GR Focus Review
A geochemical and isotopic perspective on tectonic setting and depositional environment of Precambrian meta-carbonate rocks in collisional orogenic belts
M. Satish-Kumar
a,⇑, M. Shirakawa
b, A. Imura
a, N. Otsuji-Makino
b, R. Imanaka-Nohara
a,
S.P.K. Malaviarachchi
a,c, I.C.W. Fitzsimons
d, K. Sajeev
e, G.H. Grantham
f, B.F. Windley
g, T. Hokada
h, T. Takahashi
a, G. Shimoda
i, K.T. Goto
iaDepartment of Geology, Faculty of Science, Niigata University, 2-8050 Ikarashi, Nishi-ku, Niigata 950-2181, Japan
bGraduate School of Science and Technology, Niigata University, 2-8050 Ikarashi, Nishi-ku, Niigata 950-2181, Japan
cDepartment of Geology, Faculty of Science, University of Peradeniya, Peradeniya 20400, Sri Lanka
dSchool of Earth and Planetary Sciences, Curtin University, GPO Box U1987, Perth, WA 6845, Australia
eCentre for Earth Sciences, Indian Institute of Science, Bengaluru 560012, India
fDepartment of Geology, University of Johannesburg, P.O. Box 524, Auckland Park 2006, South Africa
gSchool of Geography, Geology and the Environment, University of Leicester, Leicester LE1 7RH, UK
hNational Institute of Polar Research, Tokyo 190-8518, Japan
iGeological Survey of Japan (AIST), Tsukuba, Japan
a r t i c l e i n f o
Article history:
Received 22 November 2020 Revised 22 March 2021 Accepted 24 March 2021 Available online 23 April 2021 Handling Editor: M. Santosh
Keywords:
Meta-carbonate rocks Trace and rare earth elements C-O isotopes
Sr-Nd isotopes
EastAfrican-Antarctic Orogen
a b s t r a c t
In this contribution we review the possibility of establishing the depositional age and tectonic settings of metamorphosed carbonate rocks from continental collision zones in the East African-Antarctic Orogen.
The geochemical characteristics of regionally distributed meta-carbonate rocks from the Highland Complex (HC) in Sri Lanka are considered in detail and compared with similar occurrences in East Antarctica, India, Madagascar and Africa. The variations seen in the Highland Complex of Sri Lanka imply that carbonate deposition was younging from west to east, spanning apparent ages from Mesoproterozoic to Neoproterozoic. In the case of East Antarctica, such variations are within the Neoproterozoic, whereas in southern India, Madagascar and Mozambique they have a broader age range possibly from the Paleoproterozoic to Mesoproterozoic. There is also clear evidence that some carbonates were deposited in an open ocean surrounding volcanic islands in the Mesoproterozoic. Shale-normalized REE pat- terns have typical signatures of open ocean deposition in a passive continental margin with variable con- tinental input in platforms nearby to island arcs. In comparison to Phanerozoic equivalents, the absence of a Ce anomaly is most significant, whereas other parameters such as (Pr/Yb)SN, (Pr/Tb)SN, and (Tb/
Yb)SNwere used to evaluate relative enrichments of the LREE, MREE and HREE fractions that are charac- teristic of ambient seawater. Pronounced La, and Y anomalies with minor Eu and Gd anomalies and cor- relations of REE parameters and anomalies with carbon and oxygen isotopes,87Sr/86Sr initial ratios and eNd values are evaluated for meta-carbonate rocks in the Proterozoic collision zone. TheeNd values and Sr initial ratios suggest that basins in the western Mozambique Ocean that separated the East Gondwana from West Gondwana received contributions from Archean continental crust and ambient seawater, whereas the eastern Mozambique Ocean had REE contributions from specific cratonic conti- nents in passive margins or from continental/volcanic island arcs in active margins.
Ó2021 The Author(s). Published by Elsevier B.V. on behalf of International Association for Gondwana Research. This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/
licenses/by-nc-nd/4.0/).
Contents
1. Introduction . . . 164 1.1. Geological and tectonic significance of carbonate rocks . . . 165
https://doi.org/10.1016/j.gr.2021.03.013
1751-7311/Ó2021 The Author(s). Published by Elsevier B.V. on behalf of International Association for Gondwana Research.
This is an open access article under the CC BY-NC-ND license (http://creativecommons.org/licenses/by-nc-nd/4.0/).
⇑Corresponding author.
E-mail address:satish@geo.sc.niigata-u.ac.jp(M. Satish-Kumar).
Contents lists available atScienceDirect
Gondwana Research
j o u r n a l h o m e p a g e : w w w . e l s e v i e r . c o m / l o c a t e / g r
1.2. Trace element geochemistry. . . 165
1.3. Sr concentration and isotopic composition . . . 166
1.4. Nd concentration and isotopic composition . . . 166
2. The East African-Antarctic Orogen in Gondwana: Remnants of the extinct Mozambique ocean . . . 167
2.1. A working example from Sri Lanka . . . 168
2.1.1. Field occurrence and sample description . . . 168
2.2. The Sør Rondane Mountains, East Antarctica . . . 168
2.3. Meta-carbonate rocks from other terranes in the East African-Antarctic Orogen . . . 170
3. Analytical methods . . . 171
3.1. Screening protocol. . . 171
3.2. Mineralogy and EPMA analyses . . . 171
3.3. Trace and rare earth elements . . . 171
3.4. Carbon and oxygen isotopes . . . 172
3.5. Strontium isotopes . . . 172
3.6. Neodymium isotopes . . . 173
4. Results and interpretations . . . 177
4.1. Mineralogy, textures and mineral chemistry. . . 177
4.2. Trace and rare earth element composition . . . 177
4.3. Carbon and oxygen isotopic composition . . . 181
4.4. Strontium isotopic composition . . . 181
4.5. Neodymium isotopic composition . . . 183
5. Depositional ages and geochemical characteristics of the HC meta-carbonate rocks and tectonic implications . . . 184
5.1. Apparent depositional ages of meta-carbonate rocks . . . 184
5.2. Sr-Nd isotopic signatures in determining the proximity of a continental source . . . 187
6. Geochemical characterization of meta-carbonate rocks in Precambrian collision zones. . . 188
6.1. Preservation of seawater REE + Y patterns in meta-carbonate rocks. . . 188
6.2. Geochemical proxies for paleo-oceanography of ancient oceans . . . 188
6.3. Preservation of the C-O isotopic composition of pristine seawater . . . 192
6.4. Robustness of strontium isotopes . . . 192
6.5. Nd isotope signatures of continental margin vs. Open ocean environments of deposition . . . 197
6.6. Tectonic implications of carbonate rocks in collision zones . . . 197
7. Conclusions. . . 198
Declaration of Competing Interest . . . 199
Acknowledgements . . . 199
Appendix A. Supplementary material . . . 199
References . . . 199
1. Introduction
Carbonate rocks are the remnants of chemically precipitated sediments that serve as proxies for understanding the chemical oceanography not only of present day oceans (e.g. Grotzinger, 1986; Webb and Kamber, 2000; Zhang et al., 2017; Li et al., 2019), but also of ancient oceans, as early as the Archean period (e.g., Kamber and Webb, 2001; Van Kranendonk et al., 2003;
Bolhar et al., 2004, 2015). Although vulnerable to post- depositional alterations, if a proper screening protocol is followed, carbonate rocks with fewer siliciclastic materials are good candi- dates for constraining their depositional age, sedimentary environ- ment and the tectonic setting of sedimentary basins (Bau et al., 1997, 1999; Webb and Kamber, 2000; Van Kranendonk et al., 2003; Nothdurft et al., 2004; Bolhar et al., 2004, 2015; Zhang et al., 2017). In particular, several trace element ratios, rare earth element + yttrium (REE + Y) patterns, carbon and oxygen isotopic compositions and the radiogenic isotope systems of Rb-Sr and Sm-Nd, if considered cohesively, will be extremely effective con- straints (e.g. Keto and Jacobsen, 1987; Veizer et al., 1999;
Melezhik et al., 2001; Frimmel, 2009; Otsuji et al., 2013, 2016). This is because of their distinctive geochemical behaviors, such as con- centration in seawater, spatial distribution in oceans, proximity to a continental source, residence time, and selective incorporation in carbonate sediments under specific marine conditions (e.g. Keto and Jacobsen, 1987, 1988; Webb and Kamber, 2000; Bolhar et al., 2004). However, the validity of using geochemical parameters in carbonate rocks is also highly debated due to alteration and reset-
ting during post-depositional processes such as diagenesis (e.g.
