Source: https://mem.lyellcollection.org/content/48/1/81?ijkey=606430a628f6eb73e57efff8095f5039679fb149&keytype2=tf_ipsecsha
Timestamp: 2019-04-21 13:26:17+00:00

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Since the Early Miocene, basement uplift of the Mogok Belt in the east and the Gangaw Ranges of the upper Irrawaddy Basin in the western part of the Asian Plate, as well as the initiation of river systems, supplied orogenic detritus to the south (Allen et al. 2008; Kyi Khin et al. 2014; Robinson et al. 2014; Licht et al. 2015). Middle Miocene sandstones from both areas are rich in sub-angular monocrystalline quartz, chert and argillites with small amounts of feldspar and metamorphic lithic fragments, suggesting a relatively low supply of sediment from the Himalayas. In the early Middle Miocene (16 Ma), widespread marine transgression occurred. Post-Early Miocene dextral displacement along faults in the western part of Asian Plate possibly contributed to subsidence, which led to a subsequent rise in sea level and the accumulation of thick Middle Miocene sequences in these areas. Upper Miocene sandstones contain abundant metamorphic lithic fragments and fresh feldspars, both orthoclase and plagioclase, with volcanic rock fragments, suggesting continued orogenic unroofing in the source areas and the exhumation of younger granites in the eastern Himalayas and the Burmese Arc (Kyi Khin 2000; Kyi Khin et al. 2014).
The collisional history of the Himalayas is recorded in the sediments deposited in subsiding foreland basins to the south (Jordan 1995; Burbank et al. 1996), including the Bengal Basin of Bangladesh (Uddin & Lundberg 1998a) and the Bengal Fan (Curray 2005). Since collision in the Late Cenozoic, an arc-trench system containing sedimentary basins was developed in northwestern Myanmar. Various models have been proposed for the tectonic evolution of the Himalayas and the Tibetan Plateau, related to Late Cenozoic foreland and foredeep sedimentation (Copeland & Harrison 1990; Burbank et al. 1993; Harrison et al. 1993; Sorkhabi & Stump 1993; Beck et al. 1995; Harris 1995; Yin et al. 1999; Uddin et al. 2007; Allen et al. 2008; Licht et al. 2013, 2015; Naing et al. 2013).
The results from the present study place constraints on these models by providing important information linking syntectonic sedimentation in proximal deep-marine foreland basins, with episodic uplift and erosional fluxes from the Indian–Asian collision (Kyi Khin et al. 2014).
Regarding the timing of the Indo-Asian collision, many papers have been published from different regions: the western and eastern Himalayas (Searle et al. 1987, 2007; Beck et al. 1995; Yin et al. 1999); Siwalik Ranges and Nepal (Harrison et al. 1993; Sakai 1997; De Celles et al. 1998); Bengal Basin of Bangladesh (Johnson & Alam 1991; Uddin & Lundberg 1998a, b); Bengal Deep-Sea Fan (Curray & Moore 1974) and on two drilling legs DSDP Leg 22 and ODP Leg 116 (Ingersoll & Suczek 1979; Cochran & Stow 1989; Cochran 1990; Copeland et al. 1990; Amano & Taira 1992; France-Lanord et al. 1993); and the northern Ninetyeast Ridge on ODP Leg 121 (Klootwijk et al. 1992a, b).
However, geochemical indicators in the sedimentary record for the unroofing of the Himalaya have not been tested with even the most basic provenance and related sedimentological data from the distal part of foreland basin deposits, or by considerations of the regional tectonic evolution of Himalayan thrust belts. A more straightforward approach is to use sequence stratigraphic analysis to examine the character of deposition and the sequential evolution related to the rate of sediment accumulation in the basins containing the Himalayan detrital record.
On-land geological information from the Miocene Arakan Basin, west of the Indo-Myanmar Ranges, has not yet been well documented however, and the relationship between the sedimentary evolution of the foreland basin containing both pretectonic and syntectonic sequences, recording the unroofing history of the eastern Himalayas and Indo-Myanmar Ranges, is poorly understood. Many preliminary studies (Aung Khin & Kyaw Win 1966, 1967; Brunnschweiler 1974) have suggested that the provenance of the Miocene–Pliocene sequence was from an ancestral Brahmaputra River system to the north and from the Indo-Myanmar Ranges to the east.
The tectonics of northeastern India and northwestern Myanmar are controlled by the interaction between north–south-aligned plate convergence in the Himalayas, and east–west-aligned plate convergence in the Indo-Myanmar Ranges, a 230 km wide active orogenic belt associated with eastwards subduction of the Indian Plate beneath the Burma Plate (Le Dain et al. 1984; Sengupta et al. 1990; Johnson & Alam 1991). The northernmost extension of the Indo-Myanmar Ranges is the Chitagaung Fold Belt (Fig. 5.1), which merges with the west-trending Himalayas in the Eastern Himalayan Syntaxis. The Main Himalayan Boundary Thrust Fault, initiated in the Late Miocene or Pliocene (Le Fort 1996; Sengor & Natal'in 1996), forms the northern margin of the Himalayan foredeep which is narrow in the eastern Himalayas. The foredeep is separated from the Bengal Basin by the Shillong Plateau. After the collision the Indo-Myanmar Belt was accreted to the Indian craton margin (Packham 1996).
Tectonic divisions of western Myanmar and associated foreland basins in the eastern India region. Blue patches on the boundary between the Eastern and Western troughs are Quaternary volcanoes. Location of area studied (Fig. 5.2) shown as bold rectangular box. Dashed lines represent palaeo-shelf boundaries. Numbers represent the location of basins in the stratigraphic table Figure 5.3 (modified after Kyi Khin et al. 2014).
In their eastern part, the Indo-Myanmar Ranges represent the northern continuation of the Andaman–Nicobar arc, where the Indian Ocean floor is being subducted beneath the SE Asian continent (Le Dain et al. 1984). Tectonic evolution has included folding and westwards-directed overthrusting with a probable dextral component of movement in the western part of the Indo-Myanmar Ranges, attributed to oblique convergence between the Indian Plate and the Asian Plate (Curray et al. 1978; Hamilton 1979; Mitchell 1981). According to Mitchell (1981, 1993) thrusting was accompanied by uplift of the Indo-Myanmar Ranges, and the timing corresponds to the Eocene–Oligocene unconformity in the western trough of the Myanmar Central Tertiary Basin.
The study area is situated in the westernmost part of Myanmar, bordered by the Bay of Bengal to the west and separated from the Arakan Yoma (Indo-Myanmar Ranges) by a narrow coastal strip. The northern continuation is located in the Chittagong Hills, and lying 100 km SE are the offshore islands of Ramree (Yanbye) and Cheduba (Manaung) (Figs 5.1 & 5.2).
Location of the study area, Sittwe, Pauktaw and Baronga Islands, Arakan Coastal Ranges, western Myanmar (modified map of Earth Sciences Research Division 1977). Dashed line represents the political boundary between Bangladesh and Myanmar.
This account reports the results of an integrated study of Miocene sediments on the Baronga Islands on the eastern margin of the Arakan Basin using sequence stratigraphy and provenance studies. Data from petrography and the chemical composition of sandstones and shales are used to determine the effects of weathering and the relationship between tectonic denudation of Himalaya orogenic belts and reciprocal sedimentation related to foreland thrusting in the Himalaya–Bengal system. The sedimentary evolution of a clastic shelf-slope system, related to forced regression (Hunt & Tucker 1992, 1995; Hart & Long 1996) in the Early–Middle Miocene and Late Miocene deltaic progradation, are discussed on the basis of vertical and lateral correlation of the measured sections. These data are used to constrain the link between deltaic sedimentation in the Bengal Basin of Bangladesh and the deep-sea fan sedimentation of the Bengal Deep-Sea Fan where, since Early Miocene time, sediment was trapped in a remnant ocean basin between the Indian and Asian plates.
In order to fill the gap between geochemical and sedimentological evidence, the present study presents data from Early Miocene to Late Miocene–Pliocene siliciclastic sequences deposited in the slope and shelf of the Bengal Deep-Sea Fan System. The data comprise sequence stratigraphic analysis, including sedimentological interpretation based on the well-established age of geological units, modal petrographic data, geochemical provenance studies including the palaeoweathering index, and U–Pb–Th monazite and thorite dates from selected sandstones.
The Miocene Arakan Basin was bordered by the Arakan–Chin–Naga Ranges or Indo-Myanmar Ranges (Brunnschweiler 1974; Bender 1983) to the east and the Bengal Basin (Uddin & Lundberg 1998a; Zahid & Uddin 2005) to the west and NW. This area extends southwards into the Bay of Bengal, and the Bengal Fan continues into the NNW–SSE-trending Chittagong Hills to the immediate north and the Shillong Plateau and Himalayas to the distant north (Fig. 5.1). The Neogene units of the western Indo-Myanmar Ranges occupy the Arakan Coastal Ranges and northern portion of the Indo-Burmese wedge, which is the onshore prolongation of the Andaman–Nicobar Trench (Nielsen et al. 2004). Within the Indo-Myanmar Ranges there are north–south-running faults such as Kaladan Fault, which separates the Palaeogene accretionary complex and the Neogene accretionary prism, and additional east-dipping thrusts of an accretionary prism (Sikder & Alam 2003) with a significant dextral strike-slip component (Nielsen et al. 2004). In the eastern part of Indo-Myanmar ranges, the Kabaw Fault, which is regarded as reverse fault but with a strong right-lateral component (Pivnik et al. 1998), puts the Palaeogene metamorphic rocks into contact with the Upper Cretaceous–Tertiary sediments of the Central Myanmar Basin.
Sedimentary rocks of the Bengal Basin to the west of the study area are underlain by an oceanic basement of Late Jurassic–Early Cretaceous basaltic lava flows (Curray et al. 1982; Curray 1994). Further north the Shillong Plateau (Fig. 5.1), whose southern margin is defined by the Dauki Fault, is composed of a Precambrian basement complex of Archaean gneiss with minor greenstones and Upper Proterozoic granites (Acharyya et al. 1986; Acharyya 1998). The crystalline rocks are almost completely surrounded by Tertiary and Quaternary sedimentary deposits (Mitchell 1981; Curray et al. 1982; Chen & Molnar 1990) that dip to the south, forming a monoclinal structure. To the north, the Dauki Fault may curve into the Haflong-Disang Thrust Fault (Evans 1964; Chen & Molnar 1990; Johnson & Alam 1991). The elevation of the Shillong Plateau seems to be supported by a rigid lithosphere, as would be the case if the Indian Plate were thrust beneath it (Chen & Molnar 1990). Major uplift of the Shillong Plateau began in the Pliocene, as there is no record of significant Pliocene–Pleistocene deposition on the Shillong Plateau (Johnson & Alam 1991).
To the east, the Indo-Myanmar Ranges consist of thick deep-marine sediments, mainly indurated dark grey slaty shales, with thin interbeds of sandstone and limestone. Exotic blocks of Upper Cretaceous white, porcellaneous, lithographic limestone are found at many localities within Eocene black shales (IGCP 1978). The Indo-Myanmar Ranges are divided into two belts: the Western Belt, comprising Cretaceous–Eocene sedimentary rocks; and the Eastern Belt, consisting of schists formed from Ladinian–Carnian turbidites (Myint Lwin Thein 1972; Gramann 1974; Mitchell 1993).
The Eastern Belt is overthrust locally by serpentinized harzburgites, with pillow lavas and hornblende gabbros. In the Western Belt of the Indo-Myanmar Ranges, Miocene sediments rest unconformably on Eocene rocks along the west coast of Arakan (Yakhine) (IGCP 1978; Bannert et al. 2011).
During Oligocene time the Indo-Myanmar Ranges were periodically elevated; subsequently, the Arakan (Yakhine) to the west and the Central basins to the east were formed on the flanks of the Indo-Myanmar Ranges (Brunnschweiler 1974). The low-angle boundary between the schists (Triassic) of the Mt Victoria Dome in the eastern part of the Indo-Myanmar Ranges (Fig. 5.2) and the Triassic turbidites further to the east suggest that uplift may have been accompanied by low-angle extensional faulting, resulting in the Western Trough sequence which may be due to either Middle Eocene east-verging thrusts or faults related to extension (Mitchell 1993).
The western margin of the northern Indo-Myanmar Ranges is bounded by thrust faults along the Naga Hills in the north, and folds and thrusts in the Bengal Basin to the south (Chen & Molnar 1990). According to evidence from seismic data, especially in the northern Indo-Myanmar Ranges and uplifted marine Quaternary terraces in the Arakan (Yakhine) coastal areas, further compressional phases then narrowed the flanking basins during the Quaternary orogenic movements, which are still active. A stratigraphic scheme for the Bengal and Assam regions (Kyi Khin et al. 2014) is shown in Figure 5.3, and the Cretaceous–Holocene stratigraphy of Myanmar (Aung Khin & Kyaw Win 1966) is shown in Figure 5.4.
