Introduction

Late Cenozoic igneous rocks are widespread in Madagascar, from the southwest, between Tulear and Morondava, to the northernmost tip of the island (Fig. 1a). The most important areas covered and/or intruded by igneous rocks in northern Madagascar belonging to this cycle are the Ampasindava peninsula, the Nosy Be archipelago, the Ankaizina district, and the Massif d’Ambre stratovolcano. The Massif d’Ambre covers an area of roughly 60 km × 50 km and is the largest volcanic complex of northern Madagascar (Besairie et al. 1957; Cucciniello et al. 2011 and references therein; Fig. 1a). The age range of Bobaomby volcanism is still not well constrained: Emerick and Duncan (1982, 1983) provided K–Ar ages of 9.4–9.1 Ma for unspecified lava flows collected in the Bobaomby area, and Cucciniello et al. (2011) dated alkali feldspar separates from two phonolites with the 40Ar/39Ar method, obtaining ages of 10.14 ± 0.07 Ma and 10.56 ± 0.09 Ma. Most other occurrences in the area have roughly the same age range (cf. Cucciniello et al. 2016; Estrade et al. 2014), and so is for the volcanic rocks emplaced in the southwestern Madagascar (11–12 Ma; Cucciniello et al. 2018), strongly suggesting a regional uplift event throughout Madagascar.

Fig. 1
figure 1

a Simplified geological map of Madagascar, showing the outcrops of Cretaceous (black and blue areas) and Cenozoic (red areas) igneous rocks. b Geological setting of the Bobaomby volcanic field, with the main volcanic features, and sampling sites

The Bobaomby volcanic field is widely scattered over a triangular area of ab. 500 km2, bordered by Indian Ocean, Mozambique Sea and several bays and islets towards the south (Fig. 1b). The volcanic field is formed by large scoria cones, forming at least two prominent hills on the western side of the peninsula, associated with small, still well preserved, lava flows; a dyke swarm with two orientations (NNE–SSW and NNW–SSE), with some of the dykes cropping out also for 2 km; plugs, lava flows (e.g., at Lomoto); tuff rings, preserved in the neighborhoods of Cap d’Ambre (Tanjona Bobaomby). The volcanic rocks cross-cut a backbone of Upper Cretaceous sedimentary rocks and Paleogene nummulite- and lepidocycline-bearing limestones, (Besairie and Collignon 1972) that reach their maximum height at Windsor Castle (ca. 400 mt asl). The metamorphic basement cropping out to the south of Massif d’Ambre and Ankarana limestones, is formed by the Late Proterozoic (Pan-African age 0.76–0.71 Ga) Bemarivo belt (Jöns et al. 2006; zircon U–Pb ages in Tucker et al. 2014), which mark the latest orogenic events in this part of Gondwana.

The Cap d’Ambre (Bobaomby) volcanic rocks crop out ab.  450 km east of Mayotte, the southeastern-most island of the Comoro archipelago. The rocks of Mayotte island are mostly mafic alkaline/strongly alkaline lavas (basanites, nephelinites and olivine melilitites) and evolved trachyphonolites and phonolites (Pelleter et al. 2014; Späth et al. 1996). They are only slightly younger than the Bobaomby rocks (< 7.7 Ma; Pelleter et al. 2014 and references therein). Mayotte volcanism is currently active offshore to the south, hence very far from Karthala and la Grille currently active volcanoes at Grande Comore (e.g., Berthod et al. 2021). The origin of the Comoros archipelago in the Mozambique channel is the subject of a long controversy. This volcanic chain has been interpreted as (1) a hotspot track (e.g., Emerick and Duncan 1982); (2) a passive margin (e.g., Nougier et al. 1986); or (3) a zone of intense deformation in the Nubia-Somalia plate system (Stamps et al. 2021). Famin et al. (2020) linked the tectonics and magmatism setting of the Comoro archipelago a generalized dextral shear deformation with a NW–SE-directed maximum horizontal stress, active since at least the past million years.

The Late Cenozoic volcanism in Madagascar has been generally attributed to regional extension (intracontinental rifting) similar to that of East African Rift (Melluso et al. 2016 and references therein; Minissale et al. 2022). A continental nature of the basement beneath the Comoro archipelago, like that of northern Madagascar, has been inferred on geophysical grounds (e.g., Dofal et al. 2021).

The present manuscript presents a detailed mineralogical and geochemical study of the northernmost Madagascar volcanic rocks which increase, integrate and improve the dataset from Melluso et al. (2007), with the aim to better highlight magmatic processes which acted during the ascent in the crust. We also focus on the petrogenetic relationships and differences with other Cenozoic volcanic districts of Madagascar, particularly those of the northern Madagascar and Mayotte. Intrusive xenoliths found as lithics of spatter cones and as inclusions in plugs are detailed here for the first time. Mantle xenoliths in primitive lavas of Nosy Be, Massif d’Ambre and Bobaomby have been the subject of earlier studies (Rocco et al. 2013, 2017; Mazzeo et al. 2021; other studies are still ongoing) which indicated variable physico-chemical condition of the spinel-bearing lithospheric mantle section beneath northern Madagascar. Based on these data, we will try to highlight similarities and differences between the volcanoes in northernmost Madagascar (Massif d’Ambre, Nosy Be archipelago and Ampasindava peninsula) with the roughly coeval magmatism at Comoro archipelago (especially Mayotte, which is the closest active island of Comoros) in order to add more clues to the tectonomagmatic evolution of Mozambique Channel and Somali Basin.

Classification, petrography and mineral chemistry

The sample locations are reported in the Fig. 1b and the chemical analyses are reported in the supplementary Tables 1 and 2. The Bobaomby volcanic rocks are nephelinites, basanites, tephrites, tephritic phonolite and phonolites, on the basis of the T.A.S. diagram (Le Maitre et al. 1989; Fig. 2a), actual mineralogy and CIPW norms. The mafic rocks have mostly sodic and, less marked, potassic serial affinity (Fig. 2b). A dyke sample (M568; Na2O = 2.2 wt%; K2O = 3.7 wt%) can be considered as ultrapotassic (Fig. 2b); the mica and amphibole abundance of the sample (see below) indicates that it is a lamprophyre. It is well known that the classification in the T.A.S. diagram is not fully indicative of the distinction between basanites and nephelinites. Therefore, as shown elsewhere (e.g., in the classification of the Honolulu volcanics and the Ankaratra volcanic rocks in central Madagascar; Clague and Frey 1982; Cucciniello et al. 2017), a petrographic distinction between basanites (“ankaramites s.l.”) and nephelinites (“ankaratrites”) is pertinent in this context: basanites are petrographically distinct from nephelinites by the presence of modal plagioclase (mostly as a groundmass phase, as is alkali feldspar), whereas nephelinites have Ba-rich mica and Ba-sanidine in the groundmass, but no plagioclase (see below). This has also significant bearing for the magmatic evolution of strongly alkaline rocks.

