Abstract
The Paleoproterozoic Usagaran Orogenic Belt in East Africa is an exemplary site of preserved relics of Paleoproterozoic volcanic-arc magmas, and subducted and displaced oceanic crust. This paper describes the geochemical characteristics of displaced pillow basalts from the Usagaran Belt (Konse Group). Our data indicate that the Konse pillow basalts have tholeiitic composition and are overprinted by greenschist-facies metamorphic conditions but their primary geochemical signatures are preserved by REE and fluid-immobile elements. The (La/Sm)N ratios (0.62–1.09) and REE patterns point to normal mid-ocean ridge basalt (N-MORB) and transitional MORB (T-MORB) mantle source. The analysis of high-valency elements and trace element patterns points to a mixed signature of MORB tholeiites and island-arc tholeiites with elevation of Ba, Th, U, Eu and Sr. This composition is similar to that of Phanerozoic back-arc suprasubduction-zone ophiolites. Therefore, Usagaran Belt pillow basalts are an analogue of the Tethyan-type suprasubduction-zone ophiolite evolution and emplacement in the Precambrian. The emplacement of the Konse pillow basalts is older than the neighbouring 1920–1870 Ma volcanic-arc magmas and is probably coeval with the formation of the 2000 Ma Yalumba eclogites with MORB affinity. Thus, the Konse pillow basalts might have been displaced from their suprasubduction-zone tectonic setting of igneous construction within the Yalumba ocean basin around 2000 Ma.
Thematic collection: This article is part of the ‘Tethyan ophiolites and Tethyan seaways collection’ available at: https://www.lyellcollection.org/cc/tethyan-ophiolites-and-tethyan-seaways
The Usagaran Belt is a Paleoproterozoic orogen located at the southeastern margins of the Archean Tanzania Craton (Fig. 1; Daly 1988). The Usagaran Belt has thrust fabrics that were formed in the Ediacaran Period (650–550 Ma) when the nappes of the Neoproterozoic Mozambique Belt (eastern granulites) were transported westwards and reworked the Usagaran Belt and the Tanzania Craton to form the western granulites (Fritz et al. 2005, 2013). The Usagaran Belt hosts ancient subduction proxies (eclogites, arc-related lavas and granitoids, and ophiolite), which have emplacement and metamorphic ages dated between 2000 and 1820 Ma (Gabert & Wendt 1974; Priem et al. 1979; Möller et al. 1995; Mruma 1995; Reddy et al. 2003; Collins et al. 2004; Sommer et al. 2005; Bahame et al. 2016). Current geochronological data place the Usagaran Belt in the group of orogenic belts that evolved in the Orosirian Period. Global Archean and Paleoproterozoic orogens feature rare occurrences of preserved relics of oceanic crusts like that preserved in the Usagaran Belt (e.g. Maruyama et al. 1996; Stern et al. 2013). Therefore, the Usagaran Belt is among the few global Orosirian orogens (2.1–1.8 Ga orogenic belts; see Zhao et al. 2002), which can give information on ancient tectonic processes.
The location of the Paleoproterozoic Ubendian–Usagaran Belt and the eclogite localities with ages (Möller et al. 1995; Ring et al. 2002; Boniface & Schenk 2012; Boniface et al. 2012). Model of Neoproterozoic westward thrusting of the Mozambique Belt onto the Usagaran Belt and Tanzania Craton modified after Fritz et al. (2013).
Dilek & Furnes (2011, 2014) have provided a comprehensive review of Phanerozoic orogenic systems and ophiolite genesis. They defined an ophiolite as an allochthonous fragment of upper mantle and oceanic crustal rocks that is tectonically displaced from its primary igneous origin as a result of plate convergence. The primary igneous origins of tectonically displaced oceanic crust and upper mantle sections are broadly grouped as subduction-related (suprasubduction-zone and volcanic-arc types) and subduction-unrelated (continental margin, mid-ocean ridge and plume-type) (Dilek & Furnes 2011, 2014). Ophiolites are either derived from upper plates of subduction zones (i.e. Oman and Troodos massifs) and emplaced on passive continental margins or derived from downgoing plates and emplaced in accretionary complexes (i.e. Japanese ophiolites, ophiolites in the Franciscan Complex in California; Dilek & Furnes 2014) that mark major boundaries between amalgamated plates or accreted terranes (Dilek 2003a). Ophiolites are generally found along suture zones in both collisional-type (e.g. Alpine, Himalayan, Caledonides, Appalachian) and accretionary-type (e.g. West Pacific, North American Cordilleran) orogenic systems (Gray et al. 2000; Dilek 2003a; Wakabayashi & Dilek 2003; Lister & Forster 2009).
The pillow lava basalts of the Usagaran Belt are considered to be concomitant with 1920–1870 Ma subduction-related volcanic-arc magmatism (Whittingham 1959; Mruma 1989; Sommer et al. 2005; Bahame et al. 2016). However, it is also possible that the Konse pillow basalts belong to the same time window as the neighbouring Yalumba Hill eclogites, which represent relics of an oceanic crust that subducted at c. 2000 Ma (Fig. 1; see Möller et al. 1995; Collins et al. 2004). It should be noted that Mruma (1989, 1995) described the stratigraphy and depositional environment of the volcano-sedimentary sequences (probably an ophiolitic section) from the Usagaran Belt (Konse Group) and identified pillow lava basalts in the volcano-sedimentary sequence units (Fig. 1). However, to date there are no detailed geochemical characterizations or petrographic descriptions of the pillow lava basalts from the Konse Group of the Usagaran Belt.
