The Cenozoic India-Asia collision generated both the east-trending Himalayan orogen and the north-trending Eastern and Western Flanking Belts located along the margins of the Indian subcontinent. Although the tectonic development of both flanking belts is key to understanding mechanisms of continental deformation during indenter-induced collision, few field-based studies coupled with geochronological and geochemical methods have been applied to these tectonic domains. In this study, we investigate the lateral correlation of lithologic units between the northern Indo-Burma Ranges, the northernmost segment of the Eastern Flanking Belt, and the eastern Himalayan-Tibetan orogen by integrating field observations, U-Pb zircon geochronology, and whole-rock geochemistry. Our findings provide new quantitative constraints to interpretations that the northern Indo-Burma Ranges expose the eastward continuation of several lithologic units of the Himalayan orogen and Lhasa terrane. Our field work documents a stack of thrust-bounded lithologic units present in the study area. The northernmost and structurally highest Lohit Plutonic Complex consists of Mesoproterozoic basement rocks (ca. 1286 Ma) and Late Jurassic–Cretaceous granitoids (ca. 156–69 Ma) with positive εNd values and initial 87Sr/86Sr ratios of ∼0.705, which are correlative to the Bomi-Chayu complex and the northern Gangdese batholith, respectively. The structurally lower Tidding-Mayodia mélange complex, composed of basalt, gabbro, ultramafic rocks, and mafic schist of a dismembered ophiolite sequence, is interpreted in this study as the eastward extension of the Indus-Yarlung suture zone. Structurally below the suture zone are the Mayodia gneiss and Lalpani schist, which are interpreted to correlate with the Lesser Himalayan Sequence based on comparable metamorphic lithologies, negative εNd values, and similar Mesoproterozoic–Cambrian detrital zircon age spectra. In contrast to the above metamorphic units, the structurally lowest Tezu unit consists of siliciclastic strata that may be correlated with the Miocene–Pliocene Siwalik Group of the Himalayan orogen. Despite the above correlations, notable Himalayan-Tibetan lithologic units are absent in the northern Indo-Burma Ranges, including the Mesozoic–Cenozoic southern Gangdese batholith belt and its cover sequence of the Linzizong volcanic rocks, Xigaze forearc basin, Tethyan Himalayan Sequence, and Greater Himalayan Crystalline Complex of south-central Tibet and the central Himalaya. We interpret the absence of these lithologic units to be a result of a greater magnitude of crustal shortening and/or underthrusting of the Indian cratonal rocks than that across the Himalayan orogen to the west. This interpretation is supported by a southward decrease in the map-view distance between the active range-bounding thrust and the Indus-Yarlung suture zone in the northern Indo-Burma Ranges, from ∼200 km in the north near the eastern Himalayan syntaxis to ∼5 km in the south across a distance of ∼200–300 km.
The Cenozoic India-Asia collision generated the Himalayan collisional system, which consists of the convergence-perpendicular Himalayan orogen and convergence-parallel flanking belts located along the eastern and western margins of the Indian subcontinent (Gansser, 1964; Yin, 2006) (Fig. 1A). Although the geology and tectonic evolution of the Himalayan orogen are well documented (Le Fort, 1975; Le Fort, 1996; DeCelles et al., 2000, 2001; Hodges, 2000; Yin and Harrison, 2000; Yin, 2006, 2010; Webb et al., 2013, 2017), relatively few geologic studies have focused on the flanking belts (cf. Tapponnier et al., 1981; Ni et al., 1989; Mitchell, 1993; Haq and Davis, 1997; Haproff et al., 2018). Establishing the geologic history of the two flanking belts is crucial for understanding the holistic development of the Himalayan collisional system and differentiating the end-member models of continental deformation during the India-Asia collision (e.g., Tapponnier et al., 1982, 2001; England and Houseman, 1986; Cobbold and Davy, 1988; Dewey et al., 1988; England and Molnar, 1990; Royden et al., 1997; Zuza et al., 2019).
In this study, we investigate the divisions of major lithologic units exposed in the northern Indo-Burma Ranges, the northernmost segment of the Eastern Flanking Belt (Fig. 1). We correlate these units with the classic lithologic divisions of the Himalayan orogen and the Lhasa terrane to the west. Correlations are based on U-Pb zircon geochronology, whole-rock geochemistry, bounding Cenozoic faults, and similar lithologies. From these findings, we conclude that several lithologic units of the Himalayan-Tibetan orogen, including the Greater Himalayan Crystalline Complex, Tethyan Himalayan Sequence, Xigaze forearc basin, and Mesozoic–Cenozoic southern belt of the Gangdese batholith are absent in the study area. If these lithologic units were present at the onset of the India-Asia collision, our work implies a greater magnitude of crustal shortening and/or continental underthrusting across the northern Indo-Burma Ranges than across the Himalaya to the west.
To provide a regional context, we summarize the existing research across the eastern Himalayan orogen, the Indus-Yarlung suture zone, the Lhasa terrane, cratonal rocks of northeast India, and the northernmost segment of the Eastern Flanking Belt. The classic lithologic division of the Himalayan orogen includes from south to north: the Sub-Himalayan Sequence, Lesser Himalayan Sequence (LHS), Greater Himalayan Crystalline Complex (GHC), and Tethyan Himalayan Sequence (THS) (Yin, 2006) (Figs. 1B and 2). These Himalayan units are bounded to the north by the Indus-Yarlung suture zone, which has been shortened by two postcollisional thrust systems: the older south-directed Gangdese thrust system and the younger north-directed Great Counter thrust system (Yin et al., 1994, 1999; Harrison et al., 2000). Located north of the Indus-Yarlung suture zone, the Lhasa terrane exposes three east-trending tectonic belts from south to north: the Xigaze forearc basin and the southern and northern belts of the Gangdese batholith (Fig. 2).
Eastern Himalayan Orogen
The eastern Himalayan orogen referred to in this study extends from the Sikkim-Bhutan border at 89°E to the eastern Himalayan syntaxis at ∼95°E (Figs. 1 and 2). The aforementioned major Himalayan lithologic units are tectonically juxtaposed by the following orogen-scale faults: the Main Frontal thrust at the base of the Sub-Himalayan Sequence, the Main Boundary thrust at the base of the LHS, the Main Central thrust at the base of the GHC, and the South Tibetan detachment at the base of the THS (e.g., Jangpangi, 1974; Acharyya and Ray, 1977; Gansser, 1983; Burchfiel et al., 1992; Acharyya, 1994; Bhargava, 1995; Edwards et al., 1996, 1999; Edwards and Harrison, 1997; Wu et al., 1998; Grujic et al., 2002; Yin et al., 2006, 2010a; McQuarrie et al., 2008; Long et al., 2011a, 2011b, 2011c; Burgess et al., 2012; Webb et al., 2013; DeCelles et al., 2016) (Fig. 2). Each of these faults is not expressed as a single structure but rather as a wide, km-scale zone of deformation comprising series of brittle faults and/or ductile shear zones (Hodges, 2000; Yin, 2006).