Fig. 2
figure 2

a T.A.S. (Le Maitre et al. 1989) and b Na2O vs. K2O (Middlemost 1975) diagrams for the Bobaomby igneous rocks. The rocks of the Massif d’Ambre, Ankaratra, Itasy, Takarindiona, Nosy Be and those of Mayotte are also plotted for comparison (data from Cucciniello et al. 2011, 2017; Melluso and Morra 2000; Melluso et al. 2011, 2018 and Pelleter et al. 2014)

The phonolites (MgO < 0.4 wt%, CaO < 2.5 wt%; Fe2O3t < 4 wt%) have a peralkaline index (P.I., molar (Na + K)/Al) varying from 0.92 to 1.18.

Petrography

Thin section and backscattered electron images of the most interesting lithologies and petrographic features of the Bobaomby samples are reported in the supplementary Fig. S1.

Olivine nephelinites M634, M634N1, and nephelinite M556 have different textures. M634 is porphyritic for olivine (Fig. S1a) and rare clinopyroxene phenocrysts (Fig. S1b), set in a groundmass very rich in clinopyroxene, with additional opaques and foids. Disaggregated xenoliths composed by green clinopyroxene and amphibole (with rhönite subsolidus rims) are also noted. M634N1 and M556 are coarser-grained, and have olivine, clinopyroxene and opaques as phenocrysts, and clean nepheline as the dominant interstitial felsic phase. Biotite is also observed in the groundmass. Disaggregated xenoliths of mantle peridotite (olivine, enstatite, diopside and spinel) and xenocrysts are also found in sample M634N1 (Fig. S1c).

Basanites and tephrites. Basanites are typically porphyritic for olivine and clinopyroxene phenocrysts (Fig. S1d), set in a groundmass rich in feldspar and clinopyroxene, with additional oxides and feldspathoids (Fig. S1e). Tephrites have abundant phenocrysts of clinopyroxene and olivine, in a groundmass with feldspars, nepheline, oxides and amphiboles (Fig. S1f). More evolved tephrites (samples M576, M544, M545) have large amphibole phenocrysts (Fig. S1g) with strongly zoned clinopyroxene, in a groundmass with Ba-Sr-rich alkali feldspar and nepheline, and grade to the mineralogy of the tephritic phonolite M633. Very frequent is the presence of microxenoliths and xenocrysts (e.g., M634, M574E, M541A, M541, M755N5, M554, M554N1, M554N2, M755). A typical feature of nephelinites, tephrites and basanites is the presence of highly zoned clinopyroxene phenocrysts, having deep brown pleochroic to green Fe-rich cores (xenocrysts) and then purple, Ti-rich equilibrium rims, or also oscillatory patterns (Fig. S1h).

Lamprophyre (sample M568). The lamprophyre (this sample can be classified as tephrite based on the T.A.S.) has phenocrysts of altered olivine, with oxide inclusions, clinopyroxene, amphibole and biotite, set in a groundmass with additional feldspar, feldspathoids and oxides. A few mica crystals (possibly disaggregated xenocrysts) unexpectedly host perovskite grains with magnetite (Fig. S1k).

Tephritic phonolites. The tephritic phonolite dyke M633 is characterized by phenocrysts of amphibole and clinopyroxene, and very rare plagioclase, sanidine, magnetite and titanite, set in a trachytic groundmass rich in sanidine, magnetite, pyroxene, amphibole, rare biotite and feldspathoids (Fig. S2a).

Phonolites. The Bobaomby phonolite lavas and dykes are poorly porphyritic, with phenocrysts of alkali feldspar and nepheline, plus rare clinopyroxene, opaque oxides, titanite and corroded brown amphibole, sometimes in evident disequilibrium with the host lava (Fig. S2b). Deep green, aegirine-rich clinopyroxene and aenigmatite are more frequent as tiny microlites in the groundmass (Fig. S2c). Felsic enclaves with different structure (rich in fine-grained amphibole, and with nepheline phenocrysts, Fig. S2d), and highly reacted olivine xenocrysts are also found (Fig. S2e).

Coarse xenoliths and xenocrysts

The coarse-grained holocrystalline inclusions of Bobaomby and Massif d’Ambre lavas were classified on the basis of their petrographic characteristics and modal abundance of the minerals (supplementary Table 3). The amphibole- gabbro M541A has cumulus euhedral olivine (20%) with spinel inclusions, subhedral clinopyroxene (14%) and euhedral plagioclase (18%) (Fig. S3a) included in poikilitic amphibole (46%). The xenolith M541 is modally heterogeneous, showing patches rich in plagioclase (48%) with smaller amount of idiomorphic olivine (7%) and clinopyroxene (11%), and patches rich in olivine and clinopyroxene (25 and 24%, respectively) with minor plagioclase (13%) and large poikilitic amphibole (31%) and biotite (6–7%); opaques and apatite complete the mineral assemblages (Fig. S3b). Wehrlites have large euhedral olivine (up 78% in M775N5) and smaller opaque oxides (on average 2%) in intercumulus clinopyroxene (22%) (Fig. S3c); sample M554 is a wehrlite formed by olivine (56%) and poikilitic clinopyroxene (25%) (Fig. S3d). Olivine and clinopyroxene host small opaque crystals. A large vein of amphibole (19% of the whole thin section) cuts the xenolith (Fig. S3e); a few small interstitial amphibole crystals can be found rimming clinopyroxene. M554N2 is a wehrlite similar to M554, olivine (63%), clinopyroxene (17%) and amphibole (11%) are main forming minerals (Fig. S3f). Amphibole (2%), olivine (41%) opaques (19%) and clinopyroxene (37%) are the main phases of the xenolith M554N1 (wehrlite) (Fig. S3g). Smaller wehrlite nodules can be found in lavas as clusters of olivine and poikilitic clinopyroxene (Fig. S3h and S3k).

