About 10% of the >200 investigated intrusive and effusive CAMP basaltic rocks show gas exsolution bubble-bearing MIs, hosted mainly in clinopyroxene and occasionally in plagioclase, orthopyroxene and olivine (Supplementary Figs. 1 and 2). The studied CAMP basaltic rocks are mainly porphyritic and microcrystalline, and the principal mineral phases are labradoritic-bytownitic plagioclase, augitic (abundant) and pigeonitic (scarce) clinopyroxene, rare Mg-rich orthopyroxene, and rare and mostly altered Mg-rich olivine. As accessory mineral phases, magnetite is common, while ilmenite is rare. In effusive rock samples (from USA, Canada, Morocco and Portugal), glomerocrystic aggregates of augitic clinopyroxene and plagioclase are commonly present (Supplementary Fig. 1), and are interpreted as clots of partially crystallized mineral mush from the transcrustal magmatic plumbing system13,14. In the only studied intrusive sample (from Palisades sill, USA), olivine is abundant and usually well preserved (Supplementary Fig. 2).

MIs are nearly ubiquitous in glomerocrystic aggregates of clinopyroxene and plagioclase (Supplementary Fig. 1). The bubble-bearing MIs usually have irregular shapes, can be single- or multi-bubble MIs, and contain up to 25 bubbles per inclusion, displaying a large range of glass/bubble ratios even within the same host crystal or crystal clot (i.e., there is no proportionality between the volume of glass and the volume/number of bubbles; Fig. 2 and Supplementary Fig. 2). In detail, the estimated volume fraction of bubbles within each MI ranges from <0.1 to >0.5 approximately. Moreover, MIs present a great variability in size, approximately from 5 to 50 μm on the principal axis. Bubbles within them usually have spherical shape and generally range from 1 to 15 μm in diameter (Supplementary Fig. 2). Sometimes bubbles are aggregated in the MIs, probably due to post-entrapment coalescence (Supplementary Fig. 2). Some MIs are partially crystallized, containing μm-sized daughter minerals in addition to, or instead of, bubbles. These crystals, likely formed from the melt after the entrapment, are mainly opaque mineral phases, such as sulphides and oxides (e.g., magnetite). The MIs glass has a more silicic (mainly andesitic) and more differentiated composition compared to the host basaltic rocks, and is clearly different from typical CAMP basalts or basaltic andesites (Supplementary Fig. 3 and Supplementary Table 2). The MIs glass is generally enriched in SiO 2 and Al 2 O 3 , and depleted in FeO, MgO and CaO compared to the host rocks (Supplementary Fig. 4), and would correspond to a residual melt after fractionation of ca. 40% augitic clinopyroxene, 10% plagioclase and 5% magnetite from a typical CAMP basalt (see “Methods” section). Such differentiation can only partly be due to post-entrapment crystallization of the few tiny crystals within the MIs or of the host clinopyroxene30,31,32, which displays constant augitic composition, shows only faint chemical zonation towards the glass, and is substantially out of equilibrium with it (see “Methods” section). The most evident compositional zoning of the host clinopyroxene consists of a decrease in CaO content and a slight increase in both MgO and FeO content close to the contact with the MIs (Fig. 3 and Supplementary Fig. 5). Hence, the local thin rim around the MIs of slightly Ca-depleted and Fe ± Mg-enriched clinopyroxene suggests the probable presence of augite–pigeonite exsolution lamellae close to the boundary of MIs, which likely formed at subsolidus conditions from an intermediate composition clinopyroxene that crystallized from the entrapped melt (i.e., post-entrapment crystallization). However, the chemical disequilibrium between the MIs glass and the host clinopyroxene, and the lack of significant chemical zoning within the host clinopyroxene at the contact with MIs are not consistent with substantial diffusive re-equilibration within the host clinopyroxene and suggest a rapid cooling after melt entrapment. This indicates that a previously differentiated bubble-bearing melt was entrained between interstices of growing crystals, and rapidly cooled down, forming MIs.