In the eastern Himalaya, the Sub-Himalayan Sequence consists of a ∼4–6-km-thick section of Miocene–Pliocene sandstone and conglomerate (Gansser, 1983; Acharyya, 1994; Dikshitulu et al., 1995; Kumar, 1997; Yin et al., 2006; McQuarrie et al., 2008; Yin et al., 2010a) (Fig. 2). The LHS is composed of Neoproterozoic–Cambrian and Permian strata (Gansser, 1983; McQuarrie et al., 2008; Yin et al., 2010a; Long et al., 2011a) (Fig. 2). Various local names have been assigned for the LHS units in the eastern Himalaya, and their correlative relationships are illustrated in Figure 2. The base of the Proterozoic strata in the eastern Himalaya is marked by a meter-thick layer of quartz pebble conglomerate that was deposited atop a ca. 1.7 Ga augen gneiss (Yin et al., 2010a).
The GHC in the eastern Himalaya consists of metasediments, orthogneiss, metavolcanics, and Cenozoic leucogranites (e.g., Gansser, 1983; Yin et al., 2010a; Webb et al., 2013). In Bhutan, the north-dipping Kakhtang thrust, correlative to the Zimithang thrust to the east (Yin et al., 2006), divides the GHC into the upper and lower parts (Gansser, 1983; Swapp and Hollister, 1991; Grujic et al., 1996, 2002; Davidson et al., 1997; Daniel et al., 2003; McQuarrie et al., 2008; Long and McQuarrie, 2010; Yin et al., 2010a; Long et al., 2011a, 2011b) (Fig. 2). The upper GHC consists of upper amphibolite-facies orthogneiss, metasedimentary rocks, and Miocene leucogranites (Gansser, 1983; Swapp and Hollister, 1991; Grujic et al., 1996, 2002; Davidson et al., 1997; Daniel et al., 2003; Long et al., 2011b, 2011c; Zeiger et al., 2015). In the footwall of the Kakhtang thrust, the lower GHC consists of upper amphibolite-facies to upper greenschist-facies metasedimentary rocks and orthogneiss (Davidson et al., 1997; Daniel et al., 2003; Corrie et al., 2012; Zeiger et al., 2015). East of Bhutan, orthogneiss in the GHC yields U-Pb zircon age populations of ca. 1700 Ma, ca. 878 Ma, and ca. 500 Ma (Yin et al., 2010a) (Fig. 2), which are cut by ca. 20–18 Ma leucogranites (Aikman et al., 2012a, 2012b; Harrison and Wielicki, 2016). In the easternmost Himalaya along the Siang River Valley, the Main Central thrust splits into two thrusts referred to as the MCT-I and MCT-II by Nandini and Thakur (2011) (Fig. 2).
The THS above the South Tibetan detachment consists of pre-Triassic metasedimentary strata (Zeng et al., 2011) and Triassic–Cretaceous isoclinally folded marine strata (Yin et al., 1994, 1999; Harrison et al., 2000; Aikman et al., 2008) (Fig. 2). Folded THS strata are intruded by ca. 44 Ma granite (Aikman et al., 2008). Due to the merger of the Great Counter thrust and the South Tibetan detachment northwest of the eastern Himalayan syntaxis, the THS does not extend to the western margin of the eastern Himalayan syntaxis (Yin, 2006; Webb et al., 2013) (Figs. 1B and 2). The eastern Himalayan syntaxis is marked by the ∼90° oroclinal bend in the general strike of thrust faults including the MFT, where the east-trending Himalayan orogen transitions to the north-trending northern Indo-Burma Ranges (Fig. 1).
Indus-Yarlung Suture Zone
The Indus-Yarlung suture zone consists of tectonic mélange complexes, dismembered ophiolitic sequences, subduction-related metamorphic rocks, syntectonic conglomerate deposits, and suprasubduction-related igneous rocks (Gansser, 1964; Honegger et al., 1982; Allégre et al., 1984; Malpas et al., 2003; Ziabrev et al., 2003; Dai et al., 2011a, 2011b; Hébert et al., 2012; Cai et al., 2012; An et al., 2014; Laskowski et al., 2016; Leary et al., 2016) (Fig. 1B). The suture zone involves Early Cretaceous ultramafic rocks, volcanics, volcaniclastic strata, granitoids, and carbonate, and fossiliferous chert and flysch are intermittently exposed along the northern margin of the Himalayan orogen (McDermid et al., 2002; Malpas et al., 2003; Ziabrev et al., 2003; Dubois-Côté et al., 2005; Dupuis et al., 2005; Zhou et al., 2005; Aitchison et al., 2007; Zhu et al., 2009b). These rocks are bounded by two Cenozoic thrust systems: the south-directed Gangdese thrust system of Harrison et al. (1992) and Yin et al. (1994) and the Great Counter thrust of the Heim and Gansser (1939).
The Lhasa terrane is a ∼200–300-km-wide continental strip that trends east-west in southern Tibet and wraps around the eastern Himalayan syntaxis to a north-south trend in northern Myanmar and western Yunnan of China (e.g., Lin et al., 2013; Wang et al., 2014) (Fig. 1B). The terrane is divided into northern and southern halves by an east-trending ophiolite belt, which was interpreted to be a suture zone or a zone of rootless klippe sourced from the Bangong-Nujiang suture zone in the north (Girardeau et al., 1984; Coward et al., 1988; Hsü et al., 1995; Yin and Harrison, 2000; Kapp et al., 2003; Dong et al., 2011, Zhu et al., 2012; Zhang et al., 2014). The Lhasa terrane exposes Mesoproterozoic orthogneiss in the easternmost Bomi-Chayu Complex (ca. 1250–1350 Ma) (Xu et al., 2013), a composite Neoproterozoic (ca. 920–820 Ma) and Cambro-Ordovician (ca. 540–460 Ma) gneiss complex along the northern margin of the Lhasa terrane (Guynn et al., 2012), and fragments of Devonian orthogneiss (ca. 366 Ma) in southeast Tibet (Zhu et al., 2011) (Figs. 1 and 2).
The salient feature of the Lhasa terrane is the >2000-km-long Mesozoic to earliest Cenozoic Gangdese batholith belt of Chang and Zheng (1973) (Fig. 1B). The batholith belt can be divided into the northern and southern belts: the northern belt is characterized by Late Triassic plutons (Wang et al., 2016), Cretaceous adakite (Zhu et al., 2009a), and S-type granitoids (Wen et al., 2008), whereas the southern belt is dominated by Cretaceous to Eocene granitoids overlain by Paleocene–Eocene volcanic strata (Schärer et al., 1984; Coulon et al., 1986; Copeland et al., 1995; Ding et al., 2003; Wen et al., 2008; Zhu et al., 2008; Ji et al., 2009; Lee et al., 2009; Guan et al., 2012).
