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This is an Open Access article distributed under the terms of the Creative Commons Attribution 4.0 License (http://creativecommons.org/licenses/by/4.0/).

Nineteen large (2348–4285 m above sea level) central polygenetic alkaline shield-like composite volcanoes and numerous smaller volcanoes in Marie Byrd Land (MBL) and western Ellsworth Land rise above the West Antarctic Ice Sheet (WAIS) and comprise the MBL Volcanic Group (MBLVG). Earliest MBLVG volcanism dates to the latest Eocene (36.6 Ma). Polygenetic volcanism began by the middle Miocene (13.4 Ma) and has continued into the Holocene without major interruptions, producing the central volcanoes with 24 large (2–10 km-diameter) summit calderas and abundant evidence for explosive eruptions in caldera-rim deposits. Rock lithofacies are dominated by basanite and trachyte/phonolite lava and breccia, deposited in both subaerial and ice-contact environments. The chronology of MBLVG volcanism is well constrained by 330 age analyses, including 52 new 40Ar/39Ar ages. A volcanic lithofacies record of glaciation provides evidence of local ice-cap glaciation at 29–27 Ma and of widespread WAIS glaciation by 9 Ma. Late Quaternary glaciovolcanic records document WAIS expansions that correlate to eustatic sea-level lowstands (MIS 16, 4 and 2): the WAIS was +500 m at 609 ka at coastal Mount Murphy, and +400 m at 64.7 ka, +400 m at 21.2 ka and +575 m at 17.5 ka at inland Mount Takahe.

Supplementary material: Summary age table with locations (Table S1), Murphy age data (new data only) (Table S2), Takahe age data (new data only) (Table S3) and palaeo-ice history (Table S4) are available at https://doi.org/10.6084/m9.figshare.c.5205362

Marie Byrd Land (MBL) and western Ellsworth Land in West Antarctica have a long history of Cenozoic volcanism defined by 19 polygenetic central volcanoes and numerous smaller volcanic centres exposed above the level of the West Antarctic Ice Sheet (WAIS) (Figs 1 & 2). The central volcanoes include some of the largest composite volcanoes in Antarctica and on Earth. The history is not well known because the volcanoes are remote, mostly snow- and ice-covered, and partially or fully buried by the WAIS. The last major review of the physical volcanology and geochronology of the region was published 30 years ago (LeMasurier and Thomson 1990). This chapter updates that and includes: (1) the establishment of stratigraphic nomenclature for volcanoes on the MBL and Thurston Island tectonic blocks; (2) a general synthesis of regional volcanism; (3) updated summaries of geological histories of individual volcanoes in MBL and western Ellsworth Land; and (4) an updated analysis of the volcanic record of the WAIS. The volcano summaries include general trends in the volcanic histories based on published and unpublished fieldwork, lithofacies descriptions, 40Ar/39Ar geochronology data, and an assessment of the state of knowledge of exposed volcanoes.

Fig. 1.

Map of Marie Byrd Land Volcanic Province (MBLVP), West Antarctica and inset map of Antarctica showing the MBLVP (red box) and Thurston Island Volcanic Province (yellow box, see Fig. 2). Volcanic fields are labelled and volcanoes included in fields are within dashed line polygons and listed in Figure 3. The two large rectangular outlines show the map boundaries of Figures 5a and 5b; an inset map of Mount Siple is also shown in Figure 5a. Image maps of individual central volcanoes or volcanic fields are presented with volcano summaries. Abbreviations: VF, Volcanic Field; USAS Esc., USAS Escarpment; C.I., Cruzen Island; S.I. is Shepard Island; G.I. is Grant Island. The base image maps are derived from Google Earth Pro (image US Geological Survey; Data: SIO, NOAA, US Navy, NGA GEBCO).

Fig. 1.

Map of Marie Byrd Land Volcanic Province (MBLVP), West Antarctica and inset map of Antarctica showing the MBLVP (red box) and Thurston Island Volcanic Province (yellow box, see Fig. 2). Volcanic fields are labelled and volcanoes included in fields are within dashed line polygons and listed in Figure 3. The two large rectangular outlines show the map boundaries of Figures 5a and 5b; an inset map of Mount Siple is also shown in Figure 5a. Image maps of individual central volcanoes or volcanic fields are presented with volcano summaries. Abbreviations: VF, Volcanic Field; USAS Esc., USAS Escarpment; C.I., Cruzen Island; S.I. is Shepard Island; G.I. is Grant Island. The base image maps are derived from Google Earth Pro (image US Geological Survey; Data: SIO, NOAA, US Navy, NGA GEBCO).

Fig. 2.

Map of Thurston Island Volcanic Province, showing the Hudson Mountains and Jones Mountains volcanic fields. Inset map of Antarctica shows the map boundary (yellow box) and the boundary of the MBLVP map (Fig. 1, red box). The base image map is derived from Google Earth Pro. Abbreviations: PIB, Pine Island Bay; PIG, Pine Island Glacier.

Fig. 2.