Fig. 2: Representative bubble-bearing melt inclusions at transmitted light optical microscopy. The black arrows indicate the bubble-bearing melt inclusions. a Single-bubble MI hosted in orthopyroxene (Opx; sample NS21, Nova Scotia, Canada). b Single- and multi-bubble MIs, with very irregular shapes, hosted in augitic clinopyroxene (Cpx; sample AN137A, Morocco). c Multi-bubble MI, partially crystallized (containing also opaque mineral phases), hosted in calcic plagioclase (Pl; sample NS9, Nova Scotia, Canada). d Multi-bubble MI hosted in augitic clinopyroxene (Cpx; sample NS9, Nova Scotia, Canada). Full size image

Fig. 3: Chemical maps of glomerocrystic clinopyroxene aggregates. Backscattered electrons (BSE) image a and corresponding scanning electron microscopy with energy-dispersive X-ray spectroscopy (SEM–EDS) maps b–f of a thin section area including MIs and the hosting glomerocrystic clinopyroxene aggregates. In the BSE image the brighter portions of clinopyroxene have augitic (Aug) composition and the darker ones have pigeonitic (Pgt) composition. In the SEM–EDS maps the brighter regions correspond to higher concentrations of the analysed element. These maps were acquired on sample NEW31 (New Jersey, USA). The scale bar is shown in a. a BSE image, b Al map, c Ca map, d Fe map, e Mg map, and f Ti map. Full size image

The bubbles within MIs were investigated in all the samples by confocal Raman microspectroscopy (Supplementary Table 3), looking for carbon species (CO, CO 2 , CH 4 and elemental C), as well as for other important volatile compounds in volcanic systems (SO 2 , H 2 S and H 2 O). In almost all analysed bubbles, CO 2 (within 54 bubbles of 9 samples) or elemental carbon (within 41 bubbles of 2 samples) were detected in both single- and multi-bubble MIs of rock samples collected from all over the CAMP (Fig. 4 and Source Data 1 and 2). In detail, Raman spectra show that CO 2 in CAMP bubbles is characterized by low density (ca. 0.1 g/cm3; see “Methods” section), and elemental carbon in CAMP bubbles is characterized by low crystallinity (i.e., it is present as disordered graphite and amorphous carbon; see “Methods” section). CO 2 concentrations from 0.5 to 1.0 wt% in whole MIs (i.e., glass plus bubbles) were calculated from the density of gaseous CO 2 within the bubbles (Supplementary Table 4) and from the estimated volume fraction of these bubbles within MIs (see “Methods” section). Other volatiles such as CO, CH 4 , SO 2 and H 2 S were not detected, while H 2 O was often found within the glass of MIs (Supplementary Fig. 6), but never in the bubbles. The MIs glass, investigated through Nano-SIMS, contains about 0.5–0.6 wt% H 2 O and 30–90 ppm CO 2 (Supplementary Table 5).

Fig. 4: Bubble-bearing melt inclusions at transmitted light optical microscopy and confocal Raman microspectroscopy. Left column: transmitted light photomicrographs at optical microscope of the analysed areas, bordered by dotted lines. Right column: Raman hyperspectral maps of the corresponding areas. a, c, e Photomicrographs of elemental carbon-bearing single- and multi-bubble MIs (a and c: sample NEW31, New Jersey, USA; e: sample AN39, Morocco). b, d, f Raman hyperspectral maps of the same samples area. g Photomicrograph of an irregular-shaped CO 2 -bearing multi-bubble MI (sample NS12, Nova Scotia, Canada). h Raman hyperspectral map of the same sample area. The Raman signal of CO 2 is weak due to its low density. However, spot analyses confirmed the presence of CO 2 in all bubbles. Full size image

CO 2 in CAMP basalts

The analysed bubble-bearing MIs strongly suggest that the CAMP magmatic system was rich in CO 2 . Most of the analysed bubbles contains CO 2 or, less frequently, elemental carbon, and no detectable amounts of any other investigated volatile phase (Supplementary Note 1). In particular, confocal Raman microspectroscopy allowed to distinguish and characterize both CO 2 and elemental carbon (see “Methods” section). The Raman spectrum of CO 2 is characterized by two sharp bands, usually called Fermi diad or Fermi doublet, associated to two symmetrical weak bands, usually called hot bands33 (Fig. 5a and Source Data 1). Instead, the first-order Raman spectrum of elemental carbon is characterized by two different bands, the composite D band, activated in disordered graphite by lattice defects and typical of non-crystalline structures34,35, and the single G band, typical of graphite34 (Fig. 5b and Source Data 2). This last band is here employed in a crossplot to characterize the different types of elemental carbon, distinguishing disordered graphite and amorphous carbon (Fig. 6). Interestingly, CO 2 and elemental carbon within gas exsolution bubbles are never present together in the same samples. The elemental carbon, which is likely present as a thin film coating the inner spherical surface of the bubbles, replaces CO 2 in some samples, probably due to a change in the oxidation state within MIs, for instance related to a diffusive loss of oxygen from the bubbles to the melt, when the latter crystallized oxides during cooling. The large variability in volume and number of bubbles observed in coexisting MIs (ranging from 1 to 25 bubbles per MI, approximately occupying from <0.1 to >0.5 of the MI volume, as optically estimated in thin and thick sections) reveals heterogeneous entrapment of MIs27,36. Therefore, the bubbles within MIs are interpreted as gas exsolution bubbles, formed during exsolution of a CO 2 -rich fluid phase likely from the silicate melt prior to, or during, their entrapment. Gas exsolution within MIs after melt entrapment was probably of minor importance, particularly for MIs with large bubbles27, because the trapping of a bubble-free melt would have produced homogeneous MIs, displaying very similar glass/bubble ratios, which were not observed in this study. The volatile-saturated melt and the volatiles may have a cogenetic origin (i.e., the melt was entrapped along with volatiles immediately after, or during, gas exsolution), or may have different origins (i.e., the melt was entrapped along with volatiles exsolved from deeper magmas, or degassed and fluxed from intruded crustal rocks).