The southern part of the Lhasa terrane is marked by the Gangdese thrust system, which juxtaposes strata of the Xigaze forearc basin strata and the Gangdese batholith over the THS (Fig. 2) (Yin et al., 1994). The Xigaze forearc basin sequence consists of Cretaceous to Eocene marine deposits on top of an oceanic basement (Garzanti and Van Haver, 1988; Einsele et al., 1994; Dürr, 1996; Ding et al., 2005; Wang et al., 2012; An et al., 2014; Hu et al., 2015; Orme et al., 2015). The forearc strata are only exposed in south-central Tibet and are absent in southwest and southeast Tibet along strike. This map pattern is interpreted to be a result of postcollisional underthrusting of the Xigaze forearc basin below the Gangdese batholith (Yin et al., 1994, 1999; Harrison et al., 2000). The Cretaceous strata are intruded by a series of north-trending dikes that yield ages of ca. 18–14 Ma (Yin et al., 1994; Williams et al., 2001).
Cratonal Rocks of Northeast India
The 400-km-long, east-trending Shillong Plateau of northeast India exposes Indian cratonal rocks that are correlative to the Proterozoic metasedimentary strata and Precambrian crystalline rocks of the LHS and GHC (e.g., Gansser, 1983; Gupta and Biswas, 2000; Srivastava and Sinha, 2004a, 2004b; Srivastava et al., 2005; Richards et al., 2006; Tewari et al., 2010; Yin et al., 2010b) (Fig. 1A). Cenozoic deformation is expressed by east-striking thrust faults that bound the margins of the plateau (e.g., Clark and Bilham, 2008; Yin et al., 2010b). Active deformation within the Shillong Plateau is evident by multiple northeast-striking left-slip faults that offset Quaternary geomorphic features and sediments (Yin et al., 2010b).
Northern Indo-Burma Ranges
The research presented here is focused on the northern Indo-Burma Ranges, the northernmost segment of the Eastern Flanking Belt, which is located north of latitude 27°N (Fig. 1). Previous geological studies have established the first-order lithologic framework of the northern Indo-Burma Ranges (e.g., Wadia, 1931; Nandy, 1973; Thakur and Jain, 1975; Acharyya, 1980, 1987; Sharma et al., 1991; Singh, 1993; Misra and Singh, 2002; Gururajan and Choudhuri, 2003; Goswami, 2008, 2011; Misra, 2009; Sarma et al., 2009; Sarma et al., 2012; Goswami, 2013a, 2013b; Sharma and Sarma, 2013; Ningthoujam et al., 2015). Most recently, Haproff et al. (2018) divided the exposed rocks into six lithologic units (Figs. 3–5), which from northeast to southwest consist of: (1) igneous rocks of the Lohit Plutonic Complex (Fig. 6); (2) the Tidding mélange complex (Fig. 7); (3) metamorphic rocks of the Mayodia gneiss (Fig. 8); (4) metasedimentary rocks of the Lalpani schist (Fig. 9); (5) siliciclastic strata of the Tezu unit (Fig. 10A); and (6) metasedimentary rocks of the Sewak unit (Fig. 10B). In this paper, we further subdivide the exposures of mélange rocks into the Tidding and Mayodia mélange complexes based on their map-view separation (Fig. 3). Each lithologic unit is bounded by south- to west-directed thrust faults including (1) the Walong thrust, (2) the Lohit thrust, (3) the Tidding thrust, (4) the Demwe thrust, (5) the Lalpani thrust, (6) the Tezu thrust, and (7) the Mishmi thrust (Fig. 2). Despite extensive work in the northern Indo-Burma Ranges, disagreement remains whether these lithologic units are correlative with those of the Himalayan orogen and the Lhasa terrane (e.g., Thakur and Jain, 1975; Acharyya, 1980; Singh and Chowdhury, 1990; Gururajan and Choudhuri, 2003; Misra, 2009). Furthermore, previous lithologic correlations are solely based on comparable lithologies and remain untested with geochronologic or geochemical techniques (e.g., Gururajan and Choudhuri, 2003; Misra, 2009).
This study builds upon the work of Haproff et al. (2018), which described the structural framework of the northern Indo-Burma Ranges with a focus on the kinematics of major faults. Here we focus on the lithologies, geochemical compositions, and geochronologic ages of the thrust-bounded lithologic units and investigate how they correlate with those of the Himalayan orogen and the Lhasa terrane to the west (Fig. 2). Such an approach provides additional constraints on the possible evolutionary history of the Himalayan collisional system including both the Himalayan orogen and Eastern and Western Flanking Belts.
SAMPLING AND ANALYTICAL METHODS
A total of 65 samples were collected along the Lohit and Dibang Valleys for U-Pb zircon geochronology and whole-rock geochemistry (Figs. 3–5). Sampling locations and detailed methodologies are described below. We also present additional geologic maps of Dibang Valley (Fig. 4) and more detailed descriptions of lithologic units.
U-Pb Zircon Geochronology of the Lohit Plutonic Complex
Zircon grains from ten samples of the Lohit Plutonic Complex were analyzed for U-Pb crystallization ages using secondary ion mass spectrometry (SIMS) (Cameca ims-1270 ion microprobe) at University of California, Los Angeles (UCLA) (Table 1). Prior to analysis, zircon grains were mounted with standard AS3 (1099 Ma; Paces and Miller, 1993) on 1-inch-diameter epoxy mounts, polished with carbide paper, and coated with ∼100 Å of gold. Cathodoluminescence (CL) images were taken using the scanning electron microscope (SEM) at UCLA to observe any zonation. Zircons were sputtered using a 10–15 nA O– primary beam on ∼25-μm-diameter spots on zircon rims. U-Pb isotopic ratios (206Pb/238U and 207Pb/235U) were calculated based on a calibration curve of UO/U versus Pb/U, and corrected for common lead (Stacey and Kramers, 1975). 207Pb/206Pb ages were reported for zircon grains of >1000 Ma age. Data reduction was accomplished using the program ZIPS 3.0.3, and weighted-mean ages, concordia diagrams, and relative probability plots were generated using Isoplot/Ex (Ludwig, 1991). Data tables of all zircon analyses are shown in Tables S1 and S2 in the Supplemental Materials1. Analyses were excluded for low radiogenic lead concentrations, large analytical errors, inherited zircon grains that were significantly older than the dominant zircon-age population, and discordant ages. We calculated the weighted-mean age of the youngest population of concordant analyses to estimate the crystallization age of the plutonic samples.