In this paper, we present new geochemical data from the Usagaran Belt and compare the geochemical fingerprint of the rocks with different types of known and well-studied ophiolite tectonic settings in suprasubduction-zone and mid-ocean ridge spreading centres (e.g. Dilek et al. 1990, 2008; Wakabayashi & Dilek 2003; Metcalf & Shervais 2008; Dilek & Furnes 2011; Furnes et al. 2012). It is interesting to find out that the geochemical trends of the Usagaran Orosirian mafic pillow lavas are highly reminiscent of the Jurassic–Cretaceous suprasubduction-zone ophiolite lavas of the Tethyan ocean systems that formed during and after the Pangaea break-up (see Stampfli 2000; Bortolotti & Principi 2005; Dilek & Furnes 2009, 2011).
Geological background
The Paleoproterozoic Ubendian and Usagaran belts bind the southern margins of the Tanzania Craton. The Ubendian Belt preserves subduction-related eclogites dated between 1920 and 1860 Ma. The eclogites with a mid-ocean ridge basalt (MORB) affinity define a subduction zone that existed between 1890 and 1860 Ma in the Ubendian Belt (Boniface et al. 2012). However, relatively older kyanite–epidote eclogites dated at 1924 ± 74 Ma have been reported by Boniface & Tsujimori (2014, 2016). The subduction events dated between 1920 and 1860 Ma in the Ubendian Belt overlap with voluminous emplacement of granitoids and volcanic rocks at the margins of the Tanzania Craton, along the Ubendian Belt (e.g. basaltic andesites–trachyandesites at Ngualla). The granitoids and volcanic rocks yielded ages between 1940 and 1870 Ma (Fig. 1; Lawley et al. 2013, 2014; Tulibonywa et al. 2015; Thomas et al. 2016).
The Usagaran Belt has been regarded as a foreland of the Neoproterozoic Mozambique Belt that thrust westward onto the Tanzania Craton (see Fritz et al. 2005; Sommer et al. 2005, and references therein). The internal structural fabrics (foliation, shear zones and gneissic bands) are oriented NE–SW to east–west in its southern part (Mruma 1989; Reddy et al. 2003; Fritz et al. 2005). The Usagaran Belt is divided into three main lithologically distinct tectonic units: the Isimani Suite, Konse Group and Ndembera Group (Meinhold 1970; Mruma 1989, 1995; Fritz et al. 2005; Sommer et al. 2005; Legler et al. 2015), and the geology, geochronology and geochemical data of these units are described below in turn.
The Isimani Suite
The Isimani Suite comprises highly deformed and high-grade metamorphic rocks broadly grouped as mafic amphibolites and leucocratic orthogneisses, which form the two main units of the Isimani Suite (Meinhold 1970; Mruma 1989). The mafic amphibolites are garnet-rich and compositionally banded, including layers of high-pressure mafic granulites, eclogites and kyanite metapelites. The leucocratic orthogneiss is the most abundant association and occurs as a coarse-grained rock with in situ melting layers and inclusions of amphibolite boudins interpreted as deformed mafic dykes (Mruma 1989; Reddy et al. 2003).
Mineral assemblages from the Isimani Suite rocks point to high-grade metamorphic P–T conditions at amphibolite, granulite and eclogite facies, which are overprinted by low-grade metamorphic P–T conditions at greenschist facies localized on shear corridors and thrust zones (Mruma 1989; Reddy et al. 2003; Collins et al. 2004; Mori et al. 2018). The Isimani Suite eclogites at Yalumba Hill have a strong affinity to MORB (see Fig. 2; Möller et al. 1995; Collins et al. 2004). The eclogite-facies metamorphism reached peak metamorphic conditions at P = 18 kbar and T = 750°C at c. 2000 Ma following the subduction of an oceanic lithosphere (see Fig. 2; Möller et al. 1995; Collins et al. 2004; Fritz et al. 2005; Mori et al. 2018).
Geological map of the Usagaran Belt demonstrating location of the main tectonic units: Isimani Suite, Konse Group and Ndembera Group. The map is modified after the compilation of 56 available quarter degree sheets (QDS) by Fritz et al. (2005).
Mruma (1989) and Reddy et al. (2003) assessed the deformational fabrics in the Isimani Suite and identified five stages of deformation (D1–D5). The eclogite–granulite facies is the first deformation (D1 is compositional banding), which took place at c. 2000 ± 1 Ma. Phase one of deformation (D1) was described by Tenczer et al. (2007) as a migmatic, NE–SW-trending foliation in the basement and the predominant structural feature indicating a high-temperature tectonometamorphic event in the Usagaran Belt.
Pervasive sinistral transpressional deformation D2–D3 (D2 is mylonitic foliation and mineral elongation, and indistinguishable D3 is inclined to overturned folds) at amphibolite-facies conditions strongly overprinted the D1 event between 1991 ± 2 and 1877 ± 7 Ma related to voluminous volcano-plutonic magmatism in the Ndembera Group (Reddy et al. 2003), an event that was preceded by Usagaran basement exhumation with dextral strike-slip tectonics along west–east-trending shear planes before the deposition of the Ndembera volcanic rocks (Tenczer et al. 2007). The Paleoproterozoic structural style in the Usagaran Belt is dominated by strike-slip tectonics, with syntectonic granitoids commonly intruding along the ductile strike-slip shear zones (Fritz et al. 2005; Tenczer et al. 2007).
Finally, the greenschist-facies deformation (D4 is extensional and D5 is thrust deformation) overprinted the previously formed deformation structures and textures at 535 ± 2.5 Ma, indicating a complex thermal–tectonic reworking of the Usagaran Belt in the Neoproterozoic–Cambrian (Mruma 1989; Reddy et al. 2003, 2004; Collins et al. 2004). The Neoproterozoic reworking of the Usagaran Belt is characterized by localized thrusts and bundles of discrete west–east-trending brittle to ductile dextral strike-slip faults associated with open NE–SW-trending folding, which developed parallel to the southern margin of the Tanzania Craton (Fritz et al. 2005; Tenczer et al. 2007).