Mineral chemistry

A full data set of mineral chemistry (roughly 1500 new analyses), together with the limited dataset already available in Melluso et al. (2007) was obtained from volcanic rocks spanning the observed range of compositions, as well as the intrusive nodules. A synopsis is reported in the supplementary Table 3 and the full data are reported in the supplementary Tables 4–14.

Olivine of the mafic lavas (nephelinites, basanites and tephrites) range in composition from Fo88 in the cores of the idiomorphic phenocrysts to Fo36 in the rims of the groundmass crystals (Fo is 100xMg/(Mg + Fe) in mol). Olivine never shows reverse zoning and has the typical MnO and CaO increase towards the Fe-rich compositions. The most Mg-rich olivine cores of the volcanic facies are in equilibrium with the composition of the host rock (Fig. 3). Xenocrysts of olivine with Fo87-Fo91 composition (recognized on the basis of petrography and chemical composition) occur in basanites and nephelinites and in partially disaggregated xenoliths. Rare, highly magnesian and thus strongly resorbed xenocrysts of olivine were found also in phonolites, with composition Fo82 (Fig. 3).

Fig. 3
figure 3

Olivine Mg# plotted vs. Mg# of the host rock

The idiomorphic/subidiomorphic olivine in intrusive nodules of gabbros and wehrlites ranges in composition from Fo80 (sample M554) to Fo68–Fo87 (sample M541A) and Fo64–Fo86 (sample M541), indicating that these rocks formed from variably evolved magmas, with Mg# ranging from 66 to 30 (according to Roeder and Emslie 1970). NiO ranges from 0.17 to 0.48 wt% and from 0.12 to 0.54 wt% in olivine from gabbros and wehrlites, respectively. Olivine of gabbros has CaO in the range 0.14–0.27 wt%; olivine in wehrlites has CaO in the range 0.03–0.44 wt%. Olivine phenocrysts of host lavas are far more Fo-rich (Fo88) than olivine of the nodules, again confirming the cumulitic and xenolithic nature of the latter (Fig. 3).

Clinopyroxene phenocrysts to groundmass crystals of nephelinites, basanites and tephrites vary from diopside (Mg# = 85; where Mg# = molar Mg*100/(Mg + Fe) to titanaugite (Mg# = 70; TiO2 = 7.64 wt%; Al2O3 = 14 wt% nephelinite M556) (Fig. 4a and b) as observed in all the Cenozoic mafic/ultramafic alkaline rocks of Madagascar, suggesting the substitution (Mg, Fe) + 2Si- > Ti +  2IVAl and minor Mg + Si- > Fe3+ + IVAl. Clinopyroxene of the tephriphonolite M633 has Mg# varying from 51 to 71 and is low in TiO2, Al2O3 and Na2O (Fig. 4a and c) supplementary Table 5), which reflects crystallization from an evolved melt, and the co-crystallization with kaersutite and magnetite. The clinopyroxene of the phonolites ranges from salite, in cores and microphenocrysts, to aegirine, in the groundmass microphenocrysts and microlites (Fig. 4c; Mg# from 73 to 0; Na2O up to 13.5 wt%), with the low Al2O3 expected from an interstitial mafic mineral in a peralkaline rock, and variable concentration of TiO2, MnO and ZrO2 (supplementary Table 5). The Na-enrichment trend in Bobaomby rocks is abrupt; it appears in the groundmass compositions, and only in the phonolites (e.g., M548, M559, M572, M573; Fig. S4). The aegirine-rich compositions are characterized by low Mn and variable Ti (up to 0.15 apfu; TiO2 = 5.27 wt%). Clinopyroxene Fe-rich cores in the mafic rocks are not similar to the Ti-poor compositions of the phonolites (Fig. S4; supplementary Table 5), but have low VIAl, still indicating a low-pressure environment of crystallization.

Fig. 4
figure 4

a Clinopyroxene Mg# vs. Mg# of the host rock. b, c, d Composition of pyroxene in the Bobaomby igneous rocks

Clinopyroxene of gabbroic rocks is diopside (Ca44.8–50.3Fe6.8–14.8Mg35.6–47); clinopyroxene of amphibole-wehrlites is more rich in Ca (Ca50.8–53.5Fe9.9–15.3Mg31.6–37.9). (Fig. 4d). Al2O3 reaches the highest concentration in clinopyroxene of amphibole-wehrlites (from 7.5 to 10.1 wt%) and the lowest in clinopyroxenes of gabbros (1.4–7.2 wt%). The concentration of Na2O is very similar in the pyroxenes of the two rock types, ranging from 0.6 to 1.1 wt% for clinopyroxene of gabbros, from 0.38 to 0.8 wt% for amphibole-wehrlites. Orthopyroxene of mantle origin is found in xenoliths and as disaggregated, strongly reacted crystals in mafic lavas (Fig. S1c). Its composition, with Mg# = 87 (Ca1Mg86Fe13), broadly matches the composition of orthopyroxene in Madagascan mantle xenoliths (cf. Rocco et al. 2017; Mazzeo et al. 2021 and references therein).

Fe–Ti–Cr oxides: Cr–Al spinel is frequently hosted in the magmatic olivines, in the form of tiny idiomorphic crystals, which often show marked chemical zoning from cores to rims, with decrease of the Al concentration and increase of the magnetite + ulvöspinel components (Fig. 5). The Cr# (22–56) and Mg# (11–70) make them chemically distinct from spinels in the peridotite xenoliths found in the area (cf. Rocco et al. 2017). Cr–Al spinel of the wehrlites is richer in chromium than spinel in the gabbros (average Cr2O3 = 23 wt% vs 15 wt%; supplementary Table 6).

Fig. 5
figure 5

Chemical variation of chromium-bearing spinels in the mafic rocks of Bobaomby

Titaniferous magnetite is ubiquitous. The ulvöspinel component ranges from 26 to 82 mol%. The chemical composition of titaniferous magnetite is broadly constant throughout the lithotypes, with the exception of a higher concentration of MnO (up to 4.2 wt%) in some phonolites. Rare ilmenite (supplementary Table 7) is found in lamprophyre (M568) and a phonolite (M548) having quite similar composition (ilm96-98). The ilmenite-magnetite pair of lamprophyre (M568) equilibrated at 855–893 °C, and oxygen fugacity along the QFM buffer (− 13 to − 13.7 log units fO2). Titaniferous magnetite occurs in olivine from gabbros; its ulvöspinel composition ranges from 40 to 45 wt%.