U-Pb Detrital Zircon Geochronology
Detrital zircons from 15 metasedimentary rocks from the Sewak unit, Lalpani schist, Mayodia gneiss, and Tidding and Mayodia mélange complexes were dated via U-Pb geochronology (Table 2). Complete isotopic data of all detrital zircon analyses are shown in Tables S3–S6 (see footnote 1). Zircon age distributions were used to determine the detrital provenance, distinguish lithologic units, and test correlations with metasedimentary rocks of the Himalayan orogen to the west of the study area.
Zircon grains were separated from rocks at UCLA using standard procedures (Quidelleur et al., 1997; Schmitt et al., 2003a, 2003b). Zircons were mounted on 1-inch-diameter epoxy rounds with 91500 zircon standards (1065 Ma age; Wiedenbeck et al., 1995) and were polished with carbide paper. Zircons from two samples were analyzed by laser ablation–inductively coupled plasma mass spectrometry (LA-ICP MS) at the University of California, Santa Barbara (see Kylander-Clark et al., 2013 for detailed analytical methods). Spot sizes of ∼20 μm were each shot twice with an excimer laser at a 5 Hz pulse rate. Zircon grains from another 13 samples were analyzed using the method of Liu et al. (2018) at the Institute of Tibetan Plateau Research, Chinese Academy of Sciences in Beijing, China.
Whole-rock major- and trace-element geochemical analyses were conducted on six plutonic samples from the Lohit Plutonic Complex and nine mafic and ultramafic samples. Samples selected for whole-rock geochemical analyses were crushed using a steel mortar and pestle and were pulverized. Major-, minor-, and trace-element compositions were determined by LA-ICP MS at Activation Laboratories (Actlabs) in Ontario, Canada. Major- and minor-element compositions of nine whole-rock samples were determined at Pomona College by fusion of Li tetraborate and X-ray fluorescence (XRF) analysis using a PanAnalytical Axios wavelength-dispersive instrument. XRF instrument calibration and sample preparation methods are based on Johnson et al. (1999). Trace-element compositions were determined via ICP-MS at the Institute of Tibetan Plateau Research, Chinese Academy of Sciences in Beijing using the methods of Liu et al. (2018). Results of major- and trace-element geochemistry are shown on Tables S7 and S8 (see footnote 1). Four samples of the western Lohit Plutonic Complex belt were cut into thin sections and viewed with a petrographic microscope to determine the modal abundances of quartz, plagioclase, and alkali feldspar for granitoid classification (Fig. S4B [see footnote 1]).
Twenty-one plutonic and metasedimentary rock samples from each major lithologic unit were analyzed for 143Nd/144Nd and/or initial 87Sr/86Sr (Sri) isotopic ratios to understand the source and setting of magma genesis and to test correlation with rocks of the Himalayan orogen (Table 3). Prior to analysis, powdered samples were first placed in Teflon bombs, diluted with H2O, HNO3, and HF, and oven heated at 190 °C for 36 h. Four ml of H2O, HClO4, and HNO3 were added to solution, along with 0.1 ml Rh and Re internal standard. Final sample preparation and analyses were performed using the methods of Liu et al. (2018) via both LA-ICP-MS and thermal ionization mass spectrometry (TIMS) at the Institute of Tibetan Plateau Research in Beijing and Guangzhou, China.
LITHOLOGIC UNITS OF THE NORTHERN INDO-BURMA RANGES AND RESULTS OF U-Pb ZIRCON GEOCHRONOLOGY
Lohit Plutonic Complex
The northernmost mapped unit is the Lohit Plutonic Complex of Nandy (1973), which is divided into western and eastern belts by the north-dipping Walong thrust (Gururajan and Choudhuri, 2003) (Figs. 3, 4A, and 5). Ten samples were collected from the eastern and western belts to constrain the spatial distribution of U-Pb zircon ages.
Western Lohit Plutonic Complex Belt
The western belt of the Lohit Plutonic Complex consists of diorite, granodiorite, monzodiorite, tonalite, quartz monzonite (Fig. 6A), and younger leucogranite and mafic dikes thrust over the Tidding mélange complex (Fig. 4A). The locations of six samples collected for U-Pb zircon geochronology are shown on Figs. 3, 4A, and 5. Samples contain quartz, plagioclase, potassium feldspar, biotite, and variable hornblende and garnet. Some samples also contain epidote, clinozoisite, rutile, ilmenite, and zoned garnet. Cathodoluminescence images of representative zircons analyzed for U-Pb geochronology are shown on Figures S3A–S3C (see footnote 1). Zircon grains are euhedral, prismatic crystals with average long axes of ∼50–150 μm. Most grains display oscillatory zoning. The outermost rims of the dated zircon grains were targeted with ∼25 μm ion beam spots to date the youngest crystallization event. Complete U-Pb isotope data can be found in Table S1.
U-Pb zircon ages of the western belt are clustered at 93–115 Ma and 156 Ma (Fig. S1 [see footnote 1] and Table 1). The weighted-mean U-Pb ages (2σ error) of five granitoids from Dibang Valley are 96 ± 3 Ma (PH-1-8-13-1B), 105 ± 3 Ma (PH-1-8-13-3), 94 ± 20 Ma (PH-1-8-13-4B), 93 ± 3 Ma (PH-1-8-13-7), and 115 ± 13 Ma (PH-1-8-13-8). Lohit Valley sample PH-1-5-13-9 yields a U-Pb zircon age of 156 ± 7 Ma (Table 1).
Eastern Lohit Plutonic Complex Belt
The eastern belt of the Lohit Plutonic Complex consists of foliated diorite, garnet-bearing orthogneiss, and migmatite intermingled with marble bands, intruded by younger aplite, pegmatite, and leucogranite (Fig. 6B). Three foliated granitoid samples were collected from Dibang Valley and one from Lohit Valley (Figs. 3 and 5). Granitoids contain quartz, plagioclase, potassium feldspar, biotite, with variable calcic amphibole, garnet, rutile, epidote, and clinozoisite. Cathodoluminescence images of dated zircon grains are shown on Figure S4D (see footnote 1). Zircon grains are typically ∼50–100 μm in the longest dimension, transparent in color, and prismatic in shape. Zircon rims were targeted for dating the youngest crystallization age. Complete isotopic age data can be found in Table S2.
The eastern belt samples yield a wider range of U-Pb zircon ages compared to those of the western belt (Fig. S2 [see footnote 1] and Table 1). Sample PH-11-10-15-15 is a garnet-bearing orthogneiss from the middle part of the eastern belt in Dibang Valley (Fig. 3); the sample yields U-Pb zircon ages of 57–95 Ma with a bimodal distribution centered at ca. 70 Ma and ca. 92 Ma, respectively. The weighted-mean age of the entire U-Pb age distribution is 75 ± 9 Ma (Fig. S2A and Table 1).
