For over a century, the classic paradigm of volcanology and igneous petrology has been premised upon the existence of magma chambers, filled by crystal-free melt, forming ‘big tanks’1,2,3,4,5,6,7,8,9,10. Such magma chambers gradually lose heat and crystallize from all margins inwards and occasionally supply overlying extrusive centres (volcanoes or fissures) with magma that erupts onto the Earth’s surface1,2,3,4,5,6,7,8,9,10. This founding concept has, however, been recently challenged on the basis of observations and evidence from various disciplines. The most often-cited evidence is derived from geophysical surveys that are unable to conclusively identify any present-day magma chambers with large volumes of eruptible melt11. This is supported by thermal modelling12, which indicates that the formation of a large magma body within the upper crust is physically problematic, because it requires a magma accumulation rate that is 1–2 orders of magnitude greater than determined through geochronology11. In addition, out-of-sequence geochronology has been used to argue that the known mafic–ultramafic plutons do not require the existence of large magma chambers but could be produced as a stack of randomly-emplaced amalgamated sills13,14,15. The conclusion from these studies is that large, predominantly molten magma chambers are likely either transient16 or non-existent11,12 in the geological history of the Earth. As an alternative, it is proposed that intracrustal melt is stored within intergranular pockets of crystal-rich mushes that occupy almost the entire crust, from the Moho towards surface16,17,18,19,20,21. Periodic tectonic destabilization of the mush may produce small, discrete melt lenses that subsequently aggregate, ascend and erupt as lava. There are, however, some observations from magmatic complexes22,23,24 as well as thermal modelling constraints12 that conflict with this emerging paradigm16,17,18,19,20,21. Here we present one well-constrained example from the Bushveld Complex, indicating that the magma chamber appears to have contained, during one stage of its evolution, an enormous volume of resident melt that slowly crystallized from the base upwards to produce a continuous sequence of chemically stratified cumulate rocks.

Incremental growth of the Bushveld Complex

The 2.05 billion-year-old Bushveld Complex in South Africa is the largest mafic–ultramafic layered intrusion into the Earth’s crust. It occupies an area that most likely exceeds 100,000 km2 and extends ~ 450 km east–west and ~ 350 km north-south25,26,27,28. Despite its enormous size, this complex is merely the remaining portion of an originally much larger intrusion that has subsequently been eroded to an unknown extent by surface processes. The complex consists of several parts, of which the western, eastern and northern limbs are the largest, and is stratigraphically subdivided into five major units—the Marginal, Lower, Critical, Main, and Upper Zones, comprising a total thickness of about 7 to 9 km25. The Bushveld Complex is widely considered to be a typical example of an open-system magma chamber29. Apart from the marginal rocks, its four principal zones are attributed to major replenishing events, with numerous smaller magma recharges contributing to the formation of these zones. During this process, the magma chamber incrementally increased in size through both vertical and lateral inflation29,30,31. All major replenishing events are marked by regionally extensive magmatic disconformities, local erosive unconformities into previous strata, significant isotopic shifts, and notable changes in whole-rock and mineral compositions29,32,33,34. These relationships are best exemplified by the Main Zone (MZ) at the Tonteldoos area of the southeastern Bushveld Complex (Fig. 1), which has a much larger lateral extent than the underlying Critical Zone (CZ), as indicated by the MZ’s direct onlapping of the floor rocks in many places. The MZ is commonly attributed to a large influx of new magma that significantly expanded the chamber in both vertical and lateral extent, producing a regional disconformity29 and locally prominent unconformities with pre-existing CZ cumulates32,35. The base of the MZ is also marked by a substantial isotopic shift towards more radiogenic whole-rock 87Sr/86Sr ratios36. The laterally extensive Pyroxenite Marker in the uppermost part of the MZ indicates another major expansion of the chamber through a magma replenishment and mixing event that caused an isotopic shift towards less radiogenic whole-rock 87Sr/86Sr ratios37,38 and an increase in both the An-content of plagioclase and Mg-number of pyroxenes39,40,41,42. Successive crystallisation of the entire Bushveld Complex from the Lower and Upper Critical Zone to the Main Zone occurred within ca. 1 million years, between 2055.91 ± 0.26 Ma and 2054.89 ± 0.37 Ma43. Attempts to place stricter constraints on the crystallisation of the complex using high-precision U–Pb TIMS ages13,15 have proved problematic because zircon isotopic ages in these studies appear to be at odds with basic field relationships44.

