series of coarsening and thickening
upward sandstone units (Figs 5.48 & 5.49), separated by thin (< 5 cm) mudstone interbeds.
The sandstone units contain prominent low-angle, large scale (about 1 m ) planar cross-beds
but no basal erosion is associated with these stmctures. Reinson (1984) indicated that the
swash and backs wash mechanism is often related to the formation of distinct subparallel to low
angle, seaward-dipping, planar sedimentary stmcture in the foreshore region.
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The non-marine facies was defined particularly on the basis of the vertical sequence trends
and a mapping report by Duyfjes (1938d). Some outcrops in the Banyuurip area indicate the
presence of conglomeratic sandstone and planar cross-bedded sandstone with mudstone
interbeds.
5.8.2 Discussion
The Kabuh Formation is known as the last lithostratigraphic unit in East Java which was
deposited under the influence of marine processes (Duyfjes, 1938b,d). Deposition commenced
with a widespread development of shelf mud [green mudstone according to Duyfjes (1938d)]
which, from its textural character, appears to have been deposited as a deeper facies than the
marine-influenced deposits in the Pucangan Formation. This green mudstone facies is
recognisable on the seismic sections in Madura Strait (seismic Section 6.1) and is referred to
as the seismic unit C2a. It marked the beginning of a series of deposits (seismic unit C 2 to
C8) which are interpreted to represent frequent sea level cyclicity and rather large sea level
rises and falls. A magnetostratigraphic study by Hyodo et al. (1992) in the Perning area (Fig.
5.37, Perning section) indicated that this green mudstone is younger than the Jaramillo event.
This led to the speculation that the marine facies of the Kabuh Formation should represent the
beginning of periodic and rather extreme sea level fluctuations over the last 900 ka which are
very well recorded by the oxygen isotope record in the Pacific core V28-239 (Shackleton &
Opdyke, 1976). Unfortunately, many of the cyclic marine facies in the Kabuh Formation in
East Java are missing due to widespread erosion during the combined Late Pleistocene folding
and sea level fall during the oxygen isotope stage 6. However, some of these events are
recognised on the seismic sections from Madura Strait.
5.9 SUMMARY
The sedimentation history of the Kendeng Zone in the study area is complex. Here, the
oldest outcropping lithofacies is the Middle-Late Miocene Kerek Formation. This formation
consists of a thick (> 800 m ) turbiditic deposit which was derived from the Banyak volcanic
arc. A marl facies, the Kalibeng Formation, dominated deposition in the Early Pliocene. A
short regression occurred in the Kendeng Zone in the Middle Pliocene but deltaic deposits
(sandstone facies of the Atasangin Formation) only occurred in the eastern part of the zone.
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Since the Middle/Late Pliocene two different tectonic characters have occurred. The
central Kendeng Zone (including the Ngawi area) has undergone tectonic uplift, as indicated
by regressive vertical facies sequences in this area which ended with the deposition of the
wave-dominated unit of the Late Pliocene Sonde Formation. In contrast, in the eastern
Kendeng Zone, including Madura Strait, very rapid basin subsidence occurred, combined with
a high sedimentation rate. This resulted in the deposition of thick diatomaceous marl (marl
facies of the Atasangin Formation) and lacustrine deposits (Lidah Formation). In the Plio-
Pleistocene, activity of the Wilis Volcano commenced (Duyfjes, 1938b,d; van Bemmelen,
1949). Its clastic detritus was deposited as the marine Pucangan Formation followed by the
Kabuh Formation where only the easternmost part of the zone was in a marine environment.
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CHAPTER SIX
QUATERNARY SEISMIC STRATIGRAPHY OF MADURA STRAIT
6.1 INTRODUCTION
Madura Strait is the eastward extension of the Kendeng and Randublatung Zones. It is
a small east-west elongated basin, 130 x 65 k m in size. The basin is synclinal, bordered in
the north by the anticlinal zone of the Rembang Zone (Madura Island), and by Quaternary
volcanoes in the south (see Fig. 2.8). Discussion given in Chapter 5 indicated that the eastern
Kendeng Zone, including Madura Strait, has undergone very rapid subsidence since the
Pliocene, which m a y be due to the northward compressional stress within the larger plate
tectonic system in western Indonesia. Thick Quaternary deposits are indicated by some well
data, and seismic records provide evidence of very high sedimentation rates.
This chapter discusses the distribution and chronology of the sedimentary facies developed
in Madura Strait during the Quaternary, and briefly reviews its underlying Late Tertiary strata.
The discussion relies mostiy on the results of seismic interpretation, combined with lithofacies
analysis, including petrographic and faunal analyses, of the Late Tertiary strata. Furthermore,
this chapter can only provide a preliminary Quaternary sedimentological study, because of the
absence of reliable well data and dating. The seismic stratigraphy of Madura Strait is
interpreted by applying the sequence stratigraphic concepts developed by Vail et al (1977),
Posamentier et al. (1988) and Posamentier and Vail (1988). Sequence definition and
terminologies used in this concept have been discussed in Chapter 4.2. Interpretation of clastic
depositional facies from seismic facies is carried out based on the criteria outlined by Sangree
and Widmier (1977, Table 4.1)
The sequence stratigraphic concept was originally applied to low resolution multifold
seismic data associated with petroleum exploration and involved third order sequences or lower
(Vail et al, 1977). The present study concerns Quaternary strata, for which the sequences are
of higher order (fourth and fifth order). The thickness of such sequences is often beyond the
resolution of multifold seismic systems. However, the applicability of this concept to such
sequences is endorsed by Posamentier et al. (1988), Mitchum and van Wagoner (1991), and
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Posamentier and James (1993) who indicated that the concept of sequence stratigraphy is scale
independent.
In the last few years, a number of authors have applied sequence stratigraphic analysis to
fourth and fifth order sequences [0.1-0.2 M a and 0.01-0.02 M a respectively, Mitchum & van
Wagoner (1991)] using high resolution seismic data. O n continental shelves, successful
applications have been demonstrated by Suter et al. (1987) on the southwest Louisiana
continental shelf; Tesson et al. (1993) on the outer Rhone continental shelf (France); and
Hernandez-Molina et al. (1994) on the Spanish continental shelf. Okamura and Blum (1993)
successfully applied sequence stratigraphic analysis to high resolution seismic data from the
fore-arc region, southwest of Japan. The present study adopts the same principles as these
studies, using high resolution seismic data (from a sparker system) which have about 2 m
vertical resolution. Although this study is not intended to review the applicability of the
sequence stratigraphic concept in general, its use in Madura Strait m a y be regarded as a test
of its application to a small basin which has high subsidence and depositional rates.
Furthermore, because this basin is relatively small in lateral dimension, the trans-basinal
depositional style is able to be studied.
Five major sedimentary units can be identified on the seismic sections studied. Each of
these units is constmcted of two or more subunits which have distinct areal distributions,
thicknesses and seismic characters. The upper four major units are bounded at their base by
regional unconformities. They can be regarded as fourth order type I sequences, according to
the criteria outlined by Vail et al (1977) and Posamentier and Vail (1988), due to their
duration of less than 1 M a . The basal unconformities are associated with widespread and
relatively thin fluvial deposits which possibly represent major sea level falls during the Early
to Middle Pleistocene. The five major seismic units are referred to (from bottom to top) as
units Pre-A, A, B, C, and D. Numerical and lower case characters affixed on these letters
indicate the subunits. The seismic unit Pre-A lacks internal reflections, which does not allow
further subdivision of this unit although it is relatively thick. This unit is underlain by
Pliocene strata which in this discussion are referred to as the M u n d u Formation.
The determination of Pleistocene eustatic sea level change in Madura Strait will not be
attempted, because the basin subsidence rate is not known. In addition, recognition of coastal
onlaps in many cases is difficult due to the very low depositional dip, and the formation of the
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Late Pleistocene fold system which has disturbed the position of the sequences. The
Pleistocene sea level can only be deduced tentatively based on the position of incisions and
coastal onlaps relative to present sea level.
6.2 SEISMIC DATA
The data base for this study is drawn from seismic profiles totalling some 1500 line km.
The seismic profiling was carried out by the Marine Geological Institute of Indonesia in 1989
and 1990, using a 600 joule high frequency single channel sparker system. Figure 6.1 shows
the ship tracks during profiling in Madura Strait. Spacing of profiles was relatively regular
and the north-south profiles were run perpendicular to the general east-west structural trend.
During profiling the sparker output power was released through a three tip array. Low/high
filter settings for most of the sparker survey were 200-2000 Hz. The data were graphically
recorded in analog format. Firing and recording sweep rates were 0.5 seconds, and the vertical
recording scale was 0.5 sec, resulting in an average vertical exaggeration of 8 times on the
records. Ship positioning during the seismic profiling was accomplished with the Magnavox
M X 1157 satellite navigation system with point accuracy of 36 m.
Although the seismic system provides high resolution records, it has a maximum depth
penetration of about 350 m. In some areas where the seismic wave attenuation is high, the
penetration of the seismic system is significantly reduced. The penetration is also largely
inhibited where medium to coarse-grained sediments are interpreted to be present. With such
limitations and because the thickness of Quaternary sediments in East Java could reach 500
m, the discussion will mostly concentrate on Middle and Late Quaternary sediments in Madura
Strait.
Figure 6.2 shows the location of seismic sections used for the discussion, but isochron
(unit/sequence thickness in time units) and time structure maps presented in this discussion
were derived from all the seismic lines indicated in Figure 6.1. Due to the absence of
measured seismic velocity data, during interpretation the seismic velocity is assumed to be
1500 m/sec which is really only valid for the water column. O n the seismic profiles presented
in the discussion, the depth was calculated based on this velocity and should only be regarded
as a rough estimate.
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Analysis of seismic data used for this study includes four main steps: seismic
interpretation, seismic horizon digitising, gridding and contouring. In the seismic
interpretation, seismic horizons are recognised based on the criteria outlined by Vail et al
(1977) which was discussed in Chapter 4. However, because these criteria were derived for
use in a petroleum exploration setting, mostly on shelf margins where depositional dips are
usually obvious, a different approach will be attempted for this study. Sedimentary sequences
in Madura Strait are commonly very thin and widespread. Due to rapid sea level oscillations
and due to the very low depositional dip, recognition of seismic features such as coastal onlap
is very difficult. Here, sequence boundaries have been identified using features such as fluvial
erosion and anticline truncation.
The interpreted seismic lines were digitised by using a digitising tablet. The results from
this step were the coordinates for the picked seismic horizons, depth (z value) and horizontal
position along a line. The last value was used to obtain the absolute positions (x and y
coordinates) of picked seismic horizons by comparing it with the satellite fixed positions. The
discrete data obtained from this step are irregularly spaced along the seismic lines. Prior to
contouring, these data have to be gridded with a predefined spacing. In this study, the inverse
distance gridding method was applied, to obtain faster and reliable results. The digitising and
gridding steps were carried out on an I B M PC, using software (see Appendix F for source
codes) specially written by the author in B A S I C and C + + computer languages. The gridded
data obtained from this step were translated into the "Surfer" computer program data format
and contouring was done using this computer program. The "Surfer" computer program is in
fact able to do the gridding step; however, dealing with such a large amount of data, this
program would consume a lot of time.
6.3 WELL DATA
To date, there have been no published studies of the Tertiary and Quaternary geology of
Madura Strait. The results from two petroleum exploration wells drilled in this area, MS1-1
and M D B - 1 (Fig. 6.1), indicated a low economic potential. The current study is probably the
first attempt to discuss the Quaternary geology of this area.
Faunal analyses have been carried out on cuttings and sidewall core samples from these
petroleum wells by the Indonesian Cities Service Inc. Company. Due to the aims of the
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drilling, which were targeted on the Tertiary sequences, these wells do not provide good data
control for the present seismic study. Three disadvantages of these wells for this Quaternary
study are that: (1) they were drilled on or close to an anticlinal axis on which Quaternary
sediments are very thin due to non-deposition or Quaternary erosion; (2) faunal and lithofacies
analyses were carried out mostly on cuttings sampled at long intervals (every 30 ft or 9.1 m
penetration); and (3) no geophysical logs are available for study. With such limitations, these
wells only provide a rough estimation on the environment of deposition and age of
stratigraphic boundaries. Detailed lithofacies correlation with the seismic data is practically
impossible, and most of the lithofacies interpretation has to rely on the onshore geological data
such as from the Kendeng Zone and Madura Island.
Well MS1-1
This well was drilled on the southern flank of a Late Pleistocene anticline the crest of
which forms Kambing Island (Fig. 6.1). The projected well position is shown on seismic line
M D - M D ' (Section 6.4). Accurate projection and stratigraphic correlation are difficult to
achieve, due to the position of the well on the ramp side of an anticline. This well was drilled
in a position which has a water depth of 27.4 m and a 24° S E average dip for the strata. The
description below summarises the well report and concerns only the uppermost 700 m of
sediment in the well (measured from the rotary stage, which was 22 m above sea level).
Figure 6.3 summarises the planktonic foraminiferal analysis.
Between 49.4 and 76 m depth, the lithology is characterised by fossiliferous muddy
sandstone, with abundant benthonic foraminifers and gastropods (up to 6 5 % ) . A fairly massive
calcareous mudstone is present between 76 and 277 m depth. The Pliocene-Pleistocene
boundary (planktonic forarniniferal zones N21/N22 boundary) has been picked at the base of
this unit. It was defined by extinction of the planktonic foraminifer species Globigerinoides
obliquus and Globorotalia acostaensis. However, it should be noted that some species such
as Globigerinoides ruber, Globigerinoides quadrilobatus trilobus, Pulleniatina obliquiloculata,
Globorotalia cultrata, Orbulina universa, Hastigerina aequilateralis and Sphaeroidinella
dehiscens, which should remain until N23 (Late Pleistocene), have also disappeared. This
indicates that changes in the environment of deposition may have contributed to the extinction
of these species at this level in this area.
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The occurrence of Operculina, Amphistegina and Asterorotalia in this calcareous mudstone
suggests a fairly shallow water depth (Phleger, 1960; Murray, 1991), and a possible tidal flat
environment as Operculina commonly occurs in a slightly hypersaline region (Murray, 1991).
Stratigraphically, the calcareous mudstone is equivalent to the Pucangan Formation in the north
of Mojokerto area, which is characterised by interbeds of shallow shelf, tidal flat and non
marine deposits.
The calcareous mudstone is underlain by a sandy mudstone to a depth of 451 m. This
sandy mudstone is typified by the appearance of lignite, plant remains, quartz, pyrite and
glauconite. Planktonic and benthonic foraminifers occur sporadically. Between 451-585 m
depth, the lithology consists mostly of massive mudstone with trace amounts of glauconite,
limestone and plant remains. The consistent appearance of Ammonia beccarii indicates that
the environment of deposition was lagoonal. This mudstone is stratigraphically equivalent to
the Lidah Formation lacustrine deposit in the area north of Mojokerto which is mainly
characterised by an organic-rich, dark blue mudstone (Chapter 5.6).
Below 585 m depth, the mudstone is calcareous and fossiliferous. The upward decrease
in abundance of planktonic foraminifers in this Pliocene mudstone suggests that the
environment of deposition was gradually shoaling from open marine-influenced (below 585 m
depth) to lagoonal environments. Stratigraphically, this fossiliferous mudstone should correlate
with the M u n d u Formation marly facies (equivalent with the Kalibeng and Atasangin
Formations).
Well MDB-1
Well M D B - 1 is located 12 k m east of the easternmost seismic line. In the uppermost 488
m, the lithology is dominated by a non-calcareous mudstone, with minor interbeds of silt and
loose glauconitic sand. Below this, the facies becomes gradually more fossiliferous with depth
until 690 m, where a clastic limestone unit underlies. The Plio-Pleistocene boundary was
defined within the fossiliferous mudstone, at 673 m depth, based on analyses of planktonic
foraminifers and nanoplankton. The age of the limestone, according to the well report, is
Early to Middle Pliocene. Unfortunately, the biostratigraphic data are not available for this
study.
105
The fossiliferous mudstone in this well is correlated with the similar facies and age in the
MS1-1 well, below 585 m depth. The non-calcareous mudstone in the M D B - 1 well probably
correlates with the mudstone facies above 585 m depth in the M S 1-1 well. Correlation
between the position of the Pliocene-Pleistocene boundary and facies in both the M S 1-1 and
M D B - 1 wells suggests that the lacustrine facies was developed earlier in the M S 1-1 well
(western part).
6.4 PHYSIOGRAPHY OF MADURA STRAIT
The bathymetry of Madura Strait is shown in Figure 6.4. In general, the sea floor depth
increases towards the middle of the strait which has the form of a broad submarine valley.
Some local highs in the strait correspond to buried small anticlinal or diapiric stmctures (Fig.
6.1). The deepest portion (more than 80 m ) occurs in the easternmost part of Madura Strait.
The average east-west sea floor morphological gradient is very low (0.5 m/km) and, in the
north-south direction, it is relatively similar (2.2 m/km) on both northern and southern sides
of Madura Strait.
In Madura Strait the orientation of geological stmctures is similar to the onshore Kendeng
Zone. Seismic data indicate that anticlines, which are also developed in the onshore part,
become progressively simpler eastward and are individually separated from one another. In this
strait, the crests of anticlines are represented by two small islands (Kambing and Ketapang
Islands, Fig. 6.1). Figure 6.29a shows a three dimensional view of the sea floor morphology.
It is suspected that m u d diapirism, as found in several places in the onshore area (north of
Mojokerto area), has played an important role in the structural development in the Kendeng
Zone and Madura Strait. In this strait, it is often accompanied by normal faulting.
6.5 SEISMIC FACIES ANALYSIS AND INTERPRETATION
The discussion presented in this section refers to some representative seismic lines, and
their geographic locations are provided in Figure 6.2. However, all the surveyed lines were
used to construct the isochron, seismic facies and palaeogeographic maps. The representative
seismic sections are numbered as Sections 6.1 to 6.12.
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6.5.1 Late Tertiary units
a. Seismic facies
Because of the limited penetration of the seismic system used, the Late Tertiary units can
only be studied from the seismic sections recorded close to some anticlinal islands. The
seismic line M L - M L ' (Sect. 6.12) is representative of these lines. It was recorded above an
anticline which forms the Gili Genting and Gili Raja Islands (Figs 6.1 & 6.2). This line
displays two contrasting geological features. O n the southern synclinal side, the section is
mostly occupied by late Late Pliocene (unit Pre-A) and Quaternary (units A to D ) strata. A
continuous and rapid subsidence, combined with rapid sedimentation, has allowed an
accumulation of thick sequences that may attain more than 500 m in the synclinal axis. On
this seismic section these sequences are not well displayed due to multiple reflection
interference, but detailed discussion, based on other seismic profiles, will be presented in the
subsequent sections.
On the anticlinal portion, three distinct Late Tertiary seismic sequences can be recognised
and are referred to as units LT1, LT2 and LT3. The description below concerns these units.
The lithostratigraphic correlations are mostly tentative, based on some geological mapping
reports from the Madura, Gili Raja and Gili Genting Islands, and some direct observations
during this study on these islands.
The sedimentary facies interpretation is curtailed by the absence of exposure and well data
near seismic line M L - M L ' (Sect. 6.12), but a tentative correlation is possible through a
stratigraphic correlation with the inland geological outcrops. O n Madura Island geological
mapping was carried out by Duyfjes (1938c), Situmorang et al (1992) and Aziz et al (1993).
There is c o m m o n agreement among these authors that four major Early/Middle Miocene to
Pliocene stratigraphic units are exposed on this island, along the flanks of some east-west
trending anticlines. These are the Early/Middle Miocene paralic and clastic limestone deposits
equivalent to the Tawun and Bulu Formations respectively, and Late Miocene/Pliocene marly
facies, and reefal and clastic limestone of the locally termed Pasean and Madura Formations
respectively. The marly facies, Pasean Formation, is stratigraphically equivalent to the Mundu
and Kalibeng/Atasangin Formations in the Rembang and Kendeng Zones (of East Java)
respectively.
107
a.l Units LT1 and LT2
The lowermost unit, LT1, is topped by an angular unconformity. Seismically, this unit is
characterised by a relatively strong amplitude and a low continuity of reflectors, but some
strong and continuous reflections occur on the northern side. Unit L T 2 thickens northward,
and is typified by medium amplitude and low continuity reflectors. Some internal baselapping
patterns occur and are characterised by relatively thin, strong and continuous parallel reflection
patterns.
Units LT1 and LT2 are interpreted to correspond with the Miocene paralic deposit and
limestone respectively. This paralic deposit is dominated by organic-rich mudstone with
interbeds of quartz sandstone and clastic limestone. The occurrence of benthonic foraminifers
Ammonia sp., Brizalina sp., Triloculina sp. Quinqueloculina sp., Eponides sp. (Situmorang et
al, 1992) suggests an inner neritic environment of deposition (Phleger, 1960; Murray, 1991).
The limestone, according to Situmorang et al. (1992), is characterised by the occurrence of
larger foraminifers Lepidocyclina sp., Cycloclipeus sp., Operculina sp., which are commonly
associated with reefal deposits and calcareous algae (Moore et al, 1952; Serra-Kiel & Reguant,
1984).
The MS 1-1 well report indicated that a similar paralic facies and age as the Tawun
Formation occurred at 2627 m depth, at which the drilling was terminated. This indicates that
intensive fold development had not commenced in the Miocene, which implies the widespread
occurrence of paralic environments in the study area.
a.2 UnitLT3
Unit LT3 is characterised by a low amplitude subparallel reflection pattern with some
strong and continuous parallel reflections. The true thickness of this unit cannot be determined
on seismic line M L - M L ' (Section 6.12) due to the extensive Late Pleistocene erosion, but it
may reach more than 300 m. The uppermost part tends to show a stronger amplitude
reflection which is interpreted to represent remnants of reefal and elastic limestones. These
limestones are widely exposed along the south coast of Madura Island and in the core of Gili
Raja and Gili Genting Islands. These have been referred to as the Madura Formation by
Situmorang et al. (1992), and are equivalent to the Paciran Formation. O n these islands, the
limestone facies is underlain by a marly mudstone which is interpreted to represent the lower
amplitude subparallel reflection pattern of unit LT3. Situmorang et al (1992) assigned the
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marly facies beneath the Madura Formation limestone to the Pasean Formation which is a
lateral equivalent of the Mundu Formation. This marly facies wedges out toward the axial line
of Madura Island.
In the Kamal area, the limestone facies appears to have been deposited under the influence
of tidal currents, which resulted in southward-dipping very large cross-beds of sand waves
(Dalrymple, 1992; Fig. 6.5), and further suggests a fairly shallow shelf environment of
deposition. Duyfjes (1938c) described a limestone succession developed on the western end
of Madura Island and assigned it to the limestone facies of the Upper Kalibeng Beds.
Stratigraphically, this limestone is equivalent to the shallow water reefal facies of the Mundu
Formation. Such limestone facies is absent in MS1-1 well, suggesting that shallow marine
facies only developed along Madura Island (?southern part) and deeper marine facies occurred
along Madura Strait.
In some areas on Madura Island (east of Kamal and south of Pamekasan) and Gili Raja
Island, the limestone facies is unconformably overlain by a dark, carbonaceous mudstone
similar to the Lidah Formation found in the Mojokerto area and MS1-1 well (above the
Pliocene fossiliferous mudstone). This mudstone, based on its stratigraphic position and direct
correlation with the seismic section, is assigned to seismic unit Pre-A (lacustrine deposit).
A petrographic study carried out on four clastic limestone samples (Pasean Formation)
from the south coast of Madura, Gili Raja and Gili Genting Islands indicates the dominance
of sand-size micritised bioclasts (20 - 90%) of foraminifers, algae, brachiopods, bryozoans and
some rounded reworked limestone clasts (Appendix B.l). Rounded quartz grains are present
(up to 30%) with the grain size commonly less than 1 m m . All these clastic components are
cemented commonly by micrite. Figures 6.6 and 6.7 display the typical limestone under a
polarised microscope.
The texture and composition of all these analysed samples appear to support a shallow
shelf environment, where reef and carbonate sand bodies are commonly associated (Tucker &
Wright, 1990). Quartz grains are interpreted to have been reworked from the Tawun
Formation, which may have been exposed along the core of anticlines in Madura Island.
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b. Microfaunal study
Planktonic and benthonic foraminiferal analyses have been carried out on a number of
marly samples of the Pasean Formation from some areas on the south coast of Madura and Gili
Genting Islands. The aim was to determine the age and depositional environment of the
formation. Unfortunately, all samples represent the upper part of the formation, due to the rare
exposure of the middle and lower parts.
Planktonic foraminiferal analysis was carried out on four samples and Figures 6.8, 6.9 and
6.10 summarise the result of this analysis. The presence of Pulleniatina obliqueloculata in the
samples from Gili Genting Island (samples G G 3 & G G 7 ) suggests that the age of the upper
part of the formation is not older than Late Pliocene (planktonic foraminiferal zone N19). The
assemblages are equivalent to the Pliocene section in the MS1-1 well (Fig. 6.3).
Benthonic faunal analysis was carried out on nine samples from the Pasean Formation.
As for the petrographic study, these samples were from exposures along the south coast of
Madura Island (east of Kamal and southeast of Pamekasan), and from Gili Raja and Gili
Genting Islands (Fig. 6.1). Appendix C.2 summarises the results of the benthonic faunal
analysis.
A sample from east of Kamal indicates an inner neritic environment as shown by the
appearance of the genera Quinqueloculina and Elphidium. Six samples from the Pamekasan
area and the Gili Raja Island also suggest a similar environment, with all samples showing a
consistent appearance of the genus Elphidium with the addition of the genera Nonion and
Quinqueloculina. The high content of benthonic foraminifers in two calcareous mudstone
samples from Gili Genting Island is evidence of a littoral environment. The typical fauna for
this environment are the species Ammonia gaimardii and the genus Pseudorotalia (Phleger,
1960; Murray, 1973; 1991).