Fantle et al., 2020) and metamorphism (e.g.Melezhik et al., 2005, 2008; Satish-Kumar et al., 2008; Otsuji et al., 2013).
In this review, we aim to understand the geochemical behavior of trace and rare earth elements (REE) and multiple isotope sys- tems in pure meta-carbonate rocks from continental collision zones. As a working example, we selected the high-grade metased- imentary terrane of the Highland Complex in Sri Lanka, where there are multiple occurrences of pure meta-carbonate rocks. From the start, the existing geochemical screening protocol was tested for identifying those samples devoid of metamorphic and post- metamorphic alterations. Subsequently, we used key geochemical proxies that can deduce the sedimentary environment in which carbonate rocks are deposited. The application of C-O-Sr-Nd iso- tope geochemistry in combination with several trace element and rare earth element proxies for clarifying the interaction between paleo-ocean and continent/oceanic crust was further employed in a broader tectonic context of orogenic belts and the closure of paleo-oceans. In addition, the existing data on trace ele- ment and C-O-Sr-Nd isotopic compositions of meta-carbonate rocks from the Neoproterozoic orogens in East Antarctica, India, Madagascar and SE Africa were compiled to identify geochemical proxies for delineating their tectonic environment of deposition.
The review of a large number of carbonate geochemical data pre- sented here points to a distinct difference in ocean chemistry in the Archean, Proterozoic and Phanerozoic, perhaps reflecting the secular trends of ocean chemistry in accordance with ambient atmospheric evolution in Earth history.
1.1. Geological and tectonic significance of carbonate rocks
Carbonate sediments are distributed worldwide in shallow marine environments in low latitude areas. They are deposited in basins occurring in a wide variety of plate tectonic settings, such as active and passive continental margins, oceanic highs and islands in open oceans, and inland lacustrine systems (e.g.
Murray et al., 1990; James and Jones, 2016). They are typically composed either of biogenic or abiogenic material or a combina- tion of both, either from seawater or from freshwater. A marine system in which carbonates can deposit is either a neritic (shelf) or a pelagic (open) ocean. A variety of environmental factors, which vary from place to place, such as water temperature, salinity, trophic resources, nutrient and sediment input from continents may control the type of deposition. Therefore, it is important to clarify the depositional setting of carbonates in order to under- stand global environments in the past. Being scavengers of geo- chemical proxies of paleo-oceans, carbonate rocks have been extensively studied for understanding ancient plate tectonic envi- ronments and secular changes in seawater chemistry (e.g. Webb and Kamber, 2000; Frimmel, 2009). Many researchers have identi- fied variations in climatic conditions, ocean currents, sea level changes, and temperature from the analysis of Phanerozoic car- bonate rocks (e.g. Allan and Matthews, 1977; Grotzinger, 1986;
Dyer et al., 2019; Geyman et al., 2021;Grotzinger, 1989). On the other hand, in the Precambrian the depositional environments of carbonate deposition were supposed to be different. The composi- tion of atmosphere and oceans, especially the pCO2 and pO2 con- centrations are considered to be the critical difference. Although controlled by micorbialites, Archean and Proterozoic carbonate rocks also have distinct differences, most of the former ones lime- stones, whereas the latter are dolomites, that difference possibly caused by secular changes in ocean chemistry. Emergence of con- tinents and supercontinents had a conspicuous effect on deposi- tional settings, especially in platforms and ramps surrounding continents (Grotzinger and Knoll, 1999). Moreover, volume-wise also there was a drastic decrease in carbonate deposits from the Early Archean to the Neoproterozoic, which is attributed to changes in atmospheric pCO2 and pO2 concentrations that control the alkalinity and solubility of several metallic elements in seawa- ter (Grotzinger, 1989,1990; Kasting, 1987; Tang et al., 2021). The Mesoproterozoic is also characterized by less tectonic activities and the estimated mean elevation of continental crust was lower than late Archean, perhaps resulting in limited supply of nutrients to the ocean and less occurrence of carbonate platforms (Tang et al., 2021). Therefore, Precambrian carbonate rocks are good can- didates to gather important information about paleo-oceans and to discuss the secular changes in oceanic and atmospheric chemistry as well as the tectonic settings in which they were deposited in the Earth’s history (Grotzinger and Knoll, 1999).
Of particular interest are carbonate rocks preserved in orogenic belts that can provide key information about the ancient oceans that existed between continents (e.g.Melezhik et al., 2005, 2008;
Satish-Kumar et al., 2008; Otsuji et al., 2016). The cycle of orogen- esis starts with rifting and the opening of oceans where clastic and pelagic sediments are deposited during the life of the ocean until consumption and closure of the ocean by tectonic processes. The clastic sediments record collective (mixed) information about their provenance. Radiometric dating (e.g. U-Pb zircon) of key rocks such as volcaniclastic sediments help us to determine the age of sedi- mentation in weak- to un-metamorphosed sequences. Even in high-grade metamorphic rocks, many researchers have reported rock units with unimodal zircon age population as volcanic/depo- sitional ages, assuming that clastic rocks have a spread in detrital ages (e.g.,Goodge, 2020). This assumption may not be valid always, since a provenance area can also possess a unimodal age distribu-
tion. In contrast, the dating of detrital zircons in clastic rocks in col- lisional orogens can provide the upper limit (youngest detrital age) and the lower limit (earliest metamorphic event), which helps to constrain a window for the time of deposition (e.g. Goodge, 2020). However, accurate determination of the age of sedimenta- tion is commonly difficult in many orogenic belts. Continental col- lisions are, in general, preceded by closure of oceans, and therefore the final sutures that mark the sites of ocean closure may be marked by an abundance of clastic and carbonate sediments that were deposited in an earlier continental shelf or platform.
Several previous studies have attempted to understand the tec- tonic significance of meta-carbonate rocks in collision orogens (e.g.
Melezhik et al., 2005, 2008; Satish-Kumar et al., 2008; Otsuji et al., 2013, 2016). For example, Otsuji et al. (2013)were successful in determining the depositional age of 880–850 Ma of metamor- phosed carbonate rocks from the Sør Rondane Mountains in East Antarctica by defining the apparent depositional age. Furthermore, Otsuji et al. (2016)evaluated the tectonic setting of deposition and the nature of the source rocks based on the Nd isotopic composi- tion of meta-carbonate rocks. They suggested there were three types of settings of carbonate deposition; 1. against the East Antarctic craton, 2. along and adjacent to oceanic island arcs, and 3. in isolated oceanic islands. These studies further highlighted the possible presence of an oceanic island arc between the Mozam- bique and East Antarctic oceans. Accordingly, knowledge of car- bonate sediments and their age is vital in understanding the history and evolution of the environments of deposition in colli- sional orogens. In contrast, accretionary orogens contain thick car- bonate horizons, which are considered as remnants of reefs surrounding oceanic islands and several studies have pointed out their importance in understanding the secular changes in oceanic and climatic changes.