Figure 1
figure 1

The geological map of the southeastern part of the Bushveld Complex in the Tonteldoos area. The complex transgresses upwards through the Transvaal Supergroup over the northern 35 km of this sector and the floor of the complex steps up by approximately 6 km. The basal part of the Main Zone has three continuous markers layers termed the Lower, Middle, and Upper Mottled Anorthosites that extend along the entire area. The position of cross-sections in Figs. 2 and 6 and seven major traverses across Anorthosite Markers in Figs. 3, 4, 5 (see also Supplementary Data) are indicated. Regional geology digitised from the 1:250,000 scale geology maps 2528 Pretoria and 2530 Barberton (Geological Survey of South Africa), with additional detail from mapping by Van der Merwe49, Bevington and Hornsey (2010, Nuplats Ltd, unpublished mapping), and Latypov and Chistyakova (2021, unpublished mapping). The figure was prepared by Richard Hornsey using Micromine 2021 Release 21.5.

The resident melt column of the Bushveld Complex

The stratigraphy of the Bushveld Complex is most commonly thought to have progressively accumulated from the overlying resident melt by deposition of crystals on the chamber floor6,29,45 although there are several alternative views13,15,46. A critical unknown parameter is the volume of resident melt at any particular time during the evolution of the magma chamber26,47. We present here a potential solution through field mapping of the southeastern Bushveld Complex in the Tonteldoos area (Fig. 2), where the resident melt column thickness at the time of MZ crystallization may be assessed. The field mapping and 1:250,000 scale regional geology maps of the study area are compiled into a 3D model that includes and respects all mapped relationships between various geological units. The along-strike section in Fig. 2 is viewed down the dip of the stratigraphy and thereby preserves all the relationships between the intrusion and its host rocks, without loss of any details. It reveals that the host stratigraphy along this section dips at 11° to the west. The sedimentary host rock sequence was intruded by precursor Bushveld Complex sills that are sharply truncated by the main plutonic phase of the Bushveld Complex. The mapped geology exposes no evidence for any major syn-magmatic (e.g., the floor subsidence) or post-emplacement (e.g., fault displacement) structural deformation of the country rocks (Fig. 2), implying that the section was emplaced as shown and preserves all primary igneous field relationships.

Figure 2
figure 2

Geological along-strike section of the southeastern part of the Bushveld Complex in the Tonteldoos area. (a) The section is of the transect line AB in Fig. 1, looking at − 11° towards 270° azimuth. The section was constructed by rotating the 3D model to view the detailed geology of the complex and its immediate footwall in their true orientation. The section is not vertically or horizontally exaggerated and shows the true lithology morphology. (b) The schematic section prior to faulting highlights several important features of the complex: (1) the existence of the Roossenekal and Belfast sub-chambers with the intervening Stoffberg remnant of non-deformed host rocks, (2) the ~ 4-km and ~ 6-km thick vertical distance between the summit of the Stoffberg remnant with the Lower Mottled Anorthosite and the floor of the Roossenekal sub-chamber, respectively, (3) the concave geometry of the Pyroxenite Marker indicating that instantaneous top of the cumulate pile in the two sub-chambers was gently basinal with regards to the Stoffberg remnant. LMA, Lower Mottled Anorthosite; MMA, Middle Mottled Anorthosite; UMA, Upper Mottled Anorthosite; PM, Pyroxenite Marker; MML, Main Magnetite Layer. The CD line in (a) shows the location of the cross-section in Fig. 6. The figure was prepared by Richard Hornsey using Micromine 2021 Release 21.5.