The Ndembera Group
The Ndembera Group is composed of metavolcanic rock assemblages (basal agglomerates, andesitic tuffs and lavas, and dacitic to rhyolitic lavas and tuffs) capped with metasedimentary rocks (quartzites, quartzo-feldspathic schists and phyllites). The volcano-sedimentary sequence rests unconformably over the high-grade gneisses of the Usagaran Belt (Kürsten 1963; Mruma 1989). The Ndembera metavolcanic rocks yield crystallization ages between 1921 ± 14 and 1871 ± 15 Ma (Gabert & Wendt 1974; Sommer et al. 2005; Bahame et al. 2016), which is concomitant with the amphibolite-facies metamorphism (D2–D3) in the Isimani Suite. Geochemical and isotopic data indicate that the Ndembera metavolcanic rocks have strong affinities to calc-alkaline magmas of the volcanic-arc settings (Bahame et al. 2016).
Granitoids
The Usagaran basement is intruded by two chronological groups of granitoids: the foliated syntectonic granitoids and the weakly foliated late Usagaran granitoids (Fig. 2; Fritz et al. 2005). The foliated syntectonic granitoids intruded along ductile strike-slip shear zones between 1942 ± 89 and 1877 ± 7 Ma (Wendt et al. 1972; Reddy et al. 2003; Fritz et al. 2005; Bahame et al. 2016). The foliated syntectonic granitoids are largely metaluminous, calc-alkaline I-type granite, and have an island-arc magmatic assemblage (granodiorites, tonalites and monzogranodiorites) (Fritz et al. 2005; Bahame et al. 2016). The ages of the foliated syntectonic granitoids overlap with the ages of the Ndembera metavolcanic rocks, which were dated between 1921 ± 14 and 1871 ± 15 Ma (Gabert & Wendt 1974; Sommer et al. 2005; Bahame et al. 2016).
The weakly foliated late Usagaran granitoids (known as the Ubena granitoids to the south) are largely Hbl tonalites and Bt high-K granites connecting the Ubendian and Usagaran belts (Fig. 1; Manya & Maboko 2016). The weakly foliated late Usagaran granitoids have been dated between 1887 ± 11 and 1820 ± 9 Ma and they intrude the Ndembera metavolcanic rocks (Sommer et al. 2005; Manya & Maboko 2016).
The Ndembera metavolcanic rocks and the foliated syntectonic granitoids share similar geochemical features including very coherent REE patterns characterized by enrichment of the light REE (LREE) relative to the heavy REE (HREE) and depletion of high field strength elements (HFSE; Nb–Ta and Ti), which is a typical signature of arc magmas (Bahame et al. 2016). Therefore, a geotectonic evolution model of the Usagaran Belt proposes a continental collision with eclogite occurrences in the north and arc accretion with island-arc magmatism in the south occurring between 2000 and 1870 Ma; finally, the late Usagaran granitoids intruded the Usagaran Belt between 1880 and 1820 Ma (Fritz et al. 2005).
The Konse Group
The Konse Group comprises a volcano-sedimentary rock assemblage defining a narrow strip trending NE–SW sandwiched between the Tanzania Craton and the high-grade gneisses of the Isimani Suite (Figs 2 and 3). Quartzites, metavolcanic rocks, mica schists and crystalline dolomites form major lithological units of the Konse Group (Fig. 3). The entire group is deformed and metamorphosed at greenschist-facies metamorphic conditions (Mruma 1989). The deformational fabrics follow the regional NE–SW stretching, which is seen as the geometrical orientation of the lens-shaped crystalline dolomites. The NE–SW stretching and lens shape are also displayed in the weakly foliated granitoids, garnet–pyroxene amphibolites and gabbro-dolerites of the Isimani Suite (Fig. 2). The weakly foliated granitoids are mapped to intrude the contact between the quartzite and the metavolcanic rocks, and it should be noted that the granitoids have emplacement ages dated between 1920 and 1820 Ma (foliated granitoids 1920–1870 Ma; weakly foliated granitoids 1880–1820 Ma) (Fig. 2; Whittingham 1953; Gabert & Wendt 1974; Sommer et al. 2005).
Geological map of the study area compiled from QDS 180 and QDS 198; modified after Whittingham (1953) and Leedal et al. (1954).
A detailed stratigraphy and geological description of the Konse Group was established by Mruma (1989) and published by Mruma (1995). In summary, the Konse Group is composed of a c. 350 m thick succession that id divided into seven formations arranged from bottom to top as follows: Mkulula (quartzites); Ruaha (conglomerate); Kilimbe (mica schists); Kikuyu (metavolcanic rocks: basalts and tuff); Ihumbirisa (dolomitic marble); Mhwana (mica schist) (Fig. 4; Mruma 1989, 1995).
Stratigraphy of the Konse Group, adopted from Mruma (1989, 1995).
Despite the greenschist-facies metamorphic overprint on volcanic and sedimentary sequences of the Konse Group, stratigraphy and primary magmatic structures (pillow lavas) and sedimentary structures (cross-bedding, ripple marks and flute marks) are well preserved in the Kikuyu Formation and Mkulula Formation, respectively (Fig. 4; Mruma 1989, 1995). Owing to the presence of shallow marine sediments (marbles, conglomerates and quartzites with ripple marks), the Konse Group was interpreted as a continental margin sedimentary pile or as sediments deposited in a Paleoproterozoic inland or epicontinental sea (foreland basin) (Mruma 1995; Fritz et al. 2005).
The pyroclastic deposits and pillow basalts of the Kikuyu Formation are correlated with volcanic rocks of the Ndembera Group, and thus are assumed to have the same age (Whittingham 1959; Mruma 1989; Collins et al. 2004; Sommer et al. 2005). Therefore, the Konse volcano-sedimentary rocks are interpreted as a peripheral foreland basin formed at the margin of the Tanzanian Archean Craton during the main phase of the Usagaran deformation (Mruma 1995).