The fact that a limestone facies has developed in the upper part of the Pasean Formation,
and is in turn unconformably overlain by a lacustrine unit (Pre-A), suggests that there was a
shallowing trend in the environment of deposition. Since this faunal evidence is from the
upper part of the Pasean Formation, a deeper marine facies than littoral or inner neritic may
be present in the lower part. Thus, a transgression may have occurred after unit LT2
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deposition. The occurrence of this Pliocene sea level rise has been indicated during a sequence
stratigraphic study in the Java Sea (see Chapter 7.4.2).
6.5.2 Unit Pre-A
The deepest seismic reflections obtained by the seismic survey come from the Pre-A unit.
O n most seismic lines studied it is characterised by a lack of reflections and tends to be
chaotic. Parallel to subparallel reflection patterns, combined with medium relative amplitude,
high frequency and high continuity, do occur particularly in the upper part of the unit near its
upper boundary. Isochron maps of the unit cannot be constmcted, as its lower limit mostly
lies beyond the depth of penetration of the seismic system. The inferred thickness on some
anticlinal sites is 300 msec T W T (about 225 m ) .
Correlation with well MS1-1 suggests that this unit consists of non calcareous mudstone
featuring plant remains, lignite and pyrite. A similar facies crops out on the south coast of
Madura Island (east of Kamal and south of Pamekasan) and Gili Raja Island, and lies
unconformably on the limy facies of the Pasean and Madura Formations. Based on the
stratigraphic position and similarity in facies, unit Pre-A is interpreted as the eastern extension
of the Lidah Formation. This interpretation is also supported by an X-ray diffraction analysis
carried out on the sample from the Kamal area (Fig. 6.11b), which shows a similar character
to the diffractogram of the Lidah Formation from the Nlampin area (Fig. 5.28). The organic
hump which characterises the richness of organic materials is well displayed as well as calcite
and siderite. However, the X-ray diffractogram of a sample from Gili Raja Island (Fig. 6.11a)
shows the dominance of quartz which indicates a strong terrigenous clastic influence.
The Lidah Formation has been interpreted mostly as a lacustrine deposit (Chapter 5.6),
which was widely developed in the Rembang Zone and eastern Kendeng Zone. The extent
of this facies in Madura Strait appears to be confined to the synclinal basin of this strait. The
fact that it has an unconformable relationship with its underlying units in the Kamal area,
suggests that the deposition was later on the basin margin, and possibly followed a relative
rising of base level (?sea level) position.
6.5.3 Unit A
The beginning of unit A was marked by a depositional hiatal surface due to a relative sea
level rise which terminated unit Pre-A deposition. This sea level rise shifted the depocentre
Ill
landward, and led to sediment starvation of the basinal area. Three subunits can be identified
within unit A, these are: subunits Al, A 2 and A3, which are very apparent on the seismic
records from the easternmost area. Westward extension of these sequences is hard to trace due
to high seismic wave attenuation by the overlying sediments in the western part of Madura
Strait.
a. Subunit Al
Subunit A l is well displayed on the seismic lines from the eastern part of Madura Strait
(Sects 6.8 & 6.11). The lower boundary of this unit is prominent (bottom boundary of major
sequence A ) indicating a high contrast in physical properties with the underlying unit (Pre-A)
which has continuous and parallel reflection patterns. N o erosional features are associated with
this boundary suggesting an abmpt relative sea level rise at the end of unit Pre-A deposition.
The internal character is typified by an upward loss in acoustic transparency - almost reflection
free at the lower part which upward gradually becomes subparallel to hummocky patterns with
variable amplitude and low to high continuity. This indicates an upward coarsening of
lithofacies which m a y relate to the shoaling of the subunit (regressive). This subunit attains
35 msec T W T (about 26 m , Fig. 6.12) in thickness and occurs in the middle of the eastern part
of the strait. The thinning trend toward the basin flanks suggests that this subunit regressed
from both the north and south margins of the basin.
b. Subunit A2
At the end of subunit A l deposition, sea level rose again. The lower boundary of this unit
represents a depositional hiatus as the result of a rapid relative sea level rise which led to
sediment starvation. Almost the whole of sequence A 2 is characterised by an upward loss in
acoustic transparency (Sects 6.5, 6.8 & 6.11), indicating regressive (shoaling) deposits,
according to the criteria of Chiocci (1994). Southeast dipping sigmoid-progradational
reflections occur in the southeast area (Sect. 6.5) which indicate the presence of sedimentary
bypassing in the northern part of the region at the end of subunit A 2 deposition. Toplapping
which truncates the prograding clinoforms can be regarded as a type I unconformity. The
estimated depositional basin floor depth as indicated by the height of the prograding profile
is about 35 m.
The maximum thickness of this unit of about 55 msec TWT (about 41 m, Fig. 6.13)
occurs only locally, probably related to local subsidence. The unit thins toward the north and
112
may onlap onto unit Al, but no clear onlapping features are seen due to poor data. A n
anomalous thickness associated with the progradational clinoforms on the southeastern area is
absent, indicating that the clinoforms may relate to lateral southeastward movement of the
'shelf slope rather than a prograding delta system. The clinoform direction suggests that the
sediments were northerly derived.
c. Subunit A3
Sea level rose again after the deposition of the regressive sequence in subunit A 2 . Subunit
A 3 is almost completely eroded by the subsequent relative sea level fall. The presence of this
sequence can only be observed on Sections 6.5, 6.8, 6.9 and 6.11. The seismic facies shown
by the remainder of the sequence is a high amplitude, high continuity parallel configuration
pattern which suggests shallow marine clastic deposits.
6.5.4 UnitB
A major sea level fall occuned at the end of unit A deposition which led to the subaerial
exposure of Madura Strait, creating a type I unconformity. Unit B consists of a number of
subunits which can be observed on Sections 6.5, 6.8, 6.9 and 6.11.
a. Subunit Bla
Subunit Bla displays complex channel cut and fill features. Seismically it is characterised
by parallel to hummocky reflection patterns combined with low continuity and variable
amplitude. The maximum determined thickness is 50 msec T W T (about 37 m, Fig. 6.14). It
appears that thick deposits occur along the axial line of the syncline (Madura Strait), pinching
out northward (Sect. 6.8) and southward. These suggest that a river system was developed
along the median line of the syncline.
b. Subunit BV
A sea level rise led to the deposition of the volcanic-derived subunit B V which downlaps
onto subunit Bla in the northward direction. Seismically, subunit B V is characterised by a
nearly reflection free pattern indicating a nearly homogeneous deposit. Strong and relatively
continuous reflections occur at the top of the subunit where they are also associated with
channelling (Sects 6.8 & 6.9). The isochron map Figure 6.15 displays the limited areal extend
of this subunit. The conical form of the isochron map which thickens southeastward has
113
suggested that the deposition was related to the volcaniclastic detritus derived from the Lurus-
Ringgit-Beser Volcanic complex.
The Lurus-Ringgit-Beser Volcanic complex comprises three leucite bearing volcanoes
(Lurus, Ringgit and Beser Volcanoes). The volcanism occurred in the Early Pleistocene and
their deposits conformably overlie the neritic to littoral deposits of Plio-Pleistocene age (van
Bemmelen, 1949). These ages are consistent with the well data (MS1-1) which assigns unit
Pre-A to the Late Pliocene. A relatively steep depositional setting is shown in the seismic
sections (Sects 6.5, 6.8 & 6.9) suggesting that this subunit was deposited in a fan-delta setting
(Nemec & Steel, 1988). Stratigraphically, subunit B V is comparable to the fan-delta deposits
of the Pucangan Formation in the eastern part of the Kendeng Zone (Chapter 5.7) which were
mostly derived from the Wilis Volcano.
A similar setting, in which sedimentation was directly conelated with volcanic activity,
has been reported by Ballance (1988) who studied the Late Jurassic Huriwai delta of N e w
Zealand. H e documented that volcaniclastics were fed to a marine environment through a
braidplain delta system, and that there was evidence that the system received periodic influx
related to volcanic eruptions.
c. Subunits Bib, B2, B3 and B4
These subunits can be recognised only in the eastern part of Madura Strait (Sects 6.5, 6.8
& 6.11). They downlap on subunits Bla and B V in a southeasterly direction. In the western
part, the recognition is hindered particularly because of high sound attenuation, multiples and
only subde facies differences between the subunits. Due to the lack of erosional features on
top of each unit, these units may be regarded as parasequences (not sequences) in the sense
of van Wagoner (1985) and Posamentier and James (1993) (see Chapter 4.2 for definition).
In the eastern part of Madura Strait, although the downlap patterns of subunits Bib, B2
and B 3 are recognised, no significant form of prograding clinoforms has been developed
indicating no sedimentary bypassing. However, the subunits overall show a southeastward
progradation as indicated by the relative position of downlaps. Also apparent is that the
depositional slopes were very low and the subunits tend to be characterised by an upward loss
in acoustic transparency which indicates a shoaling upward (Sect. 6.11), The thickness
variations in these subunits is shown in Figures 6.16, 6.17 and 6.18, which all display a
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southeasterly thinning. Some localised thickenings on subunit B 3 are related to local synclinal
subsidence. These isochron maps seem to underline the importance of the straits between the
Madura, Gili Raja and Gili Genting Islands which have acted as feeder channels from the
Madura Island to Madura Strait.
6.5.5 UnitC
A relative sea level fall occurred at the end of unit B deposition and Madura Strait was
exposed subaerially (type I unconformity). Unit C consists of at least eight northwesterly
derived subunits (subunit Cl to C8) and two southerly derived subunits (included within
subunit C V ) . As for the deposition of the marine facies unit B, northwesterly derived
sedimentation appears to be dominant and shows a consistent southeastward 'shelf
progradation. Isochron maps, Figures 6.19 to 6.24, also suggest the important of the straits
between Madura, Gili Raja and Gili Genting Islands as feeder channels, in addition to
sediments derived from the west. The uppermost subunits (C6, C 7 and C8) mostly have been
eroded during the sea level fall at the end of unit C deposition. The presence of these subunits
is noted on seismic lines from the southeastern part of Madura Strait by their clinoform
patterns (Sects 6.5 & 6.8). The other subunits are mappable throughout the strait.
a. Subunits Cla and Clb
Subunit Cla was deposited unconformably on top of unit B. This subunit is characterised
by subparallel to hummocky reflection patterns with variable relative amplitude and low
continuity. Association with channelling is common. Subunit Cla is interpreted as fluvial
deposits. The isochron map (Figure 6.19) shows that this subunit is widely distributed, with
an average thickness of about 30 msec T W T (about 22.5 m ) . Locally thick deposits occur
along an east-west trending synclinal zone in the northern half of Madura Strait which
probably indicates an early development of Pleistocene folding that restricted the development
of fluvial subsystems to these synclines.
A relative sea level rise occurred after the deposition of subunit Cla. Subunit Clb
downlaps on subunit Cla (Sects 6.1,6.2 & 6.10); however, the downlapping appears to be very
subtle due to the low depositional gradient. Seismically it is characterised by parallel to
subparallel reflection patterns with variable amplitude and continuity. A relatively high
continuity with strong reflections can be found particularly in the western part of Madura Strait
(Sects 6.1 & 6.2) which is interpreted as very shallow marine deposits. The isochron map
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Figure 6.20 indicates that this subunit was only thin and was restricted to the western and
northern part of Madura Strait. This subunit has a thickness of about 15 msec T W T (about
11 m ) , and is totally absent in the southeastern part of the area.
Based on the relative position and facies difference between subunits Cla and Clb, it is
suggested that they may represent lowstand and highstand systems tracts respectively. The
absence of subunit Clb on the southern part can be regarded as due to sediment starvation,
since the sediments, which mostly derived from western and northern parts, were mostly
deposited close to the western and northern basin margins because of a sea level rise.
b. Subunit C2
Subunit C 2 has been derived from the north and western part of the basin. It was deposited
contemporaneously with subunit CV. In the eastern part of Madura Strait subunit C 2 shows
a southward downlap. Its internal reflection pattern is characterised by an upward loss in
acoustic transparency which indicates a shallowing upward regressive sequence (Sects 6.8 &
6.11). The upper part shows a subparallel reflection pattern with low amplitude, possibly
representing shallow marine elastics. The thickness attains 20 msec T W T (about 15 m ) , and
thins southward where the subunit is probably amalgamated with the southerly derived subunit
C V (Fig. 6.21).
In the western part of Madura Strait this subunit also shows a regressive sequence. The
lower part (C2a, Sect. 6.1) is characterised by a reflection free pattern suggesting a
homogeneous sediment, probably homogeneous mudstone. In the upper part, the subunit (C2b)
is channellised and associated with a low continuity, variable amplitude parallel reflection
pattern. This part is interpreted as a non-marine fluvial facies.
c. Subunit CV
Subunit C V was deposited northward, contemporaneous with the northerly derived subunit
C2. The distribution of this subunit is mostly confined to the southeastern part of Madura
Strait. The distal part is shown in Sections 6.5, 6.8 and 6.9. The proximal part has mostly
been eroded during the sea level fall at the end of unit C deposition. O n these sections this
subunit is characterised by relatively strong and high continuity reflections of oblique (Sects
6.5 & 6.8) to parallel patterns (Sect. 6.9). The isochron map (Fig. 6.21) shows that this
subunit very rapidly thickens southward (landward). As for subunit B V , the sediments are
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suspected to have been derived from volcanic deposits of the Lurus, Ringgit and Beser
Volcanoes. Both subunits B V and C V were not developed as extensively as the northerly
derived subunits. This is shown in Sections 6.5 and 6.8 where subunit B V has been overlain
by subunits Bl, B 2 and B3, and subunit C V by subunits C4, C5, C6, C 7 and C8. Apparently
the sediment supply from the south rapidly diminished soon after the deposition of subunits
B V and C V , indicating periodic as opposed to continuous volcanic activity.
d. Subunits C3 and C4
Subunits C3 and C 4 are regressive deposits. O n the seismic lines from the eastern part
of Madura Strait these subunits are characterised by an upward loss in acoustic transparency
(Sects 6.8, 6.9 & 6.11). The lower parts are characterised by a reflection free pattern
indicating homogeneous deposits of possibly mudstone. The upper parts are characterised by
parallel reflection patterns with low to high amplitudes indicating shallow marine facies with
frequent intercalations of mudstone and sandstone. Prograding clinoforms within the slope
facies are not well developed, probably due to low sediment supply. A thin oblique
progradational pattern is only shown in subunit C4.
In the western part of the strait, the upper parts of the subunits are characterised by parallel
reflection patterns with variable amplitude and low continuity. They are also associated with
channelling (Sects 6.2 & 6.10). Eastward these patterns change to a parallel pattern with low
amplitude (Sect. 6.10). These features suggest a lateral facies change from non-marine in the
west to shallow marine in the east. Section 6.10 also shows the presence of an eastward
parallel-oblique progradation pattern in subunit C4. This is interpreted as an eastward
prograding coastal facies (Sect. 6.10).
The isochron maps (Figs 6.22 & 6.23) indicate that the thickness of subunits C3 and C4
is relatively similar, averaged at 20 msec T W T (about 15 m ) . The subunits thin and downlap
southeastward, suggesting a general progradation in this direction.
e. Subunit C5
Subunit C 5 represents a type I sequence, composed of two subunits - C5a and C5b.
Figure 6.24 is an isochron map of combined C5a and C5b. Subunit C5a is a lowstand deposit
(wedge) which accumulated as sea level fell at the end of subunit C 4 deposition. The areal
distribution of this subunit is very limited and is confined to the basin floor in front of the
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prograding slope subunit C 4 which is located in the eastern part of Madura Strait (Sects 6.8
& 6.9). In Figure 6.24 subunit C5a is not represented, but can be found beneath the local
thickening in the middle east of the area. A separate isochron m a p was not constructed due
to the difficulty in defining the upper boundary of this subunit on most sections. Seismically
it is characterised by a mounded form, and probably baselaps on both subunits C V and C4.
Its internal reflection patterns, as shown on Section 6.8 are parallel and show a medium
amplitude and high continuity. Its deposition may have contemporaneously occurred with the
channelling and prograding coastal facies at the upper part of subunit C 4 (Sects 6.2 & 6.10).
As the sea level rose, deposition of subunit C5a ceased. Subunit C5b developed as a
highstand regressive deposit (Sects 6.8 & 6.10). The regressive situation apparently occurred
as the result of a combined sea level stillstand and high sediment supply. This is obvious in
the eastern part of the area where sedimentary bypassing occurred, combined with extensive
southeastward sigmoid prograding clinoforms (Sects 6.5, 6.8 & 6.9). The estimated water
depth of basin floor measured from the clinoform profile was about 70 msec T W T (about 52
m ) . Figure 6.24 shows that the thickness of the prograding lobe reaches 55 msec T W T (about
41 m ) . This southeastward prograding unit is interpreted as a prograding delta system.
Unfortunately, the recognition of the complete delta system is hindered, as its proximal facies
was eroded during the sea level fall at the end of unit C deposition. However, the direction
of clinoform dip suggests that the delta system was sourced from the northwest, and again
indicates that the strait between Madura and Gili Raja Islands acted as a feeder channel.
A much smaller delta system contemporaneous with the previous delta system probably
developed in the western part of the area as displayed on Section 6.10. It is separated from
the previous delta system by a shallow marine facies, indicated by a seismic facies with a high
amplitude and high continuity parallel pattern on Figure 6.24. In the proximal direction
(westward) the delta system is characterised by high amplitude hummocky clinoforms
associated with channelling (Sect. 6.10 & Fig. 6.24). This part is interpreted as a fluvial
facies. The delta front developed on a very low slope and profile. Seismically it is
characterised by a thin complex sigmoid-oblique reflection pattern with strong amplitude
indicating a low sediment supply and high energy depositional setting. Eastward these features
rapidly change into a low to high amplitude parallel reflection pattern of probably mudstone-
prone facies (Sect. 6.10).
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f. Subunits C6, C 7 and C 8
These subunits can only be recognised by their strong and continuous reflections on top
of prograding clinoform facies in the eastern part of the area (Sects 6.5, 6.8 & 6.9). Their
characteristic 'shelf-type' facies have mostly been eroded during the sea level fall at the end
of unit C deposition.
The stacking pattern of subunit C6 on Sections 6.8 and 6.9 suggests that the sea level rose
after the deposition of subunit C5b. This increased the accommodation space and the
sediments aggraded on top of subunit C5b. In contrast, the remnants of subunits C 7 and C8
do not permit evaluation as to whether these subunits were deposited in transgressive or
regressive settings.
Seismically subunit C6 is characterised by a parallel reflection pattern combined with low
amplitude indicating a low energy depositional setting. The internal pattern of clinoform
facies, subunits C 7 and C8, is characterised by a lack of reflections which indicates
homogeneous sediments of probably mudstone.
6.5.6 UnitD
The beginning of unit D is marked by pronounced erosion which is related to a significant
sea level fall at the end of unit C deposition. This sea level fall exposed the whole Madura
Strait subaerially, and part of unit C 6 and most of units C 7 and C 8 were eroded. The
prolonged exposure was also encouraged by the commencement of the major Pleistocene
folding event. A number of subunits can be recognised in unit D, but most of them are too
thin to be mapped. In addition, the much less pronounced subunit boundaries make then-
recognition difficult. Only the lowermost subunit (subunit Dla) was mapped seismically as
a distinct unit, and a combined isochron map of the remaining units was produced. The
present sea floor is the upper boundary of unit D.
a. Subunit Dla
O n the seismic sections subunit Dla is distinctive and characterised by variable amplitude
and poor continuity of reflections. It is interpreted as an extensive fluvial deposit. In the
upper part parallel to hummocky reflection patterns are shown, which may be associated with
ponds or floodplain deposits (Sect. 6.8). This part is associated with a river system (Sects 6.9
& 6.11) which was later abandoned during the subsequent sea level rise. Figure 6.26 displays
119
the positions of observable channel cuts which support the occurrence of an axial river system
in the synclinal basin. The flanks of the basin contributed sediment to this river system.
Toward the east, an eastward running main channel developed and exhibited a high sinuosity.
The isochron map Figure 6.25 indicates that the thickness of subunit Dla varies from 0
to 70 msec T W T (0 to 45 m ) . This m a p and the mesh diagrams in Figures 6.29b and 6.29c
clearly show that the deposition of Dla was strongly controlled by the formation of Late
Pleistocene folds. Thick deposits are confined to the synclinal areas. The westward relative
thickening of this subunit is due to the fact that the downwarping was more intensive in two
distinct centres in the western part of the basin.
Apart from the local control by the Late Pleistocene fold formation, the general
depositional style of this fluvial system is similar to subunits Bla and Cla. Miall (1982)
pointed out that river flows basically can be generalised into transverse, flowing directiy from
the uplifted area across structural grains, and longitudinal, flowing parallel to the strike of the
basin axis. The river systems in Madura Strait (Bla, Cla and Dla) are included in the second
category, where the basin controlling the flow has a relatively simple synclinal geometry.
A modem example of such a longitudinal river system is the Po River system of northern
Italy, which has been studied by Ori (1993). This river system, which supplies most of the
detritus into the Adriatic Sea, is confined within the asymmetrical subsiding Po Basin, which
is bounded to the north by the Alps mountain chain and to the south by the active thrust-
faulted belt of the Apennines. Alluvial fan and contributary river systems occur on both flanks
of the basin, and flow into the main longitudinal Po River. A complex interplay between
different subsidence rates within the basin, climatic conditions and different types of sediment
sources, has led to the difference in fluvial facies and body geometries between the north and
southern flanks of the basin. In the southern sector, where thrust faulting is active, fan and
plain systems are thin and are formed by a few depositional processes. In contrast, the
northern sector is characterised by laterally widespread fan systems formed by many
sedimentation events. Although the study area has a different setting, this model of fluvial
sedimentation may be similar to the more active alluvial fan sedimentation occurring in the
vicinity of the active Quaternary volcanism in the south.
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b. Subunit Dlb
Rapid rise of relative sea level occurred after the deposition of subunit Dla and the
depocentres shifted toward the basin margin. Drainage patterns created on top of subunit Dla
were completely abandoned. Sections 6.1 and 6.7 suggest that the top of this subunit (top of
Dl) occurs very close to 66 msec T W T (-50 m from the present sea level). While the lowest
observable previous sea level lowstand feature (fluvial channelling) on top of subunit Dla
occurs at 140 msec T W T (about -105 m from the present sea level; Sect. 6.9). Assuming very
minor structural deformation, the sea level rise from the previous lowstand was at least 74
msec or 55 m.
In the western part subunit Dlb is relatively thin (Sect. 6.1), characterised by medium
amplitude, relatively continuous reflection patterns, and is associated with small channels. It
is suspected that here, this subunit was deposited in a very flat area in a very low energy
environment (?tidal flat).
In the deeper setting (Sects 6.2 & 6.3), subunit Dlb is typified by a reflection free
character of possibly homogeneous mudstone, associated with mounded and irregular shapes
of probably patch reefs (bioherms). A very strong reflector on Section 6.3 (line M C - M C ) ,
which is possibly associated with an undulating sheet carbonate surface, has produced
pronounced multiples and diffractions.
In the topographically steeper area (Sect. 6.7) subunit Dlb is wedge shaped and shows a
relatively continuous, medium amplitude reflection pattern. Figure 6.27 illustrates the
palaeogeography during the deposition of subunit Dlb. The maximum thickness is about 21
msec T W T (16 m ) .
c. Subunit D2
Subunit D 2 comprises a number of units that resulted from repetitive small transgressions
and regressions. T w o depocentres were present along the southern and northern margins of
Madura Strait. Deposition along the southern margin was less extensive compared to the
northern margin, which was probably due to a lower sedimentation rate. O n Section 6.3 (line
M C - M C ) where the topography was relatively steep, four wedges of transgressive/regressive
facies can be identified and are referred as D2a to D2d. These wedges progressively shifted
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basinward which may suggest a progradation or gradual relative sea level lowering while
oscillating.
A similar mode of deposition, but on a larger scale, has been documented by Tesson et
al (1993) on the Rhone (France) continental shelf. They associated the wedges with
alternating episodes of coastal progradation and transgression, which are believed to occur
during the latest Pleistocene glacial lowstands.
Extensive development of subunit D2 occurred on the northern margin of Madura Strait
(Sect. 6.6). It is interpreted as a prograding wedge of delta facies, fed through a channel
system in between Madura and Gili Raja Islands. The thick deposit of the delta system forms
the main component in the combined isochron map Figure 6.28, and appears as a pronounced
bulge. The thickness reaches more than 30 m. This extensive development probably resulted
from a high sedimentation rate and the presence of folds (Fig. 6.29b) on the sides of the delta
which acted as a barrier to wave processes, and also concentrated sedimentation by funnelling
through a narrow gap.
d. Subunit D3
Subunit D 3 is the uppermost observable sequence on the seismic records. The lower
boundary is probably an erosional surface which is very prominent in the northern part of the
strait where the upper parts of deltaic deposits have been truncated (Sects 6.6 & 6.7). The
thickness of subunit D 3 may reach 15 m (Sect. 6.6). Details of this subunit are mostly
obscured by peg-leg multiples at the top of the subunit.
The lower boundary of subunit D3 is associated with strong and irregular reflections which
may have resulted in very poor seismic data quality underneath (Sects 6.6 & 6.7) and obscured
details of the earlier strata. The lateral extent of such reflections appears to follow the extent
of the previous deltaic deposit, subunit D2. A similar phenomenon has been reported by
Chiocci (1994) in association with Late Pleistocene delta systems, interpreted as the result of
interstitial gas within the delta system. However, a close examination on some sections in the
study area has indicated that such a feature occurs because of a very thin layer in between the
regressive deltaic deposit (subunit D2) and the overlying subunit (D3), which may be
associated with transgressive lag deposits over the deltaic deposits.
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6.6 D I S C U S S I O N
6.6.1 Correlation with onshore data
Quaternary sediments are widely exposed in the onshore part of the East Java area
(Kendeng Zone) along Late Pleistocene anticlinal flanks. Discussion regarding their
depositional environments has been given in Chapter 4. Correlation between these deposits
and Quaternary sediments developed in Madura Strait has been made possible due to the fact
that during the Early and Middle Pleistocene East Java and Madura Strait formed a single
elongated basin (van Bemmelen, 1949).