Two notable features illustrated by Fig. 2 provide crucial constraints. The first relates to the geometry of the chamber that has been previously subdivided and described as two large basinal structures39. We refer to them, however, as the Roossenekal and Belfast sub-chambers because they lack any associated structural deformation or down-warping of their sedimentary host rocks (Fig. 2). The floors to both sub-chambers are concave upwards and juxtapose an intervening Stoffberg remnant of non-deformed host rocks. The floor contact of the Roossenekal sub-chamber has an impressive ~ 6-km-high shelf relief (termed the Tonteldoos step; Fig. 2b) with a ~ 10° slope across ~ 35 km up to the summit of the Stoffberg remnant. This topographic relief of the floor contact constrains a sub-chamber that was emplaced along an angular discordance to its host rock stratigraphy48. From north to south, the floor of the Roossenekal sub-chamber initially overlies the Steenkampsberg Formation, then transgresses the Houtenbek Formation and finally onlaps the Dullstroom Formation. The Roossenekal sub-chamber thus attains its maximum thickness in the north and thins towards the south where the CZ eventually terminates against the intrusion floor. The MZ and UZ extend across the entire extent of the two sub-chambers. The MZ in the Roossenekal sub-chamber directly onlaps onto the sedimentary and volcanic floor rocks (Fig. 2).

The second feature relates to the igneous layering, defined by prominent layers termed the Lower, Middle and Upper Mottled Anorthosites (LMA, MMA and UMA), which occur close to the base of the MZ49. Although it is widely assumed that the MZ layering abuts the floor contact of the country rocks37,39,40, the Anorthosite Markers do not terminate but drape across the entire Roossenekal and Belfast sub-chambers, including the Stoffberg remnant49. The Anorthosite Markers have been mapped by the authors along strike at seven field traverses through the basal part of the MZ in both sub-chambers (Figs. 1, 3, 4, 5; Extended Data Fig. 1). Field mapping indicates that from north to south the Anorthosite Markers tend to decrease in thickness, without any systematic changes in their textures. Importantly, the presence of the Anorthosite Markers across the entire area indicates that the deposition of MZ cumulates occurred synchronously across the deepest (i.e., Roossenekal sub-chamber base) and shallowest (i.e., Stoffberg remnant summit) parts of the sub-chamber (Fig. 2). This occurred despite the elevation difference between the two contrasting depositional places being ~ 4 km (Fig. 2b). This relationship cannot be due to syn- or post-emplacement subsidence of the Roossenekal sub-chamber and its host stratigraphy because the local and regional geology shows no evidence for the depression of the floor rocks, either through faulting or magma emplacement (Figs. 1 and 2).

Figure 3
figure 3

Photos of field outcrops of the three Anorthosite Markers along the traverse I-I of the Roossenekal sub-chamber of the southeastern part of the Bushveld Complex. All photos were taken by Sofya Chistyakova or Rais Latypov.

Figure 4
figure 4

Photos of field outcrops of the three Anorthosite Markers along the traverse IV-IV at the Tonteldoos step of the Roossenekal sub-chamber of the southeastern part of the Bushveld Complex. All photos were taken by Sofya Chistyakova.

Figure 5
figure 5

Photos of field outcrops of the three Anorthosite Markers along the traverse VI-VI of the Belfast sub-chamber of the southeastern part of the Bushveld Complex. All photos were taken by Sofya Chistyakova.