Stratigraphic scheme of the Bengal and Assam regions (Kyi Khin et al. 2014).
Cretaceous–Holocene stratigraphy of Myanmar, after Aung Khin & Kyaw Win (1966).
The termination of the eastwards subduction of the Indian Ocean Plate is inferred to have occurred in the latest Eocene in northeasternmost Myanmar and Assam, due to the collision of the Indian continent with the Asian Plate (Mitchell 1993). This event is typically represented by underthrusting of the Middle and Upper Eocene Sylhet Limestone which was overlying on the western margin of Indian Continental Plate, beneath the Lower Eocene Disang Formation (Fig. 5.5). India collided with northern Myanmar and Tibet in the latest Eocene, along a zone south of the Triassic–Jurassic deep-marine sediments and the Yarlung Ophiolite (Mitchell 1993). Later movements of the Naga and Disang thrusts to the east suggest further convergence between the Burma and Indian plates as both moved north. India rotated anticlockwise after the collision. Stratigraphic records also indicate an Early Oligocene emplacement of the deep-marine sedimentary terrain on the Indian Plate, probably within 10 Ma of the India–Asia collision, and further northwards movement of the craton since that time (Packham 1996; Zhang et al. 2012).
Lithostratigraphic correlation of Neogene units in Arakan (Yakhine) Coastal Ranges and Surma Valley (Bangladesh), modified after Aung Khin & Kyaw Win (1966), Rao (1983) and Uddin & Lundberg (1998a).
North-trending folds, uplifted in the Chittagong-Tripura Fold Belt, plunge northwards into the Sylhet Trough (Johnson & Alam 1991) and continue into the NNW–SSE-trending Arakan (Yakhine) Coastal Ranges which occupy the western part of the Indo-Myanmar Ranges in the south. Folding within the Indo-Myanmar Ranges occurred throughout the Late Cenozoic, indicating east–west crustal shortening (Evans 1964; Brunnschweiler 1974; Mitchell & McKerrow 1975).
In the Early Miocene, oblique convergence between India and the SE Asian Plate probably sheared off the leading edge of the SE Eurasian Plate along the Sagaing Fault to form the Burma Platelet (Aung Khin & Kyaw Win 1966; Win Swe 1972; Curray et al. 1978; Mitchell 1981; Myint Thein et al. 1991; Kyi Khin & Myitta 1999; Bertrand & Rangin 2003), bounded by the Sagaing Fault to the east and either the Kabaw Fault Zone (Hla Maung 1987) and/or a foreland thrust to the west (Ni et al. 1989). Plate convergence could also have resulted from the clockwise rotation of SE Asia relative to the Indian Plate, accompanying the eastwards translation of China along the strike-slip faults described by Molnar & Tapponnier (1975). Dasgupta & Nandy (1995) termed this fold belt a ‘Neogene Accretionary Prism’ which was formed by the continuous westwards migration of the Indo-Myanmar accretionary complex. Later, compressive wrench tectonics, the result of the convergent oblique movement of the Indian Plate, also influenced the structural style of the Chittagong Fold Belt and the northern Arakan Coastal Ranges.
In Assam, the pre-Miocene sediments of the Bengal Fan and the underlying continental rise sediments of the Indian margin are considered to have been underthrust beneath the Indo-Myanmar Ranges (Curray et al. 1978; Curray 1994), where the Surma Group (Fig. 5.4) can be regarded as the oldest fan sediments deposited on ocean floor to the south. During the late Pliocene–Holocene, in response to eastwards-directed subduction and northwards interplate movement of the Indian Plate beneath the western Indo-Myanmar Ranges, elongate north–south to NNW–SSE asymmetrical anticlinal structures were formed in the Chittagong-Tripura Fold Belt and in the Arakan (Yakhine) coastal regions. The growth of these folds and associated uplift may have had a significant effect on the course of the Brahmaputra River (Lindsay et al. 1991). Because of the aseismic nature of the Indo-Myanmar Ranges and the lack of interplate events, Guzman-Speziale & Ni (1996) proposed that the Burma Plate was wedged in the Burma subduction zone and was dragged mechanically northwards together with the Indian Plate.
Since Pliocene or Pleistocene time, both the Neogene Arakan Basin and Central Basin were affected by transpression followed by a long period of extension, which led Pivnik et al. (1998) to suggest that northwards motion of the Burma sliver was resisted by the eastern Himalayas following a buttressing effect and dextral shearing, then migrated westwards, reactivating major discontinuities between Miocene and Pliocene units found in the Arakan Yoma (Beck et al. 1995). Oblique motion of the India Plate is progressively transferred from a remote fault system, which formed the Andaman Sea Basin and the eastern margin of the Central Basin (i.e. Sagaing Fault), to a close to trench fault system (Indo-Myanmar Ranges and the trench itself).
The stratigraphy of the Miocene Arakan Basin was initially established in the 1940s in the Assam and Arakan (Yakhine) oil regions by lithostratigraphic correlation with type sections in the Surma Valley, Assam and northeastern India (Sale & Evans 1940; Evans 1964). Brunnschweiler (1974) investigated selected areas of the Indo-Myanmar Ranges and considered the outer molasse basin of the Arakan coastal region from the Indo-Myanmar Ranges as a separate structural zone, adopting the stratigraphic nomenclature proposed by Evans (1964) for the Assam and Surma regions. According to Brunnschweiler (1974), the folded Eocene flysch sequences thin out beneath thick Lower Miocene mudstones in the Arakan coastal region, and the Barails Series (Oligocene) underlies the Miocene Surma Group unconformably in the Surma Basin and in the Chin Hills in northern regions (Fig. 5.3).
The Miocene Arakan Basin is the narrow southern continuation of the Assam Basin, with a lithology similar to that of the Central Tertiary Basin of Myanmar (Fig. 5.4). As proposed by Stamp (1922), these two basins may have formed as twin troughs or twin gulfs divided by the Indo-Myanmar Ranges, a mobile belt or a trench complex, which has undergone intense orogenic movements as shown by large-scale westwards overthrusts and tight folds in the flysch units. The Indo-Myanmar Ranges consist mainly of Cretaceous–Eocene pelagic sediments, overlain by thick Eocene–Oligocene turbidites and Miocene–Pliocene molasse (Brunnschweiler 1974; Ni et al. 1989; Bannert et al. 2011).
The sedimentary prism of the northern Arakan coastal region is composed of a thick section of Cretaceous–Eocene pelagic shales, claystones and greywackes (Aung Khin & Kyaw Win 1966; Brunnschweiler 1974). The section is located in the foothills of the Arakan Yoma, east of the study area, and is overlain by 4500 m of a mostly shallow- to deep-marine Miocene–Pliocene succession of sandstones, siltstones with thick pelagic shale intervals.
The previous classification and correlation of Miocene–Pliocene rocks in the Miocene Arakan Basin in coastal areas was based mainly on lithological datasets with sparse palaeontological information (Aung Khin & Kyaw Win 1966, 1967; Brunnschweiler 1974; Than Nyunt & Chit Saing 1978). A composite stratigraphic succession for the Miocene–Pliocene is shown in Figures 5.4 and 5.5. In the Miocene–Pliocene stratigraphic sequence compiled by Aung Khin & Kyaw Win (1966), five major units were defined. They include the Laung Formation (Lower Miocene), Yezaw Formation (Middle Miocene), Mayu Formation (Upper Miocene), Ngasanbaw Formation (Upper Miocene–Pliocene) and the Kyauktan Formation (Pliocene) (Figs 5.5 & 5.6). Most of the type sections are in the Mayu Range (northwestern part of Sittwe) and the Baronga Islands (Fig. 5.2).
Litho- and biostratigraphy and palaeoenvironments of Miocene sequences in Sittwe, Pauktaw and Baronga islands, Arakan Coastal Ranges, western Myanmar. Slump beds constitute folded and slumped shale containing disrupted minor thin sandstone (modified after Kyi Khin et al. 2014).
Rocks of the Laung Formation are exposed mainly in Sittwe Point, Sittwe Township and west Baronga Island, where it is made up of thick shales (about 200 m) with isolated sandstone blocks and localized conglomerate layers (only in Laychat Taung offshore island). The type section is along the Laung Chaung, a tributary of the Kaladan River (20° 39′ N, 92° 53′ E). It consists of lower (ML1), middle (ML2) and upper (ML3) members. The maximum thickness of the Laung Formation is c. 750 m. Thick shales (100–250 m) with slump beds and massive sandstones are also exposed in the core of anticline in East Baronga Island and the western part of Middle Baronga Island. Shales are blue-grey, micaceous and occasionally silty and sandstones are brown to grey, fine, medium- to coarse-grained and locally gritty in the upper part of conglomerate layers.
In the lowermost member (ML1) isolated sharp-based sandstone blocks (1.5–5.0 m thick) with dish and wet-sediment deformation structures overlie slumped shales and silty shales with a sharp contact. Shales with thin sandstone layers occur in the middle part of channelized sandstone bodies. Thin- to medium-bedded sandstones with shales are most abundant in the middle member (ML2). Bioturbations of mostly horizontal traces are abundant in these interbedded layers (Fig. 5.7). Trace fossils, mainly Lophoctenium, Radionereites, Neonereites, Dendrotichnium, Protopaleodictyon and (?)Spirodesmons, are present (Frey & Howard 1985, 1990). Most of these ichnospecies are interpreted as mud- to fine-sand dwellers in the epibathyal zone at about 200–1000 m depth (Ksiazkiewicz 1977; Seilacher 1977; Pemberton et al. 1992).
Laung Formation (Lower Miocene). (a) Basal contact features of the amalgamated sandstone bodies in the middle Laung Formation (ML2), at the southern tip of West Baronga Island. Irregularly stepped basal contact (lined) associated with deformation of immediately underlying homogenized mud unit (Hm) in the chaotic mud facies. Note the geological hammer as scale, circled. FA.3, facies association 3, chaotic beds, FA.4, facies association 4, amalgamated sandstones. (b) Slide scar (Sc) consisting of undeformed amalgamated sandstone and shale alternations, underlain by homogenized mudstones (Hm) in the upper Laung Formation (ML3) along the east coast of West Baronga Island. Note the 1 m scaled pole. (c) Channelized lenticular sandstone bodies (CS) in the thinly associated sandstone and mudstone unit, in the middle Laung Formation (ML2) at the southern tip of West Baronga Island. FA.2: facies association 2, thin alternating sandstone and mudstone. (d) Homogenized mudstone (Hm) cut by slide scars (Sc) and listric faults (arrowed). (e) Disrupted sandstone beds containing shell fragments in the chaotic mudstone, in the lower Laung Formation (ML1), Sittwe Point (East side). (f) Slumped muddy units containing localized slide scars and sandstone gutter casts. (g) Channelized sandstone bodies (CS) in the lower Laung Formation (ML1), Laychat Taung offshore island, south of Sittwe. Note the vertical beds with upwards-thinning succession (levee deposits) (AS) that overlie the channelized sand body. (h) Sandstone gutter casts containing abundant dish structures and convolute laminations overlying channel-shaped scour surface (arrowed) on the slumped, deformed thin mudstone units, near Sandawshin Pagoda, west Baronga Island. Arrows indicate direction in every photograph.
The upper member (ML3) of the Laung Formation is exposed mainly on the eastern bank of West Baronga Island and on East Baronga Island. This member consists mainly of a sandstone and shale interbedded unit, thick mudstone with slumped shales and sharp-based sandstone bodies. Thick mudstone intervals (about 200 m) include thin silty shales and sandstone layers, showing a thinning-upwards succession. These mudstones contain abundant planktonic and benthonic foraminiferal assemblages.
Planktonic foraminiferal fossils were extracted from the Laung Formation. The lowermost member contains Catapsydrax stainforthi, Globigerinella praesiphonifera, Globoquadrina tripartita, Globigerina woodi and Globigerinoides spp., assigned to planktonic foraminiferal zones N4 to N7, dated at c. 21.5–17.3 Ma. However, the middle and upper members yield a faunal assemblage containing Globoquadrina venezuelana, Globigerinoides sp., Globorotalia sp., G. bulloides, G. continuosa, G. peripheroronda, G. immaturus and Sphaeroidinellopsis sp., and are correlated with planktonic foraminieral zone N7 (17.3–16.5 Ma) of Berggren et al. (1995).
The Yezaw Formation comprises lower (MM1) and upper (MM2) members (Fig. 5.6). The lower member consists mainly of thick shale units (about 50–100 m) underlying amalgamated sandstone bodies and slump beds. The lower member is exposed mainly on the eastern coast of West Baronga Island, the western coast of Middle Baronga Island and East Baronga Island. The upper member comprises mainly thin alternating beds of grey to bluish-grey sandstone and shale with minor thick sandstone and shale intervals (Fig. 5.8).