Map of Thurston Island Volcanic Province, showing the Hudson Mountains and Jones Mountains volcanic fields. Inset map of Antarctica shows the map boundary (yellow box) and the boundary of the MBLVP map (Fig. 1, red box). The base image map is derived from Google Earth Pro. Abbreviations: PIB, Pine Island Bay; PIG, Pine Island Glacier.

The MBL Volcanic Group (MBLVG) is formally defined here to include volcanoes on the MBL and Thurston Island tectonic blocks in West Antarctica (from 156° W to 93° W) (Figs 1, 2 & 3). Specifically, the MBLVG is defined to include the MBL Volcanic Province (MBLVP) (Fig. 1) and the Thurston Island Volcanic Province (Fig. 2), the former comprising nine volcanic fields and the latter comprising two volcanic fields. Four of the volcanic fields in the MBLVP consist of central volcanoes that form linear mountain chains and have age trends along the volcano alignments; these fields bear the mountain range names (Flood, Ames and Executive Committee ranges, and Crary Mountains). Each of these volcanic fields consists of three–five large polygenetic central volcanoes. The Eastern MBL Volcanic Field includes three isolated polygenetic central volcanoes and several smaller volcanic centres. The Mount Siple Volcanic Field consists of Mount Siple, the large isolated coastal polygenetic central volcano. The McCuddin Mountains Volcanic Field includes the oldest volcanic centres, all located in central MBL. Finally, the MBLVP includes the Hobbs Coast and Fosdick Mountains volcanic fields, each composed of multiple monogenetic volcanic centres. The much smaller Thurston Island Volcanic Province in western Ellsworth Land includes the Hudson and Jones Mountains volcanic fields, both composed of monogenetic volcanoes. This study focuses more closely on the MBLVP than the Thurston Island Volcanic Province because the MBLVP is much larger and has been studied in greater detail.

Fig. 3.

Stratigraphic nomenclature for the Marie Byrd Land Volcanic Group (MBVG), including volcanic provinces, volcanic fields, and individual central volcanoes and isolated smaller centres.

Fig. 3.

Stratigraphic nomenclature for the Marie Byrd Land Volcanic Group (MBVG), including volcanic provinces, volcanic fields, and individual central volcanoes and isolated smaller centres.

The MBLVG volcanoes are exposed as mostly snow-covered nunataks protruding through the marine-based WAIS in West Antarctica. Summit elevations of the large central volcanoes range from 2348 to 4285 m above sea level (asl) and ice-sheet elevations surrounding the volcanoes range from sea level at the coast to more than 2400 m asl at inland locations near the centre of the province. Exposures of the volcano–basement rock contacts are buried at most volcanoes but, where exposed, appear to be higher in elevation towards the centre of the province (LeMasurier and Landis 1996).

The oldest known volcanism in the MBLVG is dated at 36.6 Ma but most of the documented volcanism has occurred since 13.4 Ma (LeMasurier 1990h; Wilch and McIntosh 2000). The 40Ar/39Ar ages of volcanic rocks at the large central volcanoes range from 20.46 ± 0.07 Ma (all errors ± 2σ uncertainty) at Mount Flint in the McCuddin Mountains Volcanic Field to 8.3 ± 5.4 ka at Mount Takahe in the Eastern MBL Volcanic Field (Wilch 1997). Mount Berlin in the Flood Range Volcanic Field is an active central volcano with several steaming fumarolic ice towers on the periphery of its summit caldera (LeMasurier and Wade 1968; Wilch et al. 1999). Mounts Takahe and Mount Berlin are the two central volcanoes documented to be active in the Holocene (Wilch et al. 1999), and Mount Waesche in the Executive Committee Range Volcanic Field may have had eruptions in the Holocene and be considered active because there are abundant englacial tephra layers in blue ice adjacent to the volcano (Dunbar et al. 2021). Seismic evidence suggests that an active magmatic intrusive complex exists at 25–40 km below the base of the ice sheet 55 km to the south of Mount Waesche (Lough et al. 2013; Quartini et al. 2021). Previous workers have noted geographical and temporal patterns of the large central volcanoes, with volcanoes aligned in rectilinear east–west and north–south ranges, and younger volcanoes being located towards the periphery of the province (LeMasurier and Rex 1989, 1991; Paulsen and Wilson 2010).