1.2. Trace element geochemistry
Trace element geochemistry is often used as a proxy for deter- mining the tectonic setting of deposition of carbonate rocks. For example, Y/Ho and Mn/Sr ratios are used to critically evaluate the influence of continental input on a sedimentary environment, and they also serve as tools for diagenetic alteration, in addition new isotope tracers (d44Ca,d26Mg,87Sr/86Sr) are being established to constrain the extent and rate of carbonate recrystallization (e.g.
see review byFantle et al., 2020). Sr depletion and Mn enrichment are directly related to an increasing degree of post-depositional alteration (Brand and Veizer, 1980; Derry et al., 1992). The Y/Ho ratios of modern seawater range between 60 and 90 (e.g.
Lawrence et al., 2006), whereas the upper continental crust has much lower values (c. 26,Kamber et al., 2005). Elevated Y relative to Ho in seawater reflects differences in complexation behavior of these elements that have the same charge and near-identical effec- tive ionic radii, and most rocks and minerals have ratios compara- ble to chondritic values (Nozaki et al., 1997). Holmium is scavenged by particles at roughly twice the rate of Y, leading to superchondritic Y/Ho ratios (~44–74) in seawater (Nozaki et al., 1997).
The rare earth elements (REE) and yttrium (Y) contents of car- bonate rocks have provided valuable information regarding the depositional environment and the composition of ambient seawa- ter from which they were precipitated (e.g.Bau and Dulski, 1996;
Webb and Kamber, 2000; Bolhar et al., 2004, Nothdurft et al., 2004;
Frimmel, 2009; Zhang et al., 2017; Li et al., 2019). The REE + Y pat- terns observed in modern carbonate rocks typically have a LREE depletion with characteristic positive anomalies of La, Eu, Gd, Y and a negative Ce anomaly (e,g.,Bau and Dulski, 1996; Webb and Kamber, 2000; Zhang et al., 2017). Y and Eu are common compo- nents in the continental crust, but with different redox-sensitive
relations. Ce and Eu anomalies are related to redox conditions and hydrothermal activities, respectively. Eu and Ce ions (Eu2+or Eu3+, and Ce3+or Ce4+) behave differently depending on the redox condi- tions during carbonate deposition. A positive La anomaly in car- bonates is attributed to its higher stability relative to LREEs (de Baar et al., 1985). Seawater scavenges REE + Y from riverine and aeolian sources in the continents, as well as during the interaction with seafloor (Elderfield and Greaves, 1982; Henderson, 1984;
Elderfield and Sholkovitz, 1987; Bayon et al., 2004). Biochemical reactions and partitioning effects with colloids and mineral phases control the REE + Y distribution of seawater (Henderson, 1984;
Sholkovitz et al., 1994). When normalized in PAAS (Post-Archean Australian Shale;McLennan, 1989), present-day carbonates yield seawater-like REE + Y distributions (e.g., corals: Scherer and Seitz, 1980; Sholkovitz and Shen, 1995; articulated brachiopods:
Zaky et al., 2016; foraminifera: Palmer, 1985; Osborne et al., 2017). Characteristic features of shallow oxygenated seawater are: a uniform LREE depletion, an enrichment in lanthanum, deple- tion in Ce, and a distinct positive Y anomaly (Zhang and Nozaki, 1996). Thus the REE + Y patterns of carbonates have REE + Y char- acteristics of modern seawater, and therefore carbonate rocks can be used to better understand the ambient seawater from which they were deposited. A recent review of the geochemistry of lime- stones deposited in various plate tectonic settings byZhang et al., (2017)suggested that REE ratios such as (La/Sm)SN, (Sm/Yb)SNand a Ce anomaly along with other immobile elemental ratios such as Zr/Ti and La/Sc can help in delineating the depositional environ- ment of carbonate rocks. Such geochemical proxies can be success- fully utilized to differentiate carbonate rocks deposited in active and passive continental margins, oceanic highs and islands in open oceans, and in inland basins.
However, buried carbonate rocks commonly lose their original signatures due to secondary uptake of REEs during diagenesis, especially when exposed to REEs from siliciclastic minerals or hydrothermal fluids (Haley et al., 2004; Abbott et al., 2015; Chen et al., 2015). Previous studies have shown that reefal microbialites (Webb and Kamber, 2000, 2011) and early cements (Nothdurft et al., 2004; Wallace et al., 2017) generally yield seawater-like REE distributions. These components have been used widely to evaluate paleo-seawater REE compositions in the Phanerozoic (e.g., Della Porta et al., 2015; Wallace et al., 2017; Hood et al., 2018). Furthermore, the oceanic environments of carbonate depo- sition in the Precambrian may not have been the same as those in the Phanerozoic. Especially, the redox conditions in the seawater were markedly different and therefore carbonate rocks deposited in the Archean and early Proterozoic do not display a Ce anomaly (e.g., German and Elderfield, 1999; Moffett, 1990; Bau et al., 1997, 1999; Bolhar et al., 2004; Nothdurft et al., 2004; Frimmel, 2009; Tostevin et al., 2016;Zhang et al., 2017;Li et al., 2019). Also, the variations in biological activities, composition of atmosphere, continental weathering patterns and extremely high volcanic and hydrothermal activity were important factors that have controlled seawater composition through time. Frimmel (2009) suggested that significant differences can be noted in trace element and rare earth element compositions between the proximal and distal envi- ronments from a continent to an ocean. Whereas a flat REE + Y pat- tern is common in proximal depositional sites, a depleted LREE pattern is found in distal basins, because of less input from a con- tinental source. Similar differences are also observed in meta- carbonate rocks in continental collisional zones. For example, Otsuji et al. (2013)reported various patterns in Meso- to Neopro- terozic meta-carbonate rocks from the Sør Rondane Mountains in East Antarctica, where they assigned a LREE enrichment to interac- tion with active volcanism/hydrothermal vents, and flat and LREE depleted patterns to open ocean and passive margin settings. Thus, the behavior of rare earth elements is an extremely useful indicator
in tracing carbonate depositional environments in a tectonic context.
1.3. Sr concentration and isotopic composition
One of the most common applications of carbonate geochem- istry is chemostratigraphy, which helps us to correlate sequences as well as to determine the apparent depositional ages of sedimen- tary rocks using a variety of geochemical proxies (Trønnes and Sundvoll, 1995; Veizer et al., 1999; Melezhik et al., 2001;
Halverson et al., 2010; Satish-Kumar, 2015). When compared with other elements, strontium in seawater has a long residence time (~2.4106yr;Jones and Jenkyns, 2001) relative to a short mixing time (~1000 yr;Kump, 1991; Jacobsen and Kaufman, 1999), and therefore the oceans are relatively homogeneous with respect to
87Sr/86Sr ratios (Halverson et al., 2007). Moreover, carbonate sedi- ments have high concentrations of Sr (a few hundred to thousands of ppm on average) and very low contents of Rb (ppb levels). Sr iso- tope ratios in carbonate sediments can thus directly reflect seawa- ter composition even in geologically older rocks and therefore it is well suited for chemostratigraphic applications.Halverson et al.