During the Middle Pleistocene relative sea level rose significantly in East Java. This was
marked by a widespread deposition of shelf mud (green mudstone) which signifies the
beginning of the Kabuh Formation deposition (Chapter 5.8; Figs 5.48, 5.49 & 5.50). The
stratigraphic position of this mudstone, which overlies the fluvial and coastal deposits (unit U2)
of the Pucangan Formation, has provided a good seismic marker due to the difference in
physical properties between the two formations. In the closest seismic line to the onshore area
(line M A - M A ' , Sect. 6.1), the mudstone facies is represented by seismic facies subunit C2a,
and the fluvial and coastal facies are represented by subunits Cla and Clb respectively.
A lithofacies study on the outcrops on the anticlinal flanks north of Mojokerto (Chapter
5.7; Figs 5.30, 5.37, 5.44 & 5.48) indicates that sediments below the fluvial deposits (unit U2)
of the Pucangan Formation are mostly tidally influenced deposits. Alluvial fan deposits (unit
M 3 ) occur in the section from the Perning area (Fig. 5.37) where they also contain abundant
vertebrate fossils and hominid remains. These deposits according to Duyfjes (1938b) and van
Bemmelen (1949) were sourced from the Wilis Volcano. The stratigraphic position of these
deposits, which are below unit U2, is exactly similar to the fan system (subunit B V ) developed
in the southeastern part of Madura Strait (Sects 6.5, 6.8 & 6.9). This later fan system occurred
as the result of activity of the Lums, Ringgit and Beser Volcanoes. Figure 6.30 displays the
correlation between outcrops in the area north of Mojokerto and seismic sections M B - M B '
(representing the western part) and M E - M E ' (representing the eastern part).
The lithofacies study in the northern part of the Mojokerto area indicates that the Pucangan
Formation is underlain by thick lacustrine mudstone deposits of the Lidah Formation, which
attains 500 m. Field observation in the Banyuurip area (Fig. 5.30) indicates that the facies is
almost entirely massive. The thickness and lack of bedding in this mudstone facies would
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produce a prominent, thick, reflection-free unit in the seismic sections studied. These
features occur in unit Pre-A which is widespread in Madura Strait. Well data (MS1-1)
indicate that this unit is characterised by non-calcareous mudstone occurring at a depth of 277
to 451 m , which was deposited during the Late Pliocene. Field studies in the Gili Raja Island
(Fig. 6.1) found that this same facies lies on the Pliocene marly facies (Pasean Formation)
which is equivalent to the M u n d u Formation in East Java.
6.6.2 Depositional timing
a. Units A, B and C
Absolute and relative dating have never been carried out on Quaternary sequences in
Madura Strait. In this study, the inferences on the depositional timing of Quaternary sequences
in Madura Strait are based particularly on the correlation of Madura Strait sequences with the
inland sequences (north of Mojokerto). The previous section has discussed the facies
relationship between these areas.
A recent palaeomagnetic study was carried out in the Perning area, north of Mojokerto
City (some 35 k m west of Madura Strait) by a Japan-Indonesia joint research team in 1986-
1988 (Hyodo et al, 1993; Chapter 3.5). The palaeomagnetic measurements that were carried
out on mudstone and tuffaceous mudstone samples have revealed the presence of two
Quaternary magnetic events, the Olduvai and Jaramillo. The results of these measurements are
shown on the measured section in Figures 5.37 and 6.30. The Jaramillo magnetic event is
coincident with the alluvial fan deposits (unit M 3 ) which in the seismic study is correlated with
the subunit BV.
Based on the previous correlations, unit C in the seismic sections has to be younger than
the Jaramillo magnetic event. The two important features that occurred during the deposition
of this unit were (1) the presence of two significant and widespread subaerial exposures which
may indicate significant relative sea level falls, and (2) the presence of highly repetitive
sequences. The first subaerial exposure occurred at the beginning of subunit Cla deposition
and the second occurred after the deposition of subunit C4. These features suggest that the
deposition of subunits in unit C corresponds to the relatively regular climatic cycles during the
Bmnhes Normal Epoch (Fig. 6.32), and the two significant relative sea level falls are
speculatively correlated to the glacial maximum stages 16 and 10 where the intensities of
glaciation were relatively higher.
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b. UnitD
Inferences on depositional timing of subunit D were based particularly on the assumption
that structural deformations were minor after the deposition of subunit Dla. The positions of
erosional features with respect to the present position of sea level were then compared with
the adjusted sea level curve of Huon Peninsula (New Guinea) by Chappell and Shackleton
(1986, Fig. 6.31) and the oxygen-isotope record compiled by Harland et al. (1989, Fig. 6.32).
Subunit Dla is bounded at its base and top by prominent erosional surfaces. The deepest
fluvial channelling at the top of this subunit occurs at depth of about 140 msec T W T (-105 m
from the present sea level, see Sect. 6.9). Between this fluvial channelling and the erosional
surface at the top of subunit D 2 the sequence is characterised by a number of small marine
units (Dlb, D2a, D2b, D2c, & D2d; Sects 6.3 & 6.6). These suggest that unit Dla was
deposited during the sea level lowering of oxygen isotope stage 6. Deposition was enhanced
by the development of folds, and consequent increased erosion, during this time. Subunits
Dlb, D2a, D2b, D2c and D2d on Section 6.3, and subunits Dlb and D 2 on Sections 6.6 and
6.7, which were deposited after folding, can be interpreted to have resulted from minor sea
level oscillation after the oxygen isotope stage 6 (Figs 6.31 & 6.32).
Subunit D3 is the uppermost observable unit in the study area. As shown on Sections 6.6
and 6.7, the base of this subunit is probably an erosional surface. It is interpreted that the
erosion took place during the oxygen isotope stage 2 (the last glacial maximum). Sea level
lowering during this time may have reached more than 100 m below the present sea level (Fig.
6.31; Chappell & Shackleton, 1986).
6.6.3 Sedimentary source
The study has indicated that Madura Strait has been fed from three different sources of
sediments, in the south, west and north. Seismic features and isochron maps show that the
south and westerly derived sediments were not as effective a source as northerly derived until
the Middle Pleistocene. This probably relates to the fact that sediments from the south and
western parts of the basin were volcanogenic. Their supply depended on the intermittent
volcanic activity, and lateral deposition would occur when a sea level rise provided an
accommodation space. In contrast, sediments from the north were the erosion products of
Madura Island. The absence of Quaternary marine deposits in this island (Duyfjes, 1938c;
Aziz et al, 1993; Situmorang et al, 1992) suggests that the island may have been subaerially
125
exposed since the end of the Pliocene. In addition, the Miocene and Pliocene sediments were
then folded during the Plio-Pleistocene with at least two pairs of east-west trending anticline-
synclines occurring along this island (Duyfjes, 1938c; Aziz et al, 1993; Situmorang et al,
1992). The folds have allowed the development of axial river systems which may have
become effective erosion agents and maintained a continuous sediment supply to Madura Strait.
6.7 SUMMARY AND CONCLUSION
The seismic stratigraphic study in Madura Strait revealed the presence of four major
Quaternary sedimentary units (units A, B, C and D ) . These units are bounded at their base,
above the latest Pliocene unit Pre-A, by a surface of non-deposition or erosion. Due to then-
durations being less than 1 M a , these Quaternary units can be classified as fourth order
sequences. Detailed analysis of these units has further found the presence of smaller subunits.
The lateral distribution of these subunits is strongly controlled by their depositional setting.
The southerly derived subunits, although attaining a significant thickness, are commonly
limited in their lateral extent. The direction of subunit thickening suggests that these units
were derived from the volcanic products of the Lurus, Ringgit and Beser Volcanoes. The
north and northwesterly derived subunits, in contrast, were mostly extensively developed and
show a cyclical pattern of lowstand and highstand. The two laterally extensive delta units were
related to these sources.
The depositional timing of depositional units in Madura Strait has been inferred by
correlation with the micropalaeontologically dated well M S 1-1, the palaeomagnetically dated
section in the Peming area (north of Mojokerto), the marine oxygen isotope record and the
adjusted sea level curve of the Huon Peninsula (New Guinea). A rough estimation on the
depositional timing has been obtained. The deposition of unit A occurred in the Early
Pleistocene, probably starting close to the Olduvai magnetic event. Unit B was deposited
during the late Early Pleistocene to early Middle Pleistocene. Unit C was deposited in the
middle Middle Pleistocene to late Middle Pleistocene, and unit D was deposited during the
Late Pleistocene.
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127
CHAPTER SEVEN
LATE TERTIARY AND QUATERNARY SEISMIC SEQUENCE
STRATIGRAPHY OF THE SOUTHEAST JAVA SEA
7.1 INTRODUCTION
The southeast Java Sea is the submerged part of the Sunda Shelf (Chapter 2), and lies in
a different tectonic and geological setting from Madura Strait. In Madura Strait, a north-south
compressional tectonic stress is dominant, and the sedimentary basin is simply an east-west
synclinal basin, which is bordered in the north by the anticlinal Rembang Zone (Madura
Island), and by a zone of active Quaternary volcanism in the south. In contrast, the southeast
Java Sea lies on a relatively stable continental shelf, particularly during the Quaternary, as
indicated by the present study which found less structural deformation.
Marine geological investigations in the southeast Java Sea have mostly been carried out
as part of regional studies on the Sunda Shelf. Based on regional geophysical data, Ben-
Avraham and Emery (1973) noted that Tertiary sedimentation in the southeast Java Sea
occurred in basins which were bounded mostly by northeastward trending faults. This finding
was confirmed by Sudiro et al. (1973) and Bishop (1980), who further indicated that these
basins are commonly half grabens. The present study confirms the finding by these
investigators that during the Plio-Pleistocene, marine sedimentation continued over these
basins. However, this study also finds that synclinal east trending basins occurred during the
Pliocene and Early Pleistocene. Sedimentation in these basins gradually changed basin
morphology to a relatively flat plain at the end of the Quaternary.
This chapter discusses the sedimentary facies distribution and chronology in the southeast
Java Sea during the Late Tertiary (particularly the Pliocene) and Quaternary. This discussion
is an extension of similar discussions in Chapters Four and Six which examined the northern
part of East Java and Madura Strait respectively. As in these previous chapters, the discussion
relies heavily on seismic data, which have been interpreted by applying the sequence
stratigraphic concepts developed by Vail et al. (1977), Posamentier et al. (1988) and
Posamentier and Vail (1988). In many aspects, this study faces the same problems as in
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Madura Strait due to the lack of reliable dating, as well as published geological studies. Thus,
this study represents a preliminary interpretation of the sequences in this area.
The seismic sequence stratigraphic analysis presented in this chapter is considered to be
a general overview in terms of sequence delineation. Detailed analysis cannot be carried out
on most sequences, because of three major problems: (1) the limited penetration of the seismic
system used, often resulting in seismic signals too weak to be interpreted, (2) presence of
strong multiple reflections, particularly in the area where strong reflectors, such as
unconformity surfaces, are present, (3) occurrence of complex channel cuts and fills. The first
two problems apply to the Miocene sequence and the lower part of the Pliocene sequence, and
the last applies to the Quaternary sequences.
7.2 DATA
The data base for this study is drawn from seismic profiles totalling some 3750 line km
in the Java Sea (Fig. 7.1). All geophysical data were obtained from the Marine Geological
Institute of Indonesia which ran the survey in 1989/1990. The seismic system used is similar
to that in the Madura Strait survey (Chapter 6), a single channel 600 Joule sparker system,
fired every 0.5 seconds. In addition, a 300 Joule boomer system was m n in the area near the
mouth of the strait between East Java and Madura Island. The seismic signals were not tape
recorded, but direcdy band pass filtered (200-2000 Hz) and graphically recorded in analog
format during the survey. Due to this technique, no further data processing was carried out.
The estimated vertical exaggeration on all seismic sections presented in this chapter is about
21 times for the sparker system and about 17 times for the boomer system. The ship positions
during the survey relied mostly on the Magnavox M X 1157 satellite navigation system with
point accuracy of 36 m. The profiles were oriented north-south and spaced 5 to 10 k m apart.
The isochron (thickness in time units) and time structure maps presented in this chapter
are the results of digitisation of all the seismic data following the seismic interpretation. The
computer programs used are the same as those used in the Madura Strait study, and the source
codes are provided in Appendix F. O n the seismic profiles presented, the depth was calculated
based on an assumed seismic velocity of 1500 m/sec, because of the absence of seismic
velocity data. This depth is in fact valid for the water column only, and should be regarded
as a rough estimation. Figure 7.2 shows the locations of seismic lines used during the discussion.
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Stratigraphic control for the analysis of seismic data was provided by six petroleum
exploratory wells, JS1-1, JS2-1, JS3-1, JS8-1, JS10-1 and JS16-1. A graphic summary of
biostratigraphic analyses of these wells (except for the JS1-1 well) is provided in Appendix
E. Although more well data are available for the study than in Madura Strait (Chapter Six),
the study suffers from a similar problem. Biostratigraphic and lithofacies analyses done on
these wells were based on well cuttings which were commonly sampled every 30 ft (9.1 m )
penetration. These data may not be accurate, but have narrowed the age estimation of the
stratigraphic time markers. A more serious problem is that the well data may not be applicable
to the Quaternary sequences, because most of the wells were drilled on top of anticlines where
Late Tertiary and Quaternary sediments are thin or absent due to erosion and sediments
wedging out.
7.3 STRUCTURAL GEOLOGY AND BATHYMETRY
The discussion presented in Chapter 2 indicated that northeast-trending stmctures are
prominent in the southeastern part of the Java Sea (Figs 2.2 & 7.3), and have been the major
control for the Early Tertiary sedimentation. Many of these stmctures are half grabens that
formed on the pre-Tertiary shelf (Bishop, 1980). These major features were interpreted by
Ben-Avraham and Emery (1973) as resulting from past interaction between the Eurasian and
Indian-Australian lithospheric plates, the principal ridges probably being part of an island arc
system active during the Late Cretaceous-earliest Tertiary (Bishop, 1980). Such an island arc
complex has been deduced from the occurrence of pre-Tertiary ophiolites indicating a
subduction complex cropping out in Central Java, and in Southeast Kalimantan (van
Bemmelen, 1949) possibly representing a previous subduction complex (Katili, 1989).
The Karimunjawa Arch is the dominant ridge in the eastern Java Sea which extends into
the offshore area of southern Kalimantan as a broad positive feature (Bishop, 1980). It is
capped by the Karimunjawa Islands on which pre-Tertiary quartzite and phyllitic shale, cut by
basic dykes, and probable Quaternary fissure-eruptive sheets crop out. This arch is separated
by the narrow, northeast trending West Florence Deep from the Bawean Arch. The Bawean
Arch is characterised by alkaline volcanism of the latest Neogene or Quaternary and the steeply
dipping Miocene marine strata (van Bemmelen, 1949).
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The present seismic study on the southeastern part of the Java Sea found that the Pliocene
and Early Pleistocene sedimentation was still partly controlled by the structural features
mentioned above. The sediments were derived primarily from the two positive features, and
from the growing anticline of Madura Island. The stmctures became ineffective from about
the Middle Pleistocene when the morphology became relatively flat, and the Middle and Late
Pleistocene strata are characterised by extensive channel cut and fills that have resulted from
frequent Quaternary sea level fluctuations. Intense faulting in the study area is absent, but
active faulting, at least until the Late Pleistocene (as indicated by faulted Late Pleistocene
sediments) occurs on the borders of the major structural elements in the study area (Fig. 7.3).
Figure 7.4 displays the three dimensional mesh diagrams which represent the present
situation at the top of the Miocene, Pliocene and Pleistocene (bathymetry) strata respectively.
The pronounced highs in the north represent the Karimunjawa and Bawean Arches. It is
shown that besides sedimentation occurring in the West Florence Deep (a half graben system),
the Pliocene sedimentation was strongly controlled by an east-west, asymmetrical synclinal
structure between the Karimunjawa and Bawean Arches in the north, and the east trending
ridge in the south which extends from the north of Java to Madura Island. This ridge is the
anticlinal zone termed the Rembang Zone in the previous chapters. More complex Pliocene
deposition occurred in the area north of Madura Island where a small east trending ridge was
developing during the deposition which allowed the accumulation of a locally thick Pliocene
deposit.
The present bathymetry of the study area (Figs 7.4a and 7.5) may reflect the geological
processes in the Late Pleistocene. At present the deepest water setting in the study area is
approximately 85 m. The gentle, gradual northeastward deepening may partly relate to the
growing anticlines in the northern part of East Java and Madura Island. As revealed by the
present seismic analysis, the average gradient is 0.4 m/km. In the area closer to the East Java
and Madura coasts, the bathymetry rapidly shallows due to the Late Pleistocene coastal
deposits derived from these areas. In the northwestern Rembang Zone, the deposits were
sourced mainly from the Quaternary Muria Volcano on the peninsula northwest of Rembang.
In the far north of the map (Fig. 7.5) the bathymetry shows pinch and swell stmctures which
may reflect an underlying structural feature.
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7.4 SEISMIC S T R A T I G R A P H Y
Seismic analysis indicates that the Late Tertiary and Quaternary sediments in the study
area can be subdivided into three major seismic units. These are referred to as units JM, JP
and JQ, which correspond to the Miocene, Pliocene and Pleistocene ages respectively. Ages
of these units are based on biostratigraphic data from petroleum exploratory wells in the Java
Sea provided by the Cities Service Indonesia Inc. and Pertamina Company, and outcrop
geological studies on Madura Island by Situmorang et al (1992) and Aziz et al. (1993). The
units identified are bounded at their bases by regional unconformities [type I unconformity
according to the Posamentier & Vail (1988) criteria], and each of these units consist of two
or more subunits which are characteristic in areal distribution, thickness, and seismic character.
7.4.1 UnitJM
This unit can only be observed on the structurally high areas, such as near the Bawean and
Karimunjawa Arches, and Madura Island. Its internal reflection patterns and areal distribution
are poorly defined, particularly because of the limited penetration of the seismic system used
and strong multiple reflections. The lower boundary is unidentifiable, but the upper boundary
is a regional unconformity as shown by a pronounced erosional surface on the structurally high
areas (Sects 7.1, 7.3 & 7.8). The estimated age of the unit is Miocene, based on the
correlation with the micropalaeontological data of several wells in the study area (Appendix
E). The Miocene strata in the JS8-1 and JS3-1 wells are typified by the occurrence of larger
forarninifers which are characteristic of this age, such as Lepidocyclina, Miogypsina,
Flosculinella and Alveolinella (Rahaghi, 1984; Hamaoui, 1984; Butterlin, 1984). The top of
the Miocene strata in most wells was also defined by an abmpt increase in planktonic
foraminifers. This change indicates the occurrence of a significant sea level rise at the end of
the Miocene, since larger forarninifers are commonly associated with shallow marine waters
(Moore et al, 1952; Serra-Kiel & Reguant, 1984), in contrast with planktonic foraminifers
(Murray, 1991).
The seismic lines Sections 7.1 and 7.7 suggest that unit JM can further be subdivided into
subunits JM1 in the lower part and JM2 in the upper part. The subunits appear to be separated
by an erosional surface. On the Karimunjawa and Bawean Arches units JM1 and JM2 are
characterised by a medium amplitude, continuous parallel-subparallel reflection pattern (Sects
7.1 & 7.7) which is interpreted as interbedded sandstone and mudstone facies. These deposits
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are overlain by a reefal limestone on both arches, and possibly on Madura Island as reported
by Situmorang et al (1992) and Aziz et al (1993). The mounded external form of the
carbonate buildups of unit J M 2 is clearly shown on Sections 7.3 and 7.5, located on the
southern flank of the Bawean Arch. O n the southeastern flank of the Karimunjawa Arch, the
tops of subunits J M 1 and J M 2 are characterised by a rough, undulating surface (Sect. 7.1),
which is interpreted as karst topography formed during the subsequent sea level fall of each
sequence. The lateral extent of this morphology is limited, and is not present in the deeper
setting. This suggests that the limestone formed only during sea level highstands, which
implies that the underlying siliciclastics are lowstand deposits.
Some seismic tines were ran close to Bawean Island and suggest that unit JM (?JM1) is
probably equivalent to the Miocene strata exposed on this island. A geological study by van
Bemmelen (1949) reported the occurrence of limonitic sandstone, interbedded with lignite,
marl and crystalline limestone on Bawean Island. H e indicated an age of Early to Middle
Miocene based on the occurrence of larger benthonic foraminifers Lepidocyclina, Cycloclypeus
and Alveolina bontangensis which is in agreement with the evidence from JS8-1 and JS3-1
wells (Appendix E). In the Rembang Zone, three major Middle to Late Miocene seismic
units, E M 2 , L M 1 and L M 2 , correspond to the Tawun, Wonocolo and Ledok Formations
respectively (Chapter 4.3). The Middle Miocene Tawun Formation is a paralic marine facies,
and shows a facies similarity with those developed in Bawean Island. The Late Miocene
Wonocolo and Ledok Formations are shelfal facies, and are characterised by hemipelagic marl
and tidally-influenced glauconitic sandstone respectively. The absence of the Late Miocene
deposits (Wonocolo and Ledok Formations) in the topographically high areas, such as Bawean
Island, suggests that the deposits were basinally restricted. Unfortunately the limited
penetration of the seismic system used does not permit recognition of this Late Miocene
deposit.
7.4.2 UnitJP
Unit JP is relatively thick and was deposited following unit JM2. It consists of two type-I
sequences, subunits JP1 and JP2, with subunit JP1 forming the major sequence. Correlation
between the seismic study and the micropalaeontological data from some petroleum exploratory
wells indicates that these subunits developed during the Pliocene. This age was determined
particularly by the appearance of the species Sphaeroidinella dehiscens. The first appearance
of this species (5.05 M a ; Berggren et al., 1985) has been used to indicate the boundary of
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planktonic foraminiferal zones N18/N19 (Banner & Blow, 1967; Blow, 1969). The top
boundary of unit JP is an erosional surface marking extensive subaerial exposure in the study
area at the end of the Pliocene. Subunits JP1 and JP2 may be classified as third order
sequences as their duration was no more than 1 M a (Vail et al, 1977). The sediment sources
of these subunits were mainly the Karimunjawa and Bawean Arches in the western half of the
study area (Sects 7.1 & 7.3). In the eastern half, the deposits were sourced from both the
Bawean Arch and Madura Island, but the distribution was complicated by the development of
folds (Sect. 7.8).
a. Subunit JP1
The deposition of this subunit is well described by the theoretical model of shelf margin
sequence stratigraphy by Posamentier et al (1988). T w o main systems tracts can be defined,
the lowstand and highstand systems tracts. The transgressive systems tract on most of the
seismic lines studied is absent or unidentified, probably due to a combination of low
depositional slope, low sedimentation rate and rapid sea level rise which did not permit
formation of a seismically resolvable transgressive unit. The ideal cross-sectional view for the
western half of the area is shown on Section 7.3. Figure 7.6 displays the time structure map
for the top of the Miocene strata on which unit JP1 was deposited, and Figure 7.7 displays the
isochron map of total unit JP (subunits JP1 and JP2). These maps indicate that the Pliocene
basin in the western half of the study area was still influenced by the normal fault movement
of the half graben system in the West Florence Deep. The occurrence of the deepest basin and
accumulation of the thickest Pliocene sediments in this trough (particularly along the normal
faults) has further suggested probable faster subsidence and sedimentation rates. In the eastern
half of the area, the influence of the previous structural configuration (Figs 7.3 & 7.4c) is not
really obvious. The Pliocene structural development (east-west trending folds) had more
influence on the sedimentation, as indicated by the trends of basin morphology and the
Pliocene sediment accumulation (Figs 7.6 & 7.7).
Lowstand systems tract
A sea level fall occurred after the deposition of unit JM2. This exposed most of the
positive relief areas in the study area. Karstification occurred on the Karimunjawa Arch and
possibly on the Bawean Arch (Sects 7.1 & 7.5). The estimated minimum relative sea level fall
was about 130 msec T W T (about 100 m ) , measured from the profile between the highstand
reefal facies of unit J M 2 and the nick on Section 7.3 This nick point is regarded as the
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position of sea level during the lowstand, based on the seismic features which show an
erosional surface on the shelf above it. Consequently, the position of the 'shelf edge before
erosion should be put some distance higher than this point. Figure 7.6 displays the
palaeogeographic map during the sea level lowstand, plotted on the time structure contours to
top of the Miocene strata.
Pronounced valley and fluvial incisions occurred on the Bawean Arch (Sects 7.3 & 7.7).
Most of these incisions are traceable on seismic sections and ended basinward in the lowstand
systems tract. The channel and valley incisions occur during the interval of sea level fall,
according to the sequence stratigraphic concept (Posamentier & Vail, 1988). During this active
incision, minimum stream deposition occurs and sediments are fed directly to the stream
mouths near the shelf edge and are ultimately deposited as turbidites where abrupt decreases
in slope gradient occur (Posamentier et al, 1991). The lowstand fans (basin floor fans) are
not observed in the study area due to the limited penetration of the seismic data. But some
interpreted submarine canyons which may be associated with the feeder channels are observed
beneath the lowstand wedge. The canyon on Section 7.3 in fact can be traced landward, and
is associated with the subaerial valley incision shown on Section 7.7
As sea level reached the lowest point, a seaward prograding lowstand wedge developed
(Sect. 7.3). The bottom set on most seismic lines is not observed due to weak seismic signals.
The topset is nearly horizontal (about 0.5° dip), but the parallelism and concordance of internal
strata that commonly characterise a low depositional angle (Mitchum et al, 1977) are not well
displayed, and the pattern tends to show an oblique progradation. It is suspected that the
sediment supply was relatively high during the lowstand. Wedges of lowstand deposits are
also observed on some sections (Sects 7.6 & 7.8), and are suspected to be at a similar
stratigraphic position as the above prograding deposits. O n Section 7.6, at about the same
level as the lowstand prograding clinoforms on Section 7.3, some mounded forms are
displayed. The flank of the lower lobe is rather steep (about 1°) compared with the overlying
lobes (about 0.25°).