The volcanoes of the MBLVP are mostly located on the c. 1000 × 500 km MBL structural dome on the north flank of the West Antarctic Rift System (WARS) (LeMasurier 2006). The MBL dome is associated with the MBL crustal block, one of four crustal blocks in West Antarctica (Dalziel and Elliot 1982) (Fig. 4). Bedrock elevations on the MBL dome (Fig. 4) are above sea level and rise to an elevation of c. 2700 m asl in the centre of the province (LeMasurier 1990h, 2006; LeMasurier and Landis 1996; Fretwell et al. 2013). The MBL dome has been included by some as part of the WARS based on the presence of horst and graben structures (LeMasurier 1990g), although it is more typically described as being situated on the north flank of the WARS (e.g. Paulsen and Wilson 2010). The WARS forms deep marine basins between the southern edge of the MBL dome and the northern front of the Ellsworth Mountains and Transantarctic Mountains (Fig. 4). The WARS extends from eastern MBL and the Amundsen Sea through the Ross Sea Basin to northern Victoria Land, and is one of the major continental rift systems on Earth, similar in size to the US Basin and Range Province and the East African Rift (Tessensohn and Wörner 1991; LeMasurier 2008). Large-scale asymmetrical intracontinental rifting of the nearly stationary Antarctic Plate formed the Ross Sea Basin and deep marine basins beneath the WAIS (Cooper and Davey 1985; Cooper et al. 1991; Granot and Dyment 2018). The Ross Sea Basin consists of a series of sediment-filled horst and graben structures. The intracontinental rift zone extends from the Ross Sea to beneath the WAIS (Behrendt et al. 1991; 1996). The Byrd Subglacial Basin and Bentley Subglacial Trench form deep submarine east–west-orientated basins that are typically interpreted as downfaulted graben within the WARS. The timing of extension and rifting events are not well constrained in West Antarctica, although it is generally stated that most of extension occurred in the Cretaceous (Siddoway 2008; Jordan et al. 2020). Most recognize two rifting episodes: early rifting in the Late Cretaceous; and late rifting that began in the Eocene and intensified in the late Cenozoic (Cooper and Davey 1985; Cooper et al. 1991; Jordan et al. 2020). Spiegel et al. (2016) also identified two episodes of tectonic denudation (presumably associated with uplift and rifting) at 100–60 and 20–0 Ma, based on interpretations of thermochronology data from bedrock outcrops in coastal MBL.

Fig. 4.

Map of subglacial topography of West Antarctica derived from BEDMAP2 data (Fretwell et al. 2013). Crustal blocks in West Antarctica include the Marie Byrd Land (MBL), West Antarctic Rift System (WARS), Thurston Island (TI) and the Ellsworth Mountains (EM) blocks. Most of the MBLVP central volcanoes are >2000 m asl, shown in red. The MBLVP is located on the MBL crustal block on the north flank of the West Antarctic Rift System (WARS) The WARS is dominated by deep subglacial basins including the Byrd Subglacial Basin (BSB) and the Bentley Subglacial Trough (BST). The WARS extends through the Ross Sea Embayment (RSE). The volcanoes of TIVP are situated on the Thurston Island crustal block. The Ellsworth Mountains and connected Transantarctic Mountains (TAM) form the rift shoulder on the south flank of the WARS.

Fig. 4.

Map of subglacial topography of West Antarctica derived from BEDMAP2 data (Fretwell et al. 2013). Crustal blocks in West Antarctica include the Marie Byrd Land (MBL), West Antarctic Rift System (WARS), Thurston Island (TI) and the Ellsworth Mountains (EM) blocks. Most of the MBLVP central volcanoes are >2000 m asl, shown in red. The MBLVP is located on the MBL crustal block on the north flank of the West Antarctic Rift System (WARS) The WARS is dominated by deep subglacial basins including the Byrd Subglacial Basin (BSB) and the Bentley Subglacial Trough (BST). The WARS extends through the Ross Sea Embayment (RSE). The volcanoes of TIVP are situated on the Thurston Island crustal block. The Ellsworth Mountains and connected Transantarctic Mountains (TAM) form the rift shoulder on the south flank of the WARS.

Geophysical data indicate that extended crust in the WARS is as thin as 21 km at the Bentley Subglacial Trench (Winberry and Anandakrishnan 2004). Calculated mean crustal thicknesses beneath the MBL dome from seismic data range from 28 to 33 km, 5–10 km thicker than in the WARS (Chaput et al. 2014; Ramirez et al. 2017; Shen et al. 2018). The presence of the MBL structural dome is attributed to a thermal anomaly beneath the volcanic province rather than isostatic compensation of slightly thicker crust (Winberry and Anandakrishnan 2004). LeMasurier and Landis (1996) interpreted MBL dome elevations increasing towards the centre of the MBLVP as a topographical expression of a mantle plume and suggested that the onset of uplift of the MBL dome coincided with an early pulse of volcanism at 29–27 Ma. In contrast, Spiegel et al. (2016) noted an absence of tectonic activity between 60 and 20 Ma, with acceleration of tectonic denudation beginning at 20 Ma, based on thermochronological data from MBL bedrock.