(2010) reported that the fluctuation of 87Sr/86Sr ratios from the Early Neoproterozoic to Early Cambrian displays a unidirectional, increasing trend, and they have compiled the variations in carbon isotopic compositions to demonstrate that secular trends markedly shift during global glaciation events. The strontium and carbon iso- tope ratios of carbonate rocks record the chemical compositions of coeval seawater and such secular variations can be utilized to esti- mate the apparent age of carbonate deposition. Furthermore, recently isotopic chemostratigraphy has been successfully applied both to indirect dating and to correlation of high-grade meta- carbonate sequences (e.g. Melezhik et al., 2005, 2008; Satish- Kumar et al., 2008; Otsuji et al., 2013).
1.4. Nd concentration and isotopic composition
Because of the extremely low concentration of neodymium in seawater compared with continental materials and because of its very short residence time (102-103year;Hooker et al., 1981), the Nd budget of an ocean is dominated by its continental contribu- tions. Modern seawater nearer to a continent has a non- radiogenic Nd isotope composition, i.e. negative
e
Nd values(Frank, 2002; Goldstein and O’Nions, 1981; Lacan and Jeandel, 2005; Piepgras and Wasserburg, 1980), whereas Archean seawater has a more radiogenic Nd isotope signature (i.e. positive
e
Nd(t) val-ues). Hydrothermal alteration of oceanic crust and a flux of mantle- like Nd to ambient seawater was much larger in the Archean than it is today (e.g.,Alexander et al., 2008;Bau et al., 1997; Derry and Jacobsen, 1990; Frei and Polat, 2007; Kamber and Webb, 2001;
Miller and O’nions, 1985; Viehmann et al., 2014, 2015a, 2015b).
In contrast, the growth of continents, its oxidative weathering and enhanced riverine input systematically changed ocean chem- istry through Proterozoic to Phanerozoic times. Therefore, carbon- ate sediments deposited in continental platforms can record the key local sources (Piepgras and Wasserburg, 1985; Goldstein and Jacobsen, 1987) and the Nd isotope composition of chemically pre- cipitated sediments can be used to constrain the sources of the marine REE inventory (e.g.,Bolhar et al., 2002; Alexander et al., 2008, 2009;Alibert and McCulloch, 1993; Viehmann et al., 2014, 2015a, 2015b). 143Nd/144Nd ratios from seawater differ in every basin depending on the proximity of rare earth elements (REE) from continental provenances.
By using Nd isotopic ratios from ancient carbonate rocks depos- ited in paleo-oceans, it may be possible to unravel the continental masses surrounding the basins and to make correlations between different sedimentary basins. Therefore, the Nd isotope geochem-
istry of carbonate rocks helps in understanding depositional envi- ronments and settings.
2. The East African-Antarctic Orogen in Gondwana: Remnants of the extinct Mozambique ocean
The Pangea supercontinent is considered to be one of the largest supercontinents in Earth history. It assembled most of its large continental masses during the amalgamation of Gondwana and Laurasia in the Paleozoic. It is widely recognized that the amalga- mation of Gondwana during the late Neoproterozoic to Cambrian is now reflected in the wide distribution, now fragmented, of oro- genic belts such as the East African Orogen (EAO, Stern, 1994), Kuunga Orogen (Meert, 2003), Zambesi Belt (Munyanyiwa and Hanson, 1988), Damara Belt (Lehmann et al., 2016), Dom Feliciano Belt (Philipp et al., 2016), Gariep Belt (Frimmel et al., 2002), Salda- nia Belt (Folling and Frimmel 2002), Ross Orogen (Goodge, 2020) and Delamarian Orogen and Amadeus Basin in Australia (Ward et al., 2019) (Fig. 1). Amalgamation of Gondwana in the study area depended on formation of the EAO and Kuunga Orogens in a con- tinent–continent collision setting. The EAO that resulted from the collision between East and West Gondwana, extending from the Arabian Peninsula in the north to northern Mozambique in the south (Stern 1994), where it is overprinted by the Kuunga Orogeny, which formed by the collision between North and South Gond- wana along the axis of the Damara-Zambesi-Kuunga Belt (Fig. 1).
The occurrence of Neoproterozoic carbonates contributes to the definition of the continental margins in the orogenic belts described above (Fig. 1). The basement rocks in this study are exposed in the Precambrian continental crust of East Africa, Mada-
gascar, southern India, Sri Lanka and East Antarctica, which form the roots of Neoproterozoic to Early Cambrian continental collision zones in a manner that parallels the present-day continental keel of the Tibetan plateau (Fitzsimons, 2016). Multiple collision events of smaller continental blocks might have resulted in the final amal- gamation of Gondwana and the timing of these events are still being extensively debated (e.g. Fitzsimons, 2000a, 2000b, 2003, Boger, 2011). Compared with the relatively simple amalgamation of the African and South American continents at~600 Ma in west- ern Gondwana (e.g. Trompette, 1997), the assembly of eastern Gondwana was more complicated, which is accentuated by the lack of information about the basement geology of Antarctica (e.g., Fitzsimons, 2000a, 2000b, 2003, 2016; Boger et al., 2001;
Boger and Miller, 2004; Meert, 2003; Jacobs et al., 2003, 2015;
Collins and Pisarevsky, 2005; Grantham et al., 2008, 2013, Satish- Kumar et al., 2008, 2013).
The ‘‘Mozambique Ocean” is inferred to have existed between pre-existing continents before the final amalgamation of Gond- wana. The sedimentary rocks in the EAAO are therefore considered to have been deposited in the Mozambique Ocean. The meta- carbonate rocks are ideal candidates for helping to construct the paleo-oceanography of the extinct Mozambique Ocean. The High- land Complex, in Sri Lanka, is a zone of ocean closure dominated by trench-fill sediments and thus was probably a segment of the Neoproterozoic Mozambique ocean that was connected to Antarc- tica (e.g.Kriegsman, 1995; He et al. 2016a; Malaviarachchi, 2018).
Two well-studied metasedimentary sequences of the Highland Complex, Sri Lanka and the Sør Rondane Mountains in East Antarc- tica were considered to be ideal geochemical proxies in the defini- tion of carbonate rocks applicable to continental collision zones.
Fig. 1.The Gondwana supercontinent and the distribution of Neoproterozoic to Cambrian orogenic belts and carbonate-bearing lithologies.
2.1. A working example from Sri Lanka
The Precambrian basement geology of Sri Lanka is subdivided by Nd isotope mapping and U-Pb zircon geochronology into three major terrains (Fig. 2): the Highland Complex (HC) in the centre, the Vijayan Complex (VC) in the east, and the Wanni Complex (WC) in the west (Milisenda et al., 1988, 1994;
Kröner et al., 2003; Cooray, 1994; Kehelpannala, 1997, 2003).
Prominent doubly-plunging synforms of the Kadugannawa Com- plex (KC) are located in the centre along the boundary between the WC and HC (Fig. 2). The WC, HC, and VC are considered to have independent geological histories and origins, prior to their amalgamation in the late Neoproterozoic Pan-African orogeny that led to the formation of the Gondwana supercontinent (Malaviarachchi, 2018;Malaviarachchi et al., 2019). All these ter- rains were strongly influenced by Pan-African orogenesis that took place between c. 550 and 650 Ma. The HC basement is mainly composed of meta-sedimentary rocks, whereas the WC and VC are dominated by meta-igneous rocks (Dharmapriya et al., 2015a; Osanai et al., 2016a).