Yezaw Formation (Middle Miocene). (a) Lower boundary of thin sandstone and mudstone alternation (FA.2, facies association 2, thin alternating sandstone and mudstone) showing the downlap contact (arrowed) with underlying chaotic mudstones in the upper Laung Formation (ML3), southernmost part of West Baronga Island. FA.3, facies association 3, chaotic beds. Hammer for scale circled. (b) Downlap contact (arrowed) between chaotic mudstone units and overlying thinly associated sand–mud alternation, on Sittwe Point (west side). (c) Major sequence boundary between Lower and Middle Miocene (dashed blue line). Undisturbed thickening-upwards sandstone bed (left side of photo) overlying slumped chaotic thin sand–mud units, Middle Baronga Point. Note the small scaled-minor fault aligned north–south. (d) Thick, wavy laminated thinly interbedded sand–shale unit of Middle Miocene, Yezaw Formation, Middle Baronga Island. Arrows indicate direction in every photograph.
The type section is along the Yezaw Chaung, a tributary of the Mayu River (20° 39′ N, 90° 46′ E). The maximum thickness of the formation is about 700 m in the Baronga Islands and crops out widely, forming the cores of anticlines including parts of the Mayu Ranges north of Sittwe.
The lower member is thought to include the lower Middle Miocene (c. 15.2 Ma) on the basis of the boundary between the planktonic foraminiferal zone N8 and N9 identified in the lower part of this member. The lower member yields the planktonic foraminifera Globigerina sp., Praeorbulina sp., G. baroemoenesis and Globoquaderina sp. The upper member yields mainly planktonic and benthonic foraminifers and the geological age is late Middle Miocene (c. 15–11 Ma) on the basis of two fossil-bearing shale horizons, zones N9 to zone N14 and N15. The foraminifers include Globigerinoides obliquus, G. sacculifer, G. bulloides, G. immaturus, Turborotalia siakensis, T. continuosa and Orbulina universa, indicative of the upper Middle Miocene ranging from N9 to N15.
The Mayu Formation comprises a predominantly arenaceous sequence that thickens laterally, and becomes fine-grained and more argillaceous. Thick to very thick (10–30 m) reddish brown-coloured sandstone beds crop out continuously along the strike on the Middle and East Baronga islands (Fig. 5.9). The alternating sandstones and shales are grey to greenish-grey and occasionally sandy and silty. The uppermost part of the formation is capped by a thick shale section of possible Lower Ngasanbaw Formation, and is only exposed in the core of the syncline on middle Baronga Island. The type section of this formation is on the eastern flank (20° 26′ N, 92° 40′ E) of the Mayu Range. The maximum thickness of the formation is 600 m in the study area.
Mayu Formation (Upper Miocene). (a) Thickening-upwards (yellow arrow) sandstone and shale alternation unit of lower Mayu Formation, Middle Baronga Island. (b) Sharp contact (arrowed) between heterolithic thin sandstone and shale alternation units and overlying massive- to thick-bedded sandstone with cross-laminations on the west coast of Middle Baronga Island. (c) Thick-bedded sandstone with large-scale cross-stratifications FA.5 (facies association 5 – thick bedded sandstones and shale) in the lower part of the Upper Miocene Mayu Formation near Pauktaw Township. Hammer circled for scale. (d) Heterolithic bed containing wavy to lenticular sand layers overlying mud drapes on East Baronga Island. (e) Detailed sedimentary structures of sandstone interbedded with thin shale unit of lower Mayu Formation. Note the climbing ripple laminations (middle) underlain by parallel laminations (bottom), and overlain by convolute lamination with dish structures (top). Arrows indicate direction in every photograph.
The geological age of this formation is possibly Late Miocene–Pliocene, based on the occurrence of nannofossils Discoaster sp., cf. D. broueri, Helicosphaera carteri var. carteri, Sphenolithus abies and Sphenolithus neoabies in the uppermost shale unit in the northern part of middle Baronga Island.
Evidence from palaeocurrents shows a prevailing SW–ESE transport direction although local variations are observed, including on East Baronga Island where directions shift towards the SE in the Middle Miocene units. To help understand such changes and controls on sediment flux we also conducted a detailed provenance study, using petrological and geochemical methods to constrain source-rock types and probable source regions. Palaeocurrent data summarized from the key sections in the Lower Miocene are shown in Figure 5.10. Foreset laminae in the ripple cross-laminated divisions were the most widespread pattern of palaeocurrent indicators. Sole structures are generally difficult to interpret, as the bases of most sandstone beds have been affected by load-induced deformation (Fig. 5.7h) and relatively few examples of unmodified erosional features (e.g. groove casts) were observed. Where present however, erosional sole marks are found to be parallel to the parting lineations and ripple foreset azimuths. Asymmetrical bed-base drag folds show a remarkably similar pattern of orientations to other palaeocurrent indicators. Foreset cross-laminations, channels axes and gutter-casts appear to be consistent in thick-bedded sandstones (Fig. 5.10c, d).
Palaeocurrent rose diagrams from (a, b) thin associated sandstone and shales and (c, d) the amalgamated sandstone beds of Lower Miocene (Laung Formation). Diagrams are based on: (a) foreset dip direction of cross-laminations; (b) parting lineations and sole marks of groove casts; (c) cross-laminations of sandstone beds; and (d) gutter-cast and channel directions of sandstone beds. Class interval = 10°; n = number of measurements. Cross-lamination direction (231°) perpendicular to gutter and channel directions.
Slump folds were examined in sandy and muddy slump and debris-flow deposits (Fig. 5.7d–f). Slump-fold plunge and hinge directions of the same generation of folds in the Miocene sequences are plotted on a stereonet (lower hemisphere projection) using the Hansen ‘separation-arc method’ (Hansen 1972). Directions are aligned generally nearly NE–SW (vector mean of 36 readings about 240°) and indicate a NW–SE-directed palaeoslope (Fig. 5.11a). In the lower Middle Miocene units on East Baronga Island, a southeastwards-sliding trend indicates a palaeoslope with a NE–SW-directed strike (vector mean of 21 readings about 165°; Fig. 5.11b). Evidence from palaeocurrents shows a prevailing SW–ESE transport direction although local variations are observed, including on East Baronga Island where directions shift towards the SE in the Middle Miocene units. Such a change in palaeoslope with time might be related to the shift of basin configuration, influenced by sedimentation load and tectonics in the proximal areas.
Stereograms (equal-area projection) of axial plunge directions and amounts for slump folds within the chaotic beds of Lower Miocene Laung Formation. (a) From the lower part of Laung Formation, at Sittwe and West Baronga Island from which the SW-facing palaeoslope was interpreted. (b) From the lower part of Yezaw Formation (Middle Miocene) on East Baronga Island showing SE-facing palaeoslope. All data were corrected for local structural tilt and plots were analysed by the separation angle method (Hansen 1972). Separation angles (star and dot symbols) are shown in planes approximating the fold-axis distribution (great circles). The mean downslope direction is shown by the arrow, which equally intersects the separation angle of two different groups.
The main purpose of the provenance study was to document the result of the quantitative petrographic analysis on the Miocene sediments in the Baronga Island area, and to compare these new data with previous interpretations of the uplift of the Himalaya and subsequent foreland thrusting. The exhumed sediments were mainly deposited in the palaeo-Bengal and Arakan basins and, if provenance analyses can be deciphered, provide the timing of uplift and the denudation history of the Himalayas.
In the Baronga Islands, combined petrographic and geochemical criteria for provenance determination in clastic sedimentary rocks have been applied to three formations of Lower Miocene to Upper Miocene–Pliocene age. Petrographic modal analysis carried out on 60 sandstone samples and the geochemistry of major and trace elements of 133 samples were analysed (80 sandstones, 53 shales). Geochemical criteria of discrimination based on the element ratios of trace elements (Cr/Zr, Y/Ni, Cr/V, Y/Ni and Rb/Sr) and major oxide ratios (Al2O3/TiO2, K2O/Na2O, Al2O3/SiO2-Fe2O3 + MgO, SiO2/Al2O3-K2O/Na2O, molar Na/Al-molar K/Al, Al2O3/Na2O-K2O/Na2O) allow distinction of these clastic sequences. Further, to examine the palaeoweathering trend the proxy factor chemical index of alteration (CIA), triangular A-CN-K diagrams and TiO2-Al2O3 ratio were used. All these results provide significant insights and additional constraints upon the temporal evolution of sediment supply to the Arakan Basin in Early–Late Miocene times.
Following the collision of India with Asia (c. 54 Ma; Molnar & Tapponnier 1975; Tapponnier et al. 1986; Dewey et al. 1989) the Himalayan chain experienced intense uplift and denudation; enormous quantities of terrigenous detritus accumulated not only in ancient foreland basins (e.g. Murees, Siwaliks), the Indo-Gangetic Plain (Gansser 1981) and Bengal Basin (Uddin & Lundberg 1998a, b), but also in the Bengal and Indus deep-sea fans. Main fluvial strata in these foreland basins were deposited along the flexed Indian margin (Critelli & Ingersoll 1994). From c. 21–17 Ma, an important transition in the development of southern Tibet and the Himalayas took place (Harrison et al. 1993), corresponding with the end of deposition of the Eocene Murree Supergroup and the movement of the Main Central Thrust. Furthermore, it coincides with the onset of huge sediment accumulations in the remnant ocean basins along both sides of Indian Peninsula (Curray & Moore 1974; Graham et al. 1975; Moore 1979; Curray 1994; Licht et al. 2013; Robinson et al. 2014). During and following this main tectonic uplift, the Siwalik Group was deposited between 18.3 and 4.9 Ma (Cerveny et al. 1988; Johnsson et al. 1988; Harrison et al. 1993).
Many compositional studies have considered the framework composition of sandstones from the Siwalik Group (Krynine 1937; Chaudhri 1972; Johnson & Vondra 1972; Tandon 1976; Parkash et al. 1980; Chowdhury 1982), the Makran accretionary wedge, SW Pakistan (Critelli et al. 1990), the Bengal and Surma basins (Uddin & Lundberg 1998a, b; Johnson & Alam 1991) and the modern Indus and Bengal deep-sea fans (Graham et al. 1975; Ingersoll & Suczek 1979; Moore 1979; Suczek & Ingersoll 1985; Critelli & Garzanti 1994) to obtain comparative information for the study area, where geographically and geologically connected.
Previous detailed provenance studies of individual detrital mineral species in these areas have include fission-track dating on detrital zircon (Cerveny et al. 1989), K–Ar dating of K-feldspar and muscovite (Harrison et al. 1993), heavy mineral studies from the Bengal Basin (Uddin & Lundberg 1998b; Najman et al. 2008), provenance study on Eocene–Miocene sandstones of the Rakhine Coastal Belt (Allen et al. 2008; Garzanti et al. 2013; Licht et al. 2013; Naing et al. 2013), heavy mineral constraints for Assam and Bangladesh foreland basins (Uddin et al. 2007; Allen et al. 2008; Najman et al. 2008, 2012) and Bengal Deep-Sea Fan (Copeland & Harrison 1990; Bouquillon et al. 1990; Yokoyama et al. 1990; Amano & Taira 1992). These studies concluded that, during the Neogene, sediments of Indo-Myanmar Ranges were dominantly derived from the Himalayas. The sedimentary rocks of western Myanmar are potentially important in understanding the tectonics and unroofing of the Himalayas; this study seek to find a record of Himalayan erosion preserved in the Miocene Indo-Myanmar Ranges, linking the foreland basins with the Bengal Fan.
Neogene sandstones were collected from three areas on the Arakan Coastal Ranges, from lithostratigraphic units correlated with those of the Surma and Assam areas, and from rocks that are dated by means of planktonic foraminiferal biostratigraphy. Sixty representative medium- to coarse-grained sandstone samples were selected for microscopic modal analysis on the basis of appropriate grain size and low degree of alteration. Thirty-eight samples were collected from the Lower Miocene (Laung Formation), 11 from the Middle Miocene (Yezaw Formation) and 11 from the Late Miocene–Pliocene (Mayu Formation).
Based on Vail et al. (1991) and Posamentier et al. (1992), a sequence stratigraphic approach was adopted. Petrographic and modal analyses were conducted following the Gazzi–Dickinson method (Dickinson 1970; Ingersoll et al. 1984; Zuffa 1985), counting sand-sized grains as lithic fragments to minimize the dependence of rock composition on grain size. Point counting was carried out for each sample: first, for the overall framework composition, 400–500 points per thin-section at a spacing of 0.5 mm and assigned to the petrographic categories in Table 5.1; and second, for quantifying lithic components, at least 200–300 points for each sample were counted at a spacing of 1.0 mm. Because of the potential importance of the metamorphic grade of abundant metasedimentary lithic fragments, the Ls–Lm1–Lm2 diagram (Dorsey 1988) was used to distinguish between very low- to low-grade (Lm1) and low- to intermediate-grade (Lm2) metasedimentary fragments. Counting parameters and calculations were based on the methods of Ingersoll (1983) and Dickinson (1985).