There have been significant advances in understanding of the thermal structure of the MBL lithosphere and sublithospheric mantle based on several multi-year geophysics experiments in West Antarctica. Lloyd et al. (2015) noted that in West Antarctica the slowest P- and S-wave velocities extend to 200 km below the Executive Committee Range, and Hansen et al. (2014) interpreted a low-velocity zone extending down to 800 km beneath MBL using P-wave tomography; both studies support the plume idea. Moreover, a modelling study using ambient seismic noise velocities by Heeszel et al. (2016) identified an extensive low-velocity zone down to >200 km centred on the Executive Committee Range that is again consistent with a thermal anomaly and a mantle plume. However, Heeszel et al. (2016) also pointed out that this thermal anomaly does not extend to the transition zone in the mid-mantle. Thinning of the transition zone is expected if impacted by elevated temperatures associated with a mantle plume (Reusch et al. 2008; Emry et al. 2015). In other geophysical studies, Seroussi et al. (2017) suggested that a weak mantle plume beneath MBL with a maximum geothermal heat flux of 150 mW m−2 is viable, based on modelling and assessment of the hydrology at the base of the WAIS.

In addition to the exposed volcanoes of the MBLVG, there are likely to be many subglacial volcanoes buried beneath the WAIS. Active subglacial volcanoes (Blankenship et al. 1993; Behrendt 2013) and an active intrusive magmatic system (Lough et al. 2013) have been mapped with confidence beneath the WAIS surface on the basis of geophysics, and Loose et al. (2018) inferred active volcanism beneath Pine Island Glacier based on geochemical signatures in meltwater. Extensive aeromagnetic surveys in the WARS show abundant magnetic anomalies, suggesting significant basaltic volcanism (as many as 1000 volcanic centres) beneath the WAIS (e.g. Behrendt 2013). Topographical interpretations of conical subglacial landforms beneath the WAIS, combined with pre-existing geophysical and field data, led de Vries et al. (2018) to identify 138 subglacial volcanoes in West Antarctica, including 91 previously unknown volcanoes. The inferred subglacial volcanoes are located on the MBL, Thurston Island and WARS tectonic blocks. Although these subglacial volcanoes have not been verified by samples, we suggest that the volcanoes be included in the MBLVG.

Possible causes of the MBLVG volcanism are debated by geochemists and geophysicists, and are discussed briefly here and in detail by Panter et al. (2021). Geochemically and isotopically, the MBLVG alkaline volcanic rocks have a HIMU (high μ = 238U/204Pb)-like mantle signature, typically associated with ocean island magmas (see the discussion in Panter et al. 2021). Several models to explain the origins of the HIMU-like magma and Cenozoic volcanism in MBL and the rest of the WARS have been proposed, ranging from active mantle-plume-driven volcanism since mid-Cenozoic time, to mid–late Cenozoic melting of a fossil Cretaceous mantle plume, to passive rifting and decompression melting of enriched mantle source rocks (for discussions see Storey et al. 2013; Martin et al. 2021; Panter et al. 2021; Rocchi and Smellie 2021).

Fourteen of the 19 central felsic volcanoes of the MBLVG are aligned in four linear chains (Fig. 5) (LeMasurier and Rex 1989). The linear chains show progressions in the ages of the onset of volcanic activity at the central volcanoes along each chain. In the Executive Committee Range Volcanic Field there is a north–south alignment of five polygenetic volcanoes with southward younging from Mount Hampton (13.4 Ma) to Mount Waesche (2.0 Ma). In the Ames Range Volcanic Field there is a north–south alignment of three polygenetic volcanoes with a northward younging from Mount Andrus (12.7 Ma) to Mount Kauffman (7.1 Ma). In the Crary Mountains Volcanic Field, there is north–south alignment of three polygenetic volcanoes with a northward younging from Mount Rees (9.46 Ma) to Mount Frakes (4.3 Ma). In the Flood Range Volcanic Field, there is an east–west alignment of three polygenetic volcanoes with a westward younging from Mount Bursey (10.1 Ma) to Mount Berlin (0.578 Ma). Many of these polygenetic volcanoes have two calderas with offsets that reflect the same alignments and age progressions as the central volcanoes (Fig. 5). Linear age progressions of volcanoes elsewhere (e.g. Hawaii, Snake River Plain) have been interpreted as ‘plume tracks’ recording movement of tectonic plates above mantle plumes or hotspots (Morgan 1972; Courtillot et al. 2003) but the discordant directions among the various linear trends in MBL precludes this explanation.

Fig. 5.

(a) Map of the western and central MBLVP volcanoes; and (b) a map showing the eastern MBLVP volcanoes. The inset map in (b) shows the Mount Siple Volcanic Field and central volcano. Map boundaries are delineated in Figure 1. Volcanic fields and individual central volcanoes are labelled. The approximate map view extent of the volcanoes at the level of the WAIS is shaded blue, and summit calderas of central polygenetic volcanoes are shaded yellow. Reference map with caldera age data given in Table 2.

Fig. 5.

(a) Map of the western and central MBLVP volcanoes; and (b) a map showing the eastern MBLVP volcanoes. The inset map in (b) shows the Mount Siple Volcanic Field and central volcano. Map boundaries are delineated in Figure 1. Volcanic fields and individual central volcanoes are labelled. The approximate map view extent of the volcanoes at the level of the WAIS is shaded blue, and summit calderas of central polygenetic volcanoes are shaded yellow. Reference map with caldera age data given in Table 2.