The origin of such mounded and prograding clinoform stmctures has been discussed by
several authors (e.g. Mitchum, 1985; Posamentier & Erskine, 1991), and could be associated
with submarine fan lobes. These mounded forms according to Berg (1982) can also be
associated with a fluvial-dominated delta system. But the mounded forms on Section 7.5 show
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well developed channel levees and lack evidence of surface marine processes which are
commonly shown by a flat surface (of the delta plain). Therefore, it is interpreted that these
fan systems are parts of the lowstand wedge. Posamentier et al. (1991) indicated that a
channel-levee complex (early lowstand wedge) may be formed prior to the development of the
prograding complex (late lowstand wedge). The formation of the lowstand wedge is
contemporaneous with the deposition of the channel and valley fills (Sects 7.3 & 7.7;
Posamentier & Vail, 1988; Posamentier et al, 1991).
Although the timing and related seismic features have strongly suggested the occurrence
of a shallow submarine fan system, m o d e m analogues are lacking. Walker (1992) indicated
that although turbidity currents can occur at any depth, turbidites are only preserved below
storm wave base at a minimum depth of 250-300 m, which is not the case in the study area.
The Late Pleistocene deep marine fan systems developed in the Gulf of Mexico (Weimer,
1990; 1991) and off the Amazon delta system (Rood etal, 1991) are good examples. In these
areas there is also sufficient slope length to allow turbidity currents to develop. Such slope
length and depth are absent in the study area, and thus the above interpretation has to be tested
further through detailed mapping combined with well coring.
Highstand systems tract
A rapid rise of relative sea level occurred after the deposition of the lowstand systems
tract. The estimated relative sea level rise is difficult to define due to the post-Pliocene
erosion, but the profile of the highstand clinoform on Section 7.3 suggests that it may have
reached 150 msec T W T (about 115 m ) . Since there is no transgressive systems tract developed
(within seismic resolution), the highstand systems tract appears to be directly deposited on the
lowstand systems tract. A very thin condensed section may occur between them, which may
also coincide with the first flooding surface. Most of the Pliocene deposits in the study area
in fact consist of highstand deposits.
On the Karimunjawa Arch, obvious highstand prograding clinoforms are not displayed,
probably due to very low depositional dip (about 0.1°). Here, the early highstand systems tract
is characterised by a medium to high amplitude and high continuity parallel reflection pattern
(Sect. 7.1). Sangree and Widmier (1977) suggested this type of pattern represents interbedded
strata deposited during high and low energy conditions, and commonly associated with shallow
marine clastic units deposited mainly by wave transport processes. A mounded form, which
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interfingers with surrounding marine elastics, is well displayed on Section 7.1 and interpreted
as a highstand bioherm [shelf carbonate according to Tucker & Wright (1990)].
On the Bawean Arch, the early highstand deposit is characterised by relatively steep
oblique progradational clinoforms, with the depositional dip reaching 0.75° (Sect. 7.3). It is
also associated with some probable channel stmctures (Sect. 7.5). These features are restricted
in extent to the inner shelf of the Bawean Arch, and are interpreted to indicate a highstand
prograding delta complex. Chronis et al (1991) documented that Late Quaternary deltaic
progradation in the Gulf of Patras, Greece, was greatest during the slowdown of sea level
change and the period of stable sea level. The fact that the delta system on the Bawean Arch
occurred following a sea level rise suggests a similar mode of delta progradation as in the Gulf
of Patras.
On and to the north of Madura Island, some east-trending anticlines developed during the
Pliocene. During the sea level highstand, these structural highs were flooded allowing
development of highstand carbonate facies, which shows a pronounced mounded form on
Section 7.8. O n the north flank of Madura Island, the highstand limestone is characterised by
a medium continuity and medium amplitude parallel reflection pattern (Sect. 7.10). A
landward projection to Madura Island suggests that it is equivalent to the locally known
Madura Formation reported by Situmorang et al. (1992) and Aziz et al (1993). This formation
is equivalent to the Paciran Formation in the Rembang Zone, East Java, which is the limestone
facies of the M u n d u Formation (see Chapter 3.2.1). O n Madura Island, the Madura Formation
consists of up to 250 m of interbedded reefal and sandy limestone and marl (Situmorang et al,
1992; Aziz et al, 1993). Situmorang et al (1992) reported the occurrence of benthonic
foraminifers Operculina sp., Amphistegina sp. and Gypsina sp. in this formation, which
indicate an inner shelf environment (Murray, 1973; 1991).
The late highstand tract was well developed in the western part of the study area. It
formed in response to a relatively slow sea level fall, and the basin overall experienced a
progressive shoaling. The late highstand systems tract is characterised by prograding
clinoforms in which the clinoform angles gradually decrease to nearly horizontal at the end of
the sequence. The clinoforms progressively downlap on the shelf, lowstand wedge and
possibly basin-floor fan (?). These features are shown on Section 7.3 and in the smaller
window of Section 7.2 where the pattern appears to be divergent. The change from the early
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highstand systems tract appears to be gradual with the early highstand delta being slowly
abandoned.
During deposition of the slope and basin facies of the highstand tract, water depth may
have reached 200 m (estimated from the clinoform profile), resulting in the accumulation of
planktonic foraminiferal-rich deposits. Such deposits occur in some of the petroleum
exploratory wells drilled in the study area, e.g. JS10-1 and JS8-1. The mudstone facies in the
lower Pliocene strata in both wells (between 280 m depth and the top of the Miocene)
represent the basin and slope facies whereas the overlying sandy facies represents the shelf
facies. Most of the planktonic foraminifers in fact occur in the mudstone facies.
b. Subunit JP2
Subunit JP2 appears to be the result of a smaller relative sea level fluctuation than the
previous one. The area affected by this sea level oscillation is limited and appears to be
concentrated in the deep portion (?below contour 210 msec T W T ) of the basin as indicated on
the time structure map for the top of the Pliocene (Fig. 7.8). During the deposition of this unit
a large part of the study area remained exposed. In the western part of the study area, subunit
JP2 is recognised as a thin prograding complex occurring on the erosional surface on top of
subunit JP1 (Sects 7.2 & 7.3). This erosional surface should be correlated with a lowstand of
sea level and the prograding complex with the highstand deposits.
A thick deposit of up to 80 msec TWT (about 60 m) occurs locally in the deep area north
of Madura Island. Section 7.8 shows the occurrence of a lensoidal lowstand unit, which is
interpreted as a lowstand deposit. This deposit displays some diffraction features which at
least indicate a rough surface. It is suspected that the deposit is a coarse clastic which has
been derived from the adjacent highstand deposits (unit JP1, on Madura Island and on the
anticline north of it).
A subsequent relative sea level rise led to the deposition of a highstand systems tract
which is characterised by prograding clinoforms, onlapping onto the basin margins. This
deposit has led to the shoaling of the basin prior to subaerial exposure of the study area at the
end of the Pliocene.
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7.4.3 UnitJQ
Unit JQ is a Quaternary sequence, deposited following the sea level fall which exposed
the whole study area at the end of the Pliocene. Figures 7.4a, 7.4b and 7.4c are mesh
diagrams of the present bathymetry and the present top of Pliocene and Miocene, respectively.
Roughly, these figures suggest that the basin morphology where unit J Q was deposited
appeared to be relatively flat compared with the Pliocene basin. Some local subsidence
occurred, in the western part of the area, probably related to the normal fault movement
bordering the West Florence Deep and Bawean Arch. O n the eastern part, the local subsidence
was related to the east trending folding which had been occurring since probably the Late
Miocene. O n the isochron map of unit JQ, shown in Figure 7.9, the trend of the deposition
follows the basin topography and thick sediment accumulation occurred in locally deep areas.
The seismic characters and sedimentation patterns of unit JQ differ significantly from the
preceding unit JP. They appear to be strongly influenced by extreme and rapid sea level
fluctuations. Such fluctuation during the Quaternary has been demonstrated by Emiliani (1955,
1966) and Shackleton and Opdyke (1973, 1976 & 1977) through the oxygen isotope records
of deep sea cores, which they related with the orbitally-induced fluctuations of global ice
volume. These glacio-eustatic sea level fluctuations are particularly apparent since the
Jaramillo magnetic event (0.97 M a ; Harland et al, 1989) with a relatively constant period.
Donovan and Jones (1979) have calculated that during the period of maximum Pleistocene
glacial advance the sea level was lowered by about 100 m. These studies are consistent with
the results of Bloom et al. (1974) in their work on marine terraces in the Huon Peninsula, N e w
Guinea, and the findings of Mesolella et al. (1969) on marine terraces in Barbados. Chappell
and Shackleton (1986) indicated that in the last glacial maximum some 20 ka ago, sea level
was up to 130 m lower than the present level. More recently, Fairbanks (1989) used cores
drilled from offshore Barbados coral reefs to indicate that the sea level during the last glacial
maximum was about 121 m below present level. H e also found that deglaciation toward the
present level was not monotonic, as indicated by the occurrence of a termination during 11-10
ka (Younger Dryas event).
Present water depth in the study area does not exceed 85 m and the bathymetry is
relatively flat. With respect to basin subsidence, such conditions may have persisted in most
of the study area since the end of the Pliocene, and subjected the area to two extremes in sea
level during the Quaternary. Sea level fluctuations are indicated by the occurrence of a
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repetitive sequence beginning with marine deposits and ending with extensive cut and fill
stmctures, representing the interglacial and glacial periods respectively. In terms of duration,
these sequences m a y be regarded as fourth order sequences.
Similar features have been reported by some authors in other shelf areas and have been
related to Quaternary glacio-eustatic sea level changes. Piper and Perissoratis (1991) have
demonstrated that the Late Quaternary sea level history inferred from seismic stratigraphy in
the North Aegean continental margin (Greece) correlates with the global eustatic sea level
record based on oxygen isotopic curves. In the southwest Louisiana continental shelf, Suter
et al. (1987) recognised six depositional sequences within the Late Quaternary deposits, five
of which are believed to relate to glacio-eustatic fluctuations.
The present seismic study identified nine seismic subunits in the study area, excluding
Holocene deposits. These subunits are characterised mostly by parallel to subparallel reflection
patterns or are reflection free. Each of them ended with channel cut and fill along the upper
part (Sects 7.2 & 7.3) and is interpreted to represent marine deposition and fluvial channelling
respectively. In some areas the thickness of these subunits appears to be similar (Sect. 7.2)
which may indicate constant subsidence rates and periods of sea level fluctuation. Because
these subunits have a similar seismic character and do not represent thick deposits, lateral
correlation is difficult and tends to be speculative. But they can be grouped into five subunits,
JQ1 to JQ5, based on the occurrence of widespread unconformities on top of each group.
These unconformities are commonly associated with rather deep and wide fluvial channelling.
a. Subunits JQ1 and JQ2
These subunits were deposited during the Early Pleistocene, based on their stratigraphic
position overlying Pliocene unit JP. A stratigraphic subdivision between these subunits in the
western part of the study area is rather speculative, but clear differentiation can be made in the
eastern part due to the occurrence of a relatively large sea level fall at the end of subunit JQ1
deposition (Sects 7.4 & 7.8).
Seismic features, such as rapid basinward thinning (Sect. 7.8) and pronounced anticlines
(Sects 7.2,7.3 & 7.4), and the isochron m a p (Fig. 7.10) indicate that their distribution was still
influenced by local development of folds. In the western part, the subunits asymmetrically
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thicken southward, while on the north of Madura Island local thickening occurred in the
synclinal areas between this island and Bawean Island.
Subunit JQ1 can be further subdivided into subunits JQla and JQlb based on the
occurrence of an internal erosional surface (Sects 7.2, 7.3 & 7.8). In the western part, these
subunits lie subhorizontally and onlap on the Pliocene unit JP. The thickness of subunit JQla
reaches 100 msec T W T (about 75 m ) in the deepest portion of the basin, and gradually thins
toward the basin margin. The maximum thickness of subunit JQlb is about 60 msec T W T
(about 45 m ) . The thickness variation is mainly due to local subsidence, post depositional
erosion and a gradual thinning because of the rising of the basin margin. Subunit JQlb onlaps
on subunit JQla on the southern margin, and on subunit JP when subunit JQla wedges out
(Sect. 7.3). The seismic character of these subunits is similar, a subparallel reflection pattern
with medium amplitude and medium continuity which suggests deposition in a shallow marine
environment (Sangree & Widmier, 1977). Subunits JQla and JQlb in the western part may
be regarded as the units responsible for the flatness of this area. The later subunit, JQ2, was
deposited on a flat surface and has an extensive coverage although its thickness is less than
35 msec T W T (26 m ) . Seismically, this subunit is characterised by a similar appearance to
subunits JQla and JQlb, and probably was deposited in a similar environment.
To the north of Madura Island a local deepening occurred (Fig. 7.10 & Sect. 7.8), and
subunits JQla and JQlb are characterised by northward prograding clinoform deposits,
indicating the sediments were derived from Madura Island (Sect. 7.8). The clinoform profile
provides a very rough water depth estimate for the basin centre which is about 135 msec T W T
(100 m ) . Subunit JQ2, in this area, is characterised by an upward loss in acoustic
transparency, almost reflection free at the lower part and gradually becoming subparallel to
hummocky patterns with variable amplitude and low to high continuity at the upper part. This
indicates a shoaling (regression) of the unit before finally being exposed subaerially.
b. Subunits JQ3 and JQ4
Subunits JQ3 and JQ4 can further be subdivided into three (JQ3a, JQ3b and JQ3c) and two
(JQ4a and JQ4b) respectively. These subdivisions can only be carried out in a limited area
where the subsidence and sedimentation rates were relatively high, as indicated by the thick
deposit on the combined isochron map (Figure 7.12).
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Subunits JQ3a and JQ3b are similar in seismic character, showing a medium amplitude,
subparallel reflection pattern which probably represents a shallow marine environment.
Subunit JQ3c is reflection free, indicating most probably homogeneous mudstone. In the
eastern part, these subunits are rather subtly represented, but southward prograding clinoforms
on Section 7.4 and complex channelling on Section 7.8 may be regarded as representative of
these subunits.
Subunit JQ4 is extensively distributed and characterised by an almost reflection free
character suggesting a nearly homogeneous deposit probably of mudstone. In the western part,
subunits JQ4a and JQ4b are very thin to absent which may indicate a low depositional rate
(Sects 7.2 & 7.3). A relatively thick deposit of up to 45 msec (about 34 m ) occurs in the
eastern part. Here, a further subdivision of the subunit is difficult. But parasequences,
showing northward prograding clinoforms, might be identified with subunits JQ4a and JQ4b
(Sects 7.4 & 7.8). Each of the parasequence appears to be associated with a channel notch
which shifted as the parasequence prograded.
c. Subunit JQ5
Subunit JQ5 consists of a single reflection-free sequence of possibly homogeneous
mudstone. The isochron map is shown in Figure 7.14 The maximum thickness is about 30
msec T W T (about 22 m ) with a little variation on the western part of the map. O n some parts
to the north of Madura Island this subunit is too thin to be identified, but locally thick deposits
of up to 25 msec T W T (about 19 m ) occur in a limited area (Fig. 7.14). The seismic sections
from the area near the coasts of East Java and Madura Island suggest that subunit JQ5 thickens
landward (Sects 7.9 & 7.10), and the parasequence shows a northward prograding clinoform.
A relatively thick deposit of up to 50 msec T W T (about 37 m ) occurs at the mouth of a
narrow strait between East Java and Madura Island. It is suspected that a delta system was
present at this location and this strait served as a channel which linked the Java Sea with the
sediment source area (East Java).
The base of subunit JQ5 is characterised by extensive fluvial channelling. Figure 7.13
shows the position of channels with one major channel up to 3 k m wide and 20 msec T W T
deep (about 15 m ) occurring in the northern part of the area. The occurrence of this
channelling can be related to the curve of the oxygen isotope record of deep sea cores (Fig.
7.15) and sea level position of the Huon Peninsula (Fig. 6.31). The base of subunit JQ5
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should be correlated with the sea level minimum during oxygen isotope stage 6. This age is
consistent with the occurrence of similar features found in Madura Strait occurring on top of
subunit Dla at a depth of 140 msec T W T (about 105 m ) .
Subunit JQ5 appears to be the uppermost sequence on most of the seismic sections. Its
upper boundary is rather subtle due to absence of non-marine reworking processes during the
last glacial maximum which ended the deposition of unit JQ5, although the glaciation may
have lowered the sea level to -130 m (Chappell & Shackleton, 1986). Some notches, however,
appear on some seismic sections and are interpreted to represent small scale channels during
the glacial maximum, but their lateral extent is difficult to trace. O n a regional scale, however,
Kuenen (1950) recognised drowned stream courses within the Java Sea, based on echo sounder
data (Fig. 2.7). T w o river systems with dendritic patterns were present, originating from a
topographically high area southwest of Kalimantan, flowing northward to the China Sea and
eastward to the Wallace Deep. With regard to the present bathymetry of the Java Sea, these
courses should be the most recent and should be regarded as the result of the last glacial sea
level lowering.
d. Holocene deposits
The Holocene deposits include sediments deposited after the last glacial maximum (oxygen
isotope stage 2). O n the shelf (Java Sea) these sediments are relatively thin (< 1 m ) , and in
general are not well represented on the seismic sections. They are recognised mainly from the
results of shallow sediment coring (Dewi, 1993). The thickness varies locally, and apparently
depends on the geographical position and the palaeomorphology of previous lowstand surfaces.
In the area between East Java and Bawean Island, the thickness is commonly less than 1 m,
and the sediment is characterised mostly by homogeneous greenish grey silty-sand to sandy-
silt. Near the base the sediment is commonly oolitic and contains foraminifers Ammonia
beccarii and Operculina sp. (Dewi, 1993) which are typical of a lagoonal environment
(Murray, 1973; 1991). Thus, this lower part of the sediment is thought to be due to the early
stage of flooding in the study area after the last glacial maximum.
A rather thick Holocene deposit occurs near the mouth of the Solo River and can be
regarded as a part of the Solo River delta deposit (locality shown in Fig. 7.2). Section 7.9 is
the seismic section (boomer system) taken close to the Solo River Delta system. Here, the
Holocene deposits may reach 5 m, but show a rapid thinning seaward. A detailed study of
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Holocene sedimentation in the Solo River Delta has been carried out by Hoekstra (1993). His
study indicates that the growth pattern of the delta exhibits very little lateral accretion, and the
maximum rate of deposition occurs in front of the main channel which resulted in a
prolongation of the delta. H e also found that most of the Solo River sediments, which are
mosUy mud-dominated, are deposited in the deltaic environment and only small portion escape
to the sea. Hoekstra (1993) postulated that if this condition is similar to all deltaic
environments on the northern coast of Java, most of the recent sediments in the Java Sea are
not fluvially derived from Java. This may explain the lack of Holocene vertical accretion in
the Java Sea, because most of the sediments would be derived from previous deposits or
marine organisms.
7.5 DISCUSSION
The Miocene basin configuration of the southeastern Java Sea is poorly known, and it is
suspected that the basin was still strongly influenced by the northeast-trending stmctures.
These structures are half grabens and have been the major control for the Early Tertiary
sedimentation. O n the northern part of East Java (Rembang Zone) these stmctures were still
the dominant factor in controlling the Middle and Late Miocene sedimentation (Chapter 4).
Although some elements of these stmctures were still active until the Pleistocene, their
effectiveness in controlling the sedimentation during the post-Miocene was diminished. The
Pliocene sedimentation, in general, occurred in east-trending synclinal basins which indicate
the dominance of a northward tectonic compressional stress. This continued until the Early
Pleistocene, as is indicated by some local thickening of the Early Pleistocene deposits. Since
then, further basin development appears to have ceased, and a tectonically stable condition
may have been reached.
Posamentier and Vail (1988) pointed out that depositional stratal patterns and distribution
of lithofacies are controlled primarily by relative sea level change. The above discussion
indicates that basin evolution in the study area, particularly since the Pliocene, is relatively well
understood. The regional basin subsidence occurred at a relatively slow rate. This has led to
the belief that eustatic sea level change was the only major factor in controlling the stratal
pattern and lithofacies distribution in the study area.
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7.6 SUMMARY
The seismic stratigraphic study revealed that the Late Tertiary and Quaternary sediments
in the southeast Java Sea can be subdivided into three major depositional sequences: units JM,
JP and JQ. These units, which represent the Miocene, Pliocene and Quaternary deposits
respectively, are typified by different areal distributions, thicknesses, and seismic character.
Unit JM is poorly known due to the limited penetration of the system used. However,
seismic sections from structurally high areas indicate the presence of two subunits, J M 1 and
JM2 , which are characterised by subparallel reflection patterns of possibly paralic deposits in
the lower part, and mounded and undulating forms of possibly reefal/karstified limestones in
the upper part. The age and lithofacies of this unit resemble the Tawun (paralic deposit) and
Bulu (limestone) Formations developed in the Rembang Zone.
Unit JP attains 220 msec TWT thick (about 165 m) and can be subdivided into subunits
JP1 and JP2, with subunit JP1 being the most extensively developed. In the area where the
clastic deposition was dominant, such as on the south flank of the Bawean Arch, subunit JP1
has developed systems tracts that resemble a shelf margin sequence, with sediments derived
mostly from this arch. Highstand carbonates of this unit were developed on some anticlinal
highs in the eastern part of the area and in the shelfal facies south of the Karimunjawa Arch.
They may represent a low rate of clastic deposition. Subunit JP2 was developed in a limited
area, particularly in the deeper part of the basin, which indicates that the sea level fluctuation
was much lower than during the previous sequence. Based on this difference, a tentative
correlation with the Haq et al. (1988) eustatic sea level curve suggests an age of 3.8 M a for
the boundary between subunits JP1 and JP2.
Unit JQ consists of nine small subunits, excluding the Holocene subunit, which represent
at least nine Quaternary sea level fluctuations. Each subunit is relatively thin, and tends to be
distributed widely because of deposition on a relatively flat lying area. The seismic characters
are very similar, subparallel reflection or almost reflection free patterns representing marine
deposits, topped by extensive fluvial channelling which is obvious in the western part of the
study area. This repetitive succession is thought to represent highstand and lowstand of sea
level respectively.
145
The nine subunits in unit J Q can be grouped into five subunits, JQ1 to JQ5, based on the
occurrence of relatively larger sea level falls as indicated by deeper and wider fluvial incisions.
A tentative correlation with the oxygen isotope records of deep sea cores (Fig. 7.15) can be
made with an assumption that a fluvial incision reflects a significant relative sea level fall
which exposed most of the Java Sea. Because the maximum water depth at present is about
85 m, the magnitude of sea level fall should be larger than this value. The oxygen isotope
records (Harland et al, 1989; Fig. 7.15) suggests that relatively large sea level falls occurred
after the Jaramillo Event, and correspond with oxygen isotope stages 22, 16, 12, 10, 6 and 2.
Subunit JQ5, which underlies the Holocene deposits, is bounded at the top and base by fluvial
erosional surfaces, which are interpreted to correspond with the oxygen isotope stages 2 and
6 respectively (Late Pleistocene deposition). Subunit JQ4 consists of two smaller subunits
which are separated by a fluvial erosional surface. Subunit JQ4 is inferred to have been
deposited during the interval between stages 10 and 6 (late Middle Pleistocene). The subunit
JQ3, which consists of three smaller subunits, is interpreted to have been deposited during the
interval between stages 16 and 10 (middle Middle Pleistocene). The boundaries between
subunits JQ2 and JQ1, and subunit JQ1 and unit JP are difficult to infer due to the absence on
the seismic sections of obvious features related to sea level fall.
146
147
CHAPTER EIGHT
SUMMARY AND CONCLUSIONS
8.1 INTRODUCTION
This study has involved a re-evaluation of the geological history of four aspects of the East
Java Basin including: (1) a reassessment of the onshore sedimentological history of the
Rembang and Kendeng Zones; (2) a reinterpretation of the Quaternary successions in the
eastern Kendeng Zone; (3) a seismic stratigraphic interpretation of the Quaternary succession
in Madura Strait; and (4) a seismic stratigraphic interpretation of the Late Tertiary and
Quaternary successions in the southern Java Sea. The following discussion summarises briefly
the major findings of this study.
8.2 BASIN DEVELOPMENT
During the pre-Tertiary to Early Tertiary, a series of arc and subduction complexes were
systematically developed in the western Indonesian region (Fig. 8.1). The sutures between
these complexes apparently have remained as lines of weakness which provide flexibility
within the basement of the East Java Basin. These zones of pervasive weakness have
influenced the development of the East Java Basin in the Late Tertiary and Quaternary.
Differential basement block faulting and associated uplift led to the development of a
fragmented basin with numerous small sub-basins, each characterised by contrasting
sedimentation histories. Differential block uplifts were probably caused by the oblique
orientation (northeast-trending) of the pre-Middle Tertiary sutures relative to the present
subduction zone (east-trending), which produced wrench faulting accompanied by vertical
movement (Situmorang et al, 1976). Half grabens related to this faulting influenced the
pattern of sedimentation during the Middle Miocene (Tawun Formation) in the Rembang Zone
(Chapter 4) and Pliocene (and probably earlier) sedimentation in the southern Java Sea
(Chapter 6).
Since the Late Miocene, east-trending anticlinal zones developed and were superimposed
on the previous northeast-trending stmctures. The anticlinal Rembang Zone, which extends
from the Blora area to Madura Island, became a dominant structure which controlled
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sedimentation during the Plio-Pleistocene and divided the East Java Basin into two major
basinal areas (Fig. 8.2). In the north, a basinal area formed between the Karimunjawa and
Bawean Arches and the Rembang Zone. In the south, a basinal area developed between the
Rembang Zone and the chain of volcanism along the median line of Java. The east-trending
orientation of these basins is parallel to the Neogene-Quaternary subduction zone in the south
of Java and suggests the dominance of a northward tectonic compressional stress. The smaller
sedimentary thickness and the wider dimension of the north basinal area suggest that the
compressional stress here is less than in the south basinal area.