In addition to classical Q–F–L, Qp–Lvm–Lsm, Qm–F–Lt and Lm–Lv–Ls diagrams (Dickinson & Suczek 1979; Ingersoll & Suczek 1979; Dickinson 1985), the less traditional diamond-shaped diagram (Basu et al. 1975) that discriminates plutonic v. metamorphic sources based on the relative proportions of medium-grained fragments of non-undulatory monocrystalline quartz (Qund), undulatory monocrystalline quartz (Qud) and two types of polycrystalline quartz (Qp) end-member; Qp(>3) for polycrystallinity with >3 crystal units/grain, Qp(2–3) for polycrystallinity with 2–3 crystal units/grain, and the Qm–K–P diagram were also used (Streckeisen 1976). In terms of monocrystalline light components, a Qm–P–K (where P is plagioclase and K is potassium feldspar) diagram (Dickinson & Suczek 1979; van de Kamp et al. 1994; Trop & Ridgway 1997; Uddin & Lundberg 1998a) was also used to check the relative changes in feldspar content and their alteration in sandstones.
The petrographic classification uses quartz (Q), rock fragments (RF) and feldspar (F) according to Dott (1964) and Leeder (1982). Differences in petrofacies among the sequence stratigraphic systems tracts were also interpreted in order to determine the function of variable mechanical disaggregation and hydrodynamic sorting characteristics of different systems tracts, and erosional fluxes from the hinterland.
For geochemical analysis, 133 samples (both shales and sandstones) were collected from 23 measured sections on the Baronga Islands, Sittwe Township and Pauktaw Township. Each section is stratigraphically c. 250 m thick and is well exposed in the area, with the least amount of recent weathering on wave-cut platforms and in coastal cliffs. Detailed procedures are described in the Supplementary material.
According to the petrographic classification of sandstones by Folk (1980), based on the modal proportion of quartz (Q), feldspar (F) and lithic fragments (L) medium sandstone samples from the Baronga Islands are classified predominantly into sub-arkose and subordinately lithic arenite (Fig. 5.12a). Plots on a Q–F–RF diagram by Dott (1964) and Leeder (1982) show that most of the Miocene sandstones (>15% matrix) examined are classified as feldspathic wacke (Fig. 5.12b).
Ternary plots of framework quartz–feldspar–lithic fragments for medium-grained sandstones. Classification scheme is adopted from (a) Folk (1980), (b) Dott (1964) and Leeder (1982).
The 60 samples analysed are essentially of similar lithology, with some variability in composition, fabric and diagenetic features. They are typically micaceous and quartzose sandstones, with varied quantities of feldspar and accessory minerals. Some of the sandstones are very fine to fine-grained, while others are in the coarse- to medium-grained class; all are relatively well sorted. Most of the sandstones are clean; fine silt and clay matrix material is generally absent or very sparse, although some samples have a clay matrix (Fig. 5.13a, b). Microscopic sedimentary structures include laminations on a scale of millimetres to centimetres that reflect layer-by-layer compositional differences, with a strong alignment of abundant mica plates and distortion of mica flakes due to compaction and soft sediment deformation (Fig. 5.13c, d).
Thin-section photomicrographs of sub-arkosic sandstones. (a) Densely packed, angular quartz and blue-green amphibole aggregate (centre), quartz–plagioclase aggregate (upper left centre), untwinned plagioclase and metamorphic lithic fragments (lower left) set in sparse clayey matrix. Mayu Formation, Upper Miocene. Cross-polarized light (XPL). (b) Coarse, high-grade metamorphic lithic fragments (upper right) and blue-green amphibole (centre left) with angular quartz and fresh, untwinned single feldspar grains. Mayu Formation, Upper Miocene. XPL. (c) Photomicrograph showing syntaxial overgrowth of quartz and deformed muscovite mica in sub-lithic arenitic sandstone of Laung Formation, Lower Miocene. XPL. (d) Distorted mica and monocrystalline quartz with vermicular chlorite inclusions (centre) in the sub-arkosic sandstone of Yezaw Formation, Middle Miocene. Plane-polarized light (PPL).
The detrital mineral assemblage is dominated by quartz (50–60%), mica (10–30%) and feldspar (5–30%) in most samples examined. Most of the quartz and feldspar grains are moderately sorted and angular to subrounded. Muscovite is more abundant than biotite; the latter is altered to chlorite in some samples. Feldspar is mostly clear, unaltered, twinned plagioclase (Fig. 5.14a, c, d).
Thin-section photomicrographs of sub-arkosic sandstones from the Yezaw Formation (Middle Miocene) and Laung Formation (Lower Miocene). (a, b) Blue green amphibole (centre) and fresh microcline feldspar (upper right). (c) Mafic volcanic lithic fragment (centre), polycrystalline quartz fragment (left lower centre) and abundant orthoclase feldspars (lower left) in the sub-arkosic sandstone from Yezaw Formation, Middle Miocene. (d) Mafic volcanic lithic fragments (centre) and fresh perthite feldspar (centre left) in the arenitic sandstone from Laung Formation, Lower Miocene. (a–d) XPL.
In some samples microfossils were observed embedded in the clay matrix. They are probably tests of pelagic globigerinid foraminifera but are poorly preserved, having been replaced by iron-oxide and dolomitic cement (Figs 5.15c & 5.16a). Glauconite occurs as pellets with a clastic appearance (Fig. 5.16b) and also as an authigenic mineral (Fig. 5.16d). Glauconite traces and fossils are indicators of a more open marine depositional environment (Amorosi & Centineo 1997). According to the identification of the foraminiferal tests as globigerinids, an outer shelf to slope setting is suggested. The occurrence of both autochthonous and allochthonous glaucony in Early–Middle Miocene sandstones suggests a difference in source horizons, indicating sea-level fluctuations and the reworking of extrasequential glaucony (Amorosi 1995). The intermittent occurrence of matrix-free sandstones and textural and structural features suggest deposition and winnowing in current-swept environments.
Thin-section photomicrographs of metamorphic lithic fragments in the sandstones of Laung Formation (Lower Miocene). (a) Muscovite mica schist lithic fragment and epidote grain (centre) in the sub-arkosic sandstone. (b) Mica schist lithic fragment (centre) and polysynthetic twinned feldspar (upper centre) in sub-lithic arenitic sandstone. (c) Spar-filled test of Globigerina sp. with localized iron-oxide cement in sub-arenitic sandstone. (a–c) XPL. (d) Epidote grain (upper centre), carbonate fragment (centre), high-grade metamorphic lithic fragment (quartz + mica aggregate) (upper right) and twinned feldspar (lower right). PPL.
Thin-section photomicrographs of characteristic grains in sandstones of Sittwe and Baronga Islands. (a) Myrmekitic feldspar grain (centre) in the sub-arkosic sandstone of Laung Formation, middle Lower Miocene. (b) Ferruginous mudstone lithic fragment and detrital glauconite grain (centre) in lithic arenite of Laung Formation, lowest Lower Miocene. (c) Metamorphic lithic fragments and polycrystalline quartz (centre) with orthoclase feldspar (lower right) in lithic arenitic sandstone of Mayu Formation, Upper Miocene. (d) Chert fragments (centre) in clayey matrix of sub-lithic arenite of Yezaw Formation, Middle Miocene. (e) Metamorphic lithic fragment (centre) and carbonate lithic fragment (upper right) in the sub-lithic arenite of Yezaw Formation, Middle Miocene. (a–e) XPL.
The detrital mineral assemblage is diverse and therefore somewhat immature, indicative of proximity to source and rapid depositional rates. The abundance of muscovite, complex twinned-feldspars (Gorai 1950), perthitic feldspar, microcline and quartz with chlorite inclusions (Fig. 5.13d) suggest a granitic provenance (Blatt et al. 1980) for some of the material.
Diagenetic alteration responsible for lithification and a reduction of porosity includes compaction, pressure solution and related silica-precipitation and the formation of dickite minerals (kaolinitization) (Kantorowicz 1985) (Fig. 5.17c). The main responses were: (1) fabric condensation; (2) formation of microstylolitic contacts between quartz grains; (3) overgrowth of syntaxial rims into pores from free quartz surfaces; and (4) dissolution of carbonate materials. The results are compressed grain frameworks in which densely packed quartz grains share stylolitic-suture surfaces. Ductile micas and clay minerals are deformed into adjacent intergranular voids and reduced to form thin linings along grain contacts (Fig. 5.13d). Dissolved along grain boundaries, silica is reprecipitated on free surfaces as syntaxial quartz rims that have advanced into remaining intergranular voids (Fig. 5.13c). Calcitic fossils were either dissolved or their shells replaced by less-soluble ferroan dolomite (Fig. 5.15c).
Thin-section photomicrographs of arenitic sandstones of Upper Miocene Mayu Formation. (a) Lithic arenite containing well-rounded glauconite-bearing siltstone (centre) and broken fragments of crinoid spine (lower right) along with sub-angular quartz and feldspar grains with sparse matrix support. (b) Volcanic and sedimentary rock fragments (centre) in quartz arenite. (c) Mica schist lithic fragment (lower centre) in sub-arkosic sandstone. Note irregular pores filled with numerous minute crystals of authigenic clay minerals (centre). (d) Polycrystalline quartz grains (centre left), low-grade metamorphic lithic fragments (upper centre right) and high-grade metamorphic lithic fragment (lower centre) in the sub-arkosic sandstone. (e) Various types of metamorphic lithic fragments, including quartz–mica schist lithic fragments (centre) and low-grade metamorphic lithic fragment (lower centre) in the lithic arenite. (a, b) PPL; (c–e) XPL.
Alterational overprints indicate that the post-depositional history has been compressional, arising out of burial and tectonic loading. The Baronga structures are tight folds and reverse faults. Overprints are related both to structural development by tectonic loading and/or syntectonic soft-sediment deformation, due to forced regression in a shelf-slope setting.
Sandstones in the Lower Miocene Laung Formation have compositions averaging Qt50 F20 L10 (quartz 50%, feldspar mainly of plagioclase 10–23%, micas 20%, lithic fragments 12% and accessories 5%). Feldspar grains are mostly fresh, clear and unweathered. Quartz grains are dominantly monocrystalline and exhibit both undulatory and non-undulatory extinction. Polycrystalline quartz grains include lithic fragments of chert (Fig. 5.15a) and both foliated and non-foliated metamorphic rocks. Of the sparse feldspar grains, plagioclase are both of polysynthetic twinned- and perthitic twinned types; myrmekitic feldspars are more common than potassium feldspar (P/F = 0.65).
In lower Lower Miocene samples lithic components are dominated by intermediate- to high-grade metamorphic rocks (Fig. 3.15b, d); sedimentary and low-grade metamorphic rock fragments are more abundant than in the overlying units. Although not common, volcanic lithic fragments (Fig. 5.14d) are present in the middle and upper Lower Miocene samples, in contrast to older units. Volcanic fragments form on average about 4% of lithic fragments. Mineral composition and textures of volcanic lithic fragments are mostly from mafic to intermediate volcanic rocks.
Samples from the Middle Miocene Yezaw Formation have more diverse compositions than the older units, averaging Qt55 F15 L10. However, there is an abrupt increase in quartz content in the upper Middle Miocene samples. The quartz grains are mainly monocrystalline, with undulatory extinction. Some non-undulatory quartz contains vermicular chlorite inclusions (Fig. 5.13d). Plagioclase and potassium feldspar (Fig. 5.14a) are equally abundant (P/F = 0.51). Sedimentary fragments are predominant in the lithic population, and low- to medium-grade metamorphic fragments (Fig. 5.16e) are also present in some samples. These sub-arkosic and lithic arenites are somewhat less quartzose than in the lower Middle Miocene, and more quartzose than in the Lower Miocene Laung Formation. Intermediate-grade metamorphic fragments are more abundant in the lower part of the formation, compared with the underlying Laung Formation. Volcanic fragments compose approximately 1% in volume of the lithic fragments, much less than in the Laung Formation. Micas are relatively abundant, averaging about 10% of framework grains, and are dominated by biotite.
The precursor was moderately well-sorted micaceous quartz sand; detrital grains show subrounded to angular shape. The fabric is a slightly compressed grain framework. Mica flakes and glauconite pellets are frequently distorted and deformed into interparticle space and along many grain contacts. Microstylolitic contact in quartz-to-quartz grain is rarely noted. Iron-oxide cements are found locally in interparticle voids.
