Multi-element isotope study of natrocarbonatites (1993 lava flows) from Oldoinyo Lengai volcano, Tanzania: Implications for core-mantle interactions
Introduction
The Earth's structure and composition, especially the inaccessible deep interior, is vital in better understanding of the origin and evolution of our planet. Important advances have been made recently using indirect methods to infer the nature of the Earth's deep interior, particularly the core-mantle boundary (CMB). The CMB may be a chemically very reactive region of the Earth's interior (Jeanloz, 1993; Garnero, 2000) involving interaction between the mantle and the outer core (Walker et al., 1995, 1999; Bird et al., 1999; Brandon et al., 1998, 2003; Puchtel and Humayun, 2000; Meibom and Frei, 2002; Humayun et al., 2004; Brandon and Walker, 2005) as well as the accumulation of, and the interaction with, subducted slabs (Hofmann and White, 1982; Knittle and Jeanloz, 1991; Williams and Garnero, 1996; Hofmann, 1997, 2003; Garnero, 2000; Davies, 2002). If some mantle upwellings are generated at the CMB (Weiss et al., 2016) associated with chemical interaction between the core and mantle, then magma plume products may contain a distinctive chemical or isotopic fingerprint that reflects contributions from the core. Unequivocal detection of a core constituent in plume-derived magmas could further resolve some of the major questions in geodynamics such as the depth of origin of plumes and the question of whole-mantle convection.
Previously, strong debate over core-mantle exchange has been made on the basis of Os isotope anomalies in Hawaiian picrites and Gorgonan komatiites (Walker et al., 1999; Brandon et al., 1998, 2003; Brandon and Walker, 2005 and references therein). The coupled 186Os–187Os enrichments for these rocks have been attributed to small contributions from the outer core. Such a model, however, requires relatively early crystallization of the inner core (Puchtel et al., 2005), which is in conflict with some models proposed for the timing of inner core formation (Buffett et al., 1996; Labrosse et al., 2001; Smirnov and Tarduno, 2004). It has also been suggested that Os isotope anomalies in Hawaiian picrites could result from the addition of oceanic ferromanganese sediments to the mantle source of these lavas (Ravizza et al., 2001; Baker and Jensen, 2004; Scherstén et al., 2004). However, Nielsen and co-workers (2006) using thallium isotopes in Hawaiian picrites and ferromanganese sediments cast some doubt on a sedimentary origin for the osmium isotope anomalies, and left core-mantle interaction as a possible explanation for the osmium isotope variation in Hawaiian picrites. Further concern about core-mantle exchange has come from the study of pyroxenites and base-metal sulfide grains of pyroxenites from Beni Bousera orogenic massif (Luguet et al., 2008). They concluded that the coupled 186Os–187Os enrichment observed in plume-related lavas can also have an upper-mantle origin related to source regions having experienced base-metal sulfide metasomatism by pyroxenite and/or peridotite-derived partial melts. It seems therefore, that because Os isotopes alone cannot serve as a solution to the debate on core-mantle interaction other isotopic tracers are needed to unambiguously identify the contribution of an outer core component.
The short-lived isotope system of 182Hf-182W (182Hf isotope decays to 182W with a half-life of 8.9 Ma; Vockenhuber et al., 2004) can be applied to verify the core contributions as revealed by Os isotopes. The parent nuclide of 182W (i.e., 182Hf) was extant during the first ~60 Ma of solar system history when the Earth was undergoing metal-silicate differentiation (Harper and Jacobson, 1996). Tungsten is a moderately siderophile, incompatible, and refractory element that preferentially partitions into Earth's core during core formation because of its affinity for metal. Hafnium (Hf), on the other hand, is a lithophile element so its abundance in the core is low. The accumulation of 182W in the silicate part of the Earth using a chondritic Earth model has led to the claim that the Earth had experienced a rapid accretion and an early differentiation after the birth of the solar system at 4.567 Ga (Connelly et al., 2008). The Earth appears to have undergone a protracted core formation based on the W isotope composition of the mantle, implying separation of the core either at ~30 Ma (single-stage model; Kleine et al., 2002; Yin et al., 2002) or ~11 Ma (exponential accretion model; Yin et al., 2002) after the beginning of the solar system. Early core formation of the Earth caused the radioactive 182Hf (parent