The peak development of these basins apparently occurred in the Late Miocene and
Pliocene. Sedimentation patterns, such as local thickening in the north basinal area (southern
Java Sea), suggest that basin development ceased in the Early/Middle Pleistocene. Since then,
the basin has become flatter due to infilling, promoting thin and widespread deposition of
Middle and Late Pleistocene sediments. In contrast, rapid basin subsidence and sedimentation
occurred in the south basinal area at least until the late Middle Pleistocene, when extensive
east-trending folding, accompanied by diapiric movement, gave rise to the anticlinal Kendeng
Zone.
8.3 SEDIMENTATION HISTORY
8.3.1 The onshore area during the Middle Miocene and Pliocene
A re-assessment of the Middle Miocene to Pliocene sedimentary history of the Rembang
Zone was carried out using seismic stratigraphic interpretation, while outcrop data were used
for the Kendeng Zone. The three main stratigraphic markers used to achieve a reliable lateral
correlation between these zones are: (1) the determined sequence boundaries of the Rembang
Zone strata, (2) the results of biostratigraphic analysis of previous workers (e.g.
Pringgoprawiro, 1983; van Gorsel & Troelstra, 1981), and (3) the results of the
climatostratigraphic study of van Gorsel and Troelstra (1981) on the Late Miocene and
Pliocene strata in the Kendeng Zone. These were backed up by sedimentological data such
as vertical facies sequences.
Middle Miocene strata (seismic unit MM2) in the Rembang Zone, as recognised from
seismic sections (Chapter 4), were deposited in two small basinal areas (Fig. 4.10). These
basins are half grabens that resulted from basement fragmentation. Stratal patterns revealed
that the direction of sediment transport was mainly from the north in the southern part of the
149
East Bawean Trough, and from both the south and north in the southern part of the Central
Deep (Fig. 4.10). These patterns of sedimentation were not recognised by earlier investigators
(e.g. Lemigas/BEICIP, 1969; Ardhana, 1993) who assumed that most sediments were derived
from the north (Bawean Arch).
Late Miocene sediments (Wonocolo and Ledok Formations) in the Rembang Zone were
mostly deposited in a shelfal area during sea level highstands. The seismic unit L M 1 , which
corresponds with the Wonocolo Formation, is mostly hemipelagic foraminiferal-marl, but
deltaic deposition probably occurred near the Tobo high area (Fig. 4.10). The Ledok
Formation (seismic unit L M 2 ) was deposited on a shallow shelf as a tidally-influenced deposit.
A geological reinterpretation of the Kendeng Zone was carried out based on examination
of outcrops of the Late Miocene and Pliocene strata. Here, sedimentation is the result of
volcanic activity along the southern Kendeng Zone with marine deposition occurring to the
north of the volcanic complex. The Kerek Formation, a turbiditic deposit, was formed in this
setting. The influence of volcanism diminished by the late Late Miocene, and the resultant
sediments are dominated by pelagic marls mapped as the Kalibeng Formation.
The deposition of pelagic marl became widespread in the Pliocene, and apparently
followed a global eustatic sea level rise recognised by Haq et al. (1988). In the Rembang and
Kendeng Zones, the marl is represented by the Mundu (seismic unit PL1) and Kalibeng
Formations, respectively. The Atasangin Formation, which was deposited on top of the
Kalibeng Formation, appears to result from a minor relative sea level fall in the Middle/Late
Pliocene.
A relative sea level fall occurred in the Late Pliocene and led to the development of a
stagnant lacustrine basin in a broad synclinal area south of the Rembang Zone (Fig. 8.2). This
resulted in the widespread deposition of organic-rich mudstone mapped as the Lidah
Formation.
8.3.2 The onshore area (eastern Kendeng Zone) and Madura Strait during the
Quaternary
A reinterpretation of the onshore Quaternary succession has been carried out, particularly
in the eastern Kendeng Zone, to facilitate correlation with the offshore (Madura Strait)
150
Quaternary succession interpreted from seismic data. The eastern Kendeng Zone - Madura
Strait area is characterised by syndepositional folding which formed a small east-trending
basin, located between the Rembang Zone (and Madura Island) in the north and Quaternary
volcanoes in the south. The basin is characterised by rapid subsidence and sedimentation,
which resulted in more than 200 m of Quaternary deposits.
In the eastern Kendeng Zone - Madura Strait area, deposition resulted from a complex
interplay between Quaternary volcanism along the south of the Kendeng Zone, marine
sediments derived from the exposed Rembang Zone in the north, and Quaternary relative sea
level changes. Field observations indicate that the volcaniclastic deposits are coarse-grained
and, in the seismic section from Madura Strait, they are associated with relatively steep
depositional slopes (of alluvial fan/fan delta systems). In contrast, the northerly-derived
deposits are associated with marine fossiliferous mudstone containing a small number of quartz
grains. The present detailed seismic stratigraphic study of Madura Strait indicates that the
timing of deposition of these Quaternary deposits was strongly controlled by relative sea level
changes. During a major relative sea level fall, the eastern Kendeng Zone and Madura Strait
were subaerially exposed, as indicated by thin widespread fluvial and alluvial fan deposits.
One of the major considerations of previous Quaternary geological investigations in the
eastern Kendeng Zone was the dating of hominid bearing strata (Jacob & Curtis, 1971; Hyodo
et al., 1992). The fact that vertebrate and hominid fossils occur associated with fluvial/alluvial
fan deposits (as reported by Duyfjes, 1938a,b,c,d) suggests that Quaternary sea level
fluctuations were an important factor in determining the habitat of human progenitors. This
emphasises the importance of delineating the timing of sea level fluctuation in this area. The
present study used the published Quaternary oxygen isotope curve of deep sea cores of Harland
et al (1989) to infer the depositional timing of these fluvial deposits. T w o of the three main
Quaternary fluvial horizons in Madura Strait have been correlated with two major eustatic sea
falls during the oxygen isotope stages 16 (seismic unit Cla) and 6 (seismic unit Dla)
respectively.
8.3.3 Pliocene and Quaternary sedimentation in the southern Java Sea
The Pliocene to Quaternary history of sedimentation in the southern part of the Java Sea
has been assessed through seismic stratigraphic interpretation. The limited penetration of the
seismic system used in this study did not permit a detailed evaluation of the Miocene
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sedimentation in this area. Stratal patterns, as observed on structurally high areas, such as the
Karimunjawa and Bawean Arches, indicate the occurrence of a parallel stratal pattern of
probably shallow-marine deposits associated with carbonate buildups of mounded form.
Outcrop data from Bawean Island suggest that the Miocene deposits are mainly paralic facies.
Basin development and sedimentation patterns, particularly during the Pliocene and
Quaternary, are well understood. Here, a broad asymmetrical synclinal basin developed
probably in the Late Miocene to Pliocene between the relatively stable Sunda Shelf in the north
and the anticlinal Rembang Zone in the south. During the Pliocene, sediments were primarily
derived from the stable Sunda Shelf (Karimunjawa and Bawean Arches) and the Rembang
Zone (Madura Island). The stratal patterns are mainly controlled by relative sea level changes.
For example, stratal patterns in the major Pliocene deposit (unit JP1) suggest that deposition
occurred during a slow relative sea level fall.
Pleistocene sedimentation was highly influenced by frequent sea level fluctuations. In the
southern Java Sea (north basinal area; Fig. 8.2), basin filling during the Pliocene had produced
a flat basin topography by the end of the Pliocene. Nine thin seismic subunits are recognised
in this area and are believed to represent at least nine Quaternary sea level fluctuations. Each
subunit is relatively thin, and tends to be distributed widely because of deposition on a
relatively flat-lying area. The seismic characters of these subunits are very similar; subparallel
reflection or almost reflection free patterns representing marine deposits, topped by extensive
fluvial channelling. This repetitive succession is thought to represent highstands and lowstands
of sea level, respectively, which can be related to published sea level/oxygen isotope curves
(e.g. Harland et al, 1989).
8.4 NOTES ON THE APPLICATION OF SEQUENCE STRATIGRAPHY IN THE
EAST JAVA BASIN
Sequence stratigraphy was primarily developed for basins that undergo slow and steady
differential subsidence which increases toward the basin centre. Such characteristics are
commonly displayed by shelf margin settings and divergent continental margin settings (Jervey,
1988). The widespread application of sequence stratigraphic concepts is c o m m o n in these
settings (e.g. Suter et al, 1987; Piper & Perissoratis, 1991).
152
The present study of the East Java Basin indicates that at least three types of sub-basins
occur in this area. They include (1) sub-basins showing the character of a half graben setting;
(2) sub-basins showing the character of a shelf margin setting; and (3) sub-basins showing the
character of a shelf setting. Sequence stratigraphic concepts have been applied to these basin
settings, but some remarks need to be raised related to the applicability of the concepts in these
basins.
A half graben basin setting formed during the Middle Miocene in the southern part of the
East Bawean Trough (Chapter 4.3.2a). This basin is characterised by a greater subsidence rate
toward the basin margin (Bawean Arch; Fig. 4.10). The sediments were derived mostly from
the adjacent uplifted footwall. The absence of erosional features in this basin suggests that
relative sea level falls have never exceeded the rate of subsidence. A s a result, the sequence
boundaries are difficult to place, which further emphasises the concerns of Swift et al (1987)
who suggested that sequence stratigraphic concepts are not applicable in basins characterised
by a faster subsidence rate landward, compared to rates of eustatic sea level fall. The present
study indicates that although sequence boundaries are subtle, m a x i m u m flooding surfaces
resulting from retreating stratal patterns related to transgression are observable. These surfaces
provide an important basis to correlate stratigraphic units between sub-basins.
Sub-basins showing the character of shelf margin settings include Madura Strait during the
Quaternary and the Java Sea during the Pliocene. These basins are characterised by a greater
subsidence rate toward the basin centre. The present study revealed that the development of
systems tracts within this basin setting requires further research. In many cases, transgressive
systems tracts are missing, and sequences are only represented by lowstand and highstand
systems tracts. The incomplete development of systems tracts appears to depend on (1)
palaeotopography and (2) order and magnitude of sea level change. Most parts of the basins
in Madura Strait and the Java Sea are characterised by a relatively low morphological gradient.
During a relative sea level rise, a rapid retreat (landward) of the depocentre results in sediment
starvation in the basin centre and margin (slope). This retreat appears to increase if the order
and magnitude of sea level change increases. More obviously, incomplete development of
systems tracts occurs in the shelf setting, such as the Java Sea. Where the basin morphology
is relatively flat and sea level fluctuations were rapid and of high magnitude as in the
Quaternary, this effect is particularly pronounced. The sedimentary sequences are thin and
extensive. The lowstand systems tract is only represented by fluvial channelling which was
153
later abandoned during a relative sea level rise, and the highstand systems tract is represented
by a thin marine unit.
8.5 CONCLUSION
The existence of previously uninterpreted seismic data in Madura Strait and the Java Sea,
combined with onshore seismic and outcrop data has provided a unique opportunity to study
the development of the East Java Basin. The extent of the data and the application of seismic
sequence stratigraphic concepts have allowed a comprehensive study of the timing of
deposition and deformation of the basin fill; the mode of deposition and sedimentation patterns
within the sub-basins; and the role of sea level changes on stratal patterns.
154
155
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FIGURES TO CHAPTER ONE
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FIGURES TO CHAPTER TWO
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AGE MA
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Figure 3.8. Ranges of diatoms contained in the Sonde and Lidah Formations (modified from Ninkovich & Burckle, 1978).
Figure 3.9. Lithostratigraphic correlation between the Rembang and Kendeng Zones. The Puncakwangi and Kali Sampurna measured sections are from Hasjim (1987), and the Solo River section was measured during the present study.
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Globigerina
Orbitoid Limestone
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Kujung
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Tambakromo
Mundu
Ledok
Wonocolo
Ngrayong Sand
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Lower O K
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Middle Kujung
Lower Kujung
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Selorejo Beds
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Grovergroene
Platen Complex
' ;
Prupuh Limestone
Table 3.1. Stratigraphic subdivision of the Rembang Zone according to Trooster (1937).
AGE
PLEISTOCETJ
PLIOCENE
MIOCENE
LATE
MIDDLE
EARLY
LATE
EARLY&
MIDDLE
LATE MIDDLE
LOWER
KENDENG ZONE Notopuro Breccias
Kabuh Beds
Pucangan Beds
Balanus Limestone
Sonde Marls
Klitik Limestone
Upper Kalibeng Beds
Lower Kalibeng Beds
Banyak Beds
Upper Kerek Beds
Lower Kerek Beds
Pelanu Beds
REMBANG ZONE
High Terrace accumulation
Blue Clays with marls
and limestones (Malo)
Globigerina Marls
(Mundu)
Ledok Beds
Wonocolo Beds
Lasem &.
Butak volcanic
Uplift
Karren
Limestones
y; Hia&tjs ;.
Rembang Beds
Table 3.2. Stratigraphic subdivision of the Kendeng and Rembang Zones according to van Bemmelen (1949).
AGE Blow Zone KENDENG ZONE REMBANG ZONE PLEISTOCENE N23
N22
Kabuh, Notopuro
Pucangan Formations
Dander Member
N21 Sonde Formation
PLIOCENE N20 Klitik Member
N19
MIOCENE
Atasangin Member
LATE
MIDDLE
EARLY
OLIGOCENE
N18
N17 Kalibeng Formation
N16 Banvak Formation
N15
N14
N13
N12
Nil
Kerek Formation
N10
N9
N8
N7
N6
N5
N4
N3
N2
Nl
PI9
Pelang Formation
• " • ' : , :
Malo Member
Lidah Formation
Mundu Formation
Ledok Formation
Wonocolo Formation
Paciran Formation
BifttttS :
Bulu Formation
Ngrayong Member
Tawun Formation
Hiatus
Tuban Formation
Prupuh Formation
Kujung Formation
Table 3.3. Stratigraphic subdivision of the Kendeng and Rembang Zones according to Pringgoprawiro (1983).
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The bed! above 101 metres are possibly referable to Zone N.2I but no direct evidence. The beds ore considered as beinq prior to the advent of
Globorotalia (G) truncatulinaidas (*,/)
£--Beds referred by Bolli (l96o) to his ? Equivalent of
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<&*• Gfo&oquudrino a/t/sp/ra e/t/spfro/G/o6orata//a crassafarm's
Zone,of Bolli and1 Bermudez, I9b5.
~ HIATUS •^•"Globorotalia dutBrtroi/GJobigerinoides Qbl'iQUUS vxtramus Zone
f Boomqarf (.1949) fiqured a specimen referable to < Globorotalia (ff.) multicamerata from 5 4 0 metres depth, (^therefore Zone N. 17/ Zone N. lb -boundary below this depth.
-First appearance of Giobororo/io (7T) ocostoonsis (V.s)
Last occurrence of Globorotalia (T) s/akqns/s
> Numerous small diaitems in Kawenqan Well No. 9
Last occurrence of Cassigarineiio c/it'poiansis
,213 —Last occurrence of Globigerinoides subguadrotus
-i.aoo 1,842
-t,854-4
1,914-2
-Specimens of G/oborotal/o(G^) fons/ forma typtca observed at 1,800m.
<~ HIATUS^ Zones N.9(parf^N.10 and NJifparr) misiina.
•Evolutionary first appearance of Orbulina sp.
-2,000 •TERMINATION DEPTH, 2,025 merres
n?fUrnQ4^i B^Slratrgff£hi
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Greenish-grey marl, rich in planktonic and benthonic foraminifers, interbeds of limestone and sandstone of up to 10 m thick.
Bioclastic limestone, rich in benthonic larger foraminifers and shell fragments, interbeds of glauconitic, carbonaceous quartz sandstone.
Interbedded quartz sandstone (loose, medium-grained, well sorted), bioclastic limestone and carbonaceous mudstone. Benthonic foraminifers are common.
Grey-dark grey carbonaeous mudstone, sandy, interbeds of limestone. Planktonic foraminifers are rare to absent.
Light grey-grey carbonaceous mudstone, interbeds of limestone (up to 20 m) and glauconitic quartz sandstone (up to 15 m). Planktonic foraminifers are common in the lower part.
INFERRED
ENVIRONMENT
OF DEPOSITION
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Shallow marine
Nearshore-
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—~——-
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Figure 4.5. Stratigraphic column of the BNG-1 well (Banyubang Anticline), summarised from well log and palaeontological examination reports.
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Figure 4.23. X-ray diffractogram of the Wonocolo Formation. The sample (sample no. Wl) is from the Wonocolo area, some 20 km northeast Cepu (Fig.4.1). Ca=calcite, Ar=aragonite, Qz=quartz, Mo=smectite, Ko=kaolinite.
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LITHOFACIES ENVIRONMENT
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KBJ6 Glauconitic calcarenite, planar cross-rr s-^-^Z'-^ bedded, composed mostly of foraminiferal
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Calcareous fine to medium-grained sandstone, trough cross-bedded.
Inner neritic (?)
Outer neritic
Marl, unstratified and homogenous, few interbeds of calcarenite (40 cm to 1 m).
Sandy marl, bioturbated, with thin beds (15-20 cm) of planar cross-bedded sandy calcarenites with erosional base. Calcirudite, composed of benthonic larger foraminifers, coral and volcanic rock fragments. Calcareous sandy mudstone, some beds of medium-grained, ripple-laminated sandstone. Parallel-laminated calcareous mudstone interbedded with parallel- and ripple-bedded, medium-grained sandstones, 5-20 cm bed thickness. _.
Inner neritic
Figure 5.3. Generalised vertical section of the uppermost Kerek Formation and the Kalibeng Formation, measured along the Solo River north of Ngawi.
Figure 5.4. An outcrop of the turbiditic deposits of the Kerek Formation (upper part) in the Solo River section; intercalations of calcareous sandstone and sandy mudstone.
Figure 5.5. Trough cross-bedded calcareous sandstone in the middle part of the Kalibeng Formation exposed in the Solo River section.
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I I I I I I
NORTH KABUH SECTION
LITHOFACIES ENVIRONMENT
o
400-~^~-"4-
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NKabuhlB
NKabuhl7 Marginal marine
Marl, siliceous, very well bedded, thin interbeds (15-20 cm) very fine sandstone occurred rarely.
.;- NKafauhl6
x_7*~ NKabuM5
?Open marine-influenced
_^_-r_+" NKabuhH Marl, unstratified, rich in foraminifers,
NKabuhl 4 2 conchoidal fractures.
200—i-i-:
• Nj\abuhl3
Marginal marine
Marl, siliceous, very well b e d d e d .
100-
KALIBENG FORMATION
• NKabuhl2
See Figure 5.15
lx_____xP NKabuh3
Figure 5.14. Generalised vertical section of the Atasangin Formation, measured on the area north of Kabuh (see Fig. 5.1 for the location).
UNIT
SANDSTONE FACIES, NORTH KABUH SECTION
LITHOFACIES ENVIRONMENT C! PS CS ! Sh I MS I Co
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Marl, siliceous, very well bedded.
Very coarse-grained sandstone and breccias, crudely bedded, composed of volcanic rock fragments.
Stacked fining upward units. Parallel-laminated, coarse-grained sandstone grading into parallel-laminated mudstone.
NKabuhlO
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40-
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Mudstone, interbedded with fine to coarse-grained, parallel-laminated sandstone.
Channel-fill sandstone, medium to coarse-grained, ?planar cross-bedded.
3 — iVKabuh7
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Parallel-laminated sandy marls interbedded with parallel and ripple-laminated, fine-grained sandstone, common occurrence of deformation structures.
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KALIBENG FORMATION
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Interbedded sandy mudstone and ripple-laminated sandstone, common occurrence of burrow structures and pebble-size mud clasts.
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Fine-grained sandstone, well sorted, glauconitic, structure unobserved.
Intercalation (5-15 cm) parallel-laminated calcareous sandstone and foraminiferal-rich marl.
Marl, unstratified, rich in foraminifers.
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Figure 5.19. X-ray diffractograms of the samples from A: lower (NKabuhl3) and B: upper (NKabuhl8) parts of the siliceous marl Atasangin Formation (see Fig. 5.14 for stratigraphic position). Mo=smectite, Ko=kaolinite, Cl=chlorite, Qz=quartz, Ca=calcite, Ab=albite, Si=siderite.
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UNIT
SOLO RIVER SECTION, NGAWI SH MS Co
«|«|«! LITHOFACIES ENVIRONMENT
G
U D to
Matrix supported conglomerate, non-oriented fragments. Ripple and planar cross - stratified sandstone, interbedded with massive mudstone. Planar cross-bedded conglomerate and coarse-grained sandstone.
Debris flow
Fluvial
200-
r. &r-::
iLx^.v.-.C.vfS?1
4xJX^_XX^ȣ\i
Debris flow Matrix supported conglomerate and breccia, pebble to boulder-size fragments of volcanic rocks, non-oriented fragments.
Wave-dominated Calcareous sandy mudstone, massive, contains foraminifers and pebble-size mud clasts.
Bedded calcarenite at the base, grading into reefal and calcirudite upward.
Inner neritic
Figure 5.22. Generalised vertical section of the Sonde Formation and the lower sequence of the Pucangan Formation in the Solo River, north of Ngawi (see Fig. 5.1 for the location).
Figure 5.23. An exposure of the bedded limestone Klitik Member of the Sonde Formation (Klitik Limestones, van Bemmelen, 1949) in the Solo River section.
Figure 5.24. A closer view of the bedded limestone Klitik Member in Figure 5.23.
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SUMBERRINGIN SECTION, JOMBANG
UNIT FS CS
' I ' I LITHOFACIES ENVIRONMENT
PUCANGAN FORMATION
400
300-— SRG12
o ~ VI C
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200
100
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ATASANGIN FORMATIONO
SRG9
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ParaUel-laminated, poorly sorted sandstone interbedded with sandy mudstone, abundant comminuted marine molluscan shells.
Subtidal
Bluish grey sandy mudstone, unstratified, common organic streaks. Lacustrine
Anoxic
See Figure 526
Bluish grey mudstone, unstratified, homogenous, common organic streaks.
Ripple-laminated calcareous fine sandstone (1 m thick).
Bluish grey mudstone, unstratified, homogenous, common organic streaks.
Lacustrine Anoxic
Fossiliferous, sandy_ mudstone intercalated with muddy sandstone, structure unobserved.
Mudstone, dark grey, unstratified, homogenous.
Marl, unstratified, fossiliferous, thin interbeds of medium-grained calcarenite
Subtidal
Figure 5.25. Generalised vertical section of the Lidah Formation in Sumberringin area, north of Jombang (see Fig. 5.1 for the location).
UNIT
SUMBERRINGIN SECTION, JOMBANG
Sh MS Co
?l?l?l LITHOFACIES ENVIRONMENT
O co
< s ix <
cx to
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Bluish grey, mudstone, unstratified.
Medium to coarse-grained sandstone, conglomeratic at base, fining upward into fine-grained sandstone, common planar cross-and ripple-stratifications.
Muddy sandstone fining upward into carbonaceous mudstone, unstratified, abundant clay and organic pellets.
Interbedded between massive, poorly sorted muddy sandstone and medium-sorted, ripple-laminated, medium-grained sandstone. Both contain abundant benthonic foraminifers.
Bluish grey mudstone, unstratified, homogenous, common organic streaks.
Lacustrine
Tidal/ fluvial channel
Mud flat
Subtidal
Lacustrine Anoxic
J Figure 5.26. Vertical section of the upper sandstone facies Lidah Formation in the Sumberringin area (see Fig. 5.1 for the location).
UNIT
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400"
-
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300-
200-
-
100-
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NGLAMPIN SECTION, CEPU Sh MS Co
| ?\ f| LITHOFACIES
•J =i—1,
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Light-grey calcareous mudstone, ,—. unstratified, carbonaceous and 7—' '
bioturbated in some parts, some interbeds of calcarenite (< 60 cm).
=jj4rNGLAMPINi Calcarenite interbedded with calcareous E 3 sandy mudstone, glauconitic, rich in =3 comminuted shell fragments.
-NGLAMPIN 7
Bluish grey mudstone, unstratified, carbonaceous, streaks of fine-grained quartz and pyrite, some thin interbeds (< 10 cm) of fine-grained, ripple-laminated, calcareous sandstones.
ENVIRONMENT
Tidal flat
Shoreface/ foreshore(?)
Lacustrine Anoxic
Figure 5.27. Generalised vertical section of the Lidah Formation in Nglampin area, some 30 km southeast Cepu (after Nahrowi et al. 1981; see Fig. 5.1 for the location).
3000-
= 1500-
3000-
2000 ->+
10X-
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~ m~m »_- ^~
•••••••••••••••I
60 TO
Figure 5.28. X-ray diffractograms of the samples from (a) lower (Nglampinl) and (b) middle (Nglampin7) parts of the Lidah Formation on the Nglampin area (see Fig. 5.27 for stratigraphic positions). Mo=smectite, It=illite, Ko=kaolinite, Qz=quartz, Si=siderite, Ca=calcite, Gp=gypsum.
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Figure 5.30. Geologic map of the Jombang-Mojokerto area (modified from Duyfjes, 1938b,d).
Low tide level
Figure 5.31. Depositional environments of a tidally-influenced area (after Dalrymple, 1992).
Figure 5.32. Various subenvironments in a barrier-island system (after Reinson, 1984).
UNIT Sh MS Co
Cl IFSl CSI
I I I I 11
KABUH SECTION, JOMBANG
LITHOFACIES ENVIRONMENT
150-
I00-'
50--
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JKBa
Conglomerate, massive, cobble-size fragment of andesitic rocks, normally graded into planar cross-bedded, coarse-grained sandstone, contains pebble size mud clasts, shell fragments of marine moUusc, and thin beds of calcareous mudstone. Tuffaceous mudstone
Conglomerate, massive to crudely bedded, cobble-size fragment of andesitic rocks, normally graded into planar-cross-bedded, coarse-grained sandstone, contains pebble-size mud clasts.