In sandstones of the Mayu Formation (Upper Miocene), quartz 60%, feldspar 20%, lithic fragments 10% and mica (mostly biotite) 10% occur with highly diverse accessory minerals such as garnet, hornblende, epidote, zircon, tourmaline, iron-oxide grains, rutile and sphene (Figs 5.13a, b, 5.14a, b & 5.15a, c) with a trace of glauconite. Samples reveal relatively uniform composition, except locally for quartz-rich samples in the lower Upper Miocene. Mean Qt–F–Lt values for the Mayu Formation are Qt60 F20 L20. Monocrystalline quartz is the major grain type and most of the quartz grains exhibit non-undulatory extinction. Potassium feldspar is more abundant than plagioclase (P/F = 0.3) in the middle part of the formation. Most potassium feldspar grains show no twinning (Fig. 5.13a). Lithic fragments are dominated by sedimentary types (Fig. 5.17a, b), but include a small amount of low- to high-grade metamorphic lithic fragments (Figs 5.13b, 5.16c & 5.17d, e). Volcanic fragments are also present more abundantly than in the Laung Formation (4%) (Fig. 5.17b).
The precursor was a moderately well-sorted, angular to subrounded micaceous quartz sand. The present fabric is a densely packed grain framework, cemented by calcite and minor iron-oxides. Micas are frequently distorted and deformed into interparticular spaces and along many grain boundaries. Chloritization is also seen in micas.
The results of the modal analyses of major components, monocrystalline grains and phaneritic lithic fragments (Dickinson 1985) were plotted on ternary diagrams (Fig. 5.18) to better understand compositional changes in the sandstones. In order to evaluate unroofing trends in the source regions, specific lithic components that are most likely to show useful variations in these dominantly phyllarenitic sandstones (Dorsey 1988), for example sedimentary lithic (Ls), very low- to low-grade metamorphic lithic (Lm1), and low- to medium-grade metamorphic lithic fragments (Lm2), are plotted on an Ls–Lm1–Lm2 triangular diagram (Fig. 5.18e).
(a, b) Ternary diagrams showing overall sandstone modes (Q–F–L and Qm–F–Lt), and (c) light monocrystalline components (Qm–P–K) showing means (indicated by n = number) for each stratigraphic unit. Provenance fields were adopted from Dickinson (1985). Note that Lower and Middle Miocene samples fall in the craton interior and recycled orogenic fields, whereas Upper Miocene–Pliocene sandstones fall in the recycled orogenic field. In Qm–P–K plot, shifting of quartz-rich field to moderately feldspar-rich field with increasing trend to potassium-rich feldspar in composition in Upper Miocene–Pliocene samples. (d, e, f) Ternary diagrams show lithic fragments compositions (Lm–Lv–Ls, Ls–Lm1–Lm2 and Qp–Lvm–Lsm plots). Provenance fields were adopted from Ingersoll & Suczek (1979) and Dorsey (1988). Lithic fragments suggest an unroofing trend (dotted arrows) from sedimentary to low-grade metamorphic in Lower Miocene samples, low-grade metamorphic to high-grade metamorphic in Middle Miocene, and recycled sedimentary components in Upper Miocene–Pliocene samples. Although Middle Miocene sandstones (squared) show a temporal increase in high-grade metamorphic lithic fragments derived from the orogens, the Qp–Lvm–Lsm plot shows no additional distinctions as a result of very low volcanic lithic components. However, metasedimentary lithic components contribute over time (filled arrow). (g) Qm–P–K triangular diagram (modified after Streckeisen 1976) to interpret the possible source-rock composition. (h) Point-count data derived from medium-grained detrital quartz populations, showing composite source but mostly plutonic source (after Basu et al. 1975).
The mean Q–F–L value of the Lower Miocene (Laung Formation) plots mostly within the ‘recycled orogen’ field, with some across the boundary of ‘basement uplift’ and ‘craton interior’ fields (Fig. 5.18a). On the Qm–F–Lt diagram, Laung Formation sandstones plot both on the ‘recycled orogen’ (quartzose recycled and transitional recycled orogenic) and ‘mixed’ provenance fields (Fig. 5.18b).
Sandstones of the Yezaw Formation (Middle Miocene) are mostly quartzose, and are lacking significant amounts of chert or other polycrystalline quartz grains compared to the Laung Formation; they plot as ‘craton interior’ and ‘recycled orogen’ on a Q–F–L diagram, and as ‘quartzose’ and ‘transitional recycled orogen’ and ‘mixed’ provenance on the Qm–F–Lt diagram (Fig. 5.18b). All samples from the Upper Miocene (Mayu Formation) and younger (Pliocene) sandstones plot as ‘recycled orogen’ on a Q–F–L diagram. On the Qm–F–Lt diagram, the Upper Miocene–?Pliocene Mayu Formation sandstones plot in the ‘quartzose recycled’ and, due to the slightly more feldspathic nature than the older units, the rest of the plots fall in the ‘mixed’ provenance field (Fig. 5.18b).
In terms of monocrystalline light components (Qm–P–K, where K is potassium feldspar), sands from the Baronga Islands have a diverse temporal pattern (Fig. 5.18c). Most importantly, the abundance of potassium feldspar relative to plagioclase increases systematically over time in these uppermost Lower Miocene samples. Monocrystalline quartz is abundant relative to feldspar in the Lower and Middle Miocene sandstones. The scarce feldspar in the Lower Miocene samples is mainly plagioclase, whereas feldspars in the Middle Miocene and Upper Miocene–Pliocene samples are mainly potassium feldspar. Middle Miocene and younger samples contain higher and more variable amounts of feldspar than older samples.
Volcanic lithic fragments (Fig. 5.18d–f) are generally scarce in the Baronga Islands samples, but there is a peak in the abundance of 2–5% of total framework grains (Lv/L = 0.02–0.1) in the Upper Miocene–Pliocene samples. This corresponds to abundance patterns of basaltic and ultramafic compositions in geochemical analysis of Cr/V–Y/Ni, Cr/Zr–Y/Ni diagrams, probably indicative of arc-volcanic sources as well as rare high Cr ratios, suggestive of ophiolitic sources in the orogenic belt.
Most lithic components are sedimentary and metasedimentary. These are plotted on the Lm–Lv–Ls and Qp–Lvm–Lsm diagrams (Fig. 5.18d–f). The plots show the shift from the ‘rifted continental margin’ to ‘collision orogenic’ provenances in all the Miocene samples. To interpret the unroofing trend of provenance in the entire Miocene samples, sedimentary and metamorphic lithic components were subdivided into sedimentary (Ls), very low- to low-grade metamorphic (Lm1) and low–intermediate-grade metamorphic (Lm2) lithic fragments, and plotted on the Ls–Lm1–Lm2 triangular plot (Fig. 5.18e). It suggests an unroofing trend from sedimentary to metamorphic components over time. Although the plotted ranges for the Lower Miocene samples largely overlap, the middle–upper Lower Miocene (ML2–ML3) and Middle Miocene (MM) sandstones show a temporal increase in higher-grade metamorphic lithic fragments, presumably derived from an orogenic belt.
The Lower Miocene samples are plotted within the wide field of variation such as ‘quartz-rich granitoids’ and minor ‘granite’ fields, whereas the middle Lower Miocene and late Lower Miocene samples plot in the ‘quartzolite’, ‘quartz-rich granitoids’ and ‘granodiorite’ fields. There is a distinction in the Middle Miocene samples which show the nature of the quartz-rich ‘quartzolite’ and ‘granitoids’, rather the Upper Miocene–Pliocene samples which plot in both ‘quartz-rich granitoids’ and ‘granite’ provenances due to the stratigraphic increase of feldspar content (Fig. 5.18g).
The Lower Miocene sandstones are rich in monocrystalline quartz with 35–56% of quartz grains showing non-undulatory extinction. Polycrystalline grains occupy 12–30% of the volume of total quartz. In contrast, the Middle and Upper Miocene samples show both undulatory and non-undulatory extinction in 30–65% of the total quartz population and polycrystallinity in 5–20% of the quartz grains. The Lower Miocene samples fall in the ‘plutonic’ field and ‘middle- and upper-rank metamorphic’ field of a lower triangular plot. The Middle and Upper Miocene–Pliocene samples also fall within the ‘plutonic’ field and around the ‘middle- and upper-rank metamorphic’ boundary (Fig. 5.18h). Data derived from counting the varieties of medium-grained quartz in the Miocene samples show a composite source and quartz populations produced by the erosion of plutonic and medium- and high-grade metamorphic rocks, and are easily discriminated from a detrital quartz population originating from metasedimentary sources.
In recent years, there has been an increase in the number of studies aiming to interpret the links between tectonic events and sedimentary response. In this context, much effort has been spent subdividing syntectonic successions into genetic sequences which reflect changes in parameters such as provenance, tectonically induced subsidence and uplift in sedimentary basins, and eustasy (Posamentier & Allen 1991, 1993; Posamentier et al. 1992; Burbank et al. 1993; Posamentier & James 1993; Wood et al. 1993; Ito 1994; Helland-Hansen & Martinsen 1996; Currie 1998; Dreyer et al. 1999).
In order to investigate the provenance of the sedimentary sequences, it is crucial not only to interpret the variations in sand composition but also to understand the development of depositional sequences, controlled by the interaction between the rates of sediment supply, basin tectonics and eustasy (Jervey 1988). Furthermore, the fundamental component building blocks of a depositional sequence (i.e. lowstand, transgressive and highstand systems tracts) are characterized by unique combinations of geometry, bounding surfaces and constituent lithofacies (Vail et al. 1991).
Shelf-slope to deltaic successions of Lower–Upper Miocene–Pliocene siliciclastic sequences, as much as 2000 m thick, comprise southeasterly prograded forced-regressive clastic wedges interrupted by thick transgressive Lower–Middle Miocene shale and the overlying lowstand, prograded deltaic sequences of the Upper Miocene–Pliocene. Fundamental component building blocks of depositional sequences (such as transgressive, highstand, forced-regressive wedges as slope components and lowstand prograding wedge systems tract sediments) are characterized by unique combinations of depositional processes and petrofacies (Coleman et al. 1983; Catuneanu et al. 1997).
On triangular plots of Qt–F–L, Qm–F–Lt, Qp–Lvm–Lsm and Lm–Lv–Ls (Fig. 5.19a–d), sediments from the transgressive systems tract are characterized by a high frequency of quartzose recycled petrofacies and lower feldspar content than the other systems tracts. This corresponds to the storage of terrigenous clastic sediments from the orogenic belt in a proximal position. Because the development of transgressive systems tracts is controlled by the accommodation-dominant regime condition (Thorne & Swift 1991) associated with a reduction in coarse-grained components, terrigenous sediments are trapped along river mouths during transgressive stages. The seaward area of an inner shelf lacks terrigenous sediments derived from the hinterland. The rate of quartzose and sedimentary lithic fragment supply is therefore dominant in transgressive sequences. Its preservation potential may increase during transgressive stages, in response to the reduction of active supply of feldspar and metamorphic fragments to the distal part of the basin (Fig. 5.19a, b, f).
(a–f) Ternary diagrams for sandstone petrofacies modes for systems tracts. Arrows represent general evolutionary trends. Grey shaded polygons indicate the field of deltaic sandstones from lowstand prograding wedge systems tracts. Note that TST (transgressive systems tracts) samples fall near the Q and Ls corners, and other systems tracts become progressively enriched in L and Lt and later in F. Overlying systems tracts become progressively enriched in Lm unroofing of the metasedimentary complex through time and erosion into progressively deeper (high-grade) levels in collision belts. Note backwards arrow (black in (d)) for the deltaic sandstones (LPWST) represents the compositional mixing and maturing of Ls and Lm. Note compositional enrichments of medium- to high-grade metamorphic fragments, and feldspars in HST (highstand systems tracts) in (f).
Forced-regressive wedge systems tract sediments show both quartzose and sedimentary lithic recycled petrofacies, but a minor enrichment in feldspar content compared to the transgressive systems tracts. Lowstand prograding wedge systems tract sediments and highstand systems tract sediments are distinctively characterized by a lower frequency of quartzose petrofacies and a higher frequency of feldspar and lithic components, with metamorphic and minor volcanic fragments, than the forced-regressive wedge systems tracts and transgressive systems tracts (Fig. 5.19f).
In contrast, a general decrease of the frequency of quartz in lowstand prograding wedge system tracts and highstand systems tracts (Fig. 5.19a, f) can be interpreted to be the result of the restoration of active supply of terrigenous sediments during highstand sea-level stages in response to progradation of fluvial and deltaic systems towards the distal part of foreland areas; a coastal depositional system is well developed and actively prograding during highstand sea levels (Posamentier et al. 1988; Vail et al. 1991; Cronin et al. 1998). In particular, a late highstand sea-level stage is characterized by sediment supply exceeding accommodation space (Thorne & Swift 1991). After a rapid sea-level fall with forced regression, erosion commences again in the distal part of foreland basin and cannibalizes previously deposited upper-slope and outer-shelf sediments.