Amphibole is found starting from tephrites, and varies in composition from pargasite, kaersutite and ferrokaersutite to rare alkali amphiboles in evolved lithologies or in groundmass of mafic lithologies (supplementary Table 8; Fig. S5). Mg# is always < 80, the petrography and the chemical composition indicate that amphibole is an abundant equilibrium phase of mildly evolved magmas. Nephelinites have very rare groundmass alkali amphiboles (supplementary Table 8).

Amphibole of intrusive nodules matches the composition of volcanic rocks: it is pargasite and kaersutite, with Mg# varying from 66 to 83 (Fig. S5). The concentration of TiO2 is variable, ranging from 3.1 to 5.1 wt%. The composition of amphibole in gabbros and ultramafic rocks fully overlaps that of basanites, tephrites and phonotephrites.

Mica is a rare phase, except for the lamprophyre M568. It varies from Ba-rich in the groundmass of the nephelinites (BaO up to 8.1 wt%) and some basanites (up to BaO = 6.5 wt%; sample M554) to Ba-poor compositions, including the gabbroic rocks; mica of the lamprophyre has BaO from 1 to 3 wt%. The TiO2 of mica in the nephelinites ranges from 2 to 9.6 wt%, similar to the range in the other lithologies, including the gabbroic xenoliths (TiO2 from 0.77 to 6.7 wt%) (supplementary Table 9). Mica in amphibole-wehrlites has Mg# from 77 to 55, low TiO2 (2.6–3.4 wt%) and relatively low and variable BaO (0.2–3.4 wt%) (supplementary Table 9).

Feldspar: Feldspar in basanites and tephrites varies from bytownite (An73) to oligoclase (An22), and the groundmass alkali feldspar is hyalophane and sanidine (BaO up to 7 wt%; SrO up to 2.5 wt%; Fig. 6a). The rare groundmass feldspar in the nephelinites is a Ba-rich sanidine (hyalophane, 7–10.5 wt% BaO and up to 2.3 wt% SrO), grading to K-rich sanidine (Fig. S6). Hyalophane has both oxides decreasing with increasing K2O (Fig. S6). Feldspars of the tephriphonolite M633 are sodic plagioclase (An36-28) and sodic sanidine (An2-3Ab32–48Or50–65) (Fig. 6b). Feldspar of the phonolites is a sanidine having a wide range of compositions (Or53–94), and trending to pure KAlSi3O8 (Fig. 6b). The concentration of Sr and Ba in the alkali feldspar of the phonolites is low, consistent with the low Sr and Ba of the host rocks. Anorthoclase is notably absent in the Bobaomby rocks: this is a major difference with the composition of alkali feldspar of Ankaratra and Itasy evolved rocks and the groundmass feldspar of the Ankililoaka basanites and alkali basalts, which is largely anorthoclase (e.g., Cucciniello et al. 2017, 2018; Melluso et al. 2018).

Fig. 6
figure 6

a, b Composition of plagioclase and alkali feldspar in the Bobaomby igneous rocks

Plagioclase of gabbros ranges from bytownite to andesine (An78–44), whereas plagioclase of wehrlites is labradorite (An66–50). (Fig. 6b, supplementary Table 10); rare alkali feldspar also occurs in both lithologies (Or35–57).

Feldspathoids: The groundmass nepheline of the nephelinites has low Si-excess (Ne60–76Ks9–17Sil2–12, in wt%; Fig. S7) and relatively high CaO (1.1–3.4 wt%; supplementary Table 11). On the other hand, nepheline of basanites, tephrites and phonolites has a significant Si-excess (Ne72–81Ks4–20Sil3–23), matching that of other silica undersaturated, feldspar-bearing rocks in Madagascar (e.g., Cucciniello et al. 2018; Melluso et al. 2018). The CaO in the nephelines of the basanite-phonolite sequence is lower than that in the nephelinites, though variable (0.03–2.63 wt%; supplementary Table 11).

Nepheline is also observed as minor phase in gabbro (M541A) and clinopyroxenite (M755N5).

Aenigmatite is a groundmass phase of the peralkaline phonolites. It is a solid solution of aenigmatite s.s. and limited wilkinsonite (TiO2 varies from 8.5 to 10.7 wt%, Fig. S8). The concentration of MnO from 1.3 to 2.6 wt%. MgO and Al2O3 are negligible, as expected from this phase in highly evolved peralkaline rocks. The composition of the aenigmatite of the Bobaomby phonolites is slightly different from that found in the Ankaratra and Massif d’Ambre trachytes and phonolites, being more Ti-rich and far lower in MnO that the compositions found in other trachytes and phonolites (Cucciniello et al. 2017; Lustrino et al. 2012; Melluso et al. 2014 and references therein; supplementary Table 12), and its rather high TiO2 indicates crystallization < 900 °C a relatively low fO2 (cf. White et al. 2005).

Apatite is a fluorapatite (F from 0.5 to 4 wt%), with significant Cl (from 0.16 to 1.2 wt%). The two volatile elements also show a rough negative correlation. The REE2O3 concentrations are low (La2O3 + Ce2O3 + Nd2O3 up to 2 wt%) (supplementary Table 13).

Titanite has La2O3 + Ce2O3 + Nd2O3 up to 2.7 wt%, variable, though low ZrO2, Nb2O5 and Al2O3, and is similar to the titanite compositions of the Itasy trachyphonolites (e.g., Melluso et al. 2018).

Other phases. Perovskite has been found in the lamprophyre as a corroded crystal within a phlogopite crystal (Fig. S1k). Its composition is nearly pure CaTiO3 (supplementary Table 14). The presence of this magmatic phase indicates that larnite-normative magmas were present in the feeder system of Bobaomby. Zirconolite is an accessory phase of sample M633. Sulfides are pyrite, chalcopyrite and galena. Rhönite is a typical secondary phase of mafic rocks after amphibole and has Mg# in a restricted range (Mg# = 34–58), with TiO2 from 7.3 to 10.1 wt%, similar to the rhönite around amphibole in the Massif d’Ambre, Ankaratra and Itasy. Secondary zeolites and carbonates are frequently found.