Parallel-laminated bioturbated.
mudstone, strongly
Matrix-supported conglomerate, cobble-size fragments of andesitic rocks and mud clasts, normally graded into tuffaceous, poorly-sorted sandstone, contains foraminifers.
Sandstone, medium- to coarse-grained, moderately sorted, contains plant remains, various primary structures (large scale planar, simple and trough. cross- stratifications) with well developed mud flaser and scouring.
Sandy mudstone, greyish green, rich in benthonic foraminifers.
Matrix-supported conglomerate, inverse to normally graded, pebble to boulder-size fragment of volcanic rocks and tuffaceous mudstone clasts.
Sandstone, planar cross-stratification, interbedded with bioturbated muddy-sandstone. Both contain abundant comminuted sheUof marine molluscs and foraminiferal tests.
Reworked alluvial (?)
(L6)
Tidal flat
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(L4)
Tidal flat (L3b)
Offshore facies (L3a)
Debris flow (L2)
Subtidal (LI)
Figure 5.33. Vertical section of the lower sequence of the Pucangan Formation to the north of Kabuh (see Fig. 5.30. for the location).
UNIT
KEMLAGI SECTION,MOJOKERTO
LITHOFACIES Sh MS Co
Cl I FSl CS I ENVIRONMENT
Coarse-grained sandstone, conglomeratic, structure unobserved.
Mudstone, grey, massive, bioturbated, rich in marine mollusc.
Fine-grained sandstone interbedded with mudstone, paraUel-laminated, greyish blue, highly carbonaceous.
Coarse-grained sandstone, medium sorted, conglomeratic, planar and trough cross-bedding structures, contains shells of marine mollusc and thin interbeds of lapilli tuff.
Interbedded fine-grained sandstone and mudstone, paraUel-laminated, greyish blue, highly carbonaceous.
Sandy mudstone, light grey, massive, strongly bioturbated, rich in marine mollusc fragments and foraminifers.
Mudstone, dark blue, massive and highly carbonaceous.
Interbedded fine-grained sandstoneand mudstone, paraUel-laminated, greyish blue, highly carbonaceous.
Alluvial fan ? __IU2i
Upper sequence
&
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(L6)
Lower sequence (L1-L5)
Offshore facies
Lacustrine Anoxic
Figure 5.34. Generalised vertical section of the Pucangan Formation in the area north of Kemlagi (see Fig. 5.30. for the location).
Figure 5.35. Debris flow deposits of the Pucangan Formation exposed along the Solo River section.
Figure 5.36. Similar to Fig. 5.35. This deposit has a non-erosive contact with the underlying mudstone.
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?l7l?l
PERNING SECTION, MOJOKERTO
LITHOFACIES ENVIRONMENT
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Conglomeratic sandstone, paraUel bedded and epsilon cross-bedded.
Sandstone, poorly sorted, medium to coarse-grained, conglomeratic.
Mudstone, grey, massive, bioturbated, contains marine moUusc fragments.
Sandy mudstone, very poorly sorted, highly bioturbated, rich in marine molluscs, large-scale burrows of Gastrochaenolites.
Stacked fining upward units, crudely-bedded to planar cross-bedded conglomerate, medium to coarse-grained sandstones. Each sequence separated by sandy mudstone, contains vertebrate and hominid fossils.
Tuffaceous mudstone.parallel-laminated, scoured by small-scale (15-30 cm) channels and filled by marine-fossiliferous sandstone.
Sandstone, medium-grained, medium-sorted, trough cross-bedding.
Sandstones, planar cross-bedded, medium-sorted, medium- to coarsegrained.
Sandstone, medium-grained, ripple-laminated with well developed parallel arid wavy-bedded bundled mudstone.
Sandstone, medium-grained, medium-sorted, epsilon cross-bedding, with well developed mud drapes.
Ripple-lammatedfine-grainedsandstone with interbeds of reverse and normally-graded conglomeratic sandstone.
Mudstone, parallel-laminated, interbedded with parallel to ripple-laminated fine-grained sandstone, abundant plant remains and leaf imprints between laminae.
Sandstone, medium-sorted, medium-grained, glauconitic, contains benthonic foraminifers, abundant comminuted shell of marine mollusc, and thin interbeds of carbonaceous mudstone. Coarse-grained sandstone, conglomeratic, structure unobserved.
Fluvial channel ? (U2)
Lower shoreface (UI)
Alluvial fan (M3)
Tidal flat (M2)
Subtidal (Ml)
(L6)
Figure 5.37. Vertical section of the middle sequence of the Pucangan Formation on the north of Perning (see Fig. 5.30 for the location). Palaeomagnetic dates are from Hyodo et al. (1992).
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Figure 5.43. Fine intercalation of sandstone and mudstone overlies a tidal channel deposit in the middle part of unit M2 (Perning section).
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BANYUURIP SECTION, MOJOKERTO
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Medium- to coarse-grained sandstone, planar cross-bedded.
Medium to coarse-grained sandstone intercalated with tuffaceous mudstone.
Tuffaceous sandstone.
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Conglomeratic sandstone, structure unobserved.
Medium- to coarse-grained sandstone, well sorted, planar cross-bedded.
Parallel and ripple-laminated sandstone intercalated with mudstone. Sandstone coarsens and thickens upward.
Mudstone, greenish grey, unstratified, contains tuffaceous mudstone clasts.
Interbedded poorly sorted sandstone and mudstone, strongly bioturbated, contains abundant comminuted shell and foraminifers. Sandstone, medium to coarse-grained, conglomeratic at base, planar crossbedding. Tuffaceous mudstone, parallel and wavy-laminated, carbonaceous, leaf imprints between laminae.
Tuffaceous sandy mudstone, poorly sorted, paraUel-laminated, highly bioturbated, large Gastrochaenolites burrows.
ParaUel-laminated sandy mudstone interbedded with paraUel and ripple-laminated sandstone, common plant remains and leaf imprints.
Beach ?
Shoreface
Shelf mud
Tidal flat (U2d)
Fluvial channel (U2c)
Tidal flat (U2b)
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Foreshore UI
Figure 5.44. Vertical section of the upper sequence Pucangan Formation and the lower part of the Kabuh Formation in Banyuurip area (see Fig. 5.30 for the location).
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UNIT Sh MS Ca
?l7l?l LITHOFACIES ENVIRONMENT
Foreshore Thickening and coarsening upward sandstone, planar cross-bedded, interbedded with paraUel-laminated mudstone (5 cm).
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-&• Sandstone, medium-sorted, trough y\Am.\ rhannp1(TI2d\ ^^ cross-bedding, contains comminuted ^Ciai CnanneU UZQ J
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Sandstone, medium to coarse-grained, conglomeratic at base, planar and trough cross-bedding.
Mudstone, intercalated with paraUel-laminated very fine-grained sandstone, leaf imprints between laminae. Coarse-grained sandstone, no structure, contains mud clasts (< 8 cm).
Sandstone, medium to coarse-grained, planar cross-bedding.
ParaUel-laminated tuffaceous mudstone interbedded with ripple-laminated sandstone.
Fluvial channel (U2c)
Tidal/wave influenced (U2b)
Fluvial channel (U2a)
Upper shoreface (Ulb)
Tuffaceous mudstone, intercalated with paraUel-laminated very fine-grained sandstone, 0.5-15 cm bed thickness. Lower shoreface ?
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Figure 5.48. Vertical section of the upper sequence Pucangan Formation and the lower part of the Kabuh Formation on the south of Kedamean (see Fig. 5.30 for the location).
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UNIT Sh MS Co
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Muddy sandstone, paraUel-laminated.
Medium to coarse-grained sandstone, calcareous, structure unobserved.
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Ripple-laminated calcareous sandstone interbedded with calcareous mudstone
Conglomeratic sandstone, massive.
Tuffaceous sandstone, medium- to coarse-grained, conglomeratic at base, trough and plannar cross-bedding.
Calcareous mudstone.
Fluvial
Shelf mud
Beach (?) (U2d)
Fluvial (?) (U2c)
Beach ? (U2b)
Fluvial channel (U2a)
Lower shoreface (UI)
Figure 5.50. Vertical section of the upper sequence Pucangan Formation and the lower part of the Kabuh Formation in Raci Anticline (see Fig. 5.1. for the location).
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1. Globigerinoides inns tuns 2. Globigerinoides trilobus 3. Globigerinoides sacculifer 4. Hastigeriira aequilateralis 5. Globorotalia acostaensis 6. Globorotalia tusida
Figure 6.8. Ranges of planktonic foraminiferal species from the Pasean Formation in the Kamal area (sample no. Madura-8).
1. Globigerinoides imaturus 2. Globigerinoides trilobus 3. Globigerinoides sacculifer 4. Globigerinoides ruber
Figure 6.9. Ranges of planktonic foraminiferal species from the Pasean Formation in the area south of Pamekasan (sample no. Pamekasan-8).
Blow (1969) zone I Species 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23
1. Globigerinoides imaturus 2. Orbulina universa 3. Globigerinoides elongatus 4. Globigerinoides ruber 5. Pulleniatina obliquiloculata
1. Globigerinoides imaturus 2. Globigerinoides trilobus 3. Globigerinoides sacculifer 4. Orbulina universa 5. Globigerinoides elongatus 6. Globorotalia dutertrei 1. Globorotalia tueida 8. Globigerinoides ruber 9. Sphaeroidinella dehiscens
10. Pulleniatina obliquiloculata
Figure 6.10. Ranges of planktonic foraminiferal species from the Pasean Formation in the Gili Genting Island (sample no. GG-3 & GG-7)
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Figure 6.11. X-ray diffractograms of samples from unit Pre-A exposed in (a) the Gili Raja Island, sample no. GR6; and (b) east of Kamal, sample no BH5-17. Mo=smectite, It=illite, Ko=kaolinite, Qz=quartz, Ca=calcite, Si=siderite, Cl=chlorite, Ch=chamosite, Ar=aragonite.
Figure 6.12. Seismic facies map of subunit Al superimposed on the isochron map. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec).
Figure 6.13. Seismic facies map of subunit A2 superimposed on the isochron map. Contours in milUseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec).
Figure 6.14. Seismic facies map of subunit Bla superimposed on the isochron map. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec).
Figure 6.15. Seismic facies map of subunit BV superimposed on the isochron map. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec).
Figure 6.16. Isochron map of subunit Bib. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec).
113* 113*30' 114*
Figure 6.17. Isochron map of subunit B2. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec).
Figure 6.18. Isochron map of subunit B3. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec).
Figure 6.19. Seismic facies map of subunit Cla superimposed on the isochron map. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec).
Figure 6.20. Seismic facies map of subunit Clb superimposed on the isochron map. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec). Patterned indicates partly or completely eroded.
Figure 6.21. Seismic facies map of subunits C2 (deposited southward) and CV (deposited northward) superimposed on the isochron map. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec). Arrows indicate the directions of downlaps, patterned indicates partly or completely eroded.
Figure 6.22. Seismic facies map of subunit C3 superimposed on the isochron map. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec). Arrows indicate the directions of downlaps, patterned indicates partly or completely eroded.
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Figure 6.23. Seismic facies map of subunit C4 superimposed on the isochron map. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec). Arrows indicate the directions of downlaps, patterned indicates partly or completely eroded.
Figure 6.24. Seismic facies map of subunit C5 (C5a and C5b) superimposed on the isochron map. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec). Arrows indicate the directions of downlaps, patterned indicates partly or completely eroded.
Figure 6.25. Seismic facies map of subunit Dla superimposed on the isochron map. Contours in milliseconds TWT (1 msec = 0.75 m, seismic velocity = 1500 m/sec). Patterned indicates partly or completely eroded.
Figure 6.27. Palaeogeographic map during the subunit Dlb deposition superimposed on the contour time structure map of top subunit Dla.
Figure 6.28. Palaeogeographic map during the subunit D2 deposition superimposed on the isochron map of combined subunits Dlb, D2 and D3. Arrows indicate the directions of downlaps.
BULGE DELTA SYSTEM
Figure 6.29. Three dimensional mesh diagrams of the time structures: (a) top subunit D2 or bathymetry (D3 is very thin over the area and ignored); (b) top subunit Dla or bottom subunit Dlb; and (c) bottom subunit Dla or top subunit C. Note the increasing fold complexity westward (c) and rapid filling of synclines by subunit Dla (c & b).
Figure 6.30. Correlation between geological outcrops iii the Kedamean and Perning sections, and seismic sections MB-MB' and ME-ME'.
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Figure 6.31. (A) Sea level curve in the last 240 ka for Huon Peninsula, New Guinea; (B) lsO record from east equatorial core V19-30; (C) Recalculated sea level curve after correlation with 180 record from core V19-30 (after Chappell & Shackleton, 1986).
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FIGURES TO CHAPTER EIGHT
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Figure 8.1. Lineaments of subduction zones in the western Indonesian region (after Katili, 1989).
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APPENDIX B.l. TABULATION OF PETROGRAPHIC DESCRIPTIONS
EXPLANATION
TEXTURE packing
grain size maximum grain size average sphericity sorting
QUARTZ MONOC. normal undulose
QUARTZ POLYC. FELDSPAR K-Feldspar plagioclase
ROCK FRAGMENTS volcanic metamorphic sedimentary
CARBONATE bioclastics non-bioclastics carbonate mud sparry calcite
MAFIC MINERALS pyroxene
hornblende GLAUCONITE MICA CLAY/GROUND MASS OPAQUE MINERALS POROSITY
1 (poorly), 2(moderale), 3 (well)
mm mm 1 (low) 1(rare) %
l(rare)
1(rare) %
%
l(rare) 1(rare)
> %
l(rare) 1(rare) l(rare) %
l(rare) l(rare) 1(rare) 1(rare)
- % 1(rare) 1(rare)
- % %
%
%
• %
2(moderate) , 2(common)
, 2(common) , 2(common)
, 2(common) , 2(common)
, 2(common) , 2(common) , 2(common)
, 2(common) , 2(common) , 2(common) , 2(common)
, 2(common) , 2(common)
, 3(high) 3 (abundant)
3 (abundant) 3(abundant)
3 (abundant) 3(abundant)
3(abundant) 3 (abundant) 3(abundant)
3 (abundant) 3 (abundant) 3 (abundant) 3 (abundant)
3 (abundant) 3 (abundant)
NOTE: THE RELATIVE ABUNDANCE SCALE USED HERE CONSIDERS THE INDIVIDUAL PETROGRAPHIC COMPONENT AS 100 % CORRESPONDING TO THE GROUP PERCENTAGE (IN BOLD).
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APPENDED B.2 X-RAY DIFFRACTION DATA
Key to abbreviation
Ab Ar Bo Ca Ch Cl Gp Ho It Ko Mi Mo Px Qz Si
Albite Aragonite Boehmite
Calcite Chamosite Chlorite Gypsum Hornblende Illite Kaolinite Mica Smectite Pyroxene
. Quartz : Siderite
APPENDIX B.2. X-RAY DIFFRACTOGRAMS.
3000-
2500- ••
Sample no. W l (Wonocolo Formation)
Co 2000-
1500- •
500- c ° 'co'Ca"
uJ—UflL^A-J^M
Sample no. KB9 (Kalibeng Formation)
Sample no. NKabuhl3 (Atasangin Formation)
o
10
•••••••••••••
zo ••••••••••••a
30 • • • • • • • • • • , • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • i
40 . „ SO t
60 70
Sample no. NKabuhl8 (Atasangin Formation)
CONTmUED
3000-
Sample no. Nglaminl (Lidah Formation)
o
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2500'
2000-
1500-
1000-
500-••
3000
Sample no. Nglampin7 (Lidah Formation)
Or9anie hump ^^r~>**mmmmmj-m,
rannHmnai
30 2 e
50 "T™ 60
70
<tooo-
3500-
3000
c O
Sample no. Kedamean4 (Lidah Formation)
CONTINUED
4.000 —j .
3000 —
= 2000 —
o o
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0-4
^
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10 20
z
u
02 Sample no. GR6 (Lidah Formation)
r..— 30
Ch
Mo |
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°° ' Ar I" i'
2 e 50 60 7t >
3500-
3000-
2500-
K 2000—
SOD-1
100C- ^tf"
MO MO /vy^v i
Xmm~^J ^L,, Organic hump
Ko Lf -Co
A
Sample no. BH5-17 (Lidah Formation)
X 20 30 40
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APPENDIX C.2. BENTHONIC FORAMINIFERAL DATA.
No.
1
2
3
4
5
6
7
8
9
10
11
12
13
14
15
16
17
18
19
20
21
22
23
24
25
26
27
28
29
30
31
32
33
34
35
36
37
38
39
40
41
42
43
44
45
46
47
48
49
50
51
52
53
54
55
56
57
58
59
60
61
62
FORMATION
LOCALITY
SPECIES
Ammonia sp.
Ammonia beccarii
Ammonia gaimardii
Amphicorvna sp.
Amphicorvna scalaris
Anomalina colllgera
Bolivina sp.
Bolivma albatrossi
Bolivma incrassala
Bolivmita sp.
Boltvinita quadrilatera
Bulimina aculeata
Bulimina alazamensis
Bulimina costala
Bulimina marginata
Bulimina obtusa
Bulimina spicata
Bulimina striata
Cibicides lobatulus
Cibicides sp.
Cibicides subhaidmgeril
Dentalina sp.
Discorbina sp.
Elphidium sp.
Fissurina carinata
Fissunna lacunata
Fissurina sp.
Florilus sp.
Frondicularia sagitula
Gvroidina sp.
Gvroidina neosoldanii
Lagena sp.
Lagena sulcar
Laticarinina sp.
Laticarinina halophora
Lenticulina sp.
Lenticulina cf. calcar
Lenticulma/Robulus macrodiscus
Nodosaria flintii
Nodosana sp.
Nonion sp.
Parafrondicularia sp.
Parafrondicularia helenae
Parafrondicularia vaugani
Planulina sp.
Planulina wuellerstorfi
Plectofrondicularis califomica
Polvmorphina sp.
Pvrgosp.
Ouinqueloculina sp.
Ouinquelocultna sp./Stgmollopsis schlumbergerii
Ouinqueloculina seminulum
Rectoglandulina sp.
Siphonina pulchra
Stillostomella sp.
Textularia cornea
Textularia sp.
Trifarina bradvi
Turritellella spectabllls
Uvigerina sp.
Uvigerina laevis
Uvigeiinaj>eregrina
ICBENG 1 ATASANGIN FORMATION SILICEOUS MARLY FACIES
NORTH KABUH SECTION
NK»boh3
1
1
1
1
1
1
1
1
1
1
1
1
1
1
1
1
1
1
1
1
1
2
1
1
NtXtaJii;
2
1
1
1
1
I
2
1
NKubuhn
2
NKabuJiM
1
1
1
1
1
1
1
1
1
1
NKubnhl
1
1
1
1
1
1
NKatrahK)
3
1
1
3
1
1
1
1
2
1
3
1
2
2
1
1
1
1
1
1
1
NKabuM?
3
1
1
1
1
1
1
1
JKBI0
2
3
1
1
1
2
2
1
1
2
1
1
1
I
NKibuhl8
1
1
2
1
Note:
- Relative abundace: 1 » rare; 2 = common; 3 = abundant
- The relative abundance scale considers total benthonic
species in the individual sample as 1 0 0 %
CONTINUED
No.
1
2
3
4
6
8
11
12
20
21
22
24
25
26
FORMATION
LOCALITY
SPECIES
Ammonia sp.
Ammonia gaimardii
Amphistegma lessonii
Bolivina sp.
Bulimina marginata
Cibicides praecinctus
Eponides berthelotianus
Eponides schreibersi
Pseudorotalia sp.
Ouinqueloculina auberiana
Oumaueloculma seminulum
Saracenaria sp.
Textularia sp.
Uvigerina peregrina
PUC
KABUH
JKB7
2
2
1
- 2
2
2
2
1
2
1
1
LIDAH
S'RINOrN
SR05
3
3
3
No.
1
2
3
4
5
6
7
8
9
10
11
12
13
14
15
16
17
18
19
20
21
22
23
24
25
26
27
28
29
30
31
32
33
34
35
36
37
38
39
40
FORMATION
LOCALITY
SPECIES
Ammonia sp.
Ammonia gaimardii
Amphistegma sp.
Amphistegma lessonii
Amphistegma quovii
Asterorotalia trispinosa
Bigencrma sp.
Bolivma sp. 1
Bolivma sp.
Bulimina marginata
Cancris sp.
Cibicides sp.
Cibicides praecinctus
Dentalina sp.
Discorbma australwnsis
Elphidium sp.
Elphidium advena
Elphidium craticulatum
Elphidium crispum
Eponides sp.
Fissurina orbygnvana
Florilus sp.
Lagena sp. 1
Lagena sp. cf. crenata
Lenticulina sp.
Lenticulina sp. cf submamilligera
Nodosaria flintii
Nonion sp.
Operculina sp.
Operculina ammonoides
Planorbulina larvata
Pseudomassilina sp.
Pseudorotalia sp.
Pseudorotalia schroetenana
Reusella aculeate
Oumaueloculma seminulum
Robulus sp.
Siphomna sp.
Textularia sp.
Virgulma sp.
PASEAN FORMATION
KAMAL
MADS
1
2
1
1
1
2
2
2
2
1
I
1
1
1
1
2
1
PAMEKASAN
PK1
1
1
1
1
1
PK2
1
1
1
1
PK6
1
1
2
1
2
PKS
1
1
1
2
3
I
2
GILI RAJA
DR1
1
1
2
1
3
3
1
1
1
1
OR7
1
1
1
1
QILI GENTING
GG7
2
2
1
1
2
1
1
1
1
1
1
1
1
G03
1
2
1
2
2
1
1
2
1
1
2
2
2
Note:
- Relative abundace: 1 = rare; 2 = common; 3 = abundant
- The relative abundance scale considers total benthonic
species in the individual sample as 1 0 0 %
APPENDIX D. LIST OF ROCK SAMPLES USED IN THIS STUDY
No.
1 2
3 4
5 6 7
8
9 10
11
12 13 14
15 16
17 18
19 20
21
22 23
24 25
26 27
28
29 30 31
32
33 34
35
36
37
38
39
40
41 42
43
44
45 46
47
48
49 50
Field No.
Bl
B2
B3
BGJX
BH5-17
BL13
BL14
BL16
BL2 BUI
Banyuurip 4
GG3
GG5 GG7 GR1
GR4
GR6 GR7 JKB1
JKB10
JKB12
JKB14
JKB2
JKB5
JKB6
JKB7
JKB8
JKB9
KB13A
KB 15
KB 16
KD1 KD11
KD15 KD16
KD2
KE1A KE3
KE40
KG1
KG 10 KG 11
KG 7
KG 9 Kedamean 4
L3A
L3B
L4 L5
Ml
Longihite
111*28'42.62"
11178'42.62"
11178'42.62"
11178'42.62"
112°45'54.27"
11178'42.62"
11178'42.62"
lin8'42.62"
11178M2.62"
imO'47.21"
112°33' 6.24"
113°54'55.49"
113°54'55.49"
113°54'55.49"
113°47' 9.89"
113*47' 9.89"
113°47' 9.89"
113*47' 9.89"
112*13" 8.92"
liri3' 8.92"
liri3' 8.92"
112*13' 8.92"
u r i r 8.92"
uny 8.92"
liny 8.92"
liri3' 8.92"
liri3' 8.92"
112*13' 8.92"
iino"47.2r
111*30'47.21*'
u r a w r 112*33' 6.24"
112*33' 6.24"
112*33'6.24"
112*33'6.24"
112*33" 6.24"
ni°30'47.2i"
iireo-47.21"
111*30'47.21"
112*19'55.54"
112*19'55.54"
112*19'55.54"
112*19'55.54"
ii2°i9'55.54"
112*33'6.24"
111*4673.10"
111*4673.10"
111*4673.10"
111*4673.10"
111*4673.10"
Latitute
6*5574.07"
6*5574.07"
6*5574.07"
6*5574.07"
no18.01"
6*5574.07"
6*5574.07"
6*5574.07"
6*5574.07"
7°2572.47"
7*26' 6.84"
ri2' 8.52"
T12' 8.52"
T12- 8.52"
rirsur
ri2'51.41"
ri2'51.41"
ri2'51.41"
7°23,36.04"
7^23,36.04,,
7°23,36.04"
7°23'36.04"
7°23'36.04"
7°23'36.04"
7°23,36.04"
7°23,36.04"
7^3'36.04"
7*23'36.04"
7^25•22.47"
7*25'22.47"
7°25'22.47"
776' 6.84"
7*26' 6.84"
T2616.84"
7*26' 6.84"
7*26" 6.84"
7*2572.47"
77572.47"
77572.47"
774'12.37"
7*2412.37"
774'12.37"
774'12.37"
7*24'12.37"
7*26' 6.84"
T 572.79"
T 572.79"
T 572.79"
T 572.79"
T 572.79"
Locality
Bulu
Plantungan, Blora
Bulu
Plantungan, Blora
Kamal.