The Highland Complex (HC) is a major high-grade meta- morphic belt that mainly consists of migmatitic pelitic gneisses and charnockitic gneisses that are intercalated with quartzites, meta-carbonate rocks, mafic granulites, and calc-silicate gneisses (Fig. 2). The most prominent metasedimentary rocks are layers of folded quartzites and meta-carbonates (Fig. 2; Cooray, 1994;
Madugalla, 2015). The metamorphic conditions are upper amphi- bolite to ultrahigh-temperature (UHT) granulite facies (Osanai et al., 2000; Sajeev et al., 2003; Santosh et al., 2014; Osanai et al., 2016a; Dharmapriya et al., 2015a). Rocks in the Highland Com- plex have Nd-model ages of 2000–2200 Ma (Milisenda et al., 1988; 1994), but detrital zircons in metasediments have SHRIMP U-Pb ages of 3200–2400 Ma (Kröner et al., 1987; Kröner and Williams, 1993). The HC rocks also have a pristine zircon growth age of 1940 Ma, and Pb loss events at ca.1100 Ma and 560–
550 Ma (Baur et al., 1991; Kröner and Williams, 1993; Hölzl et al., 1994; Osanai et al., 2016b; Kitano et al., 2018;
Malaviarachchi et al., 2019)
The Vijayan Complex (VC) (Fig. 2) consists predominantly of alkaline granitic gneisses, augen-gneisses and minor amphibolites possibly derived from mafic dykes. The contact between the VC and HC is a low-angle west-dipping thrust (Kehelpannala, 2003;
Krӧner et al., 2013; Malaviarachchi et al., 2021), which has trans- ported the HC eastwards over the VC. Kilppen of HC rocks overlying the VC are recognized at Kataragama (Silva et al., 1981). Ca.1000–
1100 Ma, upper amphibolite facies to granulite-facies conditions were reported from this complex (Kehelpannala, 2003; Krӧner et al., 2013; He et al., 2016b; Ng et al., 2017; Malaviarachchi et al., 2021). The VC rocks were also affected by extensive late metamorphic K-metasomatism, which transformed original tonalitic-gabbroic gneisses to K-rich granitic compositions (Kehelpannala, 2003; Krӧner et al., 2013). The Nd model ages of 1100–1800 Ma (Milisenda et al., 1988; 1994; Malaviarachchi et al., 2021) and zircon U–Pb ages of c. 510–456 Ma (Baur et al., 1991; Hölzl et al., 1994) are interpreted as the times of magmatism and metamorphism, respectively.
The Wanni Complex (WC) (Fig. 2), located to the west of the HC, is mainly composed of metamorphosed igneous rocks with minor sedimentary layers (Cooray, 1994). The Nd model ages of the WC rocks are 1100–1800 Ma (Milisenda et al., 1988; 1994;
Weerakoon et al., 2001), and the intrusive ages of orthogneisses range from 790 to 750 Ma to 1100–1000 Ma (Baur et al., 1991).
He et al. (2016b)reported a charnockite from the WC that has an emplacement age of 1000 Ma, followed by a thermal event at 570 Ma. The Kadugannawa Complex (KC) (Fig. 2), contains a com- posite unit of doubly-plunging upright folds (Almond, 1991), that
consist of hornblende- and biotite-bearing orthogneisses, gabbroic and granitic gneisses with minor metasedimentary packages (Kröner et al., 2003). The KC has Nd-model ages ranging between 1100 and 1800 Ma (Milisenda et al., 1988; 1994), and Pb–Pb zircon evaporation ages of ca.770 Ma to 1100 Ma, which were the result of multiple calc-alkaline magmatism (Kröner et al., 2003).
2.1.1. Field occurrence and sample description
Meta-carbonate rocks form conformable layers associated with gneisses of both sedimentary and igneous origin. In the Kandy area, located in the east of the KC, there are several thick folded meta- carbonate layers (Fig. 2), which are intercalated with migmatized pelitic, charnockitic and granitic gneisses. Most of the meta- carbonate layers are dolomite-rich containing only minor calc- silicate minerals, but some layers are totally free of silicates, but there are a few graphites and sulfides (Supplementary Table ST1). The layers vary in thickness from several to over a hun- dred meters, within which alternate layers of pure and impure meta-carbonates form visible beds parallel to the regional trend (Supplementary Figure SF1a). Although rare, some meta- carbonate layers have skarn zones against bordering silicate rocks, and they are included in this study in order to show how geochem- ical signatures can change under such conditions. We have studied and sampled meta-carbonate rocks throughout the HC. Altogether, we collected 104 meta-carbonate samples for detailed analyses from 48 localities from working quarries or road-cuts.
In the southeast of the HC near the boundary of the VC, (92601 inFig. 2) the meta-carbonate rocks enclose boudins and lenses of mafic granulites, the mutual contacts of which are sharp; near the contacts the calcites are coarser-grained and have variable color variations in white, light blue to light pink or yellow; indica- tive of recrystallization during intrusion or fluid-mediated metaso- matic processes (Supplementary Figure SF1b). However, in this locality 500 m away from the mafic granulite contact, the meta- carbonate rocks contain bands of pure and impure varieties (Sup- plementary Figure SF1c). Humite in association with spinel and forsterite occur in some layers, and some olivine-rich layers have been extensively serpentinized (Supplementary Figure SF1d);
these rocks were excluded from this study.
The grain size of carbonate minerals varies widely from less than a millimeter to a few centimeters. In impure silicate-bearing meta-carbonate layers, the proportions of calc-silicate minerals vary widely. The main minerals are phlogopite, olivine, diopside, spinel, apatite, humite and graphite. However, for this study we selected samples of almost pure meta-carbonate rocks, (pure white to greyish to yellowish in color) that contain more than 90% of car- bonate phases. Rarely, we sampled rocks with a higher proportion of calc-silicate minerals in order to understand the geochemical and isotopic differences between the pure and impure lithologies.
From thicker layers, multiple samples were collected perpen- dicular to the layering in order to verify possible geochemical vari- ations created during metamorphism and elemental and isotopic interactions. Salient field relations of meta-carbonate rocks from each region and mineralogical characteristics are listed inSupple- mentary Table ST1.
2.2. The Sør Rondane Mountains, East Antarctica
The Sør Rondane Mountains form part of the Meso- to Neopro- terozoic orogen in Eastern Dronning Maud Land, East Antarctica.
This region consists of highly deformed medium- to high-grade metamorphic rocks in association with various igneous rocks.
The basement rocks in the Sør Rondane Mountains comprise two terranes, the SW terrane and the NE terrane that are separated by the Main Tectonic Boundary (MTB;Osanai et al., 2013). The MTB is considered to be the collisional boundary between the
Fig. 2.Geological outline of Sri Lanka marking the sample localities of meta-carbonate rocks. The tectonic boundaries are after (Cooray, 1994) (Cooray, 1984). Meta-carbonate layers are marked in blue and sampling localities in red.
two micro-continental terranes during Gondwana amalgamation with the NE terrane being thrust over the SW terrane at ca. 650–
660 Ma, giving rise to markedly different clockwise and anticlock- wise metamorphicP–T–tpaths of the SW and NE terranes, respec- tively (Osanai et al., 2013and references therein).