Geochemistry

Major and trace-element compositions of the Bobaomby rocks are reported in the supplementary Table 1. Most of the rocks studied are relatively fresh (only a few samples have evidence for post-magmatic alteration, as indicated by the occurrence of calcite and zeolites), consistent with petrographic observations. Loss on ignition (L.O.I.) ranges from 0.8 to 5.3 wt% (supplementary Table 1). The high L.O.I. values are often linked to the presence of the volatile-rich phases’ amphibole, mica and apatite. Some of these samples host mantle xenoliths and xenocrysts (cf. Rocco et al. 2017). The Na2O–K2O diagram of the mafic rocks (Fig. 2b) indicates major oxide heterogeneity in the composition of the parental magmas: the rocks indeed vary from sodic to potassic (K2O up to 3.68 wt% in the lamprophyre M568). The MgO of the Bobaomby basanites ranges from 6.4 to 13.3 wt% (Mg# = 55–70, where Mg# = molar Mg*100/(Mg + Fe)). Major and trace-element concentrations vs. MgO and vs. Zr (used as differentiation indices) are reported in Fig. 7. The SiO2, Al2O3, Na2O and K2O concentrations generally increase with decreasing MgO. TiO2, Fe2O3t, CaO and P2O5 concentrations decrease with MgO concentration < 6 wt%. The first-group transition elements such as Cr and Ni show considerable range (Cr from 1037 to 4 ppm; Ni from 1016 to 32 ppm). Co, Sc, and V decrease with increasing Zr, whereas Zn increases (Fig. 7; supplementary Table 1). The most primitive samples (without mantle-derived xenoliths) have Mg# (66–65), Cr (645–606 ppm) and Ni (382–262 ppm) not too far from those of mantle-derived primitive melts. The tephrites and the potassic lamprophyre M568 have MgO in the range 7.0–4.4 wt% and low concentration of compatible trace elements; TiO2 ranges from 2.8 to 2.5 wt%, Nb and Zr concentrations range from 160 to 84 ppm, from 403 to 233 ppm, respectively. The phonotephrite (M633) has Rb = 210 ppm, Sr = 704 ppm and Ba = 1156 ppm and Zr/Hf = 43. The phonolites have high Rb (437–270 ppm) whereas Sr (244–46 ppm), Ba (109–9 ppm), Co (7–2 ppm), Sc (< 8 ppm), V (< 10 ppm) and Cu (32–3 ppm) are low. The concentrations of Y (26–14 ppm), Zr (792–324 ppm), Nb (162–113 ppm) partially overlap those of evolved tephrites; the Zr/Hf ratios (40–38) are significantly lower than those of the mafic rocks. The amphibole gabbros (M541, M541A), wehrlites, and amphibole clinopyroxenites (M567A, M554, M554N1, M554N2) are characterized by variable MgO (17.8–7.9 wt%), TiO2 (4.7–1.5 wt%), Ni (1016–140 ppm), Cr (122–1037 ppm), Co (71–53 ppm), Sc (48–15 ppm). The concentrations of Rb (8 ppm), Sr (332 ppm), Ba (358 ppm) and Nb (18 ppm) are low in the amphibole clinopyroxenite M554N2.

Fig. 7
figure 7figure 7

Major (wt%) and trace-element (ppm) composition of Bobaomby igneous rocks plotted vs MgO (wt%) and Zr (ppm). The composition of Massif d’Ambre and Ampasindava peninsula are taken from Cucciniello et al. (2011, 2016); the Mayotte rocks are taken from Pelleter et al. (2014)

The variation of Zr with incompatible elements such as Nb and Rb (Fig. 7) highlights that the basanites have evidence of magma differentiation trends from different parental magmas. The ΣREE concentration of the Bobaomby rocks is higher in the basanites (up to 421 ppm) than in the more evolved rocks (down to 223 ppm in sample M633; Fig. 8; supplementary Table 2). The clinopyroxenite M554N2 has the lowest concentration of REE (124 ppm). The REE patterns of basanites, nephelinites and tephrites are smooth, parallel and heavily fractionated (Fig. 8) (La/Ybn = 27–20); the REE patterns become more concave upward passing from tephriphonolite to phonolite with a decrease of MREE and Eu/Eu* (Eu/Eu* = 0.98 in M633; Eu/Eu* = 0.74 in M572). The volatile element (F, Cl and S) concentration of the Bobaomby suite varies significantly in mafic and intermediate compositions (F = 205–1564 ppm; Cl = 488–1485 ppm; S = 190–650 ppm). The concentrations of these elements are higher in the phonolites (F = 5900 ppm; Cl = 4300 ppm).

Fig. 8
figure 8

Chondrite-normalized REE patterns of the Bobaomby igneous rocks. Chondrite normalizing values are from Anders and Grevesse (1989)

The available Nd–Sr–Pb isotope data of the Bobaomby volcanic rocks are reported in the supplementary Table 2. The basanites and the tephritic phonolite M633 have 87Sr/86Sr from 0.703350 to 0.703778 and 143Nd/144Nd from 0.512735 to 0.512778 (Fig. 9a). The olivine nephelinite M634 has 87Sr/86Sr = 0.703350 and 206Pb/204Pb = 19.396. The potassic lamprophyre M568 has 87Sr/86Sr = 0.70345; 143Nd/144Nd = 0.512848 and 206Pb/204Pb = 20.065, hence plots towards the isotope composition of the HIMU mantle end-member, a very unusual feature for such a potassic/ultrapotassic volcanic rock. Such high 206Pb/204Pb values are also observed at Mayotte (206Pb/204Pb = 20.36 in a nephelinite; Pelleter et al. 2014), Anjouan and Moheli (206Pb/204Pb = 20.023 in alkali basalt sample RH-15; Späth et al. 1996; Fig. 9b), but in no other analyzed Madagascan volcanic rock (cf. Melluso et al. 2016). The Pb isotopic data plot along the Northern Hemisphere Reference Line (Hart 1984) in the 207Pb/204Pb vs. 206Pb/204Pb diagrams (Fig. 9b) and to the right of the 4.55 Ga geochron.