Orthogneiss sample PH-11-11-15-1 from the northernmost part of the eastern belt (Fig. 3) has a bimodal age distribution between 69 and 165 Ma (Fig. S2B [see footnote 1]). The weighted-mean age of the older age group is 136 ± 5 Ma, whereas the weighted-mean age of the younger group is 69 ± 9 Ma (Fig. S2B).
Sample PH-11-10-15-13 was collected from a migmatitic garnet-bearing orthogneiss unit in the southernmost part of the eastern belt (Fig. 3). Twelve zircon grains from the sample yield 207Pb/206Pb ages of 1168–1558 Ma (Tables 1 and S2 [see footnote 1]). Nine other zircon grains produce an age cluster of 1258–1343 Ma with a weighted mean of 1286 ± 14 Ma (Fig. S2C).
Tidding and Mayodia Mélange Complexes
The 8–10-km-thick Tidding and Mayodia mélange complexes are exposed in the hanging wall of the Tidding thrust (Figs. 3, 4B, and 5). Both units were originally mapped as the Tidding suture zone (Nandy, 1973; Thakur and Jain, 1975; Gururajan and Choudhuri, 2003) and the Tidding Formation by Misra (2009). In Dibang Valley, the Mayodia mélange complex is exposed as an incoherent mix of dismembered garnet mica schist, metabasite, chert, amphibolite, gabbro, and serpentinized ultramafics within the Mayodia klippe located near the foreland of the thrust belt (Figs. 4A and 7A). To the north, the Tidding mélange complex consists of amphibolite and chlorite muscovite schist (Figs. 3 and 7B). In Lohit Valley, the Tidding mélange complex consists of a single exposure of incoherently mixed amphibolite and chlorite-muscovite schist in the hanging wall of the Tidding thrust (Fig. 5). Although we mapped the Tidding and Mayodia mélange complexes as separate units based on their map-view separation, we interpret the units to originate from the same complex, which was telescoped across the orogen in the hanging wall of the Tidding thrust and subsequently folded to form an isolated klippe in the foreland (Figs. 3 and 4B).
One chlorite muscovite schist sample (PH-1-9-13-10) from the Mayodia mélange complex (Fig. 4B) yields three youngest zircon ages (±2σ error) of 940 ± 9 Ma, 1326 ± 14 Ma, and 1654 ± 15 Ma (Tables 2 and S3 [see footnote 1]). A wide age population occurs at ∼1200–1600 Ma (Fig. 11A). Garnet-mica schist sample PH-1-8-13-26 collected from the Tidding mélange complex in Dibang Valley (Fig. 3) contains age populations of ca. 500–600 Ma, ca. 800 Ma, and ca. 1000 Ma (Fig. 11A), and a youngest zircon age of 40 ± 1 Ma (three grains ≤51 Ma) (Tables 2 and S3).
The 1.5–4-km-thick Mayodia gneiss was originally mapped as the Mishmi Crystallines by Gururajan and Choudhuri (2003) and later the Mayodia Group by Misra (2009) (Figs. 3, 4C, and 5). Lithologies include paragneiss, augen gneiss, and quartzofeldspathic schist (Fig. 8A). Meter-wide syntectonic leucogranite dikes intrude the unit near its northernmost exposure (Figs. 3 and 8B). The Mayodia gneiss is in the hanging wall of the north-dipping Demwe thrust, juxtaposed against the Lalpani schist below (Figs. 3, 4C, and 5). In Dibang Valley, the unit is exposed bounding the Mayodia klippe, which originated from the root zone of the Demwe thrust (Figs. 3 and 4B).
Five metasedimentary rocks of the Mayodia gneiss were collected from the basal to uppermost structural sections of the southernmost exposure of the unit in Dibang Valley (Fig. 4C). Mylonitic paragneiss samples PH-1-9-13-19A and PH-1-9-13-19B (located ∼50 m upsection) from the basal section of the Mayodia gneiss both contain significant age populations at ca. 1100–1200 Ma, ca. 1300 Ma, and ca. 1400 Ma (Fig. 11B), and Mesoproterozoic youngest ages of 1073 ± 19 Ma (three grains ≤1181 Ma) and 1022 ± 16 Ma (three grains ≤1035 Ma), respectively (Tables 2 and S4 [see footnote 1]).
Augen gneiss sample PH-1-3-13-9, collected from the structural middle section of the unit (Fig. 4C), contains a large ca. 1100 Ma peak and smaller populations at ca. 800 Ma and ca. 1200–1300 Ma (Fig. 11B) and a youngest zircon age of 499 ± 3 Ma (three grains ≤791 Ma) (Tables 2 and S4 [see footnote 1]).
Garnet gneiss sample PH-1-3-13-8, collected from the structurally uppermost section of the Mayodia gneiss (Fig. 4C), contains a strong peak at ca. 1200 Ma, and minor populations at ca. 1100 Ma, ca. 1300–1400 Ma (Fig. 11B), and a youngest zircon age of 1069 ± 13 Ma (three grains ≤1079 Ma) (Tables 2 and S4 [see footnote 1]). Schist sample PH-1-3-13-10A, collected upsection from PH-1-3-13-8 (Fig. 4C), contains age populations at ca. 800 Ma, ca. 1100–1200 Ma, ca. 1300–1400 Ma, and ca. 1700 Ma (Fig. 11A) and a Neoproterozoic youngest zircon age of 646 ± 13 Ma (three grains ≤841 Ma) (Tables 2 and S4).
The Lalpani schist is exposed in both Lohit and Dibang Valleys (Figs. 3, 4C, and 5) and consists of quartzofeldspathic schist, paragneiss, quartzite, and carbonate, intruded in some locations by meter-wide mafic dikes (Fig. 9). The unit was originally mapped as the lower section of the Mishmi Crystallines by Gururajan and Choudhuri (2003) and later the Lalpani Group by Misra (2009).
Seven samples of the Lalpani schist were collected for U-Pb zircon geochronology (Figs. 3 and 4C). The youngest U-Pb zircon ages from the samples range from Mesoproterozoic for the basal unit to Cretaceous in the uppermost section, signaling changes in the source areas (Fig. 11A). Two quartzofeldspathic paragneiss samples, PH-1-12-13-5 and PH-1-9-13-27 from the basal section of the Lalpani schist, ∼9 km apart from one another along strike (Fig. 4C), yield nearly identical U-Pb age spectra with a significant peak at ca. 2500 Ma (Fig. 11B) and a youngest individual zircon age of 913 ± 5 Ma (three grains ≤1572 Ma) and 974 ± 6 Ma (three grains ≤1195 Ma), respectively (Tables 2 and S5 [see footnote 1]).
Schist sample PH-1-9-13-23 is structurally higher than the previous two samples (Fig. 4C) and yields zircon populations of ca. 1100–1200 Ma, ca. 1400 Ma, and ca. 1600 Ma (Fig. 11B). The youngest individual zircon age was 1054 ± 19 (three grains ≤1166 Ma) (Tables 2 and S5 [see footnote 1]).