The Konse Group shares tectonic contacts with the Tanzania Craton and the Isimani Suite high-grade gneisses defined by easterly dipping thrusts (Mruma 1989). The Tanzania Craton forms a base upon which the Konse Group tectonically sits and the high-grade gneisses of the Isimani Suite rest on top of the Konse Group (Figs 1–3). Therefore, the thrust slices of the Konse Group and Isimani Suite resting on the Tanzania Craton are regarded as the western termination of the Mozambique Belt easterly dipping thrusts characterized by greenschist-facies metamorphic thrust shear zones in the Usagaran Belt and amphibolite-facies metamorphic thrust shears in the Mozambique Belt (Reddy et al. 2003; Fritz et al. 2005; Tenczer et al. 2007).
The Usagaran Belt is bordered by the Western Granulites of the Mozambique Belt comprising orthogneisses and metasedimentary rocks (metapsammites and metapelites), which were reworked from the Tanzania Craton and the Usagaran Belt during the Neoproterozoic at c. 550 Ma (Sommer et al. 2005; Vogt et al. 2006; Fritz et al. 2009, 2013; Boniface 2019). Therefore, the Neoproterozoic westward-directed thrust faults are considered to relate to the Neoproterozoic East African Orogeny (see Fritz et al. 2013). The westward transport of the Eastern Granulite of the Mozambique Belt (meta-anorthosites) onto the Western Granulites of Tanzania and the Mesoproterozoic Unango–Marrupa Complexes of Mozambique occurred between 596 and 550 Ma during the final amalgamation of the Gondwana supercontinent (Rossetti et al. 2008; Viola et al. 2008; Fritz et al. 2013).
The tectonic model of the Paleoproterozoic Usagaran Belt based on field observations and detailed petrography and geochronology is explained by strike-slip tectonics in an island-arc regime that was overprinted by the Neoproterozoic westward thrust propagation during oblique collision of East and West Gondwana along the Mozambique Belt (Fritz et al. 2005; Tenczer et al. 2007). This is contrary to the general Paleoproterozoic thrusting model of Daly (1988).
Samples and analytical methods
Outcrops and samples
Basalt samples were collected from the Ruaha river section at the Ngengeregengere gauge in the Konse Group (Fig. 3). Six fresh rock samples from outcrops of pillow lava basalts of the Kikuyu Formation were taken for petrographic and major and trace element studies. The coordinates of the samples are given in Table 1. The outcrops occur as massive blocks or sheared mylonites and thrust slices of basalts and associated plutons (gabbro and pyroxenite) (Fig. 5a and b). In hand specimen the rocks are dark green and have an aphanitic texture. The mafic volcano-plutonic unit (basalt, gabbro and pyroxenite) is overlain by a layered dolomite–chert unit and rests on top of a quartzite units along a thrust plane dipping 40° east (Fig. 5a and c). The pillow structures are preserved in the massive blocks of basalts (Fig. 5d and e).
Geological units of the study area include quartzite, basalt (gabbro and pyroxenite, not shown) and dolomite. (a) Mafic volcano-plutonic unit occurs as thrust slices resting on the quartzite and overlain by layered dolomite. The thrust planes dip about 40° to the east. (b) Massive fresh blocks of pillow lavas at the foot of the sheared mafic volcano-plutonic rocks. (c) Close-up photograph of a layered dolomite that rests on the mafic volcano-plutonic rocks. (d, e) Preserved pillow lava structures of the massive basalt block.
Major and trace element composition of the pillow lava basalts from the Konse Group of the Usagaran Belt
Analytical methods
Six fresh rock samples were selected from more that 10 rock samples, and cut into thin sections for qualitative petrographic analysis at the University of Tohoku, Japan. All six petrographically studied samples were crushed and milled using an agate mill. About 18 g aliquots of each milled sample were packed and sent to the Activation Laboratories Limited of Ancaster, Canada for geochemical analyses. The analysis of major and trace elements was performed by the combination of lithium metaborate–tetraborate fusion whole-rock inductively coupled plasma mass spectrometry (ICP-MS) and trace element ICP-MS. Fused samples were diluted and analysed using a Perkin Elmer Sciex ELAN 6000, 6100 or 9000 ICP-MS system. Three blanks and five controls (three before the sample group and two after) were analysed per group of samples. Duplicates were fused and analysed. Details of sample preparation and analytical procedures for major and trace elements at the Activation Commercial Laboratories are as described in the Lithogeochemistry Packages manual available online or by request (www.actlabs.com; +18882285227; fusion process 4B).
Results
Petrography
Six rock samples were petrographically studied using a transmitted light microscope. The rock groundmass consists mainly of laths of plagioclase, chlorite and epidote. Hematite occurs as a minor mineral phase. The laths of plagioclase and sheets of chlorite are randomly oriented (Fig. 6a and b) or may display a slight preferred orientation (Fig. 6c and d). Plagioclase and chlorite form more than 90% of all phases (Fig. 6a–d).
Photomicrographs of pillow lava basalts showing intergranular texture. Plagioclase (Pl) laths preserve magmatic textures in the matrix of predominating metamorphic epidote (Ep) and chlorite (Chl). (a) Plane-polarized light and (b) crossed polars for sample GGG080, ×40. (c) Plane-polarized light and (d) crossed polars for sample GGG082, ×40.
Geochemistry
Major and trace element data and selected inter-element ratios of six samples of the Konse pillow lava basalt are presented in Table 1 and the geochemical characteristics of the lavas are described below.