isotope) to partition into the mantle, later decaying to 182W (daughter isotope) until it became extinct. The accumulation of 182W in the mantle consequently raised its 182W/183W and 182W/184W ratios (generally reported as ε182W or μ182W) relative to a standard, producing a deficit of ε182W = −2 units or μ182W = ~-220 ppm (Kleine et al., 2002) which have been estimated to exist between core and the bulk silicate Earth (BSE) based on the analyses of chondrites. Historically, tungsten isotopic compositions of terrestrial and extraterrestrial materials provided a robust tool to constrain the timing of metal-silicate segregation in planetary bodies and core formation in the Earth (Halliday et al., 2001; Schoenberg et al., 2002; Iizuka et al., 2010; Nebel et al., 2010). Further significant advancement in analytical techniques provided motivation to estimate the potential minor contributions (i.e., negative ε182W or μ182W) from the Earth's outer core to the overlying mantle occurring at any time during Earth's history (Mundl et al., 2017). For this purpose, several studies have been carried out on a variety of terrestrial materials including ocean island basalts (OIB; Lee et al., 1997; Takamasa et al. et al., 2009; Mundl et al., 2017; Rizo et al., 2019), mid-ocean ridge basalts (MORB; Rizo et al., 2016a; Mundl et al., 2017), other types of basalts (Iizuka et al., 2010, Willbold et al., 2011, Touboul et al., 2012; Rizo et al., 2016a; Dale et al., 2017), picrites (Scherstén et al., 2004), kimberlites (Scherstén et al., 2004), komatiites (Touboul et al., 2012), amphibolites (Willbold et al., 2015; Rizo et al., 2016b; Dale et al., 2017), ultramafic rocks (Touboul et al., 2014; Rizo et al., 2016b), gneisses (Iizuka et al., 2010; Willbold et al., 2011, 2015), and meta-sediments (Iizuka et al., 2010; Willbold et al., 2011). In spite of the uncertainly differences in the data, it is still not clear whether an anomaly exists (Willbold et al., 2011, 2015; Rizo et al., 2016a, 2019, 2016b; Touboul et al., 2012, 2014; Mundl et al., 2017) or not (Scherstén et al., 2004; Takamasa et al., 2009; Iizuka et al., 2010).
Quite a few studies have reported positive ε182W values mainly for basaltic rocks (Lee et al., 1997; Willbold et al., 2011; Rizo et al., 2016a; Dale et al., 2017), komatiites (Touboul et al., 2012), amphibolites (Willbold et al., 2011; Rizo et al., 2016b; Dale et al., 2017), ultramafic rocks (Touboul et al., 2014; Dale et al., 2017), gneisses (Willbold et al., 2015), and meta-sediments (Willbold et al., 2011). The positive values in the ε182W datasets compared to the bulk silicate Earth reference (i.e., ε182WBSE ≈ 0) hint towards the long-term preservation of tungsten isotopic compositions that were attained during early Earth-forming events (Rizo et al., 2016b; Touboul et al., 2012). It is also possible that such enrichments are result of addition of late veneer to early Earth (Willbold et al., 2015; Dale et al., 2017). Recently, coupled depletion of 142Nd and 182W recorded in 3.55 Ga old komatiites from South Africa is the first observation in terrestrial mantle materials suggesting an early mantle differentiation (Puchtel et al., 2016). More recent work shows negative ε182W anomalies in OIBs from Hawaii, Samoa, and Iceland, and these are regarded as a core contribution to the mantle plume responsible for the generation of OIBs (Mundl et al., 2017; Rizo et al., 2019).
Eruptions of natrocarbonatites from Oldoinyo Lengai (OL), Tanzania, (hereinafter referred to as OLLs) provides samples that can be used for tungsten isotopic studies. Here, we present a brief geochemical and isotopic description of the natrocarbonatites. Bell and Simonetti (see Table 1, 1996) have published the chemical compositions of the lavas used in this study including W values. One set of lavas contain silicate spheroids (i.e., SSB) and their chemical composition is different from the spheroid free material (i.e., SSF). Those without spheroids show high W, Ba, W/Pb, W/Th, W/U, W/La as well as low Pb, Th, U, and La values compared to those with spheroids (Table 1).
According to Bell and Keller (1995), natrocarbonatites are considered to be the products of a protracted periods of magmatic fractionation rarely encountered in nature. The isotopic similarities recognized between natrocarbonatites and peralkaline nephelinites from OL can be interpreted in several different ways, but the simplest explanation is one involving magma differentiation (Bell, 1998). Other explanations include generation from the same source by different melting events, or generation from different sources which have undergone similar differentiation histories.
