A 5-km-thick reservoir with &gt;380,000 km3 of magma within the ancient Earth's crust

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A 5-km-thick reservoir with >380,000 km³ of magma within the ancient Earth's crust

https://doi.org/10.21203/rs.3.rs-1104213/v1

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Several recent studies have argued that large, long-lived and molten magma chambers^1–10 may not occur in the shallow Earth’s crust^11–23. Here we present, however, field-based observations from the Bushveld Complex²⁴ that provide evidence to the contrary. In the eastern part of the complex, the magmatic layering was found to continuously drape across a ~4-km-high sloping step in the chamber floor. Such deposition of magmatic layering implies that the resident melt column was thicker than the stepped relief of the chamber floor. Prolonged internal differentiation within such a thick magma column is further supported by evolutionary trends in crystallization sequence and mineral compositions through the sequence. The resident melt column in the Bushveld chamber during this period is estimated to be >5-km-high in thickness and >380,000 km³ in volume. This amount of magma is three orders of magnitude larger than any known super-eruptions in the Earth’s history²⁵ and is only comparable to the extrusive volumes of some of Earth’s large igneous provinces²⁶. This suggests that super-large, entirely molten and long-lived magma chambers, at least occasionally, occur in the geological history of our planet. Therefore, the classical view of magma chambers as ‘big magma tanks’^1–10 remains a viable research concept for some of Earth’s magmatic provinces.

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, like ‘big tanks’^1–10. Such magma chambers gradually lose heat and crystallize from all margins inwards and occasionally supply overlying extrusive centres (volcanoes or fissures) with melts that erupt onto the Earth’s surface^1–10. This founding concept has been, however, 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 melt²⁷. This is supported by thermal modelling, 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 geochronology²⁷. In addition, out-of-sequence geochronology was used to argue that the mafic-ultramafic plutons do not require the existence of large magma chambers but сan rather be produced as a stack of randomly-emplaced and amalgamated sills^19,20,23. The conclusion from these studies is that large and mostly molten magma chambers are likely either transient¹⁴ or non-existent^11,12 in the geological history of Earth. As an alternative, it is proposed that all the melt is stored within intergranular pockets of crystal-rich mushes that occupy almost the entire crust, from its Moho towards the surface^{13–15,17,18,28}. The 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 complexes²⁹ and thermal modelling constraints³⁰ that come into conflict with this emerging paradigm^{13–15,17,18,28}. Here we present one clear-cut case from the Bushveld Complex whose 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-south^24,31–33. 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 km²⁴. The Bushveld Complex is widely considered to be a typical example of an open-system magma chamber³⁴. 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 inflation^34–36. 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 compositions³⁴. 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 its 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 disconformity³⁴ and locally prominent unconformities with pre-existing CZ cumulates³⁷. The base of the MZ is also marked by a substantial isotopic shift towards more radiogenic whole-rock ⁸⁷Sr/⁸⁶Sr ratios³⁸. 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 ratios^39,40 and an increase in both the An-content of plagioclase and Mg-number of pyroxenes^41–44.

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 floor^6,34,45 althought there are some alternative views^19,21,23. A critical unknown parameter is the volume of resident melt at any particular time during the evolution of the magma chamber^31,46. We present here a potential solution to this issue 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 10° to the west. The sedimentary host rock sequence was injected by Bushveld-related sills that are sharply truncated by the Bushveld Complex. The mapped geology exposes no evidence for any post-emplacement structural deformations, implying that the section was emplaced as shown and preserves all primary igneous field relationships.