The major element composition is consistent with the compositional range of basalts in which moderate variations are observed in SiO2 (45.94–50.10 wt%; except for one sample (081GGG) that has SiO2 = 39.90 wt%), Al2O3 (12.58–16.76 wt%), TiO2 (0.49–0.80 wt%), FeOtotal (8.21–12.84 wt%), MgO (6.60–8.34 wt%), K2O (0.07–0.38 wt%) and Na2O (1.2–4.3 wt%). The Mg# ranges between 47.82 (for highly evolved sample) and 60.11 (for the most primitive sample) (Mg# = 100Mg/(Mg + Fe2+)). All the samples display moderate variation but high loss in ignition (LOI) (6.07–8.13 wt%). The large ion lithophile elements (LILE; also known as fluid-mobile elements) are compositionally variable in all samples as observed for K (581–3155 ppm), Rb (1–10 ppm), Sr (44–126 ppm), Ba (26–672 ppm) and U (0.03–0.32 ppm). The LILE ratios are high; for example, K/Rb = 315–1660 and Sr/Rb = 6.5–126 (Table 1). Variation diagrams of selected major elements (SiO2 and K2O) and fluid-mobile elements (Ba and Sr) do not display systematic correlations with Zr contents (Fig. 7a–d). Zr was plotted against other elements in variation diagrams because it is one of the least mobile elements and, therefore, in modern volcanic rocks, Zr is used as an alteration-independent index of geochemical variations (Winchester & Floyd 1977; Pearce & Peate 1995; Polat et al. 2002).
(a–d) Zr v. selected major elements (K2O and SiO2) and fluid-mobile elements (Ba and Sr) highlighting non-correlating behaviour. (e–h) Zr v. selected fluid-immobile elements (TiO2, Sm, Nb and Gd) highlighting correlating behaviour.
The fluid-immobile trace elements display moderate variations in Ti (2888–4742 ppm), V (218–289 ppm), Th (0.14–0.31 ppm), Yb (1.26–2.37 ppm), Nb (1.1–1.9 ppm), Zr (22–36 ppm), Y (10.2–21.6 ppm), La (0.76–3.06 ppm) and Sm (0.77–2.63 ppm). The ratios between fluid-immobile elements are moderately variable and consistent with the compositional range of tholeiite basalts; for example, Zr/Ti = 0.007–0.008, Nb/Y = 0.09–0.13, Ti/V = 12.84–19.14 and (La/Sm)N = 0.62–1.09. Selected fluid-immobile elements (TiO2, Sm, Nb and Gd) display systematic correlations with Zr contents (Fig. 7e–h).
On a chondrite-normalized REE diagram, the samples display consistently flat patterns with LREE concentrations ranging between 5.58 and 9.87 times chondritic values, except for one sample (080GGG) that has depleted LREE relative to the HREE (Fig. 8a). Three samples (077GGG, 078GGG and 083GGG) display REE patterns similar to those of standard MORB glasses, which are characterized by permissively no Eu and Sr anomalies (Fig. 8a and b; e.g. Klein 2003). In contrast, two samples (081GGG and 082GGG) display positive Eu anomalies (Eu/Eu* = 1.21 and 1.19, respectively) correlating with positive Sr anomalies and one rock sample (080GGG) displays a negative Eu anomaly but positive Sr anomaly (Fig. 8a and b).
(a) Chondrite-normalized REE spidergrams highlighting the REE patterns of the Konse pillow lavas; the normalizing data are from Boynton (1984). (b) MORB-normalized trace element spidergrams; the normalizing data are from Hofmann (1988). For comparison purposes, trace element data in blue and brown patterns have been added from the Izu–Bonin–Mariana and New Britain–Manus back-arc basins respectively (data source Metcalf & Shervais 2008). (c) Nb/Yb–Th/Yb plot (after Pearce (2008).
On a normal (N)-MORB-normalized trace element diagram (Fig. 8b), the samples display elevated concentrations of slab-derived components (Cs, Rb, Ba, Th, U, K, La, Ce, Pb, Sr) that are superimposed on conservative mantle wedge-derived components (Nb, Zr, Sm, Eu, Ti, Dy, Y, Yb, Lu), therefore producing a distinctive pattern of ‘spikes’ and ‘troughs’. The patterns (e.g. strong spikes of K, Pb, Th and Sr) of the N-MORB-normalized trace element diagrams of the Konse pillow lavas are very similar to those of enriched (E)-MORB from the Izu–Bonin–Mariana and Lau–Tonga back-arc tholeiite basalts and are also comparable with those for basalts with a hybrid mixture of MORB-like and arc-like element signatures from the Jurassic Mirdita Ophiolite (Albania) (Fig. 8b; Dilek et al. 2008). Generally, the Konse pillow lavas show a pronounced enrichment in highly incompatible elements (i.e. Cs, Rb, Th and U).
The Th/Yb–Nb/Yb discrimination diagram indicates that the composition of the Konse pillow lavas straddles N-MORB and E-MORB compositions, similar to basalts that erupt in known back-arc tectonic settings (e.g. Metcalf & Shervais (2008) (Fig. 8c).
Discussion
Petrographic studies indicate that the Konse pillow lavas are intensively metamorphosed at greenschist-facies conditions. It is thus interpreted that epidote and chlorite are secondary minerals that have completely replaced magmatic counterparts, most probably pyroxene and olivine. Plagioclase occurs as laths reflecting typical magmatic textures in basalts. Therefore, plagioclase grains have preserved their original magmatic crystallization textures. The modification of the primary minerals and magmatic textures of the Konse basalt is connected to the deformation and metamorphism of the Usagaran Belt during the Neoproterozoic Era, when the tectonic nappes of the Mozambique Belt were thrusted onto the Usagaran Belt and the Tanzania Craton (Mruma 1989; Reddy et al. 2003, 2004; Collins et al. 2004; Boniface 2019, and references therein). Additionally, Reddy et al. (2003) described the earlier amphibolite-facies deformation of the Usagaran Belt to have occurred in Paleoproterozoic time during the Usagaran collisional event (c. 1990 Ma), and during the emplacement of the voluminous volcano-plutonic magmatism in the Ndembera Group (c. 1920–1870 Ma).