Fig. 5: Raman spectra of CO 2 and elemental carbon. a Raman spectrum of CO 2 , acquired on sample NS9 (Nova Scotia, Canada). The Fermi diad is represented by sharp bands, at 1285 cm−1 and at 1388 cm−1, and the hot bands are represented by symmetrical weak bands, below 1285 cm−1 and above 1388 cm−1. b Raman spectra of elemental carbon: amorphous carbon acquired on sample NEW31 (New Jersey, USA), disordered graphite acquired on sample AN39 (Morocco), and ordered graphite (detail on the G band) acquired on a common pencil. Compared to the ordered graphite Raman spectrum, our Raman spectra of disordered graphite and amorphous carbon always have one or more D peaks between 1200 and 1400 cm−1. The G band lower than 1590 cm−1 indicates disordered graphite, while the G band higher than 1590 cm−1 indicates amorphous carbon (see “Methods” section). The D band is often composed by two peaks (D1 at ca. 1350 cm−1 and D5 at ca. 1270 cm−1) for both disordered graphite and amorphous carbon. Full size image

Fig. 6: Crossplot of the Raman spectra of elemental carbon. This crossplot displays the Raman spectra of elemental carbon with the peak position of the G band ranging from ca. 1575 cm−1 to ca. 1605 cm−1. The line in correspondence of 1590 cm−1 peak position value separates the disordered graphite data (below) from the amorphous carbon data (above) according to the present study (see “Methods” section). The areas bordered by dashed lines distinguish the graphite field from the kerogens and coals field according to ref. 34. The error on PP and therefore also on FWHM is considerably smaller than the spectral resolution for the Raman spectra displayed in the crossplot (0.8 cm−1)70, and thus is much smaller than the plotted symbols. Full size image

Clinopyroxene compositions and volatile element concentrations suggest that CO 2 entrapment occurred within the deep magmatic roots of CAMP. The pressure of crystallization of host clinopyroxene crystal clots can be calculated from mineral compositions (Supplementary Table 6) using methods developed for magmatic systems37,38 (Supplementary Note 2). The geothermobarometer based on the equilibrium between clinopyroxene and a magmatic liquid37 was applied using whole rock composition as proxy for the original magmatic liquid composition, because the MIs glass is in chemical disequilibrium with the host clinopyroxene. The calculated crystallization pressure ranges from 0.1 to 0.7 ± 0.2 GPa (at temperatures from 1150 to 1230 ± 27 °C) and is consistent with previous estimates from clinopyroxene crystallization pressures (from 0.2 to 0.8 GPa) in basalts from the entire CAMP23,39,40,41. These results suggest that the crystallization of clinopyroxene in the investigated CAMP samples occurred predominantly within the middle continental crust (on average ca. 12 ± 7 km for a pressure/depth gradient of about 0.03 GPa/km; Supplementary Fig. 7).

The deep origin of MIs is consistent with observed volatile concentrations in both their glass and bubbles (Supplementary Note 3). The presence of sulphides within some MIs shows that the entrapped melt became sulphide-saturated with S concentrations likely exceeding 1500 ppm42,43. Sulphur concentrations of the same order of magnitude were estimated for CAMP basalts44. Moreover, about 0.5–0.6 wt% H 2 O was detected in the MIs glass through NanoSIMS analysis, revealing hydrated conditions for these melts. Despite the presence of H 2 O and S in the MIs glass, these volatiles were not detected in the bubbles. Hence, considering a realistic maximum primary concentration of ca. 1 wt% H 2 O and ca. 0.1 wt% SO 2 in tholeiitic within-plate basaltic melts44,45, most H 2 O and SO 2 are expected to exsolve at pressures lower than 0.1 GPa (i.e., <3 km depth)16,46. Even considering that H+ may move from the bubbles into the glass, and CO 2 from the glass into the bubbles after MI entrapment47,48, the observed distribution of volatile species between glass and bubbles within MIs suggests the dominant occurrence of gas exsolution and bubble formation at relatively high pressures from a CO 2 -rich melt.