Two notable features in Fig. 2 provide crucial constraints. The first relates to the geometry of the chamber that has been previously subdivided and called as two large basinal structures⁴¹. We refer to them here, 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 its floor contact reflects a sub-chamber that was emplaced along an angular discordance to its host rock stratigraphy⁴⁷. From north to south, the floor of the Roossenekal sub-chamber first overlies the Steenkampsberg Formation, then transgresses the Houtenbek Formation and finally onlaps the Dullstroom Basalt 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 MZ⁴⁸. Although it is widely assumed that the MZ layering abuts the floor contact of country rocks^39,41,42, the Anorthosite Markers do not terminate but drape across the entire Roossenekal and Belfast sub-chambers, including the Stoffberg remnant⁴⁸. The Anorthosite Markers have been traced by the authors along strike at seven field traverses through the basal part of the MZ in both sub-chambers (Fig. 1; Extended Data Figs. 1-4). 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 occured despite the elevational distance between these two contrasting depositional places being ~4 km (Fig. 2b). It should be noted that this relationship cannot be due to post-emplacement deformation of the host stratigraphy because the local and regional geology shows no evidence for post-emplacement structural modification, either through faulting or magma emplacement.

Igneous layering in mafic intrusions results from deposition of crystals from the overlying resident melt due to gravity settling^4,49,50 or in situ crystallization^9,51,52 onto the chamber floor. This means 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 interpretation is supported by a systematic decrease in An-content of plagioclase and Mg-number of orthopyroxene through a ~3.0-km-thick MZ stratigraphy within the Roossenekal sub-chamber, indicating internal differentiation of a resident melt column that was thicker than the crystallised sequence (Fig. 3). The transition to the overlying Pyroxenite Marker is defined by an up to 0.5-km-thick reversal towards more primitive mineral composition (Fig. 3), which has been interpreted as resulting from the mixing of a residual MZ melt with new magma entering the chamber^{34,42,44,53,54} and thereby further vertical expansion. Mass balance calculations based on Sr-isotopic data indicate that the residual melt comprised 60-70% of the resulting hybrid magma^34,39, which subsequently crystallized into a >3.0 km thick sequence overlying the Pyroxenite Marker (Fig. 2). If correct, then 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; resembling earlier estimates based upon thermal modelling of the Bushveld Complex³¹. We concur with previous studies^41,55 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 melt, as further substantiated by similar An-content of plagioclase at the base of the MZ at both sub-chambers (67.5%^41,56 and 71%^42,43, respectively). A greater thickness of MZ cumulates in the Roossenekal sub-chamber likely relates to a redeposition of crystals inside this depression and/or prevailing crystallization in the deeper parts of the magma chamber, due to a pressure-induced increase in the liquidus temperature of the melt^9,52,57.

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 subchamber towards the Belfast sub-chamber) through the 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. 4a). Our scenario is therefore substantially different from those where the melt crystallization and cumulate pile growth of the MZ occured concurrent with magma chamber replenishment³⁹. The scenario conforms better with liquidus phase equilibria predictions for a basaltic parent, like systematic changes in pyroxene assemblages (e.g., Opx-Aug through Opx-Aug-Pig to Aug-Pig)⁵⁸ and continuous decreases in both the An-content and Mg-number of cumulus plagioclase and orthopyroxene, respectively (Figs. 3 and 4b). Both the crystallization sequence and mineral compositional trends are successfully reproduced by fractional crystallization of the parental melt using the alphaMELTS software (Fig. 3; Methods, Extended Data). 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 constant⁵⁹ or irregular²² 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 is descernable within its overall decrease (Fig. 3). In fact, the appearance of the Anorthosite Markers themselves, may be associated with magma chamber replenishments⁶⁰. 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).

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. 4). The formation of such a cumulate sequence, at a typical solidification rate for large mafic intrusions (~1 cm/year^49,61) 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 in Earth’s geological history^{11,12,14,17,18}. The total volume of resident MZ melt may roughly 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 MZ³⁴, 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 km^3,³¹ 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)²⁵. It is only comparable to estimates of some of Earth’s large igneous provinces, like the Karoo (367,000 km³)⁶² and Afar (350,000 km³)²⁶. Thus, at the MZ stage 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, a current tendency in modern volcanology^14,17,18 and petrology^19,20,23 to discard the existence of such large and molten magma chambers^4–7appears 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 km³in volume⁶³. 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)^13,14,16,18, this does not automatically imply that ‘big tank’ magma chambers are absent from other regions (e.g., stable cratons with layered intrusions)^5–7. Moreover, since layered intrusions such as the Bushveld Complex are rare throughout geological time⁶⁴, 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 geophysically²⁷.