Fig. 9
figure 9

a, b Sr–Nd–Pb isotopic compositions of Bobaomby igneous rocks. The Northern Hemisphere Reference Line (NHRL) and the geochron at 4.55 Ga are shown in the 206Pb/204Pb vs. 207Pb/204Pb isotope diagram. Data for the Cenozoic Madagascar igneous rocks are from Melluso and Morra (2000), Cucciniello et al. (2011, 2017, 2018) and Melluso et al. (2016, 2018). Indian mid-ocean ridge basalt (MORB) data are from Mahoney et al. (1992)

Pressure–temperature estimates

We used mineral (i.e., olivine, pyroxene, plagioclase and amphibole) and bulk-rock compositions to constrain the pressures (P) and temperatures (T) of magmatic crystallization. However, it should be noted that P–T data are strongly dependent on the calibration of the thermobarometers employed (all these methods rely on the assumption of mineral-liquid equilibrium), and, therefore, can provide only qualitative geological information. Olivine-liquid equilibration temperatures (according to Roeder and Emslie 1970) using the most Mg-rich olivines in the olivine nephelinites range from 1178 to 1218 °C, and those of basanites and tephrites range from 1136 to 1199 °C (supplementary Table 4). Temperature estimates using clinopyroxene-only thermobarometer (Putirka 2008) range from 973 to 1204 °C for the mafic rocks (olivine nephelinites and nephelinite M556, lamprophyre M568, basanites and tephrites), and decrease to 1019 °C for the phonotephrite M633 (supplementary Table 15). Pressure estimation for clinopyroxenes (Nimis and Ulmer 1998) ranges from 0.1 to 5.9 kbar (supplementary Table 5). The amphibole gabbro M541A has a range of crystallization temperatures between 1184 and 1229 °C. Pressure estimation (Nimis and Ulmer 1998) ranges from 0.1 to 7.1 kbar. Thermobarometers using clinopyroxene-liquid equilibria (Putirka 2008) provide temperatures of 1187–1233 °C for olivine nephelinites and nephelinite M556, 1111–1276 °C for basanites and tephrites (supplementary Table S15). The amphibole thermobarometer of Putirka (2016) indicates temperatures of 1019–1070 °C and pressures of 0.5–2.2 kbar for kaersutite and pargasite in tephrites and tephriphonolite M633. Plagioclase-liquid equilibration temperatures (Putirka 2008) for basanites and tephrites range from 1132 to 1159 °C. Equilibration temperature for the tephriphonolite M633 is ~ 1078 °C. Alkali feldspar-liquid equilibration temperatures (Putirka 2008) for phonolites range from 803 to 1030 °C (supplementary Table 15). Two feldspars temperature using the method of Fuhrman and Lindsley (1988) for basanites and tephrites range from 950 to 800 °C.

Discussion

The data set provided by mineralogy, petrology and geochemistry of lavas, dykes (this paper and Melluso et al. 2007), cumulitic xenoliths (this paper) and mantle-derived xenoliths (Rocco et al. 2017) allows to highlight magmatic processes which were active in northernmost Madagascar. Mineral assemblages and bulk-rock geochemistry point out that the Bobaomby rocks formed in an environment increasingly rich in H2O and F, with low Cl and S, able to stabilize amphibole and mica in intermediate stages (tephrites), and apatite and volatile-rich feldspathoids later on. Amphibole-bearing volcanic rocks with a similar compositional range (basanites to trachyphonolites) are found in the Itasy volcanic field, central Madagascar (Melluso et al. 2018). This indicates that the mantle sources of the Cenozoic alkaline volcanism in Madagascar are rich in trace elements and volatiles (H2O, F, S, and CO2) and, therefore, must be located in an enriched mantle lithosphere. The Bobaomby rocks are distinctly more silica undersaturated and Ti-rich than the bulk of the lavas of Massif d’Ambre, where also tholeiitic rocks and rhyolites do occur (Fig. 7), and similarly silica undersaturated than the Mayotte rocks (cf. Pelleter et al. 2014; Fig. 2), including the presently ongoing eruption (Berthod et al. 2021), and to the Nosy Be sodic basanites and tephrites (Melluso and Morra 2000).

The Zr/Hf of the mafic volcanic rocks (41–50, av. 44.5 ± 1.5) and the Nb/Ta (15.5–17, av. 15.8 ± 0.35) are constant and typical of mantle-derived rocks, and Zr and Hf concentrations are high, thus neglecting any evidence of carbonatitic metasomatism in the source. The low Zr/Nb ratios of the primitive rocks (Zr/Nb = 2.1–3) are typical of melts derived from highly enriched mantle (e.g., Minissale et al. 2022 and references therein). The primitive mantle-normalized diagrams of olivine nephelinites, basanites and tephrites have relatively smooth patterns with a bell-shaped trend, peaking at Nb–Ta, with small troughs at K and Ti, common features of mantle-derived continental and oceanic magmas with sodic affinity. These geochemical features are even common in mafic volcanic rocks cropping out in the Massif d’Ambre and Ampasindava (Cucciniello et al. 2011, 2016), pointing out that mantle sources of northern Madagascar volcanoes underwent similar enrichment processes and variable degree of partial melting (from 3 to 5% in Ampasindava and Massif d’Ambre basanites up to 10% in primitive Ampasindava tholeiites) of a HFSE (Nb + Ta + Ti)-enriched source (Cucciniello et al. 2011, 2016). Melluso et al. (2007, 2016) pointed out that the most mafic basanites of Bobaomby formed after low degrees of partial melting (< 4%) of an incompatible element-enriched mantle source, similarly to the results obtained on the most primitive lavas (basanites) of the Nosy Be archipelago. The pattern of lamprophyre M568 peaks at Nb and Ta, with a peak also at K and smoothly decreasing normalized abundances from Nb to Lu with a small trough at Ti. Again, the similarities of the incompatible element patterns and isotopic composition of the mafic lavas of Mayotte are remarkable (Figs. 9, 10).

Fig. 10
figure 10

Primitive mantle-normalized incompatible element patterns for the Bobaomby igneous rocks. Primitive mantle normalizing values are from Lyubetskaya and Korenaga (2007)

Quantitative evidence of magmatic evolution

The most primitive melts followed different evolutionary paths on the route to the surface. Forsterite-rich olivine and chromiferous spinel are the liquidus phases of basanites and olivine nephelinite, followed by diopside to titanaugite clinopyroxene; plagioclase, Ba- and Sr-rich feldspar, kaersutite and Fe–Ti oxides are in the groundmass of basanites. Other batches of magma formed in the crust and underwent fractional crystallization to generate more evolved rocks from tephrites to phonotephrites. At this stage, phases in basanite groundmasses (namely plagioclase, Ba- and Sr-rich feldspar, kaersutite and Fe–Ti oxides) appear as phenocrysts of tephrites and phonotephrites.