Igneous layering in mafic intrusions results from deposition of crystals from the overlying resident melt due to gravity settling4,50,51,52 or in situ crystallization9,53,54,55,56 onto the chamber floor. This indicates that in order to blanket the topographic relief of a temporary floor of the Bushveld chamber with igneous layering (i.e., LMA), the resident melt column must have been thicker than the ~ 4.0 km height of the Tonteldoos step. The lithological interpretation is supported by a systematic decrease in An-content of plagioclase and Mg-number of orthopyroxene through the ~ 3.0-km-thick MZ stratigraphy of the Roossenekal sub-chamber, indicating internal differentiation of a resident melt column that was thicker than the crystallised sequence (Fig. 6). The transition to the overlying Pyroxenite Marker is defined by an up to 0.5-km-thick reversal towards more primitive mineral composition (Fig. 6), which has been interpreted to result from mixing of a residual MZ melt with new magma entering the chamber29,40,42,57,58, causing further vertical expansion. Mass balance calculations based on Sr-isotopic data indicate that the residual melt comprised 60–70% of the resulting hybrid magma29,37, which subsequently crystallized to form a > 3.0 km thick sequence overlying the Pyroxenite Marker (Fig. 2). If correct, the residual melt of the MZ must still have been ~ 2-km-thick prior to the Pyroxenite Marker magma influx, thereby indicating an initial ~ 5 km thickness of the MZ melt column; consistent to earlier estimates based upon thermal modelling of the Bushveld Complex26. We concur with previous studies39,59 that the instantaneous top of the cumulus pile during deposition in this region was gently basinal, with the Stoffberg remnant partially separating the Roossenekal and Belfast sub-chambers. This is best indicated by the concave geometry of the Pyroxenite Marker within the Roossenekal sub-chamber, resulting in this layer being almost 2 km stratigraphically lower at the centre of this sub-chamber, compared to the Stoffberg remnant (Fig. 2). This field evidence implies that the ~ 1.0 km and ~ 3.0 km thick MZ in the Roossenekal and Belfast sub-chambers, respectively, formed synchronously from the same interconnected resident magma, further substantiated by similar An-content of plagioclase at the base of the MZ at both sub-chambers (67.5%39,60 and 71%40,41, respectively). The greater thickness of MZ cumulates in the Roossenekal sub-chamber may be related to the greater thickness of this unit, due to redeposition of crystals inside this depression and/or prevailing crystallization in the deeper parts of the magma chamber, due to pressure-induced increase in the liquidus temperature of the melt9,54,61.

Figure 6
figure 6

Modal abundances, cumulate stratigraphy and compositional variations of minerals through the Main Zone in the Tonteldoos area of the southeastern part of the Bushveld Complex. The section is approximately along the transect line CD in Figs. 1 and 2. The lower stratigraphy is characterized by a systematic crystallization sequence Opx + Aug + Pl (A), Opx + Aug + Pig + Pl (B), and Aug + Pig + Pl (C). The uppermost stratigraphy is marked by re-appearance of cumulus Opx at the Pyroxenite Marker so that the crystallization sequence Opx + Aug + Pl (A′) and Opx + Aug + Pig + Pl (B′) is repeated. The lower stratigraphy shows systematic evolutionary trends in composition of plagioclase (decrease in An-content, 100Ca/(Ca + Na)) and orthopyroxene (decrease in Mg-number, 100 Mg/(Mg + Fe)). The transition to the Pyroxenite Marker is characterized by a gradual compositional reversal towards higher An-content of plagioclase and Mg-number of orthopyroxene. Mg-number in subzone C is for orthopyroxene that is produced by inversion of primary pigeonite. All compositional data in (a–e) are from Gruenewaldt41 and are summarized in Supplementary Data Table 1. Pl, plagioclase; Opx, orthopyroxene; Aug, augite; Pig#, inverted pigeonite. The figure was prepared by Sofya Chistyakova using CorelDRAW (version