Paragneiss sample PH-1-3-13-11B, collected from the upper-middle structural section of the Lalpani schist (Fig. 4C), contains zircon grains with Mesoproterozoic to Cambrian ages clustered at ca. 1300 Ma, ca. 800–900 Ma, and ca. 700 Ma (Fig. 11B). The sample yields a youngest age of 525 ± 3 Ma (three grains ≤668 Ma) (Tables 2 and S5 [see footnote 1]).
Samples PH-1-12-13-7 and PH-1-9-13-25 are mylonitic quartzofeldspathic paragneiss collected from the lowermost structural section of the Lalpani schist (Fig. 4C). Both samples yield identical U-Pb age spectra, with a strong peak at ca. 500 Ma (Fig. 11A) and youngest zircon ages of 158 ± 1 Ma and 177 ± 1 Ma, respectively (Tables 2 and S5 [see footnote 1]).
Schist sample PH-1-9-13-2, collected from the Hunli half-window (Fig. 3), yields a youngest zircon age of 150 ± 4 Ma (three grains ≤200 Ma), like samples PH-1-12-13-7 and PH-1-9-13-25 (Tables 2 and S5 [see footnote 1]). However, sample PH-1-9-13-2 has a larger distribution of U-Pb ages with populations at ca. 150–350 Ma, ca. 400 Ma, ca. 500–700 Ma, and ca. 800–1000 Ma (Fig. 11A).
The Tezu unit is a ∼2-km-thick sequence of non-marine, syntectonic strata including interbedded conglomerate, coarse-grained sandstone, and mudstone (Fig. 10A). The unit was observed in only one location along the range front, directly northwest of Lohit Valley (Fig. 5). Conglomeratic layers are defined by cobble-sized clasts exposed in meter-wide channels within a sandy matrix. The north-dipping Tezu unit section is oriented right-way-up, evidenced by normal grading and conglomeratic channels. The Tezu unit is in the footwall of the Lalpani thrust, juxtaposed against the Lalpani schist (Fig. 5). U-Pb detrital zircon geochronology was not performed on samples from this unit.
The Sewak unit, which was originally mapped by Misra (2009) as the Sewak Group, consists of a ∼1.5-km-thick section of low-grade metasedimentary rocks including interbedded quartzite, marble, chert, slate, phyllite, and quartzofeldspathic schist (Fig. 10B). The unit is exposed along the range front in the hanging walls of the Tezu thrust in Lohit Valley and Mishmi thrust in Dibang Valley (Figs. 3, 4C, and 5). In Dibang Valley, the Sewak unit is the southernmost lithologic unit thrust atop Quaternary alluvium by the active Mishmi thrust (Figs. 3 and 4C).
Phyllite sample PH-1-14-13-4 from the middle structural level of the Sewak unit (Fig. 4C) yields a single youngest zircon age of 27 ± 1 Ma, with three spot analyses of ≤30 Ma (Tables 2 and S6 [see footnote 1]). The sample displays prominent zircon age populations of ca. 90 Ma, ca. 130 Ma, ca. 500–600 Ma, ca. 800 Ma, and ca. 1100–1200 Ma (Fig. 11B).
RESULTS OF WHOLE-ROCK GEOCHEMISTRY
The Lohit Plutonic Complex consists of predominantly calc-alkaline diorite suite rocks featuring both I- and S-type major- and/or trace-element signatures (Figs. 12A, 13A, and S5 and Table S7 [see footnote 1]). All granitoid samples are peraluminous diorites with high Al2O3 (>19 wt%), A/CNK values (>1.58), A/NK (>2.75), and Rb (>100 ppm) (Fig. 12B and Table S7). One sample from the eastern Lohit Plutonic Complex belt (PH-11-10-15-13) is a garnet-bearing, low-K granodiorite (Fig. 12A and Table S7).
Trace-element signatures of five samples of the Lohit Plutonic Complex are consistent with volcanic arc origins, along with generation in an orogenic setting (Pearce et al., 1984) (Fig. 13B). Rare-earth element concentrations of four granitoids reflect moderate to high fractionation between heavy rare-earth elements (HREEs) and light rare-earth elements (LREEs) (La/Yb: ∼7–15) (Figs. S5 and S6 and Table S7). One granitoid intrusion of the Tidding mélange complex (PH-1-8-13-22) is a metaluminous diorite based on moderate to high Na2O (>3.75 wt%), high CaO (>5.8 wt%), low Rb (<35 ppm), and coexisting hornblende and sphene (Table S7).
Mafic and Ultramafic Rocks
Five mafic samples collected from the Mayodia mélange complex and mafic dikes in the Mayodia gneiss and Lohit Plutonic Complex yield compositions ranging from basalt and basaltic andesite on a SiO2 versus total alkalis diagram (Le Bas et al., 1986) (Fig. 14A). One sample (PH-1-9-13-7B) from the Mayodia mélange complex contains stichtite and SiO2 <45%, implying the presence of ultramafic blocks within the mélange (Fig. 14A and Table S8). Seven mafic samples predominantly fall in the mid-ocean ridge basalt (MORB) (Figs. 14B and S7) and oceanic-arc fields on trace-element discriminant diagrams (Nb/La versus La/Yb; V versus Ti; Ti versus Zr; Cr versus Y; Zr/Y versus Zr) (Pearce, 1982; Shervais, 1982; Condie, 1989) (Fig. S7 and Table S8). Mafic rock samples display relatively flat trace-element slopes (La/Yb: ∼1.5–4) on a spider diagram, in contrast to sample PH-1-9-13-7B interpreted to have an ultramafic protolith (La/Yb: ∼9) (Fig. S8 and Table S8).
Sr and Nd Isotope Geochemistry
Neodymium and strontium isotopic compositions aid in differentiating protoliths of the LHS from the GHC (e.g., Parrish and Hodges, 1996; Huyghe et al., 2001; Robinson et al., 2001; Aikman et al., 2012a) and identifying the source and evolution of igneous rocks (e.g., Kistler and Ross, 1990; Chung et al., 1998). We determined the isotopic ratios of 143Nd/144Nd (normalized to chondritic uniform reservoir [CHUR] in epsilon notation) and initial 87Sr/86Sr (Sri) for igneous rocks of the western Lohit Plutonic Complex belt (Figs. 15 and 16) and Tidding and Mayodia mélange complexes to understand source and setting of magmagenesis (Fig. 16). 143Nd/144Nd and Sri ratios were determined for the Mayodia gneiss, Lalpani schist, and Sewak unit to test correlation with the GHC and LHS of the Himalaya orogen (Fig. 16). Results of Sr and Nd isotopic analyses are summarized in Table 3.