The inferred depth of CO 2 exsolution and entrapment indicates that this volatile species has a deep origin (at least 12 ± 7 km on average). It therefore reveals that the entire CO 2 budget involved in CAMP emplacement could not have originated exclusively from assimilation and degassing of shallow intruded sediments26, because sediments in the circum-Atlantic basins only reach a thickness of 5 km in eastern North America49 and <1 km in Morocco23 and Portugal39. On the contrary, at least part of the CO 2 most probably derived from assimilation of deep- to middle-crustal metasedimentary rocks (e.g., metacarbonates or graphite-bearing amphibolites/granulites) or from the mantle source of CAMP basalts (Fig. 7), containing significant amounts of recycled sedimentary material23,39,40,50,51.

Fig. 7: Sketch of the transcrustal plumbing system of CAMP basaltic magmas from the mantle to the surface. The evolution of basaltic magmas occurs at variable depth by crystallization of minerals, which then form aggregates in crystalline mushes13,14 and entrain bubble-bearing melt, forming MIs. Different volatile species exsolve at variable depth16. In particular, CO 2 -rich fluids (white bubbles) start exsolving at great depth, whilst H 2 O-rich fluids (blue bubbles) and S-rich fluids (yellow bubbles) start exsolving at shallow depth. The black dashed arrows indicate the potential sources for the carbon in CAMP magma: the mantle, the deep crust and the Palaeozoic or Triassic sedimentary basins in which CAMP sills intruded. The carbon within the here studied MIs derives from the deep sources as demonstrated with clinopyroxene geobarometry data. Clinopyroxene crystallization pressures of this study have been calculated using ref. 37 (Supplementary Note 2). Clinopyroxene crystallization pressures of bibliography are from ref. 23 for Morocco, ref. 39 for Portugal, and ref. 40 for USA. The error (±0.2 GPa) takes into account the uncertainties from both the geobarometry model (±0.1 GPa)37 and the electron microprobe analyses (±0.1 GPa, deriving from the ±10% accuracy on measured Na concentration). Full size image

The calculated depth of entrapment (ca. 12 ± 7 km) allows an estimation of the CO 2 concentration originally present in CAMP magmas. The CO 2 saturation in basaltic melts is achieved at ca. 1000 ppm at 0.2 GPa, increasing by ca. 500 ppm for each 0.1 GPa52. Considering the calculated crystallization depths, the minimum estimate for the CO 2 concentration of CAMP magma, before gas exsolution, is between ca. 500 and 4000 ppm. Such values are consistent with the CO 2 concentrations in the MIs, calculated from CO 2 density within the bubbles, which range from 0.5 to 1.0 wt%. Moreover, starting from the minimum calculated values of the CO 2 concentration within MIs (i.e., 0.5–0.6 wt%) as representative of CAMP magma, assuming an average density of 2.90 g/cm3 for basaltic rocks53 and considering 5–6 × 106 km3 for the total volume of CAMP (in order to take into account the deep plumbing system), the total amount of degassed volcanic CO 2 during CAMP emplacement would be up to 105 Gt. Interestingly, the values estimated for the CO 2 concentration of CAMP magma (0.5–1.0 wt%) and for the total amount of degassed volcanic CO 2 during CAMP emplacement (up to 105 Gt) are consistent with those assessed in several other LIPs, using different approaches29.

Implications for the end-Triassic climatic and environmental changes

The high-volume fractions of CO 2 - and elemental carbon-bearing bubbles within CAMP MIs, along with the inferred depths of formation, reveal the high abundance (0.5–1.0 wt%) of CO 2 in the CAMP transcrustal magmatic plumbing system. The CO 2 -bearing bubbles identified in CAMP MIs can be interpreted as batches of ascending volatiles entrapped in crystalline mush shortly prior to its mobilization and prior to eruption. This evidence for CO 2 saturation in the basaltic magmas at depth can explain the pulsed eruptive style of CAMP, where CO 2 acts as propellant for magma ascent, causing rapid and violent eruptive pulses. For instance, CO 2 -rich Hawaiian basalts have been shown to rapidly rise from over 5 km depth and to cause high fountaining eruptions54.