In general, progradation of coastal depositional systems is associated with an active supply of coarse-grained, terrigenous sediments from a fluvio-deltaic system. The increase in supply of erosional fluxes from hinterland sources reduces the frequency of quartzose sediments in the upper part of highstand systems tracts and lowstand prograding wedge systems tract sediments, overlying the forced-regressive wedge systems tract (slope components). According to the triangular plots, highstand systems tract sediments show more variable petrofacies than that of lowstand prograding wedge systems tract sediments. They may be affected much more by the direct fluvial supply of terrigenous sediments causing a high sedimentation rate. The deposits have not been subjected to abrasion and destruction of unstable lithic fragments by nearshore dynamic processes.
To determine the age of the rocks in the provenance area, U–Th–Pb isotopic measurements were carried out on monazite and thorite grains contained in six selected sandstone samples from the Laung, Yezaw and Mayu formations. All U–Pb analyses were performed in the Geology Department of the National Science Museum, Tokyo. Prior to the separation of the U–Pb minerals, standard heavy liquid and magnetic separation techniques were carried out according to the analytical procedures of Parrish (1990). Six semi-consolidated sandstone samples from three horizons of Lower–Upper Miocene formations were chosen for U–Pb dating of monazite minerals. One was from the upper Lower Miocene, Laung Formation; three were from the Middle Miocene Yezaw Formation; and two were from the Upper Miocene Mayu Formation.
Modal proportions of heavy minerals in the Miocene sandstones of the Arakan Coastal Ranges area are shown in Figure 5.20. The resulting geochronological ages are shown in Figures 5.21 and 5.22.
Heavy mineral distribution of the sandstones from study area, Arakan Coastal Ranges, Western Myanmar. Note that garnet (Gar.) decreases with depth of burial. Gr., grossularite and andalusite; Alla., allanite; Ilm., ilmenite; Flo., fluorite; Tho., thorite; As., accessory minerals.
Monazite ages of (a) the sandstones from the study area and (b) Upper Miocene sandstones.
(a, b) Two distinct geological ages of the provenance constraint from the U–Th–Pb microprobe measurement on thorite in the Upper Miocene sandstone samples.
Whole-rock geochemistry of detritus is a powerful tool to aid the correlation of sedimentary successions, indicating the provenance and the palaeoweathering processes (Young & Nesbitt 1998) in sedimentary successions (Roser & Korsch 1988). Further, the chemistry is useful for understanding the plate tectonic setting of a sedimentary basin (Bhatia 1983; van de Kamp & Leake 1985; Roser & Korsch 1986, 1988). From a geochemical point of view, fields of tectonic settings on the simple bivariate Al2O3/SiO2–Fe2O3 + MgO(%), TiO2–Fe2O3 + MgO(%) and K2O/Na2O–SiO2(%) plots show the progression from an active continental margin to passive continental margin and continental collision source, with recycling of older sedimentary rocks. Characteristic trace elements, especially Rb/Sr, Zr, Nb and V, support these interpretations.
A–CN–K triangle plots and CIA values suggest granitic and granodioritic compositions in most sandstones and shales. They show the weathering trend towards the A–K join in the Early Miocene to Late Miocene–Pliocene, whereas minor enrichment of CN in the Middle Miocene sandstones may have a relationship with the marine transgression that deposited thick grey shales and buried the Lower Miocene Laung Formation.
Many workers use the average of sandstone and mudstone analyses (Crook 1974; Maynard et al. 1982; Bhatia 1983). In fact, SiO2 in sandstones can vary by as much as 15% and in combined sandstone–mudstone suites by as much as 27% (Roser & Korsch 1986; Korsch et al. 1993). Geochemical composition is therefore dependent on the grain size of the rock, as well as petrographic properties such as tectonic setting of the provenance and of the basin, transport mechanism, climatic and geomorphological conditions.
To avoid the problem of the influence of grain size on chemical composition sediments of a particular grain size should be analysed, such as medium to fine sand as recommended by many sedimentary petrologists (Dickinson & Suczek 1979; Ingersoll & Suczek 1979; Ingersoll et al. 1984). In doing so however, certain valuable information may be lost such as the chemical variability within a single suite, and comparisons with previously published geochemical data would not be available in many cases.
The ratio of Ti to Al varies markedly in primary (igneous) source rocks (Young & Nesbitt 1998). These elements are considered to be relatively immobile in most weathering regimes, so that the TiO2/Al2O3 ratio has been used to investigate the provenance of sediments and sedimentary rocks. When a plot of TiO2/Al2O3 is constructed for a suite, the resultant trend differs significantly from that shown by similar suites, providing insight into the weathering and depositional history.
Plate tectonic processes impart a distinctive geochemical signature to sediments in two different ways. First, different plate tectonic environments have distinctive provenance characteristics, and second, the environments are characterized by distinctive sedimentary processes. Three types of geochemical discrimination diagrams were used for the sandstones and shales of the Baronga Islands.
Based upon a bivariate plot of first and second discriminant functions of major-element analyses of sandstones, the discriminant function diagram (Bhatia 1983) shows fields for sandstones from passive continental margins, oceanic island-arcs, continental island-arcs and active continental margins. Most of the samples plotted in both the ‘passive continental margin’ and ‘active continental margin’ fields, except for upper Middle Miocene samples that plotted only in the ‘passive continental margin’ field (Fig. 5.23). The plots of Lower Miocene, lower Middle Miocene and Upper Miocene samples in the field of ‘active continental margin’ indicate a composite provenance and that erosional fluxes came from the orogenic belt at these times. Minor occurrences of middle and upper Lower Miocene samples in the ‘continental island arc’ field indicates that volcanic lithic fragments were derived from an arc setting. Other data, especially of upper Middle Miocene, lower Middle Miocene and Lower Miocene samples, fall well within the ‘passive continental margin’ field.
Plot of discriminant scores along Function I v. Function II, to discriminate various sandstones from the study area. The major fields of various suites are adopted from Bhatia (1983). MU–PL, Upper Miocene–Pliocene; MM2, upper Middle Miocene; MM2, lower Middle Miocene; ML3, upper Lower Miocene; ML2, middle Lower Miocene; ML1, lower Lower Miocene.
Three categories of tectonic settings, including the ‘passive continental margin’ (PM), the ‘active continental margin’ (ACM) and the ‘oceanic island arc’ (OIA), can be recognized on the K2O/Na2O v. SiO2 and SiO2/Al2O3 v. K2O/Na2O discrimination diagrams (Roser & Korsch 1986). Most of the Lower Miocene sandstones and shales plotted in the ‘active continental margin’ field near to the active continental margin/passive continental margin boundary. The plots exhibit considerable overlap in both passive margin and active margin fields (Fig. 5.24). This overlap is consistent with the interpretation, deduced from the petrography and modal analyses, that a significant proportion of the Miocene sediment was recycled from older sedimentary sequences.
Major-element compositional plots of (a, c) sandstones and (b, d) shale of the study area in tectonic setting discriminant diagrams of Roser & Korsch (1986). PM, passive continental margin; ACM, active continental margin; OIA, oceanic-island crust; A1, arc setting, basaltic and andesitic detritus; A2, evolved arc setting, felsitic–plutonic detritus. Legend as for Figure 5.23.
The provenance signature of sandstone–shale suites by using major-element contents can distinguish between sediments in which the provenance is primarily mafic, intermediate and felsic igneous, and quartzose sedimentary (Roser & Korsch 1988). The analysis is based upon major elements from 133 samples in which Al2O3/SiO2, K2O/Na2O and Fe2O3(t) + MgO are proven to be the most valuable parameters, based upon the oxides of Ti, Al, Fe, Ca, Na and K, most effectively differentiate tectonic and provenance settings (Fig. 5.25).
Discriminant function analysis classification plots of function F1 and F2 scores for (a) sandstones and (b) shales of the study area. Provenance fields after Roser & Korsch (1988). P, dominantly basaltic rocks; P2, dominantly andesitic rocks; P3, acid plutonic and volcanic rocks; P4, mature polycyclic continental sedimentary rocks. Legend as for Figure 5.23.
Most of the samples plotted in the fields of recycled and intermediate igneous provenance on these discriminant diagrams (Fig. 5.25), whereas most plots for the lower Middle Miocene and middle Lower Miocene (ML2) shale samples fall in the fields of intermediate to felsic igneous provenance (Fig. 5.25b). The discriminant function diagram for the provenance signatures of sandstone–shale suites on the basis of the ratios of the major-element contents was adopted.
The problems of biogenic CaO in CaCO3 and biogenic SiO2 is circumvented by using ratio plots in which discriminant functions are based upon the ratios of TiO2, Fe2O3(t), MgO, Na2O and K2O all to Al2O3 (Fig. 5.26). The ratio discrimination diagram is slightly more effective than that based on the raw oxides for middle Lower Miocene, upper Lower Miocene and upper Middle Miocene samples, which contain microfossils and authigenic minerals such as glauconite and clay minerals.
Discriminant function diagrams for the provenance signatures of sandstone–shale suites from the study area using major-element ratios (from Roser & Korsch 1988). Note that most of the samples fall in the quartzose sedimentary and intermediate igneous provenance fields, except upper Middle Miocene (MM2) samples which plot mostly in quartzose sedimentary provenance and the middle Lower Miocene (ML2) samples which fall in the mafic igneous provenance field. Legend as for Figure 5.23.
On these discriminant function diagrams for the provenance signature, most of the samples fall within the fields of quartzose sedimentary (recycled) and intermediate igneous provenances. Some of the middle Lower Miocene and upper Lower Miocene samples fall in the mafic igneous field, whereas lower Lower Miocene and lower Middle Miocene samples plot in the felsic igneous provenance field (Fig. 5.26).
Chemical weathering strongly affects the major-element geochemistry and mineralogy of siliciclastic sediments (Nesbitt & Young 1982; Johnsson et al. 1988; McLennan et al. 1993; Fedo et al. 1995). Quantitative parameters indicating the CIA (Nesbitt & Young 1982) are therefore potentially useful to evaluate the degree of chemical weathering. High CIA values reflect the removal of labile cations (Ca2+, Na+, K+) relative to stable residual constituents (Al3+, Ti4+) during weathering (Nesbitt & Young 1982). Conversely, low CIA values indicate the absence or scarcity of chemical alteration, and consequently might reflect cool and/or arid conditions.
The A–CN–K system is useful for evaluating fresh rock compositions and examining their weathering trends because the upper crust is dominated by plagioclase- and K-feldspar-rich rocks (Nesbitt & Young 1984, 1989) and their weathering products, the clay minerals. The A–CN–K diagram also can be used to constrain initial compositions of source rocks. In the absence of K-metasomatism, a line extended through the data points intersects the feldspar join at a point that shows the proportion of plagioclase and K-feldspar in the fresh rocks. This proportion provides a good indication of the composition of the parent rock (Fedo et al. 1995).
The influence of chemical weathering on chemical compositions of the Miocene shales and sandstones is presented on A–CN–K ternary diagrams (Fig. 5.27a, b). On these diagrams, representative mineral compositions and an estimate of average weathering trend with possible source-rock composition, in comparison to the High Himalaya crystallines, leucogranites and clay from the Bengal Deep-Sea Fan (McLennan et al. 1993), are plotted.
A–CN–K (Al2O3–(CaO + Na2O)–K2O) plots for evaluation of the weathering profile of (a) shales and (b) sandstones. Solid arrows indicate the predicted weathering trend; dashed line represents the best-fit line of the samples. Composition of the Bengal Fan sediments (IC, illite–chlorite; SK, smectite–kaolinite), High Himalaya Crystalline (HHC) (France-Lanord & Derry 1997) and Himalaya Leucogranites (Searle & Fryer 1986) are also plotted for estimation of possible source composition. Legend as for Figure 5.23.
The weathering trend intersects the feldspar join line and the A–CN tie-line intersects at about 30 mol% Al2O3, which would be the weathering trend (Nesbitt & Young 1982). This predicted weathering trend (solid line in Fig. 5.27a, b) is nearly identical in both sandstone and shale plots, except that the plot of the sandstones is slightly below this trend towards the CN apex due to the CaO problem and sorting effects (Fig. 5.27a, b) between sand and shales (Nesbitt et al. 1996). There is the minor enrichment of K in Upper Miocene–?Pliocene and upper Middle Miocene samples, which could be explained by the loss of Al2O3 or, more likely, the addition of K2O during diagenesis (Fedo et al. 1995; Bock et al. 1998).
The predicted weathering trend line intersects the feldspar join at the point that shows the proportion of plagioclase and K-feldspar of the fresh rock, and lies between granodiorite and granite compositions (Fedo et al. 1995). Moreover, lower Middle Miocene (MM1) samples fall at the lower end of line (near the CN apex), and show low CIA compared with other units. This evidence shows that the intensity of chemical weathering in the lower Middle Miocene was greater than in the other Lower Miocene samples.