The ranges in Nd-Pb-Sr isotope data of OLLs erupted in 1993 (i.e., 143Nd/144Nd = 0.51261–0.51268, 206Pb/204Pb = 19.24–19.26, 207Pb/204Pb = 15.60–15.63, 208Pb/204Pb = 39.30–39.38 and 87Sr/86Sr = 0.70437–0.70446 (Bell and Simonetti, 1996); are fairly restricted. Historically, the Nd-Pb-Sr isotope variations revealed by the east African carbonatites had been attributed to as mixing between plume and lithospheric sources (Bell and Simonetti, 1996). However, it has been later proposed, based on isotopic data from east African Rift zone, that such variations are not caused by lithospheric contributions but rather depict an isotopically inhomogeneous nature of the plume (Bell and Tilton, 2001). The143Nd/144Nd and 87Sr/86Sr ratios of OLLs not only cluster close to the chondritic uniform reservoir (CHUR)-bulk Earth intersection (see Fig. 2 in Bell and Simonetti, 1996) but also fall on a linear array defined by carbonatite samples from other east African locations (i.e., East African carbonatite line; EACL; Bell and Blenkinsop, 1987). This pattern is quite rare and defines a mixing line with varying proportions of two distinct mantle end-members (Bell and Dawson, 1995; Bell and Tilton, 2001) described as HIMU and EM1 (Hofmann, 1997). HIMU has been considered as a possible mantle source in the lower mantle on the basis of a combined Pb–Sr-Nd-Hf-Os isotope study (Hanyu et al., 2011). They proposed that OIBs from the Cook-Austral Islands represent a HIMU mantle reservoir formed by the hybridization of a subducted oceanic crust-derived melt with ambient mantle stored in the lower mantle for several billion years (Hanyu et al., 2011).
The stable isotope studies (O'Neil and Hay, 1973; Suwa et al., 1975; Hay, 1989; Keller and Hoefs, 1995; Dawson, 1993; Lee et al., 2000; Keller and Zaitsev, 2006; Halama et al., 2007) have demonstrated that the OLLs have similar isotopic compositions to other carbonatites (Deines, 1989). Volatile-bearing (CO2-enriched) mantle at high pressures (Brey and Green, 1977; Eggler, 1989; Wyllie, 1989) produces alkaline melts, such as nephelinites, alkali basalts and carbonatites depending on the degree of partial melting. There are different ways to generate natrocarbonatite melts including: (i) the evolution of a calcite-normative Ca-carbonatite melt (Guzmics et al., 2011; Weidendorfer et al., 2017) - formed by liquid immiscibility from a nephelinite or a melilite-nephelinite parent melt (Lee and Wyllie, 1998; Guzmics et al., 2012) – to a Na-carbonatite melt via continuous fractionation of calcite + apatite leading to the formation of nyerereite (i.e., Na–Ca carbonate) at the expense of calcite at ~620 °C (Weidendorfer et al., 2017), (ii) co-evolution of nephelinitic and carbonatitic melt enriched in fluorine, and alkali carbonate fluid controlled by liquid immiscibility at high temperatures (i.e., 1050 °C; Guzmics et al., 2012), which later resulted in the mixing of two carbonate phases due to CO2+H2O outgassing of alkali carbonate fluid below 850 °C and thereby producing natrocarbonatites at lower temperatures (i.e., subsolidus of nephelinite <630–650 °C; Guzmics et al., 2019), (iii) the effects of assimilation and decomposition of solid natrocarbonatites erupted in the past to produce unusual hybrid nepheline-andradite-melilite-combeite-phosphate magma (Mitchell, 2009; Mitchell and Dawson, 2012; Weidendorfer et al., 2019). In addition, their unique abundances of alkalies are attributed to enrichment of a primary olivine sövite during crystal fractionation (Gittins, 1989). However, sövite magma, characterized by having no alkalies, is neither produced as a result of liquid immiscibility – requiring substantial amount of alkalies in the immiscible carbonate melt – nor partial melting of the mantle because it would form Mg-carbonate (i.e., dolomite) rather than Ca-carbonate (i.e., calcite). Instead, sövite is considered as a fossilized carbonatite rock devoid of alkalies and it is formed by crystal accumulation (see Guzmics et al., 2011). Experimentally, it has also been observed that at low-P (~35 MPa) and undersaturated CO2 conditions, wollastonite nephelinite fractionate to peralkaline combeite-bearing nephelinite instead of the carbonate liquid exsolution (Kjarsgaard et al., 1995). Further, natrocarbonatite and nephelinite are closely associated but not conjugated (Kjarsgaard et al., 1995). Similarly, another experimental study concluded that the partition coefficients of alkalies increase significantly from anhydrous to hydrous conditions between the carbonatite and silicate melts immiscibility (Martin et al., 2013).