There are two different hypotheses about the origin of the spatial and temporal patterns in the MBLVP. LeMasurier and Rex (1989) recognized a systematic chronological pattern associated with the spatial distribution, with the oldest volcanoes near the centre of the province and progressively younger volcanoes away from the centre. They suggested that the patterns resulted from centrifugal extension and reactivation of relict fracture systems in the brittle crust caused by lithospheric doming associated with the rise of a 550–650 km mantle plume beneath the province.

Paulsen and Wilson (2010) conducted a detailed structural trend analysis of the orientations and ages of the polygenetic volcanic chains, elongate volcano edifices, elongate summit calderas and flank vents in MBL and elsewhere in Antarctica to determine stress directions during volcanism. The shape elongation measurements were assigned reliability ratings that varied from definite to indeterminate. Less reliable ratings were given in cases of low axial ratios of edifices or calderas, snow or ice cover, or erosion of part of the feature. The highest reliability data showed that overall edifice and caldera elongation directions were parallel to overall volcano alignments. The flank vent alignments were deemed mostly unreliable. The stress data combined with all available age K–Ar and 40Ar/39Ar age data showed that Miocene volcanism in MBL occurred along north–south alignments, whereas latest Miocene–Quaternary volcanism (since c. 6 Ma) occurred along east–west alignments (Paulsen and Wilson 2010). These data suggested a rapid rotation in the maximum horizontal stress field from north–south to east–west as early as 6 Ma. Paulsen and Wilson (2010) noted that the east–west orientation since 6 Ma is parallel to the absolute motion of the Antarctic Plate. They concluded that the rotation of the stress field was reflected in a change in the spatial pattern of volcanoes, and was driven by a major plate reorganization and a change in plate motion in the Pacific Basin.

Field research of the MBLVG has required major logistical support because the closest permanent research station is over 1300 km away (Table 1). The US Antarctic Program has provided logistical support for most of the volcano studies, although some field seasons included international teams of researchers and support from other international programmes (e.g. British Antarctic Survey; New Zealand Antarctic Programme). Most of the central volcanoes were discovered during overflights by Admiral Richard Byrd in the second Antarctic Expedition (1933–35) and the United States Service Expedition (1939–41). Scientists participating in several over-snow traverses from 1934 to 1959 made the first geological visits to the volcanoes. Many of the volcanoes and volcanic features are named after the expedition and traverse team members, as summarized in LeMasurier and Thomson (1990). These visits were typically limited to very few outcrops. Expeditions focused on geological mapping and sample collection began in 1960 with an investigation of the Jones Mountains in western Ellsworth Land (Craddock et al. 1964). Three major expeditions from 1966 to 1969 focused on understanding the geology of West Antarctica (Wade 1971). The 1967–68 Marie Byrd Land Survey II concentrated on the MBLVP, and all of the central volcanoes and major satellite centres of MBL were visited by helicopter reconnaissance from temporary deep field camps. Wesley LeMasurier led the 1967–68 investigations of the MBL volcanoes that resulted in the first analysis of the nature, extent and significance of the MBL volcanism (e.g. LeMasurier 1972), and laid the foundations for subsequent analysis. The 1968–69 Ellsworth Land Survey included study of the Jones and Hudson mountains in western Ellsworth Land (Wade and Craddock 1968). In 1977–78, another helicopter-based survey was conducted from a single-season deep field camp, focusing on central volcanoes in the Ames and Flood ranges, as well as isolated nunataks along the Hobbs Coast. Results from the 1967–68 and 1977–78 field seasons, and subsequent analyses on the rocks collected, are the basis of most summaries in LeMasurier (1990h) and of many subsequent papers. Most volcanoes have been revisited since this early helicopter-based fieldwork, although many outcrops have not been revisited.

Table 1.

Field expeditions to Marie Byrd Land volcanoes

Volcano19341939–411957–58195919591960–611966–671967–681968–691977–781984–85A1984–85B1984–85C1989–901990–911992–931993–941998–991999–002018–19
Mount MurphyXXX
Mount TakaheXXXX
Toney MountainXX
Mount SteereXX
Mount FrakesXX
Mount SipleXX
Mount HamptonXX
Mount CummingXX
Mount Hartigan (northern)X
Mount SidleyXX
Mount WaescheXXX
Mount FlintXX
Mount AndrusXXX
Mount KosciuskoXX
Mount KauffmanXX
Mount BurseyXXX
Mount MoultonXXXX
Mount BerlinXXXX
Kohler RangeX
Mount PetrasXXXX
Usas EscarpmentXXX
Hobbs Coast (M-P)XXX
Fosdick MountainsXX
Jones MountainsXXX
Hudson MountainsXXX
Volcano19341939–411957–58195919591960–611966–671967–681968–691977–781984–85A1984–85B1984–85C1989–901990–911992–931993–941998–991999–002018–19
Mount MurphyXXX
Mount TakaheXXXX
Toney MountainXX
Mount SteereXX
Mount FrakesXX
Mount SipleXX
Mount HamptonXX
Mount CummingXX
Mount Hartigan (northern)X
Mount SidleyXX
Mount WaescheXXX
Mount FlintXX
Mount AndrusXXX
Mount KosciuskoXX
Mount KauffmanXX
Mount BurseyXXX
Mount MoultonXXXX
Mount BerlinXXXX
Kohler RangeX
Mount PetrasXXXX
Usas EscarpmentXXX
Hobbs Coast (M-P)XXX
Fosdick MountainsXX
Jones MountainsXXX
Hudson MountainsXXX