Five granitoid samples from the western Lohit Plutonic Complex belt yield positive εNd values of 2.2–5.1 and Sri values of 0.704–0.705 (Figs. 15 and 16), comparable to those of I-type Gangdese granitoids of the southern Lhasa terrane (Fig. 15). Three samples from the Mayodia mélange complex, located near the Tidding thrust, yield positive εNd values of 5.0–7.7 and Sri of 0.704–0.706, whereas two samples of the Tidding mélange complex near the Lohit thrust contain εNd values of −7.2 and 0.8 and Sri values of 0.706–0.712 (Fig. 16). Four samples from the Mayodia gneiss yield εNd values between −17.5 and −10.3 and Sri values of 0.715–0.912 (Fig. 16). Three Lalpani schist samples yield negative εNd values of −27.8 to −16 (Fig. 16), indicative of a Lesser Himalayan affinity. The same Lalpani schist samples have Sri values between 0.709 and 0.895 (Fig. 16). One Lalpani schist sample collected from the Hunli window (Fig. 3) has a less negative εNd value of −5.5 and Sri of 0.709 (Fig. 16). One phyllite sample of the Sewak unit has a εNd value of −10.2 and Sri of 0.712 (Fig. 16).
REGIONAL CORRELATION OF LITHOLOGIC UNITS
Our results allow us to correlate the major rock units of the northern Indo-Burma Ranges with those of the Himalaya orogen and the Lhasa terrane (Fig. 2). As detailed below, several Himalayan-Tibetan lithologic units are missing in the study area, which implies either a unique tectonic process controlled the development of the northern segment of the Eastern Flanking Belt or possibly a greater magnitude of crustal shortening and/or continental underthrusting occurred across the study area compared to the Himalayan orogen to the west (Fig. 17).
The Sewak unit and Tezu units are the structurally lowest rocks exposed in the study area (Fig. 2). The Sewak unit is comparable to the Sub-Himalayan Sequence based on structural position, and the presence of young zircon ages of <30 Ma in both the Sub-Himalayan Sequence and Sewak unit (Fig. 11B) strengthens this correlation (Yin, 2006) (Fig. 2). However, we note that the presence of low-grade metamorphic rocks in the Sewak unit makes it different than the Sub-Himalayan Sequence.
The Tezu unit has a similar lithology to that of the Miocene–Pliocene Siwalik Group in the Sub-Himalayan Sequence (e.g., DeCelles et al., 1998). The structural position of the Tezu unit atop the Indian foreland basin by an active, range-bounding fault (i.e., Mishmi thrust) is also like that of the Sub-Himalayan Sequence (Yin, 2006) (Fig. 2).
Lesser Himalayan Sequence and Greater Himalayan Crystalline Complex
Because both the LHS and GHC consist of medium- to high-grade metamorphic rocks, differentiating the units without a clear structural context can be challenging. The strike-slip faults bounding the eastern Himalayan syntaxis (Fig. 1C) prevent the testing of the along-strike continuity of major thrusts between the northern Indo-Burma Ranges and Himalayan orogen. However, we compared U-Pb detrital zircon ages between lithologic units (Gehrels et al., 2011) using the nonparametric K-S statistical test for the Mayodia gneiss, Lalpani schist, LHS, and GHC. The tests were conducted using the Excel macro generated by George Gehrels at the University of Arizona LaserChron Center (https://sites.google.com/a/laserchron.org/laserchron/). In the test, the observed vertical difference (Dobs) between the cumulative probability distributions is compared to a critical value (Dcrit), determined from the number of zircon analyses per sample. In this study, a maximum Dobs value of 0.05 is established. If the Dobs is greater than Dcrit, the null hypothesis that the samples are derived from the same source can be rejected. The P value is the maximum value of the significance level at which the null hypothesis can be accepted. If the P value is greater than 0.05, then the zircon age distributions are derived from the same source at the 95% confidence level. Results of the K-S statistical test are summarized in Table S9 (see footnote 1).
The test shows that schist sample PH-1-3-13-10A of the Mayodia gneiss is similar with rocks of the Arunachal Himalaya, specifically sample AY 02-13-06-7 of the GHC (Webb et al., 2013) and sample AY9160314A of the LHS (Yin et al., 2006) (Table S9 [see footnote 1]). However, sample AY 02-13-06 from the Arunachal Himalaya is located within the Main Central thrust shear zone, which in some locations extends into the upper LHS (Yin, 2006). We interpret the Mayodia gneiss to correlate with the LHS based on the similarity of detrital zircon ages with sample AY9160314A, comparable metamorphic lithologies, and similar negative εNd values of −17.5 to −10.3 (Fig. 16 and Table 3). Furthermore, the Mayodia gneiss lacks (1) ca. 870 Ma and ca. 500 Ma orthogneiss; (2) large leucogranite sheets or laccolith bodies (e.g., Arunachal and Tsona leucogranites); and (3) significant U-Pb detrital zircon age populations at ca. 540–750 Ma and ca. 800–1200 Ma (Fig. 11B), all of which are common in the GHC (Yin et al., 2010a; Gehrels et al., 2011; Aikman et al., 2012a, 2012b; Webb et al., 2013; Harrison and Wielicki, 2016). Similarly, the Lalpani schist is correlated to the LHS based on (1) similar lithologies and metamorphic grades; (2) negative εNd values mostly between −27.8 and −16 (Fig. 16 and Table 3); and (3) the presence of the Mesoproterozoic−Cambrian detrital zircon ages (Fig. 11B).
Indus-Yarlung Suture Zone
The Tidding and Mayodia mélange complexes are interpreted to originate from the same complex, which correlates with the Indus-Yarlung suture zone. This is based on (1) the exposure of amalgamated basalts, gabbro, ultramafics, and mafic schist of a dismembered ophiolite sequence within the Tidding-Mayodia mélange complex; (2) generation of igneous rocks in a mid-ocean ridge setting (Fig. 14B); and (3) the position of the mélange complex between rocks of Indian and Lhasa terrane affinities (Figs. 3 and 5). Metasedimentary rocks containing Eocene zircons (ca. 40 Ma) (Fig. 11A) could have been part of a syntectonic basin within the suture zone.
Mesozoic Magmatic Arc and Mesoproterozoic Basement of the Lhasa Terrane
Following the interpretation of Lin et al. (2013), we suggest that the Lohit Plutonic Complex is the southeastward continuation of the Mesozoic Cenozoic northern Gangdese batholith belt based on (1) similar Jurassic–Cretaceous crystallization ages between ca. 69–156 Ma (Table 1); (2) low Th/Y and La/Yb ratios in I-type granitoids (Fig. 13A); (3) trace-element signatures indicating volcanic arc origins (Fig. 13B); (4) Sri values of ≤0.705 (Fig. 15 and Table 3); (5) positive εHf values (10–14) (Lin et al., 2013); and (6) positive εNd values (Figs. 15 and 16 and Table 3). Furthermore, similar Cretaceous zircon ages along with εNd and Sri values are found in the Wuntho-Popa and Mogok Metamorphic belts to the south (Mitchell, 1993; Mitchell et al., 2012; Wang et al., 2014). The existence of a continuous magmatic arc stretching from the southern Lhasa terrane to the West Burma block and Sibumasu plate supports the model of continuous Neotethys subduction and Andean-type magmatism along an elongated boundary (Lin et al., 2013).