Tephrites have incompatible elements ratios similar to those of more primitive rocks. Quantitative models of major and trace-element evolution between primitive and evolved rocks are reported in supplementary Table 16. The transition from basanites (e.g., M574E, M551) to tephrites (e.g., M576) was modeled through the subtraction of ca. 51% of clinopyroxene-dominated solid assemblages, resembling wehrlites or metagabbros (supplementary Table 16). The occurrence of intrusive cognate nodules seem to support this model: crystal segregation formed a series of intrusive rocks ranging from wehrlite to amphibole-rich gabbro as complementary counterpart of the volcanic rocks, as shown by phase composition, crystallization order of the intrusive rocks and bulk-rock geochemistry. We remark that the presence and the composition of amphibole (kaersutite) cannot be considered as typical of early liquidus phases, even formed in the mantle (cf. Pilet et al. 2010); indeed it is found in intrusive rocks and as phenocrysts of mildly evolved mafic alkaline magmas.

Quantitative modeling indicates that the transition from tephrite (e.g., M576) to tephriphonolite (M633) took place through 56% subtraction of an assemblage of clinopyroxene, plagioclase, kaersutite, magnetite, olivine, apatite and titanite (in order of abundance), resembling the modal mineralogy of intrusive sample M554. The paucity of mildly evolved rocks (phonotephrites) at Bobaomby mark a Daly gap; this could be related to removal of kaersutite and Fe–Ti oxides, increasing the concentration of silica and alkalies (e.g., Brotzu et al. 2007; Melluso et al. 2007).

The magmatic environment of Bobaomby likely suffered some open-system recharge, as pointed out by highly Fe-rich cores with Mg-rich rims, with reverse and oscillatory zoning, particularly in the mafic rocks; it is important to remark that the tephriphonolite M633, the only intermediate dyke so far found, has no evidence of magma mixing, given the chemical composition of the phases and the overall mineral assemblage. The magmatic evolution from tephrites (M576) to phonolites (e.g., M631, M559) involved ca. 59% subtracted solids involving clinopyroxene, amphibole, sodic plagioclase, alkali feldspar and magnetite (supplementary Table 16). Therefore, it is possible to generate phonolitic magmas after a total of 80% subtraction of various solids from basanitic melt compositions. This amount of subtracted solid is also compatible with the geochemical behavior of the highly incompatible elements Th and U, suggesting values as high as 81% of subtracted solids (f, degree of liquid fraction ≈ Thbasanites/Thphonolites = 0.19). The Bobaomby phonolites plot on the low-pressure feldspar-nepheline cotectic near the phonolite minimum (Gupta 2015; Fig. S9), as also demonstrated by idiomorphic nepheline and sanidine phenocrysts, and are peralkaline at least in the groundmass, as deduced from the presence of late aegirine and aenigmatite. Another noteworthy feature of the Bobaomby phonolites is the total lack of anorthoclase feldspar, as typical of trachytic/phonolitic rocks of sodic affinity, or typical pantellerites and comendites (e.g., Ronga et al 2010 and references therein), rather than the Italian potassic rocks (e.g., Melluso et al. 2012, 2014 and references therein). This is not a function of the crystallization temperature (which should be significantly higher in the potassic sanidine than in anorthoclase), but, rather, of the chemical affinity of the evolved rocks. The phonolites of Bobaomby resemble those of Aris and Klinghardt Mountains, Namibia (Marsh 2010), the type 1 trachytes and phonolites of Gharyan, Libya (Lustrino et al. 2012) and those of Kenya (e.g., Brotzu et al. 1983; Aurisicchio et al. 1983), Atlas (Berger et al. 2009), Antarctica (e.g., Kim et al. 2019) or Germany (e.g., Abratis et al. 2015). Many of these rocks have titanite as a notable accessory phase. Excluding aegirine and aenigmatite, there are only a few phases indicating the transition to agpaitic conditions in the groundmass of the phonolites; this may be due to the relatively low concentration of Zr in the Bobaomby phonolites (300–800 ppm), relative to other trachytes and phonolites worldwide (e.g., Ischia, Gharyan, Ilimaussaq, Brazil, Itasy in Madagascar), that probably were not sufficient to promote crystallization of groundmass F-disilicates or eudialyte. In a regional volcanic context, the Bobaomby phonolites have broad similarity with those from Massif d’Ambre and Ampasindava region (Cucciniello et al. 2011, 2016), and match the composition of the Mayotte phonolites, even in the concave REE patterns (Fig. 8). On the other hand, the Bobaomby phonolites, at the same level of magmatic evolution, have less than half the Zr concentration of some Ampasindava peninsula eudialyte-bearing phonolites and rhyolites (1700–1800 ppm Zr; Cucciniello et al. 2016; Estrade et al. 2014), which clearly reached the F-rich agpaitic stage (e.g., Guarino et al. 2021 and references therein). Removal of small amounts of titanite generate the concave REE pattern of phonolites; it could mark the transition from tephri-phonolites to phonolites, as already noted in the Itasy volcanic complex or elsewhere (e.g., Melluso et al. 2014, 2018). Amphibole could be also involved in MREE fractionation; however, such a pronounced MREE depletion cannot be accommodated by amphibole removal, also considering that partition coefficients of titanite for MREE are one order of magnitude higher than those in amphibole (Olin and Wolff 2012); moreover, kaersutite strongly prefers HREE over MREE (e.g., Melluso et al. 2018 and references therein).

The Bobaomby phonolites decrease in total iron with increasing degree of magmatic evolution (Fig. 7), differently from the typical evolution of peralkaline oversaturated rocks such as pantellerites (e.g., Ronga et al. 2010, and references therein; Macdonald et al. 2012); this suggests that removal of Fe3+-rich phases (hence, relatively high oxygen fugacity) is significant for the magmatic evolution of silica-undersaturated rocks, as already shown elsewhere (e.g., Aurisicchio et al. 1983; Lustrino et al. 2012). A comprehensive section taking into account of the various igneous rocks of Bobaomby, from mantle xenoliths through intrusive rocks and from basanites and nephelinites to phonolites is reported in the Fig. 11.