Any methods, additional references, Nature Research reporting summaries, source data, extended data, supplementary information, acknowledgements, peer review information; details of author contributions and competing interests; and statements of data and code availability are available at

Map and cross-sections constructions

The regional map was 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 Merwe (2007), Bevington and Hornsey (2010 – Nuplats Ltd, unpublished mapping), and Latypov and Chistyakova (2021). The geology was overlain on the USGS satellite digital terrane model and wireframes assigned to each lithology. The longitudinal section was drafted with the objective to accurately represent the geological relationships and unit geometries. This was achieved by rotating the map to view it in its true orientation, then digitising the lithologies as a section. This work was undertaken using Micromine™ geological modelling software. The 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.

MELTS programme, parental melts and its crystallisation sequences

The crystallization of the MZ was modeled using the MELTS algorithm for isobaric fractional crystallization at 2 kbar and FMQ oxygen buffer in steps of 5 °C. The calculations consisted of three major steps. Step 1: identified (by try-and-error) a starting composition (MZ-magma) and reproduced the A-B-C succession of cumulates (Fig. 3). This step reproduced a) the composition trend in Mg-number of orthopyroxene and An-content of plagioclase (considering the trapped liquid shift effect), and the volume proportion of minerals in the observed cumulates and b) the ratio between the volumes of cumulates observed in the stratigraphic column. Step 2: identified an orthopyroxene-saturated (O) magma by try-and-error approach, so that O-magma is more primitive than the MZ-magma, and if not mixed, it produces orthopyroxene at liquidus with a higher Mg-number. Step 3: reproduced the reversals below the pyroxenite marker (PM), the formation of PM and cumulates above it (A’ and B’) by gradual mixing of the existing volume of resident magma (R) from stage 1 (~ 30 %) with a new orthopyroxene-saturated magma (O) (see SupplTable S1). Gradually changing the mixing ratio from O₁₀R₉₀ to O₇₀R₃₀ coupled with concomitant fractional crystallization runs in steps of 1 C° reproduced the mineral assemblage and the variation of Mg-number in orthopyroxene and An-content of plagioclase as observed in zones A’ and B’ (Fig. 3, and SupplTable 5 and 4B-1 to -7, Extended Data), starting with Mg#-in-orthopyroxene and An reversals below the PM, continuing the reversal trend above the PM and turning to normal fractionation (decreasing) trend towards the top of A’ zone. The composition of MZ-, R-, and O- magmas are presented in SupplTable S1 (Extended Data), and the MELTS calculations are summarized in SupplTables S3A, S4A and S5. Detailed calculations are shown in SupplTables S2A, S2B, S2C, S3B, and S4B1 to S4B-7, Extended Data). Complete explanation of the modeling approach can be found in the extended data (Section “Thermodynamic modeling”).

Data availability

The authors declare that all relevant data are available within the article and its Supplementary Information Files.

Acknowledgements Several versions of this manuscript benefited from the constructive peer reviews by Johan Kruger, Paul Nex, Steve Prevec and Martin Klausen. The study was supported by research grants to R.L. from the National Research Foundation (NRF) and Department of Science and Innovation (DSI)-NRF Centre of Excellence for Integrated Mineral and Energy Resource Analysis (CIMERA) of South Africa. Any opinion, finding and conclusion or recommendation expressed in this contribution is that of the authors and the DSI-NRF CIMERA and NRF do not accept any liability in this regard.