1934, Byrd Antarctic Expedition II; 1939–41, US Antarctic Service Expedition; 1957–58, Marie Byrd Land Traverse Party; 1959, Byrd Station Traverse; 1959, Executive Committee Range Traverse; 1960–61, Ellsworth Land Camp Minnesota Expedition; 1966–67, Marie Byrd Land Survey, Helicopter Reconnaissance; 1967–68, Marie Byrd Land Survey II, Helicopter Reconnaissance; 1968–69, Ellsworth Land Survey, Helicopter Reconnaissance; 1977–78, Helicopter Survey (Western MBL); 1984–85A, Ice Core Tephra Source (Mount Takahe); 1984–85B, Polar Sea Helicopter; 1984–85C, Joint US–UK Expedition (Jones Mountains); 1989–90, West Antarctic Volcano Expedition (WAVE) Season 1; 1990–91, WAVE, Season 2; 1990–91, Ford Range Expedition; 1992–93, WAVE II, Season 1; 1993–94, WAVE II, Season 2; 1998–99, Mount Takahe; 1998–99, Fosdick Mountain Expedition; 1999–00, Mount Moulton Blue-ice Tephra; 2018–19, Mount Waesche Glaciation Study.

The most recent era of fieldwork involving more detailed analysis of selected volcanoes began in 1984–85 at Mount Takahe and Mount Murphy, and was carried out by small autonomous teams transported to the deep field via LC-130 aircraft. These teams accessed outcrops by snowmobiles. Subsequent field seasons in 1989–90 (Executive Committee Range), 1990–91 (Mount Murphy and Mount Hampton), 1992–93 (Crary Mountains), 1993–94 (Ames and Flood ranges, Hobbs Coast Nunataks and McCuddin Mountains), 1997–98 (Mount Waesche), 1998–99 (Mount Takahe), 1999–2000 (Mount Moulton), and 2018–19 (Mount Waesche) used this same autonomous approach augmented by transport to some sites by Twin Otter aircraft. Table 1 lists major expeditions and a fieldwork timetable for each of the central MBL volcanoes and at satellite volcanic centres.

Most of the 19 central volcanoes have a shield-like morphology, with low-angle (10°–15°) flank slopes and pronounced summit calderas (up to 8 km in diameter) (Fig. 5). Eight of the 19 volcanoes have a single summit caldera. Mount Hampton, Mount Waesche and Mount Sidley in the Executive Committee Range each have an older caldera truncated by a younger caldera. Mount Hartigan, Mount Berlin and Mount Bursey each have a summit caldera and multiple non-overlapping older calderas. Mount Moulton has a summit caldera, and a second, and possibly a third, caldera. Mount Rees and Mount Murphy are both deeply dissected and lack preserved summit calderas. If each isolated caldera were considered a distinct central volcano, there would be as many as 25 central volcanoes.

Other features of most of the central volcanoes include a lack of dissection; limited outcrops due to almost complete snow- and ice-cover; lithofacies dominated by lava rather than pyroclastic rocks; and rock compositions having largely bimodal alkaline compositions, dominated by mafic basanite and felsic trachyte and phonolite. LeMasurier (1990h) characterized the volcanoes as shield volcanoes, based on low slope angles and the dominance of lavas in outcrop and the flank slopes. Panter et al. (1994) characterized Mount Sidley with a large breached summit caldera as a stratovolcano, and early workers (e.g. González-Ferrán and González-Bonorino 1972; LeMasurier 1972) described many of the central volcanoes as stratovolcanoes. In general, the low degree of dissection precludes assessment of the internal structure of most of the volcanoes. Detailed studies of some of the more dissected volcanoes (Panter et al. 1994, 1997, 2000; Wilch et al. 1999; Wilch and McIntosh 2000) reveal complex polygenetic histories (Panter et al. 2021). Caldera wall exposures at many of the central volcanoes include explosively-erupted non-welded and welded pumiceous materials (Wilch et al. 1999; authors’ personal observations), although most early work described these exposures as lavas. Recognizing the general low angle slopes and also the complex polygenetic evolutionary patterns among the best studied volcanoes, we characterize the central volcanoes as shield-like composite volcanoes. An exception is Mount Sidley, a more typical stratovolcano.

Most of the large central volcanoes and satellite centres in the MBL and Thurston Island volcanic provinces are surrounded and partially buried by the WAIS. Many of the volcanoes and volcanic ranges are significant obstacles to flow of the WAIS, with ice on the upstream sides of the edifices as much as 800 m higher (e.g. Mount Moulton) than ice on the downstream sides. Andrews and LeMasurier (1973) evaluated rates of glacial erosion by comparing the heavily-dissected Mount Murphy and the much less dissected Mount Takahe.