The effect of greenschist-facies metamorphic overprint on the Konse pillow basalts caused the modification of primary geochemical signatures. It has been observed that all the analysed samples display large variations in major elements and LILE accompanied by non-correlating behaviour of Zr versus major elements and LILE. It has also been observed that the Konse Pillow lavas have elevated loss on ignition ranging between 6.07 and 8.13 wt%, which is most probably contributed by the abundance of chlorite and epidote (hydrous minerals).
It should be noted that major elements, LILE (K, Rb, Na, Sr, Ba, U, Pb) and LREE are prone to redistribution during metamorphism owing to their high mobility (Tatsumi et al. 1986; Polat & Hofmann 2003). In contrast, Nb, Ta, Ti, Zr, Y and HREE are relatively immobile during alteration processes (Polat & Hofmann 2003; Dilek et al. 2008; Dilek & Furnes 2009). Because of their mobility, LILE ratios of the Konse pillow basalts are not always consistent with a particular tectonic setting. For instance, in our samples, K/Rb values are relatively high, ranging between 313 and 1660; this ratio range is consistent with the composition of N-MORB. However, K/Ba ranges between unity and 99, and Sr/Rb ranges between seven and 127, showing overlapping signatures for MORB, P-type (plume-type) MORB and ocean-island tholeiite basalts (Wilson 2007). Therefore, the petrogenesis and tectonic settings of the metamorphosed Konse pillow lavas are determined by evaluating the geochemical signatures of the fluid-immobile elements (e.g. Nb, Ta, Ti, Zr, Y and HREE) (see Polat et al. 2002; Polat & Hofmann 2003; Dilek et al. 2008).
Petrogenesis of the Konse pillow basalts
The REE patterns and (La/Sm)N (0.62–1.09) of the Konse pillow basalts conform to the REE patterns of transitional (T)-MORB and N-MORB (e.g. Schilling et al. 1983 ; Sun & McDonough 1989; Klein 2003). T-MORB refers to mid-ocean ridge basalts with geochemical composition between N-MORB (with (La/Sm)N <0.7) and E-MORB (with (La/Sm)N >1.8) (see Schilling et al. 1983).
Out of the six analysed rock samples, three samples (077GGG, 078GGG and 083GGG) have REE patterns similar to that of T-MORB. However, two samples (081GGG and 082GGG) display positive Eu anomalies that correlate with positive Sr anomalies. Tang et al. (2017) explained the observation of positive Eu and Sr anomalies in MORB glasses as described below. Non-correlating negative Eu anomalies and positive Sr anomalies in one rock sample (080GGG) with strong depletion of LREE are not easily explainable.
Tang et al. (2017) analysed 72 primitive MORB glasses (MgO >8.5 wt%) from the Pacific, Indian and Atlantic mid-ocean ridges and found that they showed both positive and negative Eu (Eu/Eu* = 0.92–1.13) and Sr anomalies (Sr/Sr* = 0.81–1.98) contrary to standard MORB glasses. For a MORB source with the mineral assemblage olivine, orthopyroxene, clinopyroxene and spinel (e.g. Niu 1997), the anomalies cannot be produced by plagioclase fractionation at MgO >9 wt%, therefore disequilibrium melting in the MORB source mantle may result in Eu and Sr enrichments in the partial melts because Eu and Sr are preferentially extracted relative to Pr, Nd, Sm and Gd (Tang et al. 2017).
Although the mantle disequilibrium melting model explains the Eu and Sr anomalies in the MORB, the model cannot account for the observed elevation of Ba, Th and U, and other fluid-mobile elements. Recycled lower continental crustal materials, which have positive Eu and Sr anomalies, when partially mixed into the upper mantle could contribute to the elevation of Ba, Th and U in the MORB source (Niu 1997; Niu & O'Hara 2009; Tang et al. 2015). Therefore, we will discuss the composition of our samples in connection with subduction processes that could suitably explain the contribution of the crustal component into the mantle by subducting continental crustal materials (see Perfit et al. 1980; Shervais 1982; Gill 1987; Ikeda & Yuasa 1989; Clift 1995; Pearce 2008; Dilek & Furnes 2009, 2011; Furnes & Dilek 2017).
The analysis of high-valency (fluid-immobile) element ratios (e.g. Ti, V, Th, Yb, Nb, Zr, Y, La, Sm) and trace element patterns indicates that the Konse mafic pillow basalts display a mixed signature of a MORB tholeiite and island-arc tholeiite, which points to the composition of the Konse mafic pillow basalts being similar to that of mafic basalts of a class of subduction-related ophiolites. Subduction-related ophiolites include suprasubduction-zone and volcanic-arc types, the evolution of which is governed by slab dehydration and accompanying metasomatization of the mantle peridotites, melting of the subducting sediments and repeated episodes of partial melting of metasomatized mantle peridotites (see Volpe et al. 1987; Dilek et al. 2008; Metcalf & Shervais 2008; Dilek & Furnes 2011).
The Zr/Ti versus Nb/Yb discrimination diagram indicates that all the samples have a basaltic composition (Fig. 9a). The ratios between Ti and V (i.e. Ti/V) range between 12.8 and 19.1, which points to island-arc or back-arc basin tectonic settings (Ti/V = 10–20 for island arc; Ti/V = 10–50 for back-arc basin; Fig. 9b; Shervais 1982). The mixed signature between MORB and arc lavas is also displayed in other discrimination diagrams; for example, TiO2 (wt%) versus Zr (ppm) and the Zr/Y versus Zr (ppm) diagram (Fig. 9c and e). All these diagrams suggest a mixed geochemical signature between N-MORB and volcanic-arc tholeiites (VAT) that is not differentiated. However, the chondrite-normalized REE patterns and the N-MORB-normalized trace element patterns (Fig. 8a and b), and the trace element ratios (e.g. Th/Yb, Nb/Yb, Zr/Y, Zr/Ti, Ti/V) of the Konse ophiolitic basalts completely match the composition of back-arc suprasubduction-zone ophiolitic tholeiite basalts. Therefore, the displayed geochemical signatures in the Konse pillow basalts match the data from Izu–Bonin–Mariana, New Britain–Manus and Lau–Tonga back-arc basalts (Fig. 8f; Hawkins & Melchior 1985; Gill 1987; Ikeda & Yuasa 1989; Metcalf & Shervais 2008; Reagan et al. 2010).
