The stable isotope compositions of LCIT sulfates offer the most diagnostic clues for the formation conditions of sulfates. The unusually high δ34S (∼+49‰) of mirabilite likely arises from the dissimilatory microbial sulfate reduction (DMSR) reaction when sulfate is converted to reduced sulfur species, leaving a large enrichment of 34S in the residual dissolved sulfate. The high solubility of sodium sulfate requires a high salinity for its precipitation. Therefore, formation of mirabilite with a high δ34S would require DMSR to elevate the δ34S SO4 throughout the water body before precipitation. Melt-water ponds at the terminus of glaciers along the TAM are typically small, shallow and well aerated33 (Fig. 1), and they are therefore not suitable for DMSR. Reducing microenvironments may exist if a pond is covered by algal mats or if sufficient sediment depth and stability exists for the development of benthic communities34. However, no algae or benthic mats were observed in the ponds at the LCIT. Even with such bio-reduction, the turnover of the sulfur reservoir would be very rapid, resulting in virtually no net sulfur isotope enrichment34. On the basis of sulfate isotope signatures and direct assessment of microbial communities, DMSR in Antarctica has been inferred to occur mainly in lakes23,35. Ace Lake is a coastal, marine-derived, meromictic system in the Vestfold Hills32 where DMSR occurs and it is the only water body measured to date in Antarctica with overall highly enriched 34S (average δ34S of +42‰ (ref. 23)); it therefore appears to be a modern-day analogue for the production of mirabilite in the LCIT. DMSR is also reported for subglacial environment, but sulfate redox cycling therein produces isotope signatures distinctly different from those of the LCIT25. Alternatively, thermochemical sulfate reduction may also introduce high δ34S, but this process requires high reaction temperature (>100 °C; ref. 36) and is not known to occur in Antarctica.

The negative δ18O values (−16.9±0.3‰) of the mirabilites is indicative of the incorporation of light glacier water oxygen. As direct oxygen isotope exchange between sulfate and water is extremely slow37,38, the low δ18O values must have been acquired through incorporation of glacial melt-water oxygen during sulfate formation and/or through microbial sulfate reduction and sulfide oxidation cycles39,40. The δ34S values in Ace Lake show an apparent steady state for sulfur redox cycling23 (Supplementary Table 2), whereas the apparent Δ18O SO4–H2O at a range of +19.4 to +25.2‰ (Supplementary Table 2) approaches the predicted sulfate–water oxygen isotopic composition41. Therefore, assuming the LCIT mirabilite sulfate was also in oxygen isotopic equilibrium with water, our data imply that the δ18O of water was <−45‰ at the time the LCIT mirabilite precipitated; a value that is close to that of the present-day EAIS3,42. The uniform sulfur and oxygen isotope composition of mirabilite throughout the sampling sites also suggests they probably originated from the same water body.

The pond sulfate possesses extreme 18O depletion (−12.1 to −22.2‰) and a distinct non-mass-dependent 17O enrichment (Δ17O at +1.36 to +2.35‰) which is the only case known for modern natural water bodies. The positive Δ17O values are a clear indication that a portion of the pond sulfate derives from SAS, as this unique signature came ultimately from ozone and derivatives whose δ18O values can also be highly positive and variable in the troposphere and stratosphere43,44. However, despite having positive Δ17O values similar to the sulfates in soils of the MDVs and snow/ice cores from the EAIS45,46, the LCIT pond sulfates have much lower δ18O values (Fig. 2). Therefore, a source of sulfate in addition to SAS and with extremely low δ18O must have contributed to the pond sulfates.

It is noteworthy that the LCIT pond sulfates have δ18O values closer to sulfates in the MDVs than in ice cores. The most likely explanation is that both LCIT and DMVs sulfates have a component derived from Antarctic local weathering. Oxidation of sulfide minerals by glacial water can introduce sulfate δ18O as low as −19‰ (ref. 47). Assuming the formation pathways of SAS did not change drastically during the past few million years, the δ18O of sulfate from oxidative weathering can be calculated by treating LCIT pond sulfate as a simple mixing of SAS, oxidative weathering and minor sea-salt sulfate (SSS).

where x and y are the mole fractions of the SAS and SSS contributions, respectively. Conservatively, δ18O SAS and Δ17O SAS were set at −3‰ and +3.5‰, respectively17,27,45,46,48 (National Snow and Ice database, Fig. 2a) and Δ17O Oxidativeweathering, δ18O SSS , Δ17O SSS and y were set at −0.50‰, +10.3‰, −0.1‰ and 5%, respectively. Mass conservation calculations for isotope values show that ∼50 to 70% of the LCIT pond sulfate is SAS, and the δ18O oxidative weathering ranges from −43.5 to-58.3‰, which are much more negative than any previously recorded values. The low values put the glacier water δ18O at ∼−60 ±10‰ (ref. 26), which is also in the δ18O range of cold-based glaciers from higher latitudes of the EAIS. The accumulation of sulfate with positive Δ17O is characteristic of desert environments exposed to long periods of hyperaridity16,21. The Δ17O and δ18O of pond sulfate thus point to the existence of one or more ice-free and arid surfaces (that is, cold deserts) at high latitudes.