The central volcanoes of the MBLVG are among the largest volcanoes in Antarctica and the world (Table 2). Mount Sidley is the highest Antarctic volcano at 4285 m asl. All of the MBL central volcanoes have summit elevations >2000 m asl and 10 are >3000 m asl. Only nine central volcanoes in Antarctica outside of MBL reach elevations >2000 m asl and only four of those are >3000 m asl. Estimating the size and volume of MBL volcanoes is challenging because most of the volcanoes are buried by the WAIS, and some are likely to extend far below ice and sea level (LeMasurier 1990h; LeMasurier 2013). The amount of relief above ice level ranges from just 400 m at Mount Cumming to 2400 m at Toney Mountain. The exception is the island volcano Mount Siple along the Bakutis Coast that is separated from the continent by the Getz Ice Shelf. Mount Siple rises from sea level to about 3100 m asl. The edifice shapes and exposed bases of the central volcanoes vary from circular to very elongate, with exposed basal dimensions from <10 to 70 km in length. The exposed volumes (above ice or sea level) range from 30 km3 to 1800 km3, with the largest estimate being Mount Siple at the coast (LeMasurier 1990h) (Table 2). Mount Siple would rank third on lists of edifice volumes calculated from digital elevation models (DEMs) of more than 900 shield volcanoes and composite worldwide (Grosse et al. 2014; Grosse and Kervyn 2018). Given that many of the MBL volcanoes are likely to extend below ice level, the exposed volumes represent minimum estimates. Toney Mountain is the only volcano in MBL for which there is a geophysically-based depth estimate of the sub-ice volcano–basement contact (Bentley and Clough 1972). The contact is at 3000 m below sea level, suggesting a total relief of the volcano of c. 6600 m; with volumes estimates ranging from 2800 to 3613 km3 (LeMasurier et al. 1990d; LeMasurier 2013). Given that most of the volcanoes have portions buried beneath the WAIS and that in places the WAIS is grounded far below sea level in deep basins, the calculated exposed volumes represent minimum values. In the absence of good depth data on the volcano–basement contacts, LeMasurier (2013) estimated a likely ratio of felsic to mafic rocks in five of the central volcanoes, and used this to estimate the amount of rock buried beneath the ice-sheet surface. Using this method, calculated volumes were up to 10 times greater than the exposed volumes, with Mount Takahe having the largest estimated volume of 5520 km3 (LeMasurier 2013). This method assumes that the ratio of mafic to felsic rock is known and that the felsic volume can be estimated from limited outcrops. In this chapter, we use the more conservative estimated volumes based on the volcano volumes above ice level by LeMasurier (1990h), realizing that these are minimum volumes and may far underestimate the actual volumes (Table 2).

Table 2.