Fluid-immobile or alteration-resistant element concentrations of the Konse pillow lavas correlated with known tectonic setting. (a) Variation of Zr/Ti v. Nb/Yb points to basalt composition; diagram from (Pearce 1996). (b) Ti v. V plot of the Konse basalts (Ti/V ratios are characteristic of: 10–20 = island arcs; 20–50 = MORB; 20–30 = mixed MORB and island arc; 10–50 = back-arc basins; and boninite field; diagram after Shervais (1982). (c) TiO2 (wt%) v. Zr (ppm); diagram adopted from Pearce (1996). (d) Nb/La v. (La/Sm)N variation diagram, adopted from John et al. (2003). (e) Zr–Zr/Y plot (diagram after Pearce (1980) and modified by Furnes et al. (2007)). Suprasubduction-zone data source Furnes et al. (2012) and references therein. Suprasubduction-zone (SSZ) setting: FA, forearc; BA-FA, back-arc to forearc; CBA, continental back-arc; OBA, oceanic back-arc; SSOC, Solund–Stavfjord ophiolite complex (for ophiolite samples from Norway, see Furnes et al. 2012). WPB, within-plate basalt; IAT, intra-arc tholeiite. (f) Nb/Yb–Th/Yb plot comparing data from the Izu–Bonin–Mariana and New Britain–Manus back-arc basins; IOB, ocean island basalt (for data source see Metcalf & Shervais 2008); diagram after Pearce (2008).
The genesis of suprasubduction-zone ophiolites involves the initiation of subduction, followed by rapid slab rollback leading to extension, seafloor spreading in the upper plate and generation of melt with MORB-like compositions in the proto-arc to back-arc setting in the early phase of melting (Gill 1987; Dilek & Furnes 2011). However, the subsequent phases of melting are strongly influenced by slab dehydration and related mantle metasomatism, melting of subducting sediments, repeated episodes of partial melting of metasomatized peridotites, and mixing of highly enriched liquids from the lower fertile source with refractory melts in the melt column beneath the extending protoarc–forearc region (see Shervais 2001; Dilek 2003b; Hawkins 2003; Pearce 2003; Metcalf & Shervais 2008; Dilek & Furnes 2011; Saccani et al. 2018). Suprasubduction-zone ophiolites generally display a characteristic, sequential evolution of MORB to island-arc tholeiite to boninitic igneous activity, which manifests itself in a vertically and laterally well-developed chemostratigraphy of the Phanerozoic ocean systems, such as the Izu–Bonin–Mariana arc system (Dilek & Furnes 2009; Reagan et al. 2010; Furnes et al. 2014).
Preserved Orosirian MORB and volcanic-arc magmas
Our geochemical data and the previously published geochemical, petrological and geochronological data indicate that the Usagaran Belt is an exemplary site of preserved Paleoproterozoic relics of volcanic-arc magmas, fragments of subducted oceanic crust and allochthonous island arc systems (Möller et al. 1995; Reddy et al. 2003; Collins et al. 2004; Sommer et al. 2005; Bahame et al. 2016). In the Isimani Suite (NE Usagaran Belt), the Usagaran Belt preserves eclogites with MORB affinity that represent a 2.0 Ga subducted oceanic crust (Fig. 10a; Möller et al. 1995; Reddy et al. 2003). In the Ndembera Group (SW Usagaran Belt), the Usagaran Belt preserves calc-alkaline volcanic rocks with geochemical affinity to arc magmas that were emplaced above the subduction zone between 1920 and 1870 Ma (Fig. 10b; Gabert & Wendt 1974; Reddy et al. 2003; Fritz et al. 2005; Bahame et al. 2016). The entire Usagaran Belt is intruded by suites of granitoids with affinity to arc magmas, which are concomitant with arc volcanic rocks dated between 1920 and 1870 Ma (Gabert & Wendt 1974; Reddy et al. 2003; Sommer et al. 2005; Bahame et al. 2016).
(a) Sketch map and evolution model of the Usagaran Belt with profile X–X′ through the Isimani Suite and profile Y–Y′ through the Ndembera Group adopted from Fritz et al. (2005). (b) The Isimani Suite and northern Usagaran Belt evolved by subduction of an oceanic crust at c. 2000 Ma and collision at c. 1990 Ma (Möller et al. 1995; Collins et al. 2004). (c) The Ndembera Group and southern Usagaran Belt evolved between 1920 and 1870 Ma in a volcanic-arc system. (d) Our new geochemical data indicate that the pillow lava basalts of the Konse Group formed in a back-arc environment and the basalts have a geochemical signature similar to the ophiolites of the Jurassic–Cretaceous suprasubduction zones of the Tethyan ocean systems.
The Konse volcano-sedimentary rock assemblages point to a derivation from a suprasubduction-related ophiolitic terrane that was dismembered and accreted at the margins of the Tanzania Craton in the Paleoproterozoic Era (Fig. 10b). The pillow lava basalts of the Konse Group have affinity to N- and T-MORB and have well-preserved trace element ratios and trace element and REE patterns indicating the injection of a subduction fluid component, the composition of which is generated only in suprasubduction-zone back-arc systems. Therefore, the Konse pillow basalts are hereby interpreted as relics of subduction-related magmas that erupted at a spreading centre in a back-arc tectonic setting. The intrusion of 1920–1820 Ma granitoids into the Konse Group indicates that the minimum age of the Konse Group is 1920 Ma.