As the mirabilite and pond water sulfates have distinct isotopic compositions, the mirabilite evaporites could not have come from the pond water, or vice versa. Instead, the isotope composition is consistent with an Ace Lake-type reservoir for the formation of the mirabilite sulfates and a hyperarid desert with oxidative weathering for the pond sulfates. These conditions closely resemble the key geochemical and climatic characteristics of present-day MDV, and to an extent, the Vestfold Hills. The cold deserts could have existed in the vicinity of the LCIT and/or a region further towards the interior of the EAIS along the glacier flow path. As the LCIT geological settings cannot generate mirabilite sulfate with high δ34S (+49.8‰) or accumulate a significant amount of SAS with high Δ17O (+2.1‰), it is very likely that the mirabilite and pond sulfates are terrigenous, having been originally transported from deeper within the Antarctic interior.

Therefore, we argue that the LCIT sulfates originated from a time when the ice cap covered a smaller area than present, along the flow paths of the Beardmore Glacier2,9 that includes some of the topographic lowlands along the Queen Alexandra Range and the Pensacola Basin located between the TAM and the East Antarctic Polar Plateau. We suggest that the distribution of cold deserts may be patchy among relatively large ice-free lands, but it is unlikely that a few ice-free outcrops existed within thick ice caps. This once MDV-like environment enabled SAS to accumulate to high levels, glacial melt remnant water-fed lakes allowed for extended periods of DMSR and a relatively warm ambient temperature allowed for oxidative weathering of sulfur-bearing bedrocks. As the climate cooled and glaciers advanced, mirabilite beds may have been formed as a result of salt concentration by freezing of lakes. The mirabilite beds, together with regolith containing SAS and weathering-produced sulfates, would have been disrupted and carried by glaciers and eventually stranded at the LCIT moraine (Fig. 3). The exact location and the size of the cold deserts are yet to be determined. The provenance or geological terrains of these sediments may be traced using radiogenic isotopes and more data from other localities along the TAM4.

The initial age of the mirabilite sulfate salt formation would help determine its origin. Unfortunately, evaporite formation are difficult to date and inferences about age must be gleaned from elsewhere. The ice sheet producing the Beardmore Glacier was suggested to have been overriding the TAM in the Quaternary2,9, indicating that the cold deserts would be older than the Quaternary. Glacier palaeotemperature can be estimated based on a robust relationship between δ18O of precipitation and the local temperature49,50. The δ18O of the glacier water responsible for the LCIT sulfate δ18O is comparable, or lower, than those of the Dome C and Vostok ice cores that represent the past hundreds of thousands of years, suggesting that the temperature before emergence of the hypothesized interior cold deserts was as low as it is today49. However, the Miocene ice cap would have had a much higher δ18O than the present-day glacial δ18O because: first, the temperature of the Antarctic interior was not expected to be as low as today in the Miocene51; and second, the Southern Ocean provided relatively more moisture than it presently does, thus reducing the latitudinal effect for the vapour movement. The δ18O of the EAIS was estimated to be higher than −35‰ in the Oligocene and Miocene52,53. This is also supported by a recent leaf-wax study that showed that the δD of precipitation along the Antarctic coast was as high as −50‰ during 20 to 15.5 Myr (ref. 54), which indicates a much higher coeval δ18O of precipitation than today. Our sulfate oxygen isotope data therefore point to formation of LCIT sulfates with coeval waters similar to modern-day glacier waters. This suggests that the LCIT sulfates did not form during the early stages of Antarctic ice cap formation. Instead, the hypothesized interior cold deserts must have formed during a time of relative warmth but a period long after the global hydrological cycle was established similar to that of today. The Pliocene epoch is therefore the most feasible geological time for this occurrence. If we assume the interior cold deserts had 10% of the sulfate concentration as present-day upland soils in the MDVs, the interior cold deserts would have persisted for at least >1 Myr (ref. 16).

Our data support the existence of an ice-free, warmer but dry environment underneath the interior of the high EAIS in the past, likely during the Pliocene warmth. The implied massive ice contraction associated with the emergence of interior cold deserts is supported by recent evidence that during the Pliocene warmth, ∼500 km of ice retreated in the Wilkes Basin, and the ice-sheet basal thermal regime transitioned from polythermal to cold4,55. The geological processes leading to the formation of the interior cold deserts following glacier melt are not clear. However, the environment is topographically favourable, with lowlands (for example, Pensacola basin) surrounded by a high-altitude barrier to prevent glacier ice flow and katabatic winds to prevent moisture entry from the sea; conditions that are similar to today’s MDVs.