The ratio of Ti to Al varies greatly in primary source rocks of igneous origin (Young & Nesbitt 1998). These elements are considered to be relatively immobile in most weathering regimes, so that TiO2:Al2O3 ratios have been used to investigate the provenance of sediments and sedimentary rocks (Young & Nesbitt 1998). Under extreme weathering conditions however, the upper parts of profiles commonly show a marked increase in the Ti:Al ratio that is thought to be due to preferential translocation of Al-rich phases. Moreover, the Ti:Al distribution reflects mixing of Al- and Ti-enriched fine-grained materials with sands that are depleted in these elements due to physical weathering and sorting (Young & Nesbitt 1998). Comparison of these results with data from sedimentary rock suites suggests that trends shown by plots of TiO2 v. Al2O3 may provide insight into weathering and depositional history, in addition to their use as a provenance indicator.
The plots of TiO2 v. CIA (Fig. 5.28a, b) show higher TiO2 and CIA values in the Lower and Upper Miocene samples than some Middle Miocene samples. The similarity in the chemical behaviour of Fe and Ti is remarkable. These elements would be expected to behave differently during weathering because the solubility of Fe is largely controlled by the redox condition, whereas that of Ti is not. The similar behaviour of Fe and Ti is thought to indicate that both are effectively immobile, and the variations in Ti:Al ratio in the weathering profile of the samples (Fig. 5.29) probably indicate that Al has been mobilized rather than Ti.
TiO2 v. chemical index of weathering (CIA). (a) Increasing TiO2 and CIA in Lower Miocene (ML1) and Upper Miocene–Pliocene (MU–PL) sandstones (upper right); the arrow represents the weathering trend. (b) Most of the shale samples yield higher TiO2 contents than the sandstones and, apart from samples from the upper Middle Miocene, do not exhibit increasing TiO2 but show the opposite trend (dotted arrow). Legend as for Figure 5.23.
Plots of TiO2 v. Al2O3 (Young & Nesbitt 1998) for (a, b) sandstones and (c, d) shales from the Baronga Islands area. Note that the shale samples show higher values of both Ti and Al than the sandstone samples. (b, d) Distribution of Ti: Al ratio trends; most of the samples fall within a fairly constant field between 1:20 and 1:10 lines and, as weathering becomes extreme, show a marked increase in both Ti and Al. Bengal Fan Clay (IC, illite–chlorite; SK, smectite–kaolinite), High Himalaya Crystalline (HHC) from France-Lanord & Derry (1997). Legend as for Figure 5.23.
Plots of TiO2 v. Al2O3 for sandstones and shales show that the profiles of the Lower–Upper Miocene–Pliocene samples display fairly straight, linear trends (Fig. 5.29) and fall within a fairly constant field between the 1:20 and 1:10 lines. As weathering becomes more extreme, it shows a marked increase in both Ti and Al. As a result, the relative abundance of TiO2 and Al2O3 in siliciclastic sequences of the study area can be used as a possible indicator of the nature of the source terrains. In contrast, a suite of well-sorted sediments from the studied samples shows a roughly linear distribution, sub-parallel to the lines of equal Ti:Al ratio.
Strontium isotope ratios (87Sr/86Sr) were analysed on the calcareous planktonic and benthonic foraminifers taken from 29 shale samples of the Lower–Upper Miocene Laung, Yezaw and Mayu formations. The foraminiferal samples consisted of approximately 300 specimens of benthonic species (Uvigerina spp., Bulimina spp., Globobulimina spp., Globorotalia spp., Pullenia bulloides and Dentalina spp.) and planktonic species (Globigerinoides spp., Globoquadrina spp. and Paragloborotalia spp.) which were extracted from shales and mudstone from eight stratigraphic sections.
These data were plotted in stratigraphical order and the fluctuations examined, along with the chemical index of alteration (CIA) and geochemical parameters (Fig. 5.30). All of the values fall in the range 0.709183–0.711221. Both of these samples are from the Middle Miocene Yezaw Formation. There is a major fluctuation in lower Middle Miocene, coinciding with episodes of sediment influxes from hinterlands (Fig. 5.30). This is the same trend as that of the TiO2 v. CIA (Fig. 5.28a, b). Sr isotope data and the palaeoweathering index (CIA) yield results that are essentially identical to those obtained from major oxide and trace-element ratios.
Summary of composite sedimentary sequences with Sr isotope ratio, CIA, heavy minerals, and K-feldspar and plagioclase grains in columns A–D, respectively. G, garnet; Z, zircon; E, epidote; T, tourmaline; A, amphiboles; (*) U–Pb monazites, thorite age dated samples.
The Lower Miocene sandstones analysed from the Baronga Islands, Myanmar are quartzose, although very few in number, and contain relatively large amounts of low- to intermediate-grade metamorphic lithic fragments and a moderate amount of unweathered feldspar (plagioclase), pre-dating a significant input of collisional detritus (Fig. 5.31a, b). The modal analysis of the Lower Miocene sandstones of the Laung Formation documents compositions that are dominated by feldspar grains and lithic fragments, including low- to intermediate-grade metamorphic lithic fragments. This suggests the onset of uplift with erosional unroofing in the eastern Himalayas, and the generation of fluvial systems which supplied orogenic detritus to the south.
(a) Plot of molar Na/Al v. K/Al showing a trend for shales from the Baronga Islands study area that intersects the Bengal Fan Clay (illite–chlorite) and PAAS. Note samples from the episodes (Middle Miocene: MM1, MM2) fall near to the composition of HHC. (b) Plot of SiO2/Al2O3 v. K2O/Na2O (wt%) also shows a mixing line with two end-members (Bengal Fan Clay: Bengal IC, SK and Middle Miocene samples of the Baronga Islands area). (c) Companion plot to (b), plotting the ratio of the denominators (Al2O3/Na2O) against one of the ratios (K2O/Na2O) after Langmuir et al. (1978). The data define a straight line, which is a good test of internal consistency for mixing relationship. Bengal Fan Clay, High Himalaya Crystalline (HHC) from France-Lanord & Derry (1997). Legend as for Figure 5.23.
The abundance of sub-angular to subrounded, coarse monocrystalline quartz with fresh plagioclase grains in the sandstones of the lower part of Laung Formation and the Upper Miocene Mayu Formation reflect little mechanical weathering, whereas the conspicuous change in lithic type, such as chert and metasedimentary lithic fragments increasing in the Upper Yezaw Formation, suggest a low input from orogenic belts and a high proportion of reworking from older sedimentary deposits. Authigenic quartz, dickite minerals and glauconite suggest a slight reducing environment in the middle Lower Miocene.
The dominance of monocrystalline, sub-angular to subrounded quartz grains, the presence of plagioclase feldspar with minor potassium feldspar, and small amounts of volcanic (basaltic) lithic fragments all suggest there were erosional fluxes from continental sources rather than collisional orogenic terrains during late Early–Middle Miocene time. These units were possibly derived from the adjacent Indian Craton, as well as from deposits on the northeastern passive continental margin of India and the Indo-Myanmar Ranges. The decrease of fresh metamorphic lithic input during late Early–late Middle Miocene time shows that the supply from the eastern Himalaya was exceeded by the sedimentary supply from the uplifted Indo-Myanmar Ranges, probably due to higher rainfall and exhumation in the east coinciding with interrupted marine transgressions (Kyi Khin et al. 2014).
Johnsson et al. (1988) and Uddin (1998) discussed the provenance of the Eocene–Oligocene sandstones of the Bengal and Assam basins, northeastern India, as these sequences contrast strongly in composition with sandstones from the Bengal Basin of Bangladesh, indicating quite distinct provenance histories for these two eastern Himalayan foreland basins. Sandstones from the Eocene–Lower Oligocene units are texturally immature. They are composed mainly of quartz (both mono- and polycrystalline), sedimentary and low-grade metamorphic lithic fragments, including abundant chert and plagioclase.
Sandstones of the overlying Middle–Upper Oligocene formations from Assam are similar in lithic composition, but also contain significant amounts of volcanic and higher-grade metamorphic detritus. The pre-Miocene sandstones are clearly derived from an orogenic source, exposing and eroding sedimentary and low-grade metamorphic rocks, followed by increasing contributions from volcanic and higher-grade metamorphic rocks during the deposition of the Miocene sandstones. In contrast, the Eocene–Oligocene strata from the neighbouring Surma Basin of Bangladesh contain mature quartzose sandstones that probably originate from a cratonic source. The Bengal Basin was apparently protected from orogenic sedimentation during Eocene–Oligocene time, but clearly orogenesis had begun to the north (Sorkhabi & Stump 1993; Yin et al. 1999; Uddin et al. 2007; Allen et al. 2008; Licht et al. 2013, 2015; Naing et al. 2013).
The abrupt increase in feldspar grains, lithic fragments and heavy minerals (Fig. 5.30) in the lower Lower–lower Middle Miocene samples marks the first clear signal of erosion fluxes from the orogenic terrains into the Arakan Basin, Myanmar based on our stratigraphic sampling and heavy mineral analysis. However, it records the initiation or simply migration into this area of stream systems that supplied orogenic detritus to the Bengal and Arakan basins; this may represent the actual onset of collisional uplift and erosional unroofing in the eastern Himalayas.
Extremely thick Oligocene sequences, including both marine and continental deposits overlain by Miocene fluvial sequences, lithostratigraphically correlated with the Tipam Group of the Surma Basin, NE Bangladesh, are documented in Assam by Rao (1983). This suggests that the orogenic activity had begun in the eastern Himalayas by Oligocene time, and shifted southwards to the Indian crust as it developed large stream systems that evolved to funnel voluminous detritus into the Neogene Surma and Arakan remnant basins by earliest Miocene time. It appears that the bulk of deltaic accumulation migrated from Assam in Oligocene time and through Surma, Bangladesh into the Arakan Basin during the Early Miocene and through Middle Miocene time.
The upper Middle Miocene sandstones are rich in sub-angular monocrystalline quartz grains, chert and argillite with a relatively small amount of feldspar and metamorphic lithic fragments, suggesting a low supply from the Himalayan terrain, associated with a widespread marine transgression. The Upper Miocene–Pliocene sandstones containing abundant argillitic and low- to intermediate-grade metamorphic lithic fragments (Fig. 5.16d, e) and feldspar grains suggest continued orogenic unroofing. Potassium feldspar is dominant in these younger sandstones relative to the plagioclase-rich Lower–Middle Miocene sandstones suggesting a granitic source, probably in the Lesser Himalayas and the High Himalayan Crystalline terrain. The ratio of potassium feldspar to plagioclase increases in the Upper Miocene (Mayu Formation). This is probably not due to changes in the intensity of chemical weathering alone, because the abundance of potassium feldspar and heavy minerals (blue green amphiboles) (Figs 5.13a, b & 5.14a, b) is high relative to the older units.
The greater content of volcanic lithic fragments and the occurrence of blue-green amphiboles in Upper Miocene samples suggest the erosion of arc rocks from the Himalayan and Indo-Myanmar Ranges. Most of the palaeocurrent directions show transport from NE to SW however, with subordinate NW to SE transport (Fig. 5.10). Erosional fluxes from the Himalaya are present in strata of similar age in the Siwalik Sandstones, derived from the Pakistan Himalaya (Johnson et al. 1985; Cerveny et al. 1989; Critelli et al. 1990; Critelli & Garzanti 1994), Nepal (Critelli & Ingersoll 1994), the Surma Basin (Uddin & Lundberg 1998a), and are also found in cores from ODP Leg 116 sites on the distal Bengal Deep-Sea Fan (Ingersoll & Suczek 1979; Suczek & Ingersoll 1985; Yokoyama et al. 1990; Amano & Taira 1992).
Most detritus has been transported by the Indus and Ganges/Brahmaputra river systems to their respective deltas, with much of the fluvial sediment bypassing the deltas to be further transported by turbidity currents into the deep-water Indus and Bengal/Nicobar fans (Graham et al. 1975). Isotopic data on the Bengal Fan (Copeland & Harrison 1990; Amano & Taira 1992) confirm rapid uplift and unroofing of the tectonic units of the High Himalayan crystalline rocks. During Oligo-Miocene time, sediment derived from the Himalayan Belt was probably transported westwards and eastwards into the Makran and offshore Sumatra (Nias Island) subduction zones (Moore 1979; Critelli 1993). This suggests a sedimentary–metasedimentary–volcanic source terrain in the Himalayas to the north, draining through the Bengal Fan and reaching the subduction complex.
Miocene sandstones from the Arakan Basin (average Qt59 F26 L15) are very similar in composition to sandstones of the same age in the Bengal Fan and the Surma Basin. Comparison of the Q–F–L, Qp–Lvm–Lsm and Lm–Lv–Ls plots suggests that most of the sediments were derived from the collision and foreland uplift of recycled orogen provenances, with varying proportions of feldspar and lithic populations in times that were related to erosional fluxes from Himalayan uplift and relative sea-level change.