Fig. 11
figure 11

Sketch model for the origin of the Bobaomby igneous rocks. It summarizes the mantle and crustal processes that influenced the magmatic evolution in the Bobaomby area: melting in the lower lithosphere of an enriched source gave rise to basanites and rare nephelinites, sometimes hosting spinel-lherzolites disrupted from the uppermost lithospheric mantle (e.g., Rocco et al. 2017). Emplacement in the crust of the magmas gave rise to crystallization of intrusive/cumulitic rocks, locally generating evolved tephrites, phonolitic tephrites and phonolites, finally poured out as lavas, dykes and pyroclastic rocks. The intrusive rocks are locally extracted by successive magma batches and found as xenoliths in the diatreme and plug facies. The thin section of the Madagascan basement is taken from Tucker et al. (2008). The figure is not to scale

Conclusions

A comprehensive study of the igneous rocks of the Bobaomby area provided the following constraints to the magma feeder system:

  1. (1)

    The age of the magmatism of the Bobaomby area (10–11 Ma) constrains uplift processes in this part of Madagascar, and broadly matches the age ranges of the central and southwestern volcanic rocks (Cucciniello et al. 2017, 2018).

  2. (2)

    Different types of alkaline magmas (olivine nephelinites/nephelinites, various basanite types, potassic lamprophyres) were erupted at the surface in the form of lavas, dykes and pyroclastic rocks; mantle-derived debris (xenoliths, xenocrysts) is observed in the primitive samples. Particularly interesting is the presence of potassic and nearly ultrapotassic volcanic rocks, indicating independent magma batches, expected in such a scattered monogenetic lava field.

  3. (3)

    Some of the basanitic magmas had the time to differentiate in ephemeral crustal reservoirs towards peralkaline phonolites, through prolonged removal of olivine, clinopyroxene, spinel, amphibole, and feldspar, together with minor accessory phases such as apatite and titanite. The magmas reached liquid compositions with cotectic potassic sanidine and nepheline, and allowing crystallization of groundmass aegirine and aenigmatite in a peralkaline residual system; the solid assemblages of the intrusive xenoliths broadly match the subtracted mineral assemblages at various steps of fractional crystallization of basanites and tephrites; there is no mineralogical, geochemical or geological trace of evidence for phonolites being mantle-derived melts in the same way as are the Mg-rich basanites and olivine nephelinites.

  4. (4)

    The sources of the Bobaomby primary magmas are heterogeneous, and rich in hydrous phases such as amphibole and/or phlogopite, and, possibly, carbonates, in a mantle volume where these phases are stable along the peridotite solidus, hence with geothermal gradients compatible with the deepest parts of the continental lithospheric mantle in northern Madagascar. There is no trace of carbonatitic metasomatism, nor evidence of subduction-related old enrichment events.

  5. (5)

    The compositional characteristics of the Bobaomby volcanic rocks, from the mafic lithotypes to the phonolites, closely match those of the Mayotte rocks, implying similar genesis, mantle source and subsequent crystallization history. The effects of titanite removal at Bobaomby (and Mayotte, or elsewhere in Madagascar) are easily visible in the HFSE and REE patterns of more evolved phonolites.

Analytical techniques

The quantitative classification of intrusive xenoliths was determined on the basis of the modal abundance of the minerals (supplementary Table 1) by image-analysis software Leica QwinPlus. New bulk rock major and trace-element analyses of the Bobaomby rocks (supplementary Table 1) have been obtained (on pressed powder pellets) by X-ray fluorescence XRF at University of Naples (Dipartimento di Scienze della Terra, dell’Ambiente e delle Risorse) utilizing an Axios Panalytical instrument, equipped with six analyser crystals, three primary collimators, and two detectors (flow counter and scintillator), operating at different kV and mA for each analyte with techniques already described in Guarino et al. (2021) and by ICP-MS at Actlabs (cf. www.actlabs.com for analytical procedures, and supplementary Tables 1 and 2 for reference standards). Further analytical details can be found in Guarino et al. (2021).

The mineral compositions were obtained at DiSTAR, University of Napoli Federico II, utilizing a JEOL JSM-5310 microscope and an Oxford Instruments Microanalysis Unit, equipped with an INCA X-act detector and operating at 15 kV primary beam voltage, 50–100 mA filament current, variable spot size and 50 s net acquisition time. Measurements were done with an INCA X-stream pulse processor. The following standards were used for calibration: diopside (Mg), wollastonite (Ca), anorthoclase (Al, Si), albite (Na), rutile (Ti), almandine (Fe), Cr2O3 (Cr), rhodonite (Mn), orthoclase (K), apatite (P), fluorite (F), barite (Ba), strontianite (Sr), zircon (Zr, Hf), synthetic Smithsonian orthophosphates (REE, Y, Sc), pure vanadium, niobium and tantalum (V, Nb, Ta), Corning glass (Th and U), sphalerite (S, Zn), sodium chloride (Cl). The Kα, Lα, Lβ, or Mα lines are used for calibration, according to either element (standards in supplementary Table 17). Backscattered electron (BSE) images were obtained with the same instrument (cf. Melluso et al. 2014 and Guarino et al. 2021 for further details).

Sr and Nd isotope analyses were obtained at DiSTAR and the Istituto Nazionale di Geofisica e Vulcanologia (Osservatorio Vesuviano), Naples. Sample powders were dissolved in a HF–HNO3–HCl mixture. Sr and Nd fractions were separated following standard chromatographic techniques. The total procedural blank at Osservatorio Vesuviano was < 220 pg for Sr and < 100 pg for Nd, making blank correction negligible. Mass spectrometric analyses were performed on a Thermo Finnigan Triton-Ti® thermal ionization mass spectrometer equipped with nine movable faraday cups.

Pb isotope analyses were performed at the School of Ocean and Earth Science and Technology, University of Hawaii. Sample chips were leached with HF and dissolved successively in a HF–HNO3–HCl mixture. Pb was then separated into both spiked and unspiked splits in a two-step elution with mixed HBr–HNO3 solutions on 100 μl anion exchange columns. Mass spectrometric measurements were performed on a VG Sector multicollector thermal ionization mass spectrometer. Total procedural blank for Pb was < 19 pg. Further details can be found in Cucciniello et al. (2013). The analytical data obtained at the Osservatorio Vesuviano follow those already described by Mazzeo et al. (2021).