Section snippets
Geological background
OL is the only active carbonatite volcano that has erupted alkali-rich carbonatite lavas (i.e., natrocarbonatites). Situated in the East African Rift (EAR) valley in northern Tanzania and it remained relatively unknown until first recognized in 1960 (Dawson, 1962). Details about general geology and historical volcanic activity of this volcano are reported elsewhere (Dawson, 1989; Bell and Keller, 1995; Neukirchen et al., 2010). The lava flows are mainly composed of non-stiochiometric minerals
Sample description
A suite of OLLs are taken from a collection of C. Shrady who collected fresh lava flows from the 1993 eruption. Three out of 7 samples have contained 8–10 wt% of up to 2 mm size spheroids made up of mainly silicate material, a feature of OLLs first pointed out by Dawson et al. (1994). We divided the samples into SSB (i.e., silicate spheroid bearing) and SSF (i.e., silicate spheroid free) subgroups. OLLs are chemically different from all other known carbonatites (see Simonetti et al., 1997 and
Analytical techniques
The details of the tungsten purification method and isotope measurements using a multi-collector inductively-coupled mass spectrometer (MC-ICPMS) are described elsewhere (Sahoo et al., 2006). A brief description of the methodology is given here. Aliquots of digested samples were loaded onto the ion exchange chromatographic column to separate tungsten from major and trace elements by eluting different acidic solutions (see Table 2 in Sahoo et al., 2006). Solutions of the purified tungsten
C–O isotope data
The C and O isotopic data of OLLs are given in Table 2 and plotted in Fig. 1. The δ13C values range from −5.9 to −6.9‰ and 9.5‰–10.9‰ for δ18O. Our δ13C values are identical to the previous data; however, our δ18O data are heavier compared to those reported in other studies (Table 2).
W isotope data
The ε182W values are calculated using 182W/183W ratios measured for OLLs. Average data of each sample are shown in Table 3 and plotted in Fig. 4. The values for SSF and SSB subgroups range from −0.01 to −0.05
C–O isotopic compositions
Our data fall within the range of primary igneous carbonatites (PIC box; Fig. 1) established by Taylor et al. (1967) and Hoefs (1997). However, previously published OL data (Keller and Hoefs, 1995; O'Neil and Hay, 1973; Hay, 1989; Javoy et al., 1988; Dawson et al., 1995a) occupy a restricted area (OL box; Fig. 1) within the domain of primary igneous carbonatites (PIC box; Fig. 1). The discrepancy between our data and the previously reported values could possibly be the result of prolonged
Summary
The W isotopic compositions of fresh natrocarbonatite lava flows collected from the 1993 eruption at Oldoinyo Lengai, Tanzania have been measured in order to assess core-mantle interactions. To our knowledge, this is the first W isotope dataset obtained for OLLs. The ε182W values of these lava flows are in close proximity to the value assigned to the bulk silicate Earth (i.e., ε182WBSE ≈ 0). The absence of ε182W anomaly in these samples suggests that the ultimate source of the carbonated magma
Declaration of interest statement
This is to certify that the authors have shown a consent on the paper being submitted. It is an original work produced in Japanese laboratories at The University of Tokyo and Okayama University. This research work (title: Multi-element isotope study of natrocarbonatites (1993 lava flows) from Oldoinyo Lengai volcano, Tanzania: Implications for core-mantle interactions) has not been submitted for publication elsewhere and we have no conflict of interest to declare.
Acknowledgements
We are thankful to the editor Dr. Read Brown Mthanganyika Mapeo and two anonymous reviewers for their constructive comments that helped to significantly improve the quality of this manuscript. Authors acknowledge the help of Dr. Keith Bell for providing the samples and valuable suggestions. The support given by Dr. Minoru Kusakabe during the use of stable isotope facility at Institute for the Study of Earth's Interior (ISEI), Misasa, Japan is gratefully acknowledged. T. Nogi is thanked for
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