The SW terrane, which is located in the south and southwest of the Sør Rondane Mountains is divisible into a meta-tonalitic unit and a meta-sedimentary unit. Kamei et al. (2013)proposed that two types of tonalite-tholeiitic- calc-alkaline rocks have distinct magmatic zircon ages. Tholeiitic tonalite formed at 998–995 Ma, its magma being generated by partial melting of low-K basalt at lower crustal depths. In contrast to the calc-alkaline tonalite rocks, which formed at 945–920 and 772 Ma being derived by partial melting of subducted oceanic crust. North of the main shear zone, the major rocks are greenschist to granulite facies meta- sedimentary rocks, accompanied by minor mafic rocks, meta- pelites, marbles and calc-silicate rocks. These rocks have igneous zircon and monazite detrital ages of 1190–960 Ma and metamor- phic ages of 690–520 Ma (Shiraishi et al., 2008; Osanai et al., 2013). Meta-carbonate rocks form thin conformable layers whose strontium isotopes suggest an apparent depositional age of ca.
880–850 Ma (late Tonian) (Otsuji et al., 2013).
The NE terrane is dominated by meta-sedimentary and meta- igneous rocks, which attained peak metamorphicP–T conditions of ca. 800 °C and 7–8kbar followed by amphibolite-facies retro- grade metamorphism and regional rehydration at ca. 530–580°C and 5.5 kbar. Previous studies reported a wide range of U–Pb ages between 3250 and 640 Ma for igneous zircons and 780–530 Ma for metamorphic zircons from the NE terrane, and 1010–1000 Ma igneous and 740–510 Ma metamorphic events from the northern domain of the SW terrane (Osanai et al., 2013). Meta-carbonate rocks from these two regions were likely deposited at 820–
790 Ma (early Cryogenian) and at ca. 1000–900 Ma, respectively (Otsuji et al., 2013).
2.3. Meta-carbonate rocks from other terranes in the East African- Antarctic Orogen
In addition to our detailed studies in the Highland Complex, Sri Lanka and the Sør Rondane Mountains in East Antarctica, we stud- ied representative meta-carbonate rocks from neighboring Gond- wana continents, such as the Mozambique belt in central Malawi and N Mozambique (Melezhik et al., 2006, Fritz et al., 2013), south- ern Madagascar (Collins, 2006), the Trivandrum and Madurai Blocks in southern India (Satish-Kumar et al., 2001a, 2001b), the Lützow Holm Complex in East Antarctica (Satish-Kumar et al., 2006) and at Sverdrupfjella in Western-Dronning Maud Land (Grantham et al., 1991, 1995; Groenewald et al., 1995; Elvevold and Ohta, 2010) in Eastern Antarctica (Fig. 1). With this material we aim to compare the depositional environments of meta- carbonate rocks throughout a wide extent of Gondwana.
The meta-carbonate samples from the Mozambique belt were collected from the Eastern granulite – Cabo Delgado Complex (Fritz et al., 2013) in NW Mozambique and central Malawi and sup- plement the data from Melezhik et al., (2006)from NE Mozam- bique from the Nampula Terrane and Namuno Terranes (Grantham et al., 2008). The marbles, which are poorly exposed in road cuts, form layers up to 100s of metres thick that are con- formable with the gneissic foliation. They are typically coarse- grained and pure with subordinate calc-silicate minerals (diopside, olivine, phlogopite).
The lithological framework of southern Madagascar resembles that in southern India and the Mozambique belt except that Mada- gascan rocks have a greater range of minerals and assemblages.
The meta-carbonate rocks are high-grade calc-silicate marbles that form layers that are up to 15 m wide, and extend along strike for
many kilometres. They are intercalated with mainly meta- sediments andpara-gneisses. The mineral assemblages in the mar- bles vary in different areas. At Vohidava they contain spinel, clino- humite, diopside, forsterite and phlogopite, whereas in the Tranomaro area they contain hibonite and corundum (Razakamanana et al., 2010) and dolomitic marbles have clinohu- mite, humite and chondrodite (Pradeepkumar and Krishnanath, 2000). In the Vohibory block in SW Madagascar dolomite-rich metacarbonate layers are up to 5 km thick and extend for at least 80 km, intercalated with pillow-bearing amphibolites, gabbros, peridotites, quartzites, jasper cherts, pelitic schists and orthog- neisses (Jöns and Schenk, 2008) interpreted by Wiindley et al.
(1994)as a shelf sequence interthrust with ophiolites and orthog- neisses. The Bekily area comprises diopside marbles (samples BW- 86.3, BW-86.4, BW-86.5) intercalated with layers of sillimanite quartzite, diopsidite, two-pyroxene amphibolite, sillimanite- cordierite-garnet paragneiss, and sapphirine-kornerupine rocks (Wiindley et al., 1994). In the Ampandrandava area the meta- carbonate layers are intercalated with layers of garnet-magnetite quartzite, phlogopite-spinel-scapolite diopsidite (up to 100 m wide), orthopyroxene-bearing charnockite, graphite-sillimanite schist, and cordierite-spinel-quartz-biotite-garnet-sillimanite mig- matitic paragneiss (Pierdzig, 1992). The diopsidites contain > 90%
diopside, but also, commonly in pegmatitic pockets, phlogopite (crystals 1–2 m long), calcite, scapolite, anhydrite, anorthite and fluorite. From carbon and oxygen isotope data Boulvais et al.
(1998)concluded that Tranomaro marbles were derived by decar- bonation of impure siliceous marbles and dolostones with fluids released by granitic intrusions. The marble layers are isoclinally folded with the other lithologies. An important development of southern Madagascar was the formation of many, strike-parallel, vertical shear zones (up to 25 km-wide) dated as 530–500 Ma (Martelat et al., 2000), which formed after the peak, 590–530 Ma, granulite facies metamorphism (Martelat et al., 2000), and which acted as channels for the transport of fluids enriched in CO2, H2O and Boron (Pili et al, 1997). Zircons in diopside-calcite veins that cross-cut marbles in the Vohidava area have a U-Pb age of 523 ± 5 Ma at 850-650C and 6–4 kbar (de Grave et al., 2002).
The meta-carbonate rocks in South India occur as thick layers in the Trivandrum and Madurai blocks of the Southern Granulite Ter- rain; we collected nine marble layers. The local geological setting and petrography of seven meta-carbonate layers were described with their calcite-graphite carbon isotope thermometry by Satish-Kumar et al. (2002). Two additional localities close to the Palghat Cauvery shear zone, where meta-carbonate rocks are mined for the cement industry (Madukkarai and Pudussery), are also included in this study. The meta-carbonate layers are gener- ally several kilometers in strike length and a few to hundreds of meters thick. They are mostly associated with pelitic gneisses and charnockites (Satish-Kumar et al., 2001a, 2001b, 2002).
Although pure meta-carbonate layers lack evidence of deforma- tion, the effects of intense deformation are expressed by boudi- nage, disruption and rotation of marly or pelitic bands within the marbles. In fresh calcite-rich meta-carbonate samples with few sil- icate grains, the calcite grain size varies from a few hundred microns to more than 5 mm. Many hand specimens from each locality were studied to identify their silicate mineral assemblages.
The meta-carbonate rocks in the Lützow Holm Complex occur as layers up to 100 m thick, which are intercalated with aluminous metapelites and commonly separated from them by decimeter- scale, coarse-grained skarn minerals, such as scapolite, clinopyrox- ene, phlogopite, amphibole and spinel. One marble layer (~100 m thick) we examined is closely associated with a ~ 200 m-thick meta-pelitic layer. The marble comprises medium to coarse- grained, dolomitic or calcite-rich layers with minor forsterite, phl- ogopite, spinel and apatite. Calc-silicate rocks that separate the
marble and metapelite are distinctly zoned with a phlogopite/
amphibole-rich zone near the metapelite and a scapolite/
diopside-rich zone in contact with the marble; these zones reflect extensive metasomatic activity along the lithological boundaries.