Rather than implying that the entire ~ 5.0 km thick column formed from a single large magma influx, it is proposed that the intrusion progressively grew to its final size (from the Roossenekal sub-chamber towards the Belfast sub-chamber) through emplacement of numerous magma influxes, yet over a much shorter time scale than solidification. During this period of repeated injections, each replenishment effectively mixed with the resident melt in the chamber, thereby delaying or impeding the onset of crystallization. Thus, crystallization commenced within a completely filled, large and homogenized magma chamber, which can thereby be modelled as having crystallized as a ‘single pulse’ of magma (Fig. 7a). Our model is therefore substantially different from those where the melt crystallization and cumulate pile growth of the MZ occurred concurrent to magma chamber replenishment37. Our model conforms better to liquidus phase equilibria predictions for a basaltic parent, such as systematic changes in pyroxene assemblages (e.g., Opx-Aug through Opx-Aug-Pig to Aug-Pig)62 and continuous decreases in both the An-content and Mg-number of cumulus plagioclase and orthopyroxene, respectively (Figs. 6 and 7b). Both the crystallization sequence and mineral compositional trends are reproduced by fractional crystallization of the parental melt (Extended Data, Extended Data Fig. 2) using the alphaMELTS software. It should be noted, however, that the modelling results are not unique (e.g., they are greatly dependent on the choice of a parental melt composition) and cannot, therefore, be considered as solid evidence of the model developed in this study. Such unidirectional evolutionary trends are also inconsistent with models that attribute the formation of the MZ to externally-derived crystal-rich slurries, because these would result in either constant63 or irregular64 trends in mineral compositions through the MZ stratigraphy. It is quite conceivable, however, that during protracted fractionation of the MZ resident melt, the chamber may have been further replenished by additional minor magma pulses with or without phenocrysts since minor local reversals in the An-content of plagioclase are discernable within the overall decreasing trend (Fig. 6). In fact, the Anorthosite Markers themselves may be associated with magma chamber replenishments65. Such occasional and relatively small magma chamber recharges do not, however, modify our major conclusion with respect to the initial resident melt column thickness, prior to the onset of crystallization at the first anorthosite marker (i.e., LMA in Fig. 2). It was only after a protracted and relatively quiet period of continuous crystallization of the MZ, that a major influx of orthopyroxene-saturated magma incrementally mixed with the magma chamber’s resident melt, while concomitantly crystallizing a succession of cumulates both below and above the Pyroxenite Marker (see Thermodynamic modeling in Extended Data).

Figure 7
figure 7

Model for the proposed crystallization history of the Main Zone in the Tonteldoos area of the southeastern part of the Bushveld Complex. (a) At onset of crystallization, the MZ resident melt column had a total thickness of about 5 km. This resulted in simultaneous deposition of the Lower Mottled Anorthosite along the entire extent of the Roossenekal and Belfast sub-chambers, including the Tonteldoos step and Stoffberg remnant. (b) Internal crystallization and differentiation of the MZ resident melt from the base upwards produced a continuous sequence of chemically-stratified cumulate rocks. The instantaneous top of the cumulate pile in the chamber was gently basinal with the Stoffberg remnant separating two sub-chambers. (c) Mixing of the MZ residual melt with a new magma influx resulted in the formation of the MZ hybrid melt that produced the laterally extensive Pyroxenite Marker with an associated reversal in mineral compositions. The total volume of resident melt before onset of crystallization is estimated at > 380,000 km3 which allows to classify the Bushveld chamber as a ‘big tank’ open-system within the Earth’s crust. The figure is not to scale and is for illustrative purposes. The figure was prepared by Rais Latypov using CorelDRAW (version