Th–Pb ion microprobe measurements made on monazite and thorite grains of six selected sandstone samples from the three Miocene horizons, showing episodic fluxes (Fig. 5.30) from granitic and metamorphic sources, are interpreted to indicate that these sources crystallized at 1100, 500–600, 55–65 and 15–25 Ma (see Figs 5.21 & 5.22). Most of the U–Pb isotopic results on the monazite grains demonstrate important distinctions of the Greater Himalayan metasedimentary provenance that included a major source of 0.8–1.0 Ga zircons, implying a Late Proterozoic age (Windley 1983; Hodges et al. 1996; Parrish & Hodges 1996). The provenance age of 500–600 Ma can be interpreted as derived from Lesser Himalayan thrust sheets, which are composed of Lower Palaeozoic rocks, mainly low-grade metamorphics and granites (Windley 1983; Mohan et al. 1989; Upreti & Le Fort 1999).
Another provenance age of 55–65 Ma represents a source in the Transhimalayan (Gangdese) plutonic belt of the southern Tibetan Plateau, the Lhasa Block, defined by the Indus–Zangpo suture zone (Windley 1983; Scharer et al. 1984; Xu et al. 1984; Yin et al. 1994). The prominent occurrence of 15–25 Ma provenance ages suggests a major source in the High Himalayas Crystalline and younger granitic bodies (France-Lanord et al. 1993; Sorkhabi & Stump 1993; Noble & Searle 1995; Hodges et al. 1996; Edwards & Harrison 1997; Guillot et al. 1999). The major source for detrital material has therefore been the equivalent of the modern High Himalaya Crystallines. The Lesser Himalayas and the Transhimalayan and Greater Himalayan metasedimentary sequences must have been subordinate sources compared with the High Himalaya Crystallines of the eastern Himalayas. The primary conclusion from the isotopic data is that an equivalent of the High Himalayan Crystalline series has been the dominant source of sediment supply to the Arakan Basin since before 17 Ma.
The underlying assumption of the geochemical discrimination diagrams for sedimentary rocks is that there is a close link between plate tectonic setting and sediment provenance (Rollinson 1993). This is largely true, and the chief successes of the technique are with immature sediments containing a significant volume of lithic fragments from which provenance, and hence tectonic evolution with episodic erosional fluxes, may be identified.
Sandstones and shales of the studied area have intermediate to high SiO2 contents, ranging over 52.4–86 wt% but, as expected, the sandstones have higher SiO2 and correspondingly lower Al2O3. An exception is lower Middle Miocene shale samples with low K2O and Al2O3 contents (Fig. 5.31), but high carbonate content. These analysed samples are not only shales but also sandstones, enriched in metamorphic lithic fragments and heavy minerals (Zr = 1200 ppm; Cr = 250 ppm).
Miocene sandstones of the Baronga Islands show a strong negative correlation on a plot of molar Na/Al v. molar K/Al ratio (Fig. 5.31a). The trend intersects the K/Al axis at about 0.27, which indicates an illitic end-member (cf. Argast & Donnelley 1987). This illite could be either detrital or authigenic. There are two distinctive zones with respect to the ratio of Na/Al, whereas late Lower Miocene samples plot in the upper zone and fall in the field of high Na/Al, but with lower K/Al ratios than other samples.
Samples from the Baronga Islands follow a hyperbolic trend on a SiO2/Al2O3 v. K2O/Na2O plot (Fig. 5.31a) that may be interpreted as due to mixing. To ensure that this fit is not fortuitous, a companion plot of the ratio of denominators (Al2O3/Na2O) v. one of the ratios (K2O/Na2O) (after Langmuir et al. 1978) is also shown (Fig. 5.31c). On this diagram, the data fall along a straight line as predicted if the observed trend was due to the mixing of provenances subject to different palaeoweathering conditions (cf. McLennan et al. 1990, 1993).
The K2O/Na2O ratios of the sandstones lie between 1.1 and 2; however, K2O/Na2O ratios of less than 1 are observed for the lower Middle Miocene sandstones. The high K2O/Na2O ratios for the sandstones of the active margin-related samples probably reflect derivation from recycled sedimentary sources that have the characteristic signature of an extended weathering history (McLennan et al. 1993).
Geochemical variation within the Lower–Upper Miocene–Pliocene sequences is shown in Figure 5.30. CIA values for each stratigraphic subdivision are plotted with the other major oxides and trace elements. In these sections there are episodic anomalies in several horizons coinciding with variations in CIA and major-element composition, indicative of erosional fluxes from hinterland sources (especially from the Himalayan orogenic belts).
The 87Sr/86Sr ratios of dissolved Sr in the seawater from the measurements of planktonic and benthonic foraminiferal tests exhibit fluctuations in the Lower–Upper Miocene succession of the Baronga Islands. The period of the most rapid change in the Sr isotopic ratio is 20–15 Ma (Fig. 5.30), which is also a period of exceptionally high sediment accumulation in the Lower Miocene succession. The relationship between erosion and Sr flux allows the use of the Sr isotopic evolution of seawater to reconstruct a history of erosion since collision in the hinterland regions.
The flux of dissolved Sr carried by rivers originating in the Himalaya–Tibet region, on the other hand, is presently a significant fraction of the global Sr budget (cf. Richter et al. 1992). In support of this, both the timing of collision and the periods of most rapid unroofing of the Himalaya and Tibetan Plateau correspond closely to periods of rapid increase in seawater 87Sr/86Sr and the total quantity of material removed by erosion from the uplifted areas. Furthermore, the period of particularly rapid unroofing of Himalaya and Tibetan Plateau of 20–15 Ma, indicated by new geochronological data (England & Houseman 1986; Richter et al. 1990; Copeland & Harrison 1990), occurred at exactly the same time as the maximum change in the Sr isotopic composition of seawater.
High seawater 87Sr/86Sr ratios reflect continent–continent collisions where old basement rocks are uplifted, partially remobilized and exposed to extensive erosion (Richter et al. 1992). A prominent feature in the timing of the Sr flux change, the short-lived pulses of high CIA starting at 22–15 Ma (Fig. 5.30), correlates in time almost exactly with a period of exceptionally high erosion in at least some portions of the Himalayas and Tibet. There was a further increase in Sr ratios at 10 Ma (see Fig. 5.30), which shows a nearly 10 Ma time lag between the initial collision of India and Asia and increased erosion. This could reflect either a delay in generating the high topography needed for enhanced erosion and weathering, or the fact that a substantial thickness of carbonates (possible Tethys Himalayas) needed to be stripped from the emerging highland before exposing the more radiogenic rocks. In either case, the sedimentary input as a function of time since Early Miocene (see Fig. 5.32) contains important clues in respect of the erosional fluxes consequent on the unroofing of the Himalayas and the timing of weathering, both physical and chemical, in the Neogene sequences of the Arakan Coastal Ranges, Myanmar (Fig. 5.33).
Sediment accumulation curves from the Assam–Bengal foreland basins, based on the relative time control. Overall accumulation rates (equal to the slope of the accumulation curves and shown in the inset) decrease after 15 Ma (early Middle Miocene), and remain steady, or lower, during the subsequent Late Miocene interval. Dashed lines show constant accumulation for comparison. CH, Chittagaung Fold Belt; Sch, Schuppen belt; Ng, Naga Hills; As, Assam basin; Ak, Arakan Basin (Baronga Islands area). Numbers represent the location of the basins (see Fig. 5.1). The change in sedimentation rates corresponds to a major change in sedimentation pattern in proximal areas (onset of fluvial sedimentation of Siwalik Group) and delta prograding to the distal part of the foreland areas.
Correlation of Neogene Arakan Basin sedimentation with Bengal and Indus fans evolution and tectonics events of Tibet–Himalayas. Modified from Nishi & Sakai (1997). STDS, South Tibetan Detachment System; MBT, Main Boundary Thrust; N–S, north–south-trending faults; MCT, Main Central Thrust.
Sequence stratigraphic and provenance studies, including trace-element geochemistry and model compositions of the Miocene siliciclastic sequences preserved in the Baronga Islands, the Arakan Coastal Range, western Myanmar, have yielded the following significant results.
On the basis of regional stratigraphic correlations and geological age determinations from planktonic foraminiferal zonation of the siliciclastic sequences in the Baronga Islands, the Laung and Yezaw formations were deposited in deep-marine slope and shelf environments during Early–Middle Miocene time (21.5–11 Ma). The age of the lower and upper bounding units suggest that the Mayu Formation was deposited in a southwards-prograding shelf-delta environment during Late Miocene–Pliocene time.
Early Miocene underthrusting along the Himalayan front is well documented by the forced-regressive sedimentation patterns in the slope and shelf systems in the Bengal–Arakan basins.
The sequential evolution of the Miocene clastic sequences shows that forced-regressive wedged systems tracts, in which the slope evolved into bypassing, slumping and deep-marine channel infilling, and began to accumulate an increased sediment load due to the rapid fall in base level during Early–early Middle Miocene time.
Provenance data from the Laung Formation and the lower Yezaw Formation demonstrate an overall upsection enrichment in feldspar and medium- to high-grade metamorphic lithic fragments, at the expense of quartzose grains and low-grade metasedimentary and sedimentary lithic grains. The proportion of K-feldspar grains also increases abruptly in the lower part of Mayu Formation, suggesting that High Himalayan granitic rocks became exposed widely at 11 Ma.
Abundant sub-angular to subrounded coarse monocrystalline quartz grains reflect little mechanical weathering and relatively brief transport, whereas the conspicuous change in lithic type, such as chert and metasedimentary lithic fragments, which increase in the Yezaw Formation, suggest a low input from orogenic belts and increased reworking of older sedimentary deposits.
Sandstones of the late Middle Miocene are rich in sub-angular monocrystalline quartz grains, chert and argillite with relatively small amounts of feldspar and metamorphic lithic fragments, suggesting a low supply from the Himalayan terrain associated with a widespread marine transgression.
Sandstones of the Late Miocene–?Pliocene (Mayu Formation) contain abundant argillitic and low- to medium-grade metamorphic lithic fragments and feldspar grains, suggesting continued orogenic unroofing. Potassium feldspar is dominant in these younger sandstones, relative to the plagioclase-rich Early–Middle Miocene sandstones, suggesting a granitic source, probably in the Lesser Himalayas and the High Himalayan Crystalline terrain.
There are two distinct episodes, mainly in the Laung Formation and lower part of Yezaw Formation, which show the enrichment of acidic source materials and low chemical alteration. Lower CIA values are suggestive of retention of composition of source material in the sediments. Most of these episodes coincide with late transgressive highstand sequences. There are higher CIA values and high TiO2, MgO, Fe2O3 and Rb/Zr, Rb/Sr ratios in lower and middle horizons, representing early lowstand (forced-regressive periods). This anomaly can be interpreted as due to sedimentary reworking and/or low input from the hinterlands.
There is a period of rapid change in the Sr isotopic ratio during 20–15 Ma. This is also a period of exceptionally high sedimentation in part of the Early Miocene succession. The relationship between erosion and Sr flux allows the use of the Sr isotopic evolution of seawater to reconstruct a history of erosion since collision in the hinterland regions. There is a further increase at 10 Ma, which shows that there was a nearly 10 Ma time lag between the initial collision of India and Asia and increased erosion. This could reflect either a delay in generating the high topography needed for enhanced erosion and weathering, or that a substantial thickness of carbonates (possible Tethys Himalayas) was stripped from the emerging highland before the more radiogenic rocks were exposed.
Th–Pb ion microprobe measurements on monazite and thorite grains from the sandstones of the late Early–Late Miocene formations were made to determine the geochronological ages of the source rocks in the areas of provenance. The data show that a close analogue of the High Himalaya Crystalline was already sub-aerially exposed to active erosion during Miocene time.
Based on the geochemical and provenance data proxy for the unroofing of Himalayas, episodes are related to the interactions between the timing of Himalayan thrusting, tectonic denudation and palaeoweathering, migration of foreland fluvial system and the direction and rate of base-level changes in the Miocene Arakan Basin of Myanmar.
We would like to express our gratitude to Professor H. Okada, Professor H. Sano of Kyushu University and Professor H. Sakai of Kyoto University, Dr Wonn Soh of Japan Marine Science and Technology Center (JMSTEC), Professor H. Nishi of Hokkaido University for identification of foraminifers and helpful suggestions during this study, and to Associate Professor T. Ikeda, Kyushu University for heavy mineral identification in thin-sections. The first author would like to express sincere thanks to the Department of Geology, Fukuoka Education University and Dr Yasuji Saito of National Science Museum, Tokyo for carrying out U–Pb microprobe analyses, Dr Takanori Nakano of Institute of Geosciences, University of Tsukuba for Sr isotope analyses in their laboratory and the Ministry of Education (Monbusho) of the Japanese Government for financial support. Three periods of fieldwork could not have been completed without support in many ways from colleagues (U San Win, U Hla Min and Dr Myo Myint) and the local people from the villages of Sittwe and Pauktaw townships, Rakhine State, Myanmar. We also thank Dr H. Smyth, Professor A. Carter, Dr A.J. Barber and an anonymous reviewer for their comments, which helped to improve the manuscript.
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