Bulu
Bulu
Bulu
Bulu
Solo River
Kedamean
Gili Genteng Island
Gili Genteng Island
Gili Genteng Island
Gili Raja Island
Gili Raja Island
Gili Raja Island
Gili Raja Island
Kabuh
Kabuh
Kabuh
Kabuh
Kabuh
Kabuh
Kabuh
Kabuh
Kabuh
Kabuh
Solo River
Solo River
Solo River
Kedamean, Mojokerto
Kedamean, Mojokerto
Kedamean, Mojokerto
Kedamean, Mojokerto
Kedamean, Mojokerto
Solo River
Solo River
Solo River
Kemlagi, Mojokerto
Kemlagi, Mojokerto
Kemlagi, Mojokerto
Kemlagi, Mojokerto
Kemlagi, Mojokerto
Kedamean
Kawengan, Cepu
Kawengan, Cepu
Kawengan, Cepu
Kawengan, Cepu
Kawengan, Cepu
Formation
Bulu
Bulu
Bulu
Tawun, Lower
Lidah
Tawun
Tawun
Bulu
Bulu
Kerek
Pucangan
Pasean
Pasean
Pasean
Pasean
Pasean
Pasean
Pasean
Kabuh
Atasangin
Pucangan
Atasangin
Kabuh
Pucangan
Pucangan
Pucangan
Pucangan
Pucangan
Kalibeng
Kalibeng
Kalibeng
Pucangan
Pucangan
Pucangan
Pucangan
Pucangan
Kerek
Kerek
Kerek
Lidah
Lidah
Lidah
Lidah
Lidah
Lidah
Ledok, Middle
Ledok, Middle
Ledok, Top
Ledok, Top
Mundu, Lower
Age Mid Miocene
Mid Miocene
Mid Miocene
Mid Miocene
Early Pleistocene
Mid Miocene
Mid Miocene
Mid Miocene
Mid Miocene
Late Miocene
Early Pleistocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Mid Pleistocene
Late Pliocene
Early Pleistocene
Late Pliocene
Mid Pleistocene
Early Pleistocene
Early Pleistocene
Early Pleistocene
Early Pleistocene
Early Pleistocene
Pliocene
Pliocene
Pliocene
Early Pleistocene
Early Pleistocene
Early Pleistocene
Early Pleistocene
Early Pleistocene
Late Miocene
Late Miocene
Late Miocene
Early Pleistocene
Early Pleistocene
Early Pleistocene
Early Pleistocene
Early Pleistocene
Early Pleistocene
Late Miocene
Late Miocene
Late Miocene
Late Miocene
Pliocene
Description
Limestone
Limestone
Limestone
Organic shale
Mudstone
Quartz sandstone
Quartz sandstone
Calcarenite
Quartz sandstone
Calcirudite
Sandy mudstone
Marl
Marl
Marl
Calcarenite
Calcarenite
Black mudstone
Black mudstone
Sandstone
Marl
Sandstone
Medium sandstone
Sandy tuff
Tuffaceous sandstone
Sandstone
Sandstone
Sandstone
Sandstone
Fine sandstone
Calcarenite
Calcarenite
Tuffaceous sandstone
Medium sandstone
Sandy mudstone
Muddy sandstone
Tuffaceous mudstone
Sandstone
Calcareous sandstone'
Calcirudite
Sandstone
Sandstone
Mudstone clast
Sandstone
Sandstone
Bluish grey mudstone
Calcarenite
Calcarenite
Glauconitic sandstone
Glauconitic sandstone
Fos.marl
51
52
53
54
55
! 56
57
58
59
60
61
! 62
63
64 65
66
67
68
69
70
71
72
73 74
75 76
77
78
79 80
Madura 18
Madura 8
N.Kabuh 1
N.Kabuh 11
N.Kabuh 12
N.Kabuh 13
N.Kabuh 14
N.Kabuh 16
N.Kabuh 17
N.Kabuh 18
N.Kabuh 3
N.Kabuh 4
N.Kabuh6
N.Kabuh 7
N.Kabuh 9
N B I 1
Nglampin 1
Nglampin 7
PK1
PK2
PK5
PK6
PK8
Perning 2
SI
S. Ringin5
SR2
SR5 SR6 Wl
113°0'54.49"
112*46' 6.88"
112*12'12.72"
liri2'12.72"
112°12'12.72"
112*12'12.72"
112°12'12.72"
112*12'12.72"
112*12_V2.72"
112*12" 12.72"
112°12'12.72"
112°12'12.72"
112*12'12.72"
112°12'12.72"
112°12'12.72"
112°15'31.39"
111'40'43.45"
111"40"43.45"
113°31'37.56"
113°31'37.56"
113*31'37.56"
113°31'37.56"
113*31'37.56"
11276'43.88"
111*3672.60"
112*1577.66"
112*1577.66"
112*1577.66"
112*1577.66"
111*3672.60"
riO'50.07"
7*8'56.83"
772'32.15"
7*22"32.15"
772'32.15"
772'32.15"
772'32.15"
772'32.15"
772'32.15"
7*22'32.15"
772'32.15"
7*22'32.15"
772'32.15"
7*22'32.15"
772'32.15"
ri6'48.06"
77078.67"
77078.67"
7*12'49.52"
7°12'49.52"
7°12'49.52"
7°12'49.52"
7*12'49.52"
777' 2.51"
T 572.79"
773'57.34"
r23'57.34"
773'57.34"
773'57.34"
T 572.79"
Blega, Madura
Sukolilo, Kamal
North Kabuh
North Kabuh
North Kabuh
North Kabuh
North Kabuh
North Kabuh
North Kabuh
North Kabuh
North Kabuh
North Kabuh
North Kabuh
North Kabuh
North Kabuh
Ngimbang
Nglampin, Cepu
Nglampin, Cepu
Pamekasan
Pamekasan
Pamekasan
Pamekasan
Pamekasan
Perning, Mojokerto
Ledok, Cepu
Sumbeningin, Kabuh
Sumberringin, Kabuh
Sumberringin, Kabuh
Sumberringin, Kabuh
Ledok. Cepu
Pasean
Pasean
Atasangin
Atasangin
Atasangin
Atasangin
Atasangin
Atasangin
Atasangin
Atasangin
Kalibeng
Atasangin
Atasangin
Atasangin
Atasangin
Mundu
Lidah
Lidah
Pasean
Pasean
Pasean
Pasean
Pasean
Pucangan
Mundu
Lidah
Pucangan
Pucangan
Pucangan
Wonocolo
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Early Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Early Pliocene
Early Pleistocene
Early Pleistocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Late Pliocene
Early Pleistocene
Pliocene
Early Pleistocene
Early Pleistocene
Early Pleistocene
Early Pleistocene
Late Miocene
Conglomerate
Calcarenite
Marl
Sandy marl
Marl
Marl
Marl
Mad
Marl
Fine sandstone
Sandy marl
Calcareous sandstone
Medium sandstone
Tuffaceous marl
Marl
Calcarenite
Sandy mudstone
Sandy mudstone
Mudstone
Calcareous mudstone
Calcareous mudstone
Calcareous mudstone
Calcareous mudstone
Sandstone
Fos. sandstone
Sandy mudstone
Muddy sandstone
Sandstone
Sandstone
Fos. calcarenite
APPENDIX E.l. GRAPHIC SUMMARY OF PETROLEUM EXPLORATORY WELL REPORT FOR JS2-1 WELL.
MIOCENE T
PLIOCENE in
o o
o o
u o o
to
o o
o o
I I I I I I I
I I I I I I I ! I i
i i i i i i i i i r I I I I I I I I I I
I I I I I I I I I I I I I I I I I I I I
I I I I I I I I I I J I I I I I I I I I
Epoch
Depth (m)
Est. Lithology
CO
o O
8£ 3 3 m 9
1 » O i • 3
I - s . I e».rr
Planktonic Forarninifera zones
Cycloclypeus cf. postlndopaclflcus
Asterorotalla spp.
Pseudorotalia spp.
Globigerinoides trilobus s.l.
Orbulina s.l.
Globigerinoides obllquus
Qloboquadrlna altlspira
Sphaeroidinellopsis subdehiscens
Globigerinoides extremis
Globorotalia tumida
Globorotalia plesiotumida
Globorotalia dutertrel
Pulleniatina prlmalis
Globorotalia humerosa
Globorotalia cf. flexuosa
Globorotalia menardlllcultrata
Globorotalia scltula
Sphaeroidinellopsis semlnullna
Globigerinoides conglobatus
Hestlgerlna sp.
Globorotalia pseudoplma
Globigerinoides ruber
Globigerinoides canimarensls
Globigerina bulloldes
Pulleniatina praecursor
Globorotalia merotumlda
Sphaeroidinella dehiscens
Globorotalia cf. exllis
Globorotalia cf. ungulate
Globigerinoides elongatus
Globorotalia cf. subcretacea
Globorotalia cf. Inflata
Tl O —l
0)
2 —H CD — I
00
APPENDIX E.2. GRAPHIC SUMMARY OF PETROLEUM EXPLORATORY WELL REPORT FOR JS-3-1 WELL.
o o a. UJ
UJ
z UJ
O O
Q.
D
100-
UJ z UJ
o O
200-
300
O O
tn UJ
Foraminifers Benthonic Planktonic
ra i-,
0) •*— c
E ra k.
o u. o c o "5 w
! • 2. o a. N
Globorotalia tosaensis
Sphaeroidinella dehiscens
400-
500-
600-
T~T 1 1 1
55
Globorotalia plesiotumida
Recrystallized
Globigerinoides trilobus
APPENDIX E.3. GRAPHIC SUMMARY OF PETROLEUM EXPLORATORY WELL REPORT FOR JS8-1 WELL.
APPENDIX E.4. GRAPHIC SUMMARY OF PETROLEUM EXPLORATORY WELL REPORT FOR JS10-1 WELL.
u o Q. UJ
9
o
z HI
o g
200-
o o
tn
UJ
300
UJ
z UJ
o o
400-
o <p
c
E o \-
u.
c CO
E 8
500
Globorotalia tosaensis
Planktonic^ Foraminifers
0) a -o o —
• X U-x Q. to to
-S> to rj • ^ . - - ^
~- o P c
O X X <B o co l~ •-QJ -Q
tn c-—
tO 10 a 0" to X. •— W 1» — to -; OJ <0 Q.--•Q -£3. -CJ to — O O 73 -~ to
X x> 6 to >x rj. to to ix
* u a » B T3 to T3 tO \7- to -— c to O -~ B O •» C C -~ c c x -i
v. *J • v. v. ta oj OJ O OJ OJ 3 •-. Co v. Co o> rj- c — o — — Q 01
3 -O -Q +J -Q 47j -~ •X O J3 to — I
o o to o o -~-— x to • 3
cjorjjcjTjarcjTjrjQ.
to to to
is ° T3 *-Xl QJ
doji W Q.-C
to to to •- to -~ — c — to — to O QJ O X t3> X rj -~ o •5 * -Q
o o o
to TJ
§ Xi to mt •-
tD UI C
to tu 3 X 73 to _& E -~ e to -Q tO T3 O O -~
E tO to 73 m rn x> 13 a
to
•— -~ to X X Xi OJ OJ a Co CO X •-. •- Q •O -Cl J7j
O O O
•c .
3 .-•O to to to
XJ
to to tO -X
~ o OJ (0
, c • "
T3 to
X o 10 X 73
u tu to X
to
to — 73
to to o
O o X X oj o 10 J7J •c o
U 5 cj 3 3 5 <J3" 5
to to _ c to U OJ to o O O S 3
'- ~- to •- ;< >• OJ •- Q. OJ U 3 -c o — ^ Cr o> -o>*. ID -~. TD 3 C ~~ OJ . Q) -TJ to to *-E O ;- Q. O to to 0) to to -_-- c c •» -to -x "rj to to •X to •— Xi XJ
o -~ o o a X C X X X o oj to o o •O. •— to -Q -JO o ~~ -c o a -- 3 Q.-~ — csci.iJycjC3ttC3<acat3C3crj
to to
to 53) O •X C H-to O to — — to 3 OJ to D) to C to X 3 7JJ U
a to •-•- c — to — to Xi X Xi O OJ o X O) X o — ro -Q -O 47J
o o o
to
i3 -H -x Tj
to to tg •-
i-c-x3! ••~ E xi
i<.' P ° Oj S- XJ —
O O M x oj
*- to OJ —
u a; s n o
to — to to •^ Q -«. c ~- -~ to'-- to to Xi X X» +J rj oj o o k- tj) X X rj-~ o o •S3 -D « JTJ O rj o o
Sphaeroidinella dehiscens
6oa
Globorotalia tumida
rM
^ n ^ D l X E-5' GRApHIC SUMMARY OF PETROLEUM EXPLORATORY WELL REPORT FOR JSI 6-1 WELL
JZ
u o a UJ
MIOCENE
! PLIOCENE
E JC
a ©
a 300-
400-
500-
600-
o> o o sz -J to Ui
T T —
-i- -i- -r
T T T
_T T_-r
-r -r -T-
T — I — r
-r- -r -r
-r- -r -r
T T- T
-I- T- T
-r -i r
— —
:;
Plan
kfon
ic F
orarninifera
zone
s
. Foraminifers Benthonic . Planktonic
Pseu
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sacculifer
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APPENDIX F.l. COMPUTER PROGRAM TO DIGITISE SHIP POSITION ON MAP
1 REM Written by Susilohadi 1992 10 20 30 40 50 60 70 80 90 100 110 120 130 140 150 155 160 170 180 190 200 210 220 280 290 300 310 320 330 340 350 360 370 380 390 400 410 420 430 440 450 460 470 480 490 500 510 520 530 540 550 560 570 580 590 600 610 620 630 640 650 660 670 680 690 700 710 715 720 730 740 750 760
CLEAR "PROGRAM
using BASIC
TO DIGITISE UNKNOWN POINT ON MAP"
The and
program maximum
CLS PRINT PRINT PRINT " PRINT " PRINT REM Digitise reference points INPUT "File name of data : ", OPEN NF$ FOR APPEND AS #2 PRINT "Put map on digitiser tablet DIM X(5), Y(5), P#(5), Q#(5)
1 TO 5 "Reference point I 1160 2 THEN 2 90
needs at least two reference points, of five points"
NF$
and begin to digitise the reference points"
FOR I = PRINT PRINT GOSUB IF C =
no.
A: Y(I) = B
324000# -Longitude
(3600 * DER)
X(I) = BEEP PRINT INPUT Q#(D INPUT P#(D
NEXT I IF I <> BEEP: PRINT FOR J = 1 TO CLS : RUN Z = I - 1: OZ = REM Measure the FOR I = 1 TO Z
FOR J = (I + OZ = OZ + JPET# = ( JDIG# = ( PJ# = PJ# IF (Y(J)
Enter the coordinate :" Latitude (degree, minute, second)
3600 * DER) + (60 * MIN) (degree, minute, second) + (60 * MIN) + DET
", DER, MIN, DET DET) : ", DER, MIN, DET
THEN 330 PRINT " 600: NEXT
0
NOT ENOUGH REFERENCE POINT"
distance ratio (pj) and angle (alpha)
1 1 (P# (X + -Y -X
TO Z
ID + (Q#(D - Q#(J)) (Y(I) - Y(J))
P#(J)) A
I) - X(J)) A 2 (JPET# / JDIG#) I))=0 THEN SDTDIG#=0 I))=0 THEN SDTDIG#=1.5707963#
2) 2)
- Y(I)) / (X(J) THEN SDTPET#=0 THEN SDTPET#=1.5707963#
IF (X(J) GOTO 450 SDTDIG# = ATN((Y(J) IF (Q#(J)-Q#(D)=0 IF (P#(J)-P#(D)=0
' GOTO 490 SDTPET# = ATN((Q#(J) - Q#(D) / (P#(J) ALPHA# = ALPHA# + (SDTPET# - SDTDIG#)
NEXT J NEXT I PJ# = PJ# / OZ ALPHA# = ALPHA* / OZ CLS
Coordinate reference points are :"
X(I)))
P#(D)
PRINT " PRINT PRINT " FOR I = A$ = B$ = C$ = PRINT OT# = PRINT OT# = PRINT
NEXT I PRINT
X y 1 TO Z " ###### ###### "###° ##'##.###""
A$; X(I); Y(I) GOSUB 1220
B$; DJ; GOSUB
C$; DJ;
longitude latitude"
USING P#(I): USING Q#(I): USING
MT; 1220 MT;
DT;
DT
PRINT REM Determination of arbitrary points PRINT " Begin to digitise arbitrary points GOSUB 1160 IF C = 2 THEN 1130 K = K + 1 LAT# = 0: LON# = 0 FOR I = 1 TO Z JTTK# = (((A - Xdl IF B = Y(I) AND A =
) * 2 + (B - Y(D) : X(I) THEN SDTTK# =
, 2) A .5) * 0: GOTO 780
PJ#
CONTINUED 770 SDTTK# = ATN((B - Y(I)) / (A - X(I))) 7 80 SPTTK# = SDTTK# - ALPHA# 790 DLAT# = ABS(JTTK# * SIN(SPTTK#)) 800 DLON# = ABS(JTTK# * COS(SPTTK#)) 810 IF A <= X(I) AND B > Y(I) THEN 910 820 IF A <= X(I) AND B <= Y(I) THEN 850 830 IF A > X(I) AND B > Y(I) THEN 940 840 IF A > X(I) AND B <= Y(I) THEN 880 850 OLAT# = Q#(I) - DLAT# 8 60 OLON# = P#(I) - DLON# 870 GOTO 960 880 OLAT# = Q#(I) - DLAT# 890 OLON# = P#(I) + DLON# 900 GOTO 960 910 OLAT# = Q#(I) + DLAT# 920 OLON# = P#(I) - DLON# 930 GOTO 960 940 OLAT# = Q#(I) + DLAT# 950 OLON# = P#(I) + DLON# 960 LAT# = LAT# + OLAT# 97 0 LON# = LON# + OLON# 980 NEXT I 990 LAT# = LAT# / Z 1000 LON# = LON# / Z 1010 S# = 324000# - LAT# 1020 R# = LON# 1030 PRINT "Coordinate point no. : "; 1040 PRINT K; 1050 PRINT "is : " 1060 OT# = R#: GOSUB 1220 1070 PRINT USING B$; DJ; MT; DT; 1080 OT# = S#: GOSUB 1220 10 90 PRINT USING C$; DJ; MT; DT; 1100 PRINT 1120 'PRINT #2, USING »#######.###,#######.###,\ 1121 PRINT #2, USING "#######.### #######.###"; R#; 1125 GOTO 710 1130 CLOSE #2 1140 END 1150 REM Digitiser control 1160 OPEN "COM1:9600,N,7,2,CS,DS" FOR RANDOM AS #1 1165 PRINT "Press red button when finish" 1170 PRINT #1, "P"; 1180 INPUT #1, A, B, C 1200 CLOSE #1 1210 RETURN 1220 REM Change to degree format 1230 DJ = INT(OT# / 3600) 1240 MT = INT((OT# - (DJ * 3600)) / 60) 1250 DT = OT# - (3600 * DJ) - (60 * MT) 12 60 RETURN
APPENDIX F.2. COMPUTER PROGRAM TO DIGITISE SEISMIC SECTION CLS : CLEAR DIM in(4) PRINT *** COMPUTER PROGRAM TO DIGITISE SEISMIC SECTION ***"
Written by Susilohadi in 1992, using QBASIC" FUNCTION OF BUTTONS ON THE DIGITISING PUCK
> Button no. > Button no. > Button no.
(red) to digitise seismic horizon" (blue) to change seismic sector to be digitised" (white) to quit"
PRINT " PRINT " PRINT PRINT " PRINT " PRINT " PRINT INPUT " INPUT " INPUT- " PRINT PRINT " PRINT PRINT 1400 REM Subrouting data entry
REM variables: jsi=left horizontal distance REM jsa=right horizontal distance REM xas=absis at zero point REM yas=ordinat at zero point REM js=average horizontal distance REM s=ship time travel PRINT " DIGIT SEISMIC BORDERS :" 1410 GOSUB 1300
File name of data : ",nf$ Interval of digitising (hour,minute,second) : Moving to the left (i) of right (a) : ",tb$
PUT SEISMIC SECTION ON THE DIGITISING TABLET"
",jtb,mtb,dtb
CLS PRINT " BEGIN TO DIGITISE :" 1420 GOSUB 1100
IF z = 2 OR z = 3 THEN 1410 IF z = 4 THEN END nv = y - (mv * x) xas = ( (nv - n) / (mh - mv) ) yas = (mh * xas) + n jsi = ((((xas - xai) A 2) + ((yas - yai) A 2)) A
jsa = ((((xaa - xas) A 2) + ((yaa - yas) A 2)) A
jsi = jsi + dw# jsa = -jsa + dk# js = INT((jsi + jsa) / 2) PRINT " Point position (hour:minute:second) : ";
jj = INT(js / 3600) mm = INT((js - (jj * 3600)) / 60) dd = INT(js - (jj * 3600) - (mm * 60))
.5)
.5) * jh * jh
PRINT jj; PRINT ":"; PRINT mm; PRINT ":"; PRINT dd
PRINT " Time travel (TWT) : "; s = (( ((xas - x)A2 + (yas - y)A2)A.5) * jv)
PRINT USING "####.##";s; PRINT " msec."