An earlier study bySatish-Kumar et al. (1998)of carbon and oxy- gen stable isotopes in a pure marble indicated an extremely large (up to 21‰) oxygen isotope heterogeneity at a sub-millimeter scale. This was interpreted by grain boundary migration of aqueous fluids with a meteoric component. A micrometer-scale stable iso- tope profile across the grain boundaries suggests that the outer rims of zoned calcites were isotopically modified by a dissolution-reprecipitation mechanism, whereas during cooling an isotopic exchange between the relict cores and newly precipi- tated outer rims produced a diffusion profile. Nevertheless, other calcite-rich marble layers in the region were less altered during metamorphism, and became potential candidates for Sr isotope chemostratigraphic studies, which suggested a possible age of deposition sometime within the period 730–830 Ma (Satish- Kumar et al., 2008).
In Western Dronning Maud Land, East Antarctica, meta- carbonate rocks were collected from the Fuglefjellet Gneiss Complex in eastern Sverdrupfjella (Grantham et al., 1991, 1995; Groenewald et al., 1995; Elvevold and Ohta, 2010). The meta-carbonates form layers up to ca. 100 m thick, interlayered with quartzo-feldspathic paragneisses, calc-silicates and rare meta-conglomerates. The thick- est continuous sequence is at the type locality at Fuglefjellet with further extensive exposures at Kvitkjolen and Skarsnuten. At Kvitk- jolen carbonate layers are duplicated by thrust-imbricates. The car- bonates are typically white to pale yellow, pink or grey. Grain sizes vary from fine (<1 mm) to coarse grained (ca. 1 cm); fine-grained varieties are relatively competent in contrast to coarse, crumbly coarse-grained marbles (Grantham 1992). Minor accessory silicates are tremolite, talc, phlogopite and serpentine. The meta-carbonates are limited to eastern Sverdrupfjella where they are inferred to be part of an allochthonous mega-nappe that overlies a footwall con- taining the ca. 1140 Ma authochthonous Jutulrora Complex that has tonalite-trondhjemite gneisses (Elvevold and Ohta, 2010;
Grantham et al., 2019; Bumby et al., 2020).
3. Analytical methods
Multiple samples were collected from each locality. All meta- carbonate rocks were investigated in detail for their mineralogy, their calcite-dolomite textures, and alterations. Thin sections were prepared for all samples and petrographic examinations were car- ried out with a polarizing microscope. In order to select samples that were geochemically least altered during metamorphism, an initial screening was carried out based on their mineralogy. The presence of minor amounts of silicate minerals in meta- carbonate rocks indicates that the original carbonate sediments contained siliciclastic components. Previous studies have demon- strated that even a low volume of siliciclastic components can cause a pronounced modification of the original carbonate REE + Y signal (e.g.Li et al., 2019). Considering this, pure meta- carbonate rocks that contained more than 90% carbonate minerals were selected and subjected to further rigorous screening using various geochemical tools, such as carbon and oxygen isotopes and trace elements. Sr and Nd isotopic compositions were analyzed only for those samples that were selected using the screening pro- tocol described below.
3.1. Screening protocol
Firstly, meta-carbonate samples were selected based on their field occurrence and mineralogy. Initial sample selection was
based on mineralogy, i.e. to visually select samples with > 90% car- bonate minerals. Previous studies have suggested that oxygen and carbon isotope values can be used as a parameter for screening and selecting samples that are least affected by post-depositional alter- ations (e.g.Otsuji et al., 2013). In the case of the meta-carbonate samples in the Highland Complex (HC), the d13C value of 1‰ and d18O value of 20‰ were fixed as cutoff thresholds (Otsuji et al., 2013), and samples falling below that were not considered for further analysis.
Mn/Sr ratios are also useful as a selection parameter, because of their sensitivity to the environment of carbonate deposition.
Denison et al., (1994)reported that concentrations of Mn, Sr and Fe vary systematically with the Sr isotope ratio in seawater. They argued that samples containing < 300 ppm of Mn yield the best Sr isotope ratios and those with > 700 ppm of Mn give altered Sr isotopic ratios. Thus, carbonate rocks with a Mn concentration of < 300 ppm along with Sr/Mn higher than 2 are the best to retain their original seawater ratio (Jacobsen and Kaufman, 1999, Halverson et al., 2010). The meta-carbonate rocks from the HC dis- play large variations in Sr isotopic compositions in relation to their Mn/Sr ratios (Supplementary Table ST7). In this study, following Kaufman and Knoll (1995), we considered that a conservative Mn/Sr ratio of 6 is an optimum cut-off value, beyond which any samples were considered to have been affected by post- sedimentary alterations. However, some samples gave high Sr iso- tope ratios irrespective of their lower Mn/Sr ratios. Therefore, in addition to the Mn/Sr ratios, we considered other geochemical parameters.
3.2. Mineralogy and EPMA analyses
The petrography of polished thin sections of all meta-carbonate samples used for the geochemical analyses were studied with a polarizing microscope. Carbonate minerals in selected samples were analyzed for major elements using a wavelength dispersive electron microprobe (JEOL JXA733; Shizuoka University). For the carbonates measurement conditions were a 12nA, 15 kV accelerat- ing voltage and a defocused beam (~10
l
m). Analytical results are presented inSupplementary Table ST2.3.3. Trace and rare earth elements
Whole-rock powder or calcite/dolomite mineral separates were analyzed for trace and rare earth element concentrations. Whole- rock powders were prepared from slabs that were initially crushed to small pieces with a tungsten mortar, and then to a fine powder in an agate mortar. For analysis of carbonate minerals, the required amount of powder was collected by drilling with a dentist micro- drill under a binocular microscope. Aliquots from the powders weighing between 20 and 50 mg was put into vials, which con- tained 1 ml distilled water and 1 ml 1 M CH3COOH. After drying, a mixture of acid (nitric acid (37%), hydrochloric acid (0.4%), and hydrofluoric acid (0.08%)) was used to dissolve the powder and dried repeatedly three times. CaF2 crystallization was not observed. The solution was diluted to 10,000 times when measur- ing REE, and 50,000 times when measuring for Mn. All analyses were carried out with an ICP-MS at Niigata University, Japan, the detailed procedure of which is reported inNeo et al. (2009). Stan- dards used were BHVO-2, W-2a, and JB-2. In order to reduce the scatter caused by low concentration of trace elements, we did not consider samples that were too close to detection limits. Fol- lowing the criteria given inLi et al. (2019)we set the cutoff value of 0.9 ppm for total REE contents. The analyzed trace and rare earth element compositions of all meta-carbonate rocks (208 samples) in this study are presented inSupplementary Tables ST3–ST6.
Measured raw REE + Y concentrations were normalized to the Post-Archean Australian Shale (PAAS) (McLennan, 1989) for easy comparison with the seawater composition. Yttrium (Y) is shown in the REE graphs between Dy and Ho on the basis of effective ionic radius (Bau and Dulski, 1996). The ratios of (Pr/Yb)SN, (Pr/Tb)SN, and (Tb/Yb)SNwere used to evaluate relative enrichments of the LREE, MREE (from Sm to Dy), and HREE fractions. La, Ce, Eu and Gd, anomalies represented by (La/La*)SN, (Ce/Ce*)SN, (Eu/Eu*)SN
and (Gd/Gd*)SNwere calculated geometrically following the equa- tions given below (afterLawrence et al., 2006). An Yttrium anom- aly (Y/Y*)SNwas calculated followingShields and Stille (2001).