Author Contributions M.M. has done an original mapping of the Anorthosite Markers across the entire study area. R.L., S.C. and R.H. undertook additional field work in this area, including traversing, small-scale mapping and sampling of outcrops as well as conceptualized the original idea and wrote a draft of the paper. R.H. constructed a 3-dimensional geology model of the Eastern limb from which the geological map and longitudinal sections of the study area were produced. G.C. performed thermodynamic modelling and participated in interpretation of data, editing the paper as well as in improving clarity of figures. All co-authors discussed the results and problems and contributed to producing a final draft for peer reviews.

Supplementary Information is available in the online version of the paper.

Competing interests The authors declare no competing interests.

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LatypovfieldobservationsUTM36S.txt
Latypov_field_observations UTM36S.kml
HornseyBevingtonfieldobservationsUTM36S.zip
Hornsey_Bevington_field_observations UTM36S.kmz
LatypovChistyakovafieldobservationsUTM36S.zip
Latypov_Chistyakova field_observations UTM36S.kmz
20211005LatypovetalFieldXWPTGeologyObservationsUTM36S.xlsx
Latypov et al Field X WPT Geology Observations UTM36S
ExtendedData22Nov2021.docx
Extended Data_22Nov_2021.docx
SupplTableS1CompositionsofmodeledMZRandOmagmas.xlsx
SupplTable S1_Compositions of modeled MZ- R- and O-magmas.xlsx
SupplTableS2AFractionalcrystallizationofdryMORBtholeiiteat2kbar.xlsx
SupplTable S2A_Fractional crystallization of dry MORB tholeiite at 2 kbar.xlsx
SupplTableS2BFractionalcrystallizationofwetMORBtholeiiteat2kbar.xlsx
SupplTable S2B_Fractional crystallization of wet MORB tholeiite at 2 kbar.xlsx
SupplTableS2CRmagmamixedwithMORBwetfractcrystalat2kbar.xlsx
SupplTable S2C_R-magma mixed with MORB(wet)_fract_crystal_at 2 kbar.xlsx
SupplTableS2DB2magmafractionalcrystallizationat2kbar.xlsx
SupplTable S2D_B2 magma fractional crystallization at 2 kbar.xlsx
SupplTableS3AMineralassemblagefractcrystalat2kbarsMZmagma.xlsx
SupplTable S3A_Mineral assemblage, volume, Mg# in orthopyroxene and An in plagioclase resulted by fractional crystallization at 2 kbars of MZ-magma
SupplTableS3BDetailedfractionalcrystallizationofMZmagmaat2kbar.xlsx
SupplTable S3B_Detailed fractional crystallization of MZ-magma at 2 kbar.xlsx
SupplTableS4ADetailedfractionalcrystallizationofOmagmaat2kbar.xlsx
SupplTable S4A_Detailed fractional crystallization of O-magma at 2 kbar.xlsx
SupplTableS4B1FractionalcrystallizationofthemixO10R90.xlsx
SupplTable S4B-1_Fractional crystallization of the mix O10R90.xlsx
SupplTableS4B2FractionalcrystallizationofthemixO20R80.xlsx
SupplTable S4B-2_Fractional crystallization of the mix O20R80.xlsx
SupplTableS4B3FractionalcrystallizationofthemixO30R70.xlsx
SupplTable S4B-3_Fractional crystallization of the mix O30R70.xlsx
SupplTableS4B4FractionalcrystallizationofthemixO40R60.xlsx
SupplTable S4B-4_Fractional crystallization of the mix O40R60.xlsx
SupplTableS4B5FractionalcrystallizationofthemixO50R50.xlsx
SupplTable S4B-5_Fractional crystallization of the mix O50R50.xlsx
SupplTableS4B6FractionalcrystallizationofthemixO60R40.xlsx
SupplTable S4B-6_Fractional crystallization of the mix O60R40.xlsx
SupplTableS4B7FractionalcrystallizationofthemixO70R30.xlsx
SupplTable S4B-7_Fractional crystallization of the mix O70R30.xlsx
SupplTableS5SummaryofPMmodellingresults.xlsx
SupplTable S5_Summary of PM modelling results.xlsx

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A 5-km-thick reservoir with >380,000 km³ of magma within the ancient Earth's crust

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