Implications for the ‘big tank’ magma chamber paradigm

The 5-km-thick resident melt column that produced the chemically stratified MZ cumulate sequence, including its three prominent Anorthosite Markers, indicates that during this stage the Bushveld chamber was exceptionally large and entirely molten (Fig. 7). The formation of such a cumulate sequence, at a typical solidification rate for large mafic intrusions (~ 1 cm/year50,66) would take ~ 300,000 years indicating that the intrusion was also long-lived. These conclusions are at odds with an emerging paradigm that such ‘big tank’ magma chambers were ephemeral at any given time throughout Earth’s geological history16,19,20,67,68. The total volume of resident MZ melt may be estimated as follows. The MZ melt column varied between ~ 5.0 km in the thicker and, at least, 1.0 km thick in the thinner parts of the intrusion. Based on reconstructions of the lateral extent of the MZ29, it is estimated that the thicker areas occupied ~ 70% of the MZ, and the remaining ~ 30% were thinner zones. The estimated Bushveld Complex area is approximately 100,000 km3,26,28 and therefore the total volume of the MZ resident melt is ~ 380,000 km3 (5 km * 70,000 km2 + 1.0 km * 30,000 km2). This volume is several orders of magnitude larger than the largest ignimbrite/tuff super-eruptions in Earth’s history (e.g., Bishop tuff—600 km3 and Youngest Toba eruption—up to 13,200 km3)69. It is only comparable to estimates of some of Earth’s large igneous provinces, such as the Karoo (367,000 km3)70 and Afar (350,000 km3)71. Thus, during emplacement of the MZ, the Bushveld magma chamber was a repository of an enormous volume of resident melt and may be regarded as a ‘big tank’ open-system within the Earth’s crust. Therefore, the current tendency in modern volcanology16,19,20 and petrology13,14,15 to discard the existence of such large and molten magma chambers4,5,6,7 appears to be premature. There is also no compelling reason to believe that ‘big tank’ magma chambers, such as the 2.05 Ga Bushveld Complex, are restricted to a long-forgotten past of our planet, such as the Precambrian. This is indicated by the 55 Ma Skaergaard intrusion in Greenland whose spectacular chemical stratigraphy indicates that before the onset of crystallization it was a ‘big tank’ of crystal-free tholeiitic parent magma up to 4 km in thickness and up to 300 km3 in volume72,73. Another example is the 1.3 Ga Kiglapait intrusion in Labrador with 3500 km3 of magma in a > 8 km thick magma chamber that shows a continuous differentiation sequence with little or no magma recharge74,75. It is therefore conceivable that such magma chambers have developed throughout the entire Earth’s evolution. Even if some regions of the Earth’s crust may behave as giant crystal mushes (e.g., mid-ocean ridges or deep roots of continental arcs)16,17,20,76, this does not automatically imply that ‘big tank’ magma chambers are absent from other regions (e.g., stable cratons with layered intrusions)5,6,7. Moreover, since layered intrusions such as the Bushveld Complex are rare throughout geological time77, it is not surprising that there are currently no active examples of large and molten magma chambers in Earth’s crust which can be detected geophysically11.


Map and cross-sections constructions

The maps were created by digitally capturing the 1:250,000 geological survey maps (2528 Pretoria, and 2530 Barberton, Geological Survey of South Africa) into Micromine software ( Additional details come from mapping by Van der Merwe49, Bevington and Hornsey (2010, Nuplats Ltd, unpublished mapping), and Latypov and Chistyakova (2021, unpublished mapping). Micromine is a commercially available geological software package widely used for exploration and mine planning to model geological datasets in 3d space. The version used for the latest iteration of the model was Micromine 2021 Release 21.5. The maps were digitally captured as polygons, which were then draped onto a digital terrain model downloaded from the USGS Earth Explorer website ( The terrain data were used to create a digital terrain model wireframe (DTM) that was cut into separate entities using the geological polygons and attributed according to the lithological unit. To create the geological section, the geological model was rotated using Micromine to enable visualisation of the data as a section looking down the plunge of the Bushveld Complex. The detailed geology was then manually digitised from the 3-dimensional section as a planar section that was then attributed to show the lithological units using the same colour scheme as for the plan. The imagery was exported from Micromine as formatted plans and sections. These were then imported into MSPowerpoint software, within which the labelling and legends were added. The final product was then exported as an image for inclusion into the research paper. The geological section therefore illustrates the entire stratigraphy and relationships between the Bushveld Complex and its host stratigraphy showing the “as-mapped” relationships and geometries in their correct spatial location. The maps are all in WGS84 Datum, UTM zone 36 South.