The eastern Lohit Plutonic Complex belt is distinct in having both Cretaceous (ca. 69–136 Ma) and Mesoproterozoic (ca. 1286 Ma) zircons (Table 1). The ca. 1286 Ma orthogneiss correlates with the Bomi-Chayu Complex (1276–1342 Ma), which is exposed surrounding the eastern Himalayan syntaxis (Xu et al., 2013) (Fig. 1C). The Bomi-Chayu Complex is notable for being the oldest rocks exposed in the Lhasa terrane (Xu et al., 2013), and our observations suggest that these rocks extend southeast to the northern Indo-Burma Ranges in the hanging wall of the Walong thrust (Fig. 1C).
Absence of Cenozoic plutonic and volcanic rocks (e.g., analogous to the Linzizong volcanics) of the southern Gangdese batholith belt in the study area raises the question of whether Cenozoic magmatism was occurring along the entire southern Lhasa margin during the Cenozoic. Existing research has shown that ca. 50–55 Ma mafic rocks of Gangdese affinity are exposed in the Mogok Metamorphic Belt of northwest Yunnan to the south of the study area (e.g., Wang et al., 2014). Furthermore, Eocene–Oligocene granitoids and Miocene–Quaternary volcanic flows related to Neotethys subduction are exposed throughout the Wuntho-Popa belt in present-day Myanmar (Mitchell, 1993; Mitchell et al., 2012; Lee et al., 2016). The existence of Cenozoic subduction-generated rocks both to the northwest and south of the study area suggests Cenozoic magmatism did occur in the northern Indo-Burma Ranges, but these igneous rocks may have been shortened and eroded and/or underthrusted during development of the orogen (Fig. 17).
Correlation of the Lalpani schist and Mayodia gneiss with the LHS (Fig. 2) suggests that prior to the Cenozoic India-Asia collision, rocks comprising the northeast Indian cratonal sequence extended to the east and southeast of the present-day location of the eastern Himalayan syntaxis. Rocks of the Indian passive continental margin (i.e., THS) and Indian cratonal sequence (i.e., GHC and LHS) would have existed along the northeast margin of the Greater Indian continent. The existence of Late Jurassic–Cretaceous granitoids within the study area that are equivalent to the northern Gangdese batholith belt implies that subduction of Neo-Tethys oceanic lithosphere and arc magmatism occurred along a continuous belt encompassing both the southern Lhasa terrane and Eastern Flanking Belt (e.g., Lin et al., 2013). Following initial India-Asia collision, Cenozoic magmatism (i.e., southern Gangdese batholith belt) also occurred along an elongated belt from the southern Lhasa terrane to the Eastern Flanking Belt. An accompanying Mesozoic–Cenozoic forearc basin was likely present along the entire Neo-Tethys subduction margin.
We interpret the absence of these major Himalayan-Tibetan units in the northern Indo-Burma Ranges to result from a greater magnitude of (1) Cenozoic shortening and erosion compared to the Himalayan orogen to the west and/or (2) continental underthrusting. Haproff et al. (2018) showed that the south- to west-directed northern Indo-Burma thrust belt accommodated clockwise rotation of Himalayan-Tibetan rocks around the eastern Himalayan syntaxis. Progressive clockwise rotation of the thrust belt from an originally east-trending configuration implies a spatial gradient in shortening strain, such that the magnitude of crustal shortening increases with distance from the pole of rotation at the eastern Himalayan syntaxis. This interpretation is supported by the dramatic southward decrease in the map-view distance between the active foreland basin and Tidding mélange complex (= IYSZ), which is ∼33 km along Dibang Valley and ∼5 km along Lohit Valley farther to the southeast. For comparison, the map-view distance between the Main Frontal thrust and IYSZ across the western Arunachal Himalaya, perpendicular to the strike of the orogen, is ∼200 km (Yin et al., 2010a). This width of the thrust belt decreases to ∼5 km over a map-view distance of ∼250 km to the south of the eastern Himalayan syntaxis (Fig. 1). Upper-plate shortening at the longitude of the northern Indo-Burma Ranges may also have been coeval with the underthrusting of complete sections of Greater India and southern Lhasa terrane. In this case, the Tethyan Himalayan thrust belt, GHC, Xigaze forearc basin, and southern Gangdese batholith belt were underthrusted together with the Indian continent below the southeastern Tibetan Plateau (Fig. 17).
Our study presents the results of U-Pb zircon geochronology and whole-rock geochemistry of lithologic units exposed in the northern Indo-Burma Ranges, located east to southeast of the eastern Himalayan syntaxis. Geochronological and geochemical data were coupled with lithologic and structural observations to test the lateral continuity of rocks of the easternmost Himalayan collisional system. We infer that lithologic units exposed in the study area are the easternmost continuations of the Cenozoic Sub-Himalayan Sequence, Lesser Himalayan Sequence, and Indus-Yarlung suture zone. The eastern Lohit Plutonic Complex contains Mesoproterozoic basement rocks of the Lhasa terrane (i.e., Bomi-Chayu Complex), whereas the western Lohit Plutonic Complex belt correlates with the Mesozoic northern Gangdese batholith belt. Based on comparable U-Pb zircon ages and geochemical composition between the Gangdese batholith and Lohit Plutonic Complex, subduction of Neotethys oceanic lithosphere, Mesozoic arc magmatism, and forearc basin sedimentation occurred along a continuous belt encompassing the southern Lhasa terrane and Eastern Flanking Belt along the southern margin of Asia. Prior to the early Cenozoic India-Asia collision, the northeast continental margin of Greater India was likely a continuous sequence involving rocks of the Lesser Himalayan Sequence, Greater Himalayan Crystalline Complex, and the Tethyan Himalayan Sequence. Once collision initiated, rocks of the Himalayan orogen and southern Lhasa terrane experienced coeval clockwise rotation and shortening around northeast India; this rotation was accommodated by a south- to west-directed thrusting. During progressive development of the northern Indo-Burma thrust belt, a greater magnitude of crustal shortening and/or continental underthrusting occurred as compared to the Himalayan orogen to the west, resulting in local complete erosion and/or underthrusting of several Himalayan-Tibetan lithologic units.
This work is supported by a grant from the Tectonics Program of U.S. National Science Foundation. We thank editor Raymond Russo, Peter Zeitler, and an anonymous reviewer for their constructive comments.