OPEN nf$ FOR APPEND AS #2 PRINT #2, USING "####### ####.##";js, s
CLOSE #2 GOTO 1420
1100 REM subroutine to pick values from digitiser REM variables: x=absis; y=ordinat; z=button number OPEN "coml:9600,n,8,1" AS #1
PRINT #l,chr$(0) PRINT #1,"B" FOR i=0 to 4
in(i)=asc(INPUT$(l,#D) next i z=7 and in(0) x=128*(127 and in(2))+(127 y=128*(127 and in(4))+(127
CLOSE #1 RETURN
1300 REM subroutine to change sector to be digitised REM (actified by button 2 or 3 on the digitiser puck) REM variables earlier border : jw=hour; mw=minute; dw#=second
and in(l)) and in(3))
CONTINUED REM xai=absis minimum; yai=ordinat minimum REM xbi=absis maximum; ybi=ordinat maximum REM variables later border: jk=hour; mk=minute; dk#=second REM xaa=absis minimum; yaa=ordinat minimum REM xba=absis maximum; yba=ordinat maximum REM other variables: jh=horizontal unit; jv=vertikal unit REM nk=minimum scale value; nb=maximum scale value IF z = 2 and tb$ = "a" or z = 2 and tb$ = "A" THEN
jw = jk:mw = mk:dw = dk jk = jk + jtb:mk = mk + mtb:dk = dk + dtb IF dk >= 60 THEN mk = mk + 1:dk = dk - 60 IF mk >= 60 THEN jk = jk + l:mk = mk - 60 GOTO 1320
END IF IF z = 2 and tb$ = "i" or z = 2 and tb$ = "I" THEN
jk = jw:mk = mw:dk = dw jw = jw + jtb:mw = mw + mtb:dw = dw + dtb IF dw >= 60 THEN mw = mw + l:dw = dw - 60 IF mw >= 60 THEN jw = jw + l:mw = mw - 60 GOTO 1320
END IF 1310 INPUT " Type in hour,minute,second left side of the section : ", jw, mw, dw
INPUT " Type in hour, minute, second right side of the section : ", jk, mk, dk 1320 PRINT " Batas kiri : ";jw;":";mw;":";dw
PRINT " Batas kanan : ";jk;":";mk;":";dk dw# = dw + (3600 * jw) + (60 * mw) dk# = dk + (3600 * jk) + (60 * mk)
PRINT " Digit TWT minimum and maximum left side" GOSUB 1100:BEEP
xai = x: yai = y GOSUB 1100:BEEP
xbi = x: ybi = y PRINT " Digit TWT minimum and maximum right side"
GOSUB 1100.-BEEP xaa = x: yaa = y
GOSUB 1100:BEEP
mh = mh = mh = mv = n =
xba (yaa mh + mh / -1 / ( (yai
= x: yba = y - yai) ((yba • 2 mh - mh *
/ (xaa • - ybi) /
xai) + INPUT " TWT minimum : ", INPUT " TWT maximum : ", jv = Jv = jv = jh = jh = jh =
(nb -jv + jv / (dk# jh + jh /
RETURN
- nk) / ((nb -2 - dw#) ((dk# -2
- xai) (xba - xbi))
(yaa - mh nk nb
(((((xbi - xai) ' nk) / (((((xba -
/ (((((xaa - xai) - dw#) / (((((xba
* xaa
v 2) + xaa)
A 2) - xbi
)) / 2
((ybi * 2) +
- yai) * ((yba -
+ ((yaa ) A 2) +
- yai) ((yba
2))) * yaa) "
A 2))) - ybi)
.5) 2))) A
A .5) A 2)))
.5))
v -5)
APPENDIX F.3. COMPUTER PROGRAM TO CONVERT GRIDDED DATA TO 'SURFER* COMPATIBLE FORMAT (ASCII) DIM z(13000) PRINT "PROGRAM TO CONVERT GRIDDED DATA" PRINT "FROM XYZ FORMAT (ASCII) TO 'SURFER' (ASCII) COMPATIBLE FORMAT" PRINT "Written by Susilohadi, 1992, usinq QBASIC" PRINT 'INPUT "Name of file to convert : ", nfu$ 'INPUT "New file name : ", nfh$ OPEN nfu$ FOR INPUT AS #1 INPUT #1, x INPUT #1, y INPUT #1, zz
CLOSE #1 xspc = X yspc = y OPEN nfu$ FOR INPUT AS #1 spacel: INPUT #1, x INPUT #1, y INPUT #1, zz IF xspc = x AND (NOT EOF(l)) THEN GOTO spacel xspc = ABS(xspc - x)
CLOSE #1 OPEN nfu$ FOR INPUT AS #1 space2: INPUT #1, x INPUT #1, y INPUT #1, zz IF yspc = y AND (NOT EOF(l)) THEN GOTO space2 yspc = ABS(yspc - y)
CLOSE #1 OPEN nfu$ FOR INPUT AS #1 INPUT #1, xmin INPUT #1, ymin INPUT #1, zmin zmin = -zmin zmax = zmin
CLOSE #1 i = 1 OPEN nfu$ FOR INPUT AS #1 WHILE (NOT EOF(l)) INPUT #1, x INPUT #1, y INPUT #1, z(i)
PRINT USING1"#######.## #######.## ####.##"; x; y; z(i) IF xmin >= x THEN xmin = x ELSE xmin = xmin IF ymin >= y THEN ymin = y ELSE ymin = ymin IF zmin >= z(i) THEN zmin = z(i) ELSE zmin = zmin IF xmax <= x THEN xmax = x ELSE xmax = xmax IF ymax <= y THEN ymax = y ELSE ymax = ymax IF zmax <= z(i) THEN zmax = z(i) ELSE zmax = zmax i = i + 1
WEND CLOSE #1 PRINT xmin, xmax PRINT ymin, ymax PRINT zmin, zmax xnum = ((xmax - xmin) / xspc) + 1 ynum = ((ymax - ymin) / yspc) + 1 OPEN nfh$ FOR OUTPUT AS #2 'PRINT HEADING ON SCREEN PRINT "DSAA" PRINT USING "#### ####"; xnum; ynum PRINT USING "######.## ######.##"; xmin; xmax PRINT USING "######.## ######.##"; -ymax; -ymin 'PRINT USING "####.## ####.##"; xspc; yspc PRINT USING "####.## ####.##"; zmin; zmax 'SAVE HEADING PRINT #2 "DSAA" PRINT #2, USING "#### ####"; xnum; ynum PRINT #2, USING "######.## ######.##"; xmin; xmax PRINT #2, USING "######.## ######.##"; -ymax; -ymin 'PRINT #2, USING "####.## ####.##"; xspc; yspc
CONTINUED PRINT #2, USING "####.## ####.##"; zmin; zmax
'PRINTING AND SAVING FOR k = ynum TO 1 STEP -1 FOR 1 = 1 TO xnum PRINT USING "####.## "; z((xnum * (k - 1)) + 1); PRINT #2, USING "####.## "; z((xnum * (k - 1)) + 1 ) ; IF (1 / 10) - INT(1 / 10) = 0 THEN PRINT : PRINT #2, ;
NEXT 1 PRINT : PRINT PRINT #2, PRINT #2,
NEXT k CLOSE #2 BEEP END
APPENDIX F.4. COMPUTER PROGRAM FOR GRIDDING SEISMIC DATA
//Written by Susilohadi 1992, using C++ language #include <stdio.h> #include <math.h> #include <conio.h> #include <iostream.h> #include <ctype.h> #include <stdlib.h> ttinclude <dos.h> unsigned int
innrsnnm . lat[6000];
int depth[6000]
a, b, c;
float xmin, xmax, ymin, ymax, zmin, zmax, x_spc y_spc power xmine ymme
float di
/Vlpjigitude saved, as integer, of .(xaxis-xmin) * 1.0 //latitude saved as integer of (yaxis-ymin)*10
//depth saved as integer of //number of data point //dummy variables
//longitude minimun on file //longitude minimun on file //latitude minimun on file //latitude minimun on file //depth minimum on file //depth maximum on file //grid spacing on x-axis //grid spacing on y-axis //power, weight st,
st; stsort[64], zsort[64];
(zaxis-zmin)*10
char inpfile[50] , outfile[50];
//file name of data to be retrieved //file name of gridded data to be saved
FILE *file_in, *file out;
//for ascii x,y,z file to be retrieved //for ascii x,y,z grided data to be saved
//subroutines void heading(void); void subheadingl(void); void subheading2(void); void getfilename(); void readmaxmin(); void readdataO; void prtmaxmin(); void grdspc(); float distance(float, float, float, float); float xrange(float, float, float, float); float yrange(float, float, float, float); float truex(int); float truey(int); float truez(int); void readmaxmin(); void saving(float, void rectangular(float, void ellipsoidal(float, void area(); void closestpoint(); void freefind(int); void quadrantfind(int); void octantfind(int); void sorter(int, int); void notes(void); void beep(void);
float, float); float, int); float, float)
void main() {
int selectmenu; start: clrscr();
CONTINUED heading(); menuopt: selectmenu = getche(); switch(selectmenu) {
case '1': area() ; break;
case '2': closestpoint(); break;
case '3': exit(1); default : beep()_;
goto menuopt; } gotoxy(15, 25); printf("CALULATION COMPLETED, PRESS ANY KEY"); getch(); goto start;
} //estimate value based on points in the given area void area() {
int search, azim; float mayor, minor; clrscr (); subheadingl(); getfilename(); readmaxmin(); prtmaxmin(); grdspc(); gotoxy(53, 16); search = toupper(getche());
// printf(search); gotoxy(44, 17) cin » minor; gotoxy(44, 18) cin >> mayor; gotoxy(44, 19) cin » azim; readdata (); if (search == 'R') {
rectangular(mayor, minor, azim); } else {
ellipsoidal(mayor, minor, azim); } return;
//estimate z based on rectangular area of searching void rectangular(float mayor, float minor, int azim) {
int pointh = 0; float x_est, //longitude of estimeted point x_point, //longitude of known point y_est, //latitude of estimated point y point, //latitude of known point dist_P, //distance between estimated and known point dist_R, //maximum distance of known point allowed x, y, truelon, truelat, truedepth, vdepth, wdepth, value, distance_P, distance_R[5], xe, ye, m_P, me[4], x_R, y_R, n_P, n_horiz, n_vert, ne[5];
me[l] = me[2] = tan(azim * 3.14159/180); me[3] = me[4] = tan((azim+90) * 3.14159 / 180); forty = ymin; y <= ymax; y += y spc) {
for(x = xmin; x <= xmax; x += x spc) { ++pointh;
CONTINUED vdepth = 0; wdepth = 0; value = 0; switch(azim) {
case 0: break; case 90: break;
case -90: break; default:
n_horiz = y - (me[l] * x) ; n_vert = y - (me[3] * x); ne[l] = n_horiz - (minor / cos(azim * 3.14159 / 180)); ne[2] = n_horiz + (minor / cos(azim * 3.14159 / 180)); ne[3] = n_vert - (mayor / sin(azim * 3.14159 / 180)); ne[4] = n_vert + (mayor / sin(azim * 3.14159 / 180));
for (a = 1; a != i; ++a) { gotoxy(17, 21); printf("%d",a);
truelon = truex(a); truelat = truey(a); truedepth = truez(a); distance_P = distance(x, truelon, y, truelat); if (distance_P == 0) {
vdepth = truedepth; wdepth = 1; break;
} if (truelon != x)
{ m_P = (truelat - y) / (truelon - x); n_P = y - (m_P * x);
} switch(azim) {
case 0: if (truelon == x)
{ distance_R[l] = minor;
} else
distance R[l] = mayor / cos(atan(m_P)); distance~R[2] = minor / cos((90 * 3.14159 /180) - atan(m_P));
if (distance_R[l] >= distance_R[2]) distance_R[l] = distance_R[2];
} if (distance_P <= distance_R[l]) vdepth += truedepth/(pow(distance P, power));
wdepth += 1/(pow(distance_P, power)-); } break; case 90: if (truelon == x)
{ distance_R[l] = mayor;
} else
distance R[l] = minor/ cos(atan(mP)); distancejl[2] = mayor/ cos ((90 * 1.14159 /180) - atan(m_P));
if (distance R[l] >= distance_R[2]) { distance R[l] = distance_R[2]; }
} if (distance P <= distance R[l])
vdepth += truedepth/(pow(distance_P, power));
CONTINUED wdepth += 1/(pow(distance_P, power));
} break; case -90: if (truelon == x)
{ distance R[l] = mayor;
} else {
distance_R[l] = minor/ cos(atan(m P)-) ; distance_R[2] = mayor/ cos((90 * 1.14159 /180) - atan(m_P));
if (distance R[l] >= distance R[2]) { distance R[l] = distance R[2]; }
} if (distance P <= distance R[l]) {
vdepth += truedepth/(pow(distance P, power)); wdepth += 1/(pow (distance P, power J") ;
} break; default : if (truelon == x)
{ distance_R[l] = mayor / cos((90-azim) * 3.14159 /180); distance_R[2] = minor / cos(azim * 3.14159 / 180); if (distance R[l] >= distance R[2]) { distance R[l] = distance R[2]; }
} else {
for (b = 1; b <= 4; ++b) { x_R = xrange(m_P, n_P, me[b], ne[b]); y R = yrange(m_P, n_P, me[b], ne[b]); ~distance_R[b] = distance(x, x_R, y, y_R); if (distance R[l] >= distance R[b]) { distance R[l] = distance R[b];
} }
} if (distance P <= distance R[l]) {
vdepth += truedepth/(pow(distance_P, power)); wdepth += 1/(pow(distance_P, power));
} if (vdepth == 0 & wdepth == 0) value = 0; else value = vdepth / wdepth; gotoxy(14, 23);
printf("%d",pointh); gotoxy(26, 23); printf("%7.2f", x); gotoxy(42, 23); printf("%7.2f", y); gotoxy(58, 23); printf("%5.2f", value); saving(x, y, value); }
} } _ //estimate z based on ellipsoidal area of searching
CONTINUED void ellipsoidal(float mayor, float minor, float azim)
int pointh = 0;
disttRdiSt-P' //distance between estimated and known point
x v~ rmoinn //maxim™ distance of known point allowed ' ' U l ° ! truelat, truedepth, vdepth, wdepth, value, ?iwf -P? dlstance_R' x_R, y R, m P, n P; tor(y = ymm; y <= y^^. y + = y_spcy
for(x = xmin; x <= xmax; x += x_spc) ++pointh;
vdepth = wdepth = value = 0-for (a = 1; a <= i; ++a)
truelon = truelat = truedepth = 0; gotoxy(17, 21); printf("%d", a); truelon = truex(a); truelat = truey(a); truedepth = truez(a); distance_P = distance(x, truelon, y, truelat); if (distance P == 0) {
vdepth = truedepth; wdepth = 1; break;
} if (truelon == x) {
y_R = minor * sin ((90 - azim) * 3.14159 / 180); x_R = mayor * cos((90 - azim) * 3.14159 / 180);
else {
m_P = (truelat - y) / (truelon - x); x_R = mayor * cos((azim * 3.14159 / 180) + atan(m P) y_R = minor * sin((azim * 3.14159 / 180) + atan(mj?)
distance_R = distance(0, x_R, 0, y_R) ; if (distance P <= distance R ) {
vdepth += truedepth/(pow(distance_P, power)); wdepth += 1/(pow(distance_P, power));
if (vdepth == 0 & wdepth = 0) value = 0; else value = vdepth / wdepth; gotoxy(14, 23);
printf("%d",pointh); gotoxy(2 6, 23); printf("%7.2f", x); gotoxy(42, 23); printf("%7.2f", y); gotoxy(58, 23); printf("%5.2f", value); saving(x, y, value); }
} } //estimate z based on closest point values void closestpoint() {
int numpoint, search; float dist P, //distance between estimated and known point x, y, trueTon, truelat, truedepth, vdepth, wdepth, value; clrscr();
CONTINUED subheading2(); getfilename(); readmaxmin(); prtmaxmin(); grdspc (); repeatcalll: gotoxy(52, 16); search = toupper(getche ());
/* if(search != 'F' I search != 'Q' I search != '0') {
beep(); goto repeatcalll;
} */ repeatcall2:
gotoxy(37, 17) ; cin » numpoint; if(numpoint >= 64 & search == 'F') {
notes(); beep(); goto repeatcall2;
} if(numpoint >= 16 & search == 'Q') {
notes (); notes(); goto repeatcall2;
} if(numpoint >= 8 & search == '0') {
notes (); beep (); goto repeatcall2;
} readdata (); switch(search) {
case 'Q': quadrantfind(numpoint); break;
case '0': octantfind(numpoint); break;
default : freefind(numpoint); }
} void freefind(int numpoint) {
int pointh = 0; float truelon, truelat, truedepth,
x, y, dist_P, vdepth, wdepth, value; —numpoint;
for(a = 0; a <= numpoint; ++a) { truelon = truelat = truedepth = 0; truelon = truex(a); truelat = truey(a); zsort[a] = truez(a); distsortfa] = distance(x, truelon, y, truelat);
} sorter(0, numpoint);
forty = ymin; y <= ymax; y += y spc) {
for(x = xmin; x <= xmax; x += x spc) { ++pointh; vdepth = wdepth = value = 0; for(a = 1; a <= i; ++a) { truelon = truelat = truedepth = 0; gotoxy(17, 21); printf("%d", a); truelon = truex(a); truelat = truey(a);
CONTINUED truedepth = truez (a) ; ^!li-P = distance (x, truelon, y, truelat); if(dist_P ==0)
vdepth = truedepth; wdepth = 1;
break; } _ if(dist_P <= distsort[numpoint])
distsort[numpoint] - dist_P; zsort[numpoint] = truedepth;
sorter(0, numpoint);
if(dist P != 0) { for(a = 0; a <= numpoint; ++a)
vdepth += zsort[a]/(pow(distsort[a], power)); wdepth += 1/(pow(distsort[a], power));
} value = vdepth / wdepth; gotoxy(14, 23); printf("%d",pointh); gotoxy(26, 23); printf("%7.2f", x); gotoxy(42, 23); printf("%7.2f", y); gotoxy(58, 23); printf("%5.2f", value); saving(x, y, value); }
} return;
}
II-
//estimate z based on quadrant method void quadrantfind(int numpoint) {
int sect[4]; int num, numr, numint; int pointh = 0; float phi4 = 90; float truelon, truelat, truedepth,
x, y, dist P, vdepth, wdepth, value; float deltax, deltay; float angle, numfloatl, numfloat2;
—numpoint; for(y = ymin; y <= ymax; y += y_spc) {
for(x = xmin; x <= xmax; x += x_spc) { ++pointh; vdepth = wdepth = value = 0; sect[0] = sect[l] = sect[2] = sect[3] = 0; for(a = 1; a <= i; ++a) { truelon = truelat = truedepth = 0; gotoxy(17, 21); printf("%d", a); truelon = truex(a); truelat = truey(a); truedepth = truez(a); dist P = distance(x, truelon, y, truelat); if(dist P == 0) {
vdepth = truedepth; wdepth = 1; break;
} deltax = (truelon - x); deltay = (truelat - y); if(deltay >= 0)
(
CONTINUED int sortl, sort2; float opsortd, opsortz;
for(sortl = begin; sortl <= end; ++sortl) {
for(sort2 = sortl; sort2 <= end; ++sort2) { if(distsort[sortl] >= distsort[sort2]) { opsortd = distsort[sortl]; opsortz = zsort[sortl]; distsort[sortl] = distsort[sort2] ; distsort[sort2] = opsortd; zsort[sortl] = zsort[sort2]; zsort[sort2] = opsortz;
} }
} return;
} //print heading void heading(void) {
printf(" printf(" printf(" printf(" printf(" printf(" printf(" printf(" printf(" printf(" printf(" printf(" printf(" printf(" printf(" printf(" printf(" printf("\n"); printf("
}
GRIDDING PROGRAM USING INVERSE
SEARCHING METHOD:
1. Area 2. Closest Points
3. Quit
Note: <> Data must be in ASCII form ->
DISTANCE METHOD
X Y Z.
Press the option key ")
//print subheadingl (area) void subheadingl(void) {
gotoxy(20, 1); printf("INVERSE DISTANCE
METHOD (AREA SEARCH)\n\n")
printf(" File name of data to be retrieved : \n"); printf(" File name of gridded data : \n\n"); printf(" The program found that maximum of: \n"); printf(" X = and Y =\n"); printf(" minumum of:\n"); printf(" X = and Y =\n\n"); printf(" Grid spacings on x-axis : , and on printf(" Top left coordinate: X = printf(" Bottom right coordinate: X = printf(" Number of point to be estimated =\n"); printf(" Weighting, power of distance :\n"); printf(" Searching pattern (R)ectangular or (E)llipsoidal printf(" Length of minor axis in unit :\n"); printf(" Length of mayor axis in unit :\n"); printf(" Azimuth of minor axis (± 90°, integer):\n\n"); printf(" Calculation ->\n"); printf(" Results:\n"); printf(" point no. = ; X = ; Y =
y-axis :\n"); Y =\n"); Y =\n");
\n" \n" \n" \n" \n" \n" \n" \n" \n" \n" \n" \n" \n" \n" \n" \n" \n"
An");
} //====
Z =");
//print subheading2 (closest point) void subheading2(void) (
gotoxy(20, 1); printf("INVERSE DISTANCE METHOD
(CLOSEST POINT)\n\n")
u = 0;
CONTINUED v = 0; for(q = ymin; q <= ymax; q += y_spc)
++v; U = 0; for(p = xmin; p <= xmax; p += x_spc)
++u; ++pointnum; }
} gotoxy(37, 14); printf("%d", pointnum); gotoxy(33, 15); cin » power; return;
} float xrange(float m_xrange, float n_xrange, float me_xrange, float ne_xrange) return((n_xrange - ne_xrange)/(me_xrange - m_xrange));
float yrange(float m_yrange, float n_yrange, float me_yrange, float ne_yrange)
return(((((n_yrange - ne_yrange)/(me_yrange - m yrange)) * me yrange) + ne_yrange)); } //===========================================================================
//read file names void getfilename(void) {
gotoxy(37, 3); gets(inpfile); gotoxy(37, 4); gets(outfile);
} //save results' void saving(float absis, float ordinate, float height) {
file_out = fopen(outfile, "a"); fprintf(file_out,"%7.2f %7.2f %5.2f\n", absis, ordinate, height); fclose(file_out); return;
} //calculate the real values of x float truex(int idx) {
float valuex; valuex = (lon[idx] / 10) + xminest; return(valuex);
} // //calculate the real value of y float truey(int idx) {
float valuey; valuey = (lat[idx] / 10) + yminest; return(valuey);
} // //calculate the real value of z float truez(int idx) {
float valuez; valuez = (depth[idx] / 10) + zmin; return(valuez);
//sort array to ascending order void sorter(int begin, int end)
ymin = ymin;
CONTINUED fscanf(file_in,"%f\n", Szaxis); if (zmax <= zaxis) zmax = zaxis ; else zmax = zmax ; if (zmin >= zaxis) zmin = zaxis ; else zmin = zmin;
} xminest = xmin; yminest = ymin; fclose(file_in); return;
} _ void readdataO {
float xaxis, yaxis, zaxis, mayor; gotoxy(4, 20); printf("Extend area to : " ) ; cin » mayor; i = 1; file_in = fopen(inpfile, "r") ; while(!feof (file in)) { fscanf(file_in,"%f\n",&xaxis); fscanf(file_in,"%f\n",&yaxis); fscanf(file_in,"%f\n",&zaxis); if (xaxis <= (xmax + mayor) & yaxis <= (ymax + mayor) & xaxis >= (xmin - mayor) & yaxis >= (ymin - mayor)) { lon[i] = (xaxis - xminest) * 10; lat[i] = (yaxis - yminest) * 10; depth[i] = (zaxis - zmin) * 10; i++;
} } fclose(file_in); return;
} _ //print maximum and minimum values on screen void prtmaxmin(void) {
gotoxy(8, 7); printf("%7.2f", xmax); gotoxy(28, 7); printf("%7.2f", ymax); gotoxy(8, 9); printf("%7.2f", xmin); gotoxy(28, 9); printf("%7.2f", ymin);
} //=
//read and calculate grid spacing void grdspcO {
int pointnum; int u, v; float p, q; gotoxy(28, 11) cin » x_spc; gotoxy(51, 11) cin » y spc; gotoxy(3T, 12) cin >> xmin; gotoxy(47, 12) cin » ymin; gotoxy(31, 13) cin » xmax; gotoxy(47, 13) cin » ymax; pointnum = 0;
CONTINUED
else
{
{ if(dist_P <= distsort[num*8+numpoint])
distsort[num*8+numpoint] = dist P; zsort[num*8+numpoint] = truedepth;
sorter(num*8, num*8+numpoint) ;
}
} " if(dist P != 0) { for(num = 0; num <= 7; ++num) { for(a = 0; a <= sect[num]; ++a)
vdepth += zsort[num*8+a]/(pow(distsort[num*8+a] , power)); wdepth += 1/(pow(distsort[num*8+a], power));
} } value = vdepth / wdepth; gotoxy(14, 23); printf("%d",pointh); gotoxy(26, 23); printf("%7.2f", x); gotoxy(42, 23); printf("%7.2f", y); gotoxy(58, 23); printf("%5.2f", value); saving(x, y, value); }
//sub-routine for distance calculation float distance(float xl, float x2, float yl, float y2) {
float x = x2-xl; float y = y2-yl; float xy = pow(x,2) + powfy, 2); return(sqrt(xy));
} //find maximum and minimum values of data void readmaxmin() {
float xaxis, yaxis, zaxis; file_in = fopen(inpfile, "r"); xmin = 648000; ymin = 648000; while(!feof(file in)) { fscanf(file_in, "%f\n",&xaxis); if (xmax <= xaxis) xmax = xaxis; else xmax = xmax ; if (xmin >= xaxis) xmin = xaxis ; else xmin = xmin ;
fscanf(file_in, "%f\n", &yaxis); if (ymax <= yaxis) ymax = yaxis ; else ymax = ymax ; if (ymin >= yaxis) ymin = yaxis ; else
{
CONTINUED vdepth += zsort[num*16+a]/(pow(distsort[num*16+a], power)); wdepth += 1/(pow(distsort[num*16+a], power));
} }
} } value = vdepth / wdepth;
// gotoxy(51,1); // printf("vdepth=%f; wdepth=%f",vdepth,wdepth);
gotoxy(14, 23); printf("%d",pointh); gotoxy(2 6, 23); printf("%7.2f", x); gotoxy(42, 23); printf("%7.2f", y); gotoxy(58, 23); printf("%5.2f", value); saving(x, y, value); }
} } //estimate z based on octant method void octantfind(int numpoint) {
int sect[8], num; int pointh = 0; float phi8 = 0.78539; float truelon, truelat, truedepth,
x, y, dist_P, vdepth, wdepth, value; float tann; double angle;
—numpoint; for(y = ymin; y <= ymax; y += y spc) {
for(x = xmin; x <= xmax; x += x spc) { ++pointh; vdepth = wdepth = value . = 0; for(a = 1 ; a <= i; ++a) { truelon = truelat = truedepth = 0; gotoxy(17, 21);
printf("%d", a); truelon = truex(a); truelat = truey(a); truedepth = truez(a);
dist P = distance(x, truelon, y, truelat); if (dist P == 0)
{ vdepth = truedepth; wdepth = 1; break;
} if(truelon == x) {
angle = phi8; } else {
tann =(truelon - x) / (truelat - y); angle = atan(tann);
} for(num = 0; num <= 7; ++num) { if(angle >= (num * phi8) & angle < ((num+1) * phi8)) if(sect[num] < numpoint) { distsort[num*8+sect[num]] = dist_P; zsort[num*8+sect[num]] = truedepth; ++sect[num]; }
(
CONTINUED
if(deltax >= 0) {
angle = fabs((float)atan(deltay/deltax));
else {
angle = 180 - fabs((float)atan(deltay/deltax));
} else {
if(deltax >= 0) {
angle = 180 + fabs((float)atan(deltay/deltax)) ; else
{ angle = 360 - fabs((float)atan(deltay/deltax)) ;
} // angle = angle * 180 / 3.14158; // gotoxy(1,1); // printf("angle=%f",angle) ;
for(num = 0; num <= 3; ++num) {
if(angle >= (num * phi4) & angle < ((num+1) * phi4)) // printf("dist=%5.2f; angle=%5.2f",dist_P,(float)angle); // gotoxy(1,1); // printf("sectnum: %d",sect[num]);
if(sect[num] < numpoint) { distsort[num*16+sect[num]] = dist_P; zsort[num*16+sect[num]] = truedepth;
// printf("dist=%f; z=%f",dist P,truedepth); // printf("dist=%f; z=%f", distsort[num*16+sect[num]],zsort[num*16+sect[num] ] ) ;
++sect[num]; } else { sorter(num*16, num*16+numpoint) ;
// for(numr = 0; numr <= numpoint; ++numr) // { // numfloatl = fabs(distsort[num*16+numr] - dist P) ; // // if(dist_P <= distsort[num*16+numpoint])
if(dist_P <= distsort[num*16+numpoint]) {
distsort[num*16+numpoint] = dist_P; zsort[num*16+numpoint] = truedepth;
// gotoxy(1,1); // printf("dist=%f; z=%f",dist_P,zsort[num*16+numpoint]); // gotoxy(l,num+2); // printf("%3.1f-%3.If: angle: %f; D: %f; Z: %f",(num * phi4),((num+1) * phi4), // (float)angle,dist_P,truedepth); } // sorter(num*l6, num*16+numpoint);
} }
} } if(dist P != 0) { for(num = 0; num <= 3; ++num) {
for(a = 0; a <= sect[num]; ++a) {
// gotoxy (1, num+1) ; ^r nv
// printf("D: %f; Z: %f",zsort[num*16+a],distsort[num*16+a]); // printf("arr: %d", num*16+a);
if (zsort[num*16+a] != 0 & distsort[num*16+a] != 0)
CONTINUED printf(" File name of data to be retrieved : \n"); printf(" File name of gridded data : \n\n"); printf(" The program found that maximum of: \n"); printf(" X = and Y =\n"); printf(" minumum of:\n"); printf(" X = and Y =\n\n"); printf(" Grid spacings on x-axis : , and on y-axis :\n"); printf(" Top left coordinate: X = Y =\n"); printf(" Bottom right coordinate: X = Y =\n"); printf (•"-Number ,of"-p'0~iTjtr *to~ be estimatech -An-"')7 —•-•••• printf(" Weighting, power of distance :\n"); printf(" Searching pattern: (F)ree, (Q)uadrant, (O)ctant :\n"); printf(" Number of point in each sector :\n\n\n\n"); printf(" Calculation ->\n"); printf(" Results:\n"); printf(" point no. = ; X = ; Y = ; Z =");
} //print notes on error void notes(void) {
gotoxy(5, 17); printf("Maximum number of point on:"); gotoxy(5, 18); printf("Free search is 64, Quadrant search is 16 and Octant search is 8"); delay(1500); gotoxy(5, 17); printf(" " ) ; gotoxy(5, 18); printf(" " ) ;
} //sound a beep on error void beep(void) {
sound(880); delay(lOO); nosound(); delay(lOO); sound(440); delay(lOO); nosound();
Volume 2 of this thesis contains seismic sections that are too large to be scanned. The
seismic sections can be viewed in the Archives of the University of Wollongong Library.