La=La
ð ÞSN¼ LaSN=PrSNx Prð SN=NdSNÞ2
ð1Þ Ce=Ce
ð ÞSN¼ CeSN=ðPrSN ðPrSN=NdSNÞÞ ð2Þ Eu=Eu
ð ÞSN¼ EuSN= Sm2SNx TbSN
1=3
ð3Þ Gd=Gd
ð ÞSN¼ GdSN= Tb2SNx SmSN
1=3
ð4Þ Y=Y
ð ÞSN¼ 2 YSN=ðDySNþ HoSNÞ ð5Þ Table 1present the range of values for trace element ratios and REE parameters for all studied terranes andTable 2shows the sali- ent geochemical parameters of the screened samples from the Highland Complex in Sri Lanka.
3.4. Carbon and oxygen isotopes
All meta-carbonate samples were analyzed for carbon and oxy- gen isotopic compositions. Polished slabs were stained with Ali- zarin red-S to distinguish between calcite and dolomite.
Comparison of the same samples with and without staining with Alizarin red-S has shown that staining does not affect C and O iso- tope ratios (Wada et al., 1984). Sampling was made with a knife edge or micro-drill under a microscope. Powdered samples for C–
O isotope analyses were taken from different domains in each slab to identify whether there is heterogeneity or not. Dolomite or cal- cite powders were placed in small steel thimbles and dropped into a pot containing phosphoric acid at 60°C for calcite and 100°C for dolomite in vacuum to produce CO2gas. The CO2gas produced was
cleaned to remove impurities like H2O with a pentane slush and collected by using a liquid nitrogen cold trap (Wada et al., 1984).
Stable isotope measurements of carbon and oxygen were carried out with a Finnigan MAT-251 mass spectrometer at Niigata Univer- sity, Japan.d13C values relative to V-PDB andd18O values are com- pared with V-SMOW and reported in per mil (‰). A total of 189 carbon and oxygen isotope analyses of dolomite or calcite were carried out on 124 samples and presented in Supplementary Table ST7with selected data inTable 2. The positions of analytical spots on stained slabs along with their isotope data of representa- tive samples are shown inFig. 3and information about all studied samples are shown inSupplementary Figure SF2.
3.5. Strontium isotopes
For the strontium isotope analysis samples were selected with a basic criterion of an oxygen isotopic composition of > 20‰. New data for 54 meta-carbonate samples from Sri Lanka and 40 samples from other terranes of East Gondwana are presented here. A few samples with lower oxygen isotope ratios were included in order to confirm the effect of metamorphic fluid-rock interaction. Car- bonate samples were firstly decomposed in a mixed solution of distilled water and 1 M CH3COOH in a TeflonÒvessel. Only carbon- ate minerals will dissolve in acetic acid, which restricts the effect on minor silicate minerals, (especially detrital) on strontium iso- topes. After dissolution the vial was centrifuged and the super- natant was separated. Distilled water was added once again to the residue and a supernatant liquid was collected, which was evaporated to dryness and 6 M HCl was added and evaporated to dryness again. 1 ml of 2.5 M HCl was added to the dried sample and Ca–Rb–Sr and REE were separated by 1st column separation (BioRad, AG50W-X8 cation-exchange resin). Purification of Sr was made with a 2nd column separation using 2 M HNO3(Eichrom resin for Sr). The Sr isotope analyses were carried out following the procedures of Miyazaki and Shuto (1998) and Takahashi et al., (2009)using a Thermal Ionization Mass Spectrometer on a Finni- gan MAT 262 Mass Spectrometer at Niigata University, Japan. Usu- ally a Ta (sample) and Re (ionization) filament combination was used, however in some cases a Re-single filament method with a tantalum activator (Birck, 1986) was adopted to obtain higher ion- ization. This method is effective for samples that contain very low concentrations of Sr. The NIST 987 standard was analyzed before
Table 1
Range of values for various geochemical parameters of metacarbonate rocks from the East African-Antarctic Orogen.
Western Dronning Maud Land, East Antarctica
Sor Rondane Mountains, East Antarctica
Lutzow Holm Complex, East Antarctica
Highland Complex, Sri Lanka Southern Granulite Terrane, India
Madagascar Mozambique (All data) Localities Belt
close to Wanni Complex
Localities in the central HC
Localities close to Vijayan Complex Geochemical
parameters
n= 11 n= 47 n= 4 n= 108 n= 17 n= 5 n= 8 n= 21 n= 8 n= 3
(Pr/Yb)SN 0.36–2.77 0.05–8.78 0.58–3.0 0.23–9.78 0.23–4.94 0.36–0.60 0.30–4.09 0.27–0.93 0.12–1.71 0.45–1.13 (Pr/Tb)SN 0.53–2.28 0.07–3.32 0.49–2.26 0.23–3.36 0.23–3.36 0.31–0.55 0.42–2.23 0.32–0.88 0.24–1.43 0.43–1.14 (Tb/Yb)SN 0.69–1.55 0.34–2.64 0.86–1.33 0.71–4.33 0.74–1.86 0.88–1.18 0.71–1.83 0.59–1.31 0.50–1.42 0.99–1.14 (Nd/Yb)SN 0.42–2.29 0.07–7.11 0.65–2.57 0.01–9.25 0.27–4.14 0.35–0.63 0.01–3.40 0.30–1.06 0.13–1.55 0.49–1.04 RREE 3.33–85.03 1.48–72.54 3.45–28.95 0.56–88.62 0.56–33.73 0.64–6.50 0.63–48.39 2.07–54.34 3.74–102.5 1.3–5.1 Y/Ho 23.9–38.0 25.0–57.9 27.2–32.1 29.4–66.7 40.1–65.3 47.0–49.5 37.5–59 25.9–49.9 38.8–63.4 30–5-37.0 Mn/Sr 0.02–1.04 0.02–9.03 1.23–6.62 0.07–56.94 0.11–4.94 0.69–4.27 0.31–3.28 0.003–1.90 0.29–7.73 0.24–1.30 Ce/Ce* 0.86–1.10 0.64–1.30 0.92–1.20 0.58–1.46 0.80–1.4 0.67–1.05 0.83–1.24 0.47–1.23 0.76–1.02 0.90–1.12 Eu/Eu* 1.09–1.93 0.58–2.37 0.97–1.49 0.49–3.95 0.71–2.20 1.22–1.45 0.38–1.56 1.01–1.97 0.89–2.20 0.94–1.42 La/La* 0.88–1.30 0.70–1.52 0.98–1.20 0.66–2.06 0.83–2.13 0.63–1.25 0.81–1.79 0.88–1.75 0.73–1.20 0.92–1.24 Gd/Gd* 1.07–1.25 0.91–1.95 1.13–1.22 0.82–1.63 0.83–1.53 0.87–1.28 1.13–1.38 1.12–1.22 1.09–1.30 0.99–1.22 Lu/Lu* 0.92–1.54 0.57–1.47 0.83–1.25 0.64–5.67 0.96–2.66 0.92–1.87 0.92–1.41 0.95–1.25 1.03–1.86 0.93–2.34 (Y/Y*)SN 0.81–1.35 0.87–2.19 0.97–1.19 1.07–2.73 1.58–2.67 1.81–2.08 1.54–2.54 0.93–1.84 1.36–2.61 1.05–1.52