The geometric configuration of the Usagaran Belt indicates that the Konse Group was emplaced between the Archean Tanzania Craton and the 2000 Ma high-grade Isimani Suite (subduction–collision terrane) (Fig. 10). The low-grade (greenschist-facies) Konse Group represents a back-arc terrane derived from a suprasubduction-zone geodynamic setting and might be coeval with the 2000 Ma high-grade Isimani Suite. Therefore, the Konse suprasubduction-zone related tectonic domain was emplaced onto the Tanzania Craton during the 2000 Ma collision following the subduction of the Yalumba oceanic crust.
Our geochemical data indicate that the Konse back-arc terrane related pillow lavas are highly reminiscent of the Jurassic–Cretaceous suprasubduction-zone ophiolite lavas of the Tethyan ocean systems that formed during and after the Pangea break-up (see Stampfli 2000; Bortolotti & Principi 2005; Dilek & Furnes 2009, 2011). Therefore, the Konse pillow lava basalts and the associated chert and limestone (dolomite) probably represent the upper part of a suprasubduction-zone type columnar ophiolitic section (see Dilek & Furnes 2011, 2014), which was displaced onto the Tanzania Craton. However, the geometric relations of the Konse pillow lavas and the sedimentary cover (chert and dolomite), along with other parts of the ophiolitic section, have not yet been described.
Other researchers consider that the Konse Group formed in a younger ocean basin that is concomitant with the 1920–1870 Ma Ndembera volcanic suites (Whittingham 1959; Mruma 1989; Collins et al. 2004; Sommer et al. 2005). The 1920–1870 Ma Ndembera calc-alkaline volcanic rocks and the associated calc-alkaline granitoids seem to represent arc magmas that preceded a subduction that existed at c. 1920 Ma. This means that the Ndembera volcanic rocks might represent orogenic magmatism at the active continental margins that followed an undocumented subduction phase postdating the 2.0 Ga subduction. Recently, Boniface & Tsujimori (2014, 2016) have reported kyanite–epidote eclogites with a subduction age of 1924 ± 74 Ma at the SW margins of the Tanzania Craton, which indicates that the ocean basin persisted at the margins of the Tanzania Craton until around 1860 Ma (e.g. Boniface et al. 2012). Therefore, the Ndembera calc-alkaline volcanic rocks and the associated calc-alkaline granitoids are probably the manifestation of pre-collisional orogenic events at the active continental margins following the c. 1920 Ma subduction process that is recorded in the Ubendian Belt (see Fig. 10; Fritz et al. 2005; Liégeois et al. 2013; Boniface & Tsujimori 2014, 2016; Bahame et al. 2016). It is also possible that the Konse back-arc terranes accreted onto the Tanzania Craton at c. 1920 Ma during the initiation of arc magmatism of the Ndembera Group.
Conclusions
The Usagaran Belt in Tanzania is an exemplary site of preserved remnants of a subducted Paleoproterozoic oceanic crust, allochthonous back-arc pillow lava basalt and volcanic-arc magmas of the Orosirian Period.
The Orosirian Usagaran oceanic crust remnants are metamorphosed at greenschist-facies conditions; consequently, major elements and LILE (e.g. K, Rb, Na, Sr, Ba, U, Pb) are mobilized. However, fluid-immobile elements (e.g. Nb, Ta, Ti, Zr, Y and HREE) preserve the primary geochemical characteristics of the Konse pillow basalts.
The geochemical trends of the Usagaran Orosirian pillow basalts are highly reminiscent of the composition of the Jurassic–Cretaceous suprasubduction-zone ophiolites of the Tethyan ocean systems that formed during and after the Pangea break-up.
The rock assemblage (mafic pillow lava basalt, chert and dolomite) in the Konse Group probably represents the upper part of a Paleoproterozoic suprasubduction-zone columnar ophiolitic section. The geometric relations of mafic pillow lava basalt, chert and dolomite in the Konse Group are not yet established, along with other parts of the columnar ophiolitic section.
The Konse suprasubduction-zone related pillow basalts were emplaced onto the Tanzania Craton probably during the 1990 Ma collision following the subduction of the 2000 Ma Yalumba oceanic crust. Therefore, the Konse ophiolite might have been emplaced at c. 2000 Ma from the Yalumba ocean system.
The recorded arc magmas and the associated granitoids in the Usagaran Belt are dated between 1920 and 1870 Ma, pointing to the presence of subducting oceanic crust at c. 1920 Ma. The 1920 and 1870 Ma granitoids have been mapped to intrude the Konse Group. Therefore, the Konse pillow basalts might be the relics of an oceanic crust that accreted before the onset of the Ndembera arc volcanism at c. 1920 Ma.
The Konse pillow lava basalts have a tectonic contact with the high-grade Isimani suites and the Konse lavas were metamorphosed at greenschist-facies conditions, which might imply that the lavas were not part of the 2.0 Ga subduction and 1.99 Ga collisional events. Establishing zircon crystallization ages will precisely determine the stratigraphic position of the Konse lavas.
Acknowledgements
We acknowledge Y. Dilek, A. Polat and H. Furnes for their constructive reviews that significantly improved the paper, along with editors who handled the submissions efficiently. We are also grateful to the University of Dar es Salaam for the support in various forms during field visits. Special thanks go to A. Mruma and M. Maruvuko of the Geological Survey of Tanzania for their support. We appreciate the TANESCO management for granting permission to access the sampling site at the Mtera Dam. We thank R. Kinyaiya for driving us safely in a Tokyo LandCruiser 70 in difficult but beautiful terrains.
Funding
This research was partly supported by the Center for Northeast Asian Studies, Tohoku University and partly supported by grants from the MEXT/JSPS KAKENHI (JP24403010, JP15H05212 and 18H01299) to T.T. and for a visiting professorship to N.B. at Tohoku University.
Scientific editing by Yildirim Dilek
- © 2019 The Author(s). Published by The Geological Society of London. All rights reserved
References
Pillow lava basalts with back-arc MORB affinity from the Usagaran Belt, Tanzania: relics of Orosirian ophiolites