MBL Volcanic Province central volcanoes

Volcanic fields (VF), central volcanoes (1–19), secondary calderasLatitude (°)Longitude (°)Summit elevation* (m asl)Elevation above ice level (m)Upstream ice elevation* (m asl)Downstream ice elevation* (m asl)Estimated volume (km3)CalderasCaldera diameter (km)Edifice-building age range
age ± 2 SD (Ma)
nPost-edifice-building age range
age ± 2 SD (Ma)
n§
Eastern MBL VF
1. Mount Murphy−75.3496−110.71012703200800200580No caldera9.46 ± 0.105.18 ± 0.18233.65 ± 0.160.609 ± 0.0275
2. Mount Takahe−76.2817−112.10933460210014001400780Summit caldera80.194 ± 0.0060.008 ± 0.00560.105 ± 0.0280.007 ± 0.013
3. Toney Mountain−75.8072−115.86003595240016001200550Summit caldera39.6 ± 1.011.0 ± 0.40.29 ± 0.12
Crary Mountains VF20001600400
4. Mount Rees−76.6744−118.0732270990020001600n.d.No caldera29.46 ± 0.247.62 ± 0.06157.01 ± 0.216.91 ± 0.262
5. Mount Steere−76.7250−117.78123558180020001600n.d.Summit calderanone8.66 ± 0.045.81 ± 0.0430
6. Mount Frakes−76.8078−117.69943654190020001800n.d.Summit caldera3 × 2.54.26 ± 0.0533.93 ± 0.030.032 ± 0.0105
Boyd Ridge−76.9503−116.8107237580018001600n.d.No calderanone2.17 ± 0.321.29 ± 0.033
7. Mount Siple VF−73.4313−126.764131003100001800Summit caldera4.50.230 ± 0.0080.171 ± 0.00520.746 ± 0.0360.008 ± 0.0882
ECR VF
8. Mount Hampton−76.4937−125.784733238002500250070Summit caldera6.5 × 5.511.43 ± 0.048.6 ± 1.0511.4 ± 1.210.7 ± 0.82
Whitney Peak−76.4577−125.97673003600Whitney Peak calderaindet.13.36 ± 0.054
9. Mount Cumming−76.6786−125.8061261230024002400naSummit caldera4.5 × 3.510.4 ± 1.010.0 ± 1.023.0 ± 0.41
10. Mount Hartigan−76.8196−126.02632811600n.d.naBoudette Peak caldera3.58.50 ± 0.666.02 ± 0.505
Tusing Peak−76.8717−126.07422652450n.d.Tusing Peak caldera3.58.36 ± 0.827.57 ± 0.603
11. Mount Sidley−77.0570−126.13464181220026002000250Summit caldera4.54.43 ± 0.0614.37 ± 0.064.24 ± 0.083
Weiss Peak−77.0352−126.0428Weiss Peak caldera2.55.77 ± 0.124.87 ± 0.0654.66 ± 0.104.51 ± 0.025
12. Mount Waesche−77.1686−126.89383292120024002000160Summit caldera1.51.0 ± 0.2<0.1 ± 05
Chang Peak−77.1201−126.76482920700Chang Peak caldera102.01 ± 0.101.09 ± 0.103
McCuddin Mountains VF
13. Mount Flint−75.7223−129.059026959002000200052Summit caldera2 × 320.46 ± 0.079.67 ± 0.203.75 ± 0.063
Ames Range VF252
14. Mount Andrus−75.8083−132.34742978140018001600115Summit caldera4.512.71 ± 0.0611.18 ± 0.19511.60 ± 0.80<0.13
15. Mount Kosciusko−75.7161−132.20002909130018001600107Summit caldera3.5 × 59.2 ± 1.31
16. Mount Kauffman−75.6308−132.376623648001800160030Summit caldera?indet.7.1 ± 1.51
Flood Range VF
17. Mount Bursey−76.0260−132.51542787100022002000290Summit caldera56.04 ± 0.4810.49 ± 0.121
Koerner Bluff−76.0078−132.9713Koerner Bluff caldera6 × 8.510.12 ± 0.349.31 ± 0.7438.56 ± 0.060.25 ± 0.032
18. Mount Moulton−76.0373−135.126230781700240017000325Summit caldera5.54.03 ± 0.1411.04 ± 0.041
Prahl Crag−76.0513−134.6589Prahl Crag caldera4.5 × 75.95 ± 0.051
Kohler Dome−76.0459−134.2753Kohler Dome caldera?5n.d.
19. Mount Berlin−76.0547−135.86553478210019001400200Summit caldera1.50.0279 ± 0.00640.01 ± 0.00537
Merrem Peak−76.0378−135.9588Merrem Peak Caldera2.5 × 10.578 ± 0.0090.143 ± 0.006160.214 ± 0.0186
Volcanic fields (VF), central volcanoes (1–19), secondary calderasLatitude (°)Longitude (°)Summit elevation* (m asl)Elevation above ice level (m)Upstream ice elevation* (m asl)Downstream ice elevation* (m asl)Estimated volume (km3)CalderasCaldera diameter (km)Edifice-building age range
age ± 2 SD (Ma)
nPost-edifice-building age range
age ± 2 SD (Ma)
n§
Eastern MBL VF
1. Mount Murphy−75.3496−110.71012703200800200580No caldera9.46 ± 0.105.18 ± 0.18233.65 ± 0.160.609 ± 0.0275
2. Mount Takahe−76.2817−112.10933460210014001400780Summit caldera80.194 ± 0.0060.008 ± 0.00560.105 ± 0.0280.007 ± 0.013
3. Toney Mountain−75.8072−115.86003595240016001200550Summit caldera39.6 ± 1.011.0 ± 0.40.29 ± 0.12
Crary Mountains VF20001600400
4. Mount Rees−76.6744−118.0732270990020001600n.d.No caldera29.46 ± 0.247.62 ± 0.06157.01 ± 0.216.91 ± 0.262
5. Mount Steere−76.7250−117.78123558180020001600n.d.Summit calderanone8.66 ± 0.045.81 ± 0.0430
6. Mount Frakes−76.8078−117.69943654190020001800n.d.Summit caldera3 × 2.54.26 ± 0.0533.93 ± 0.030.032 ± 0.0105
Boyd Ridge−76.9503−116.8107237580018001600n.d.No calderanone2.17 ± 0.321.29 ± 0.033
7. Mount Siple VF−73.4313−126.764131003100001800Summit caldera4.50.230 ± 0.0080.171 ± 0.00520.746 ± 0.0360.008 ± 0.0882
ECR VF
8. Mount Hampton−76.4937−125.784733238002500250070Summit caldera6.5 × 5.511.43 ± 0.048.6 ± 1.0511.4 ± 1.210.7 ± 0.82
Whitney Peak−76.4577−125.97673003600Whitney Peak calderaindet.13.36 ± 0.054
9. Mount Cumming−76.6786−125.8061261230024002400naSummit caldera4.5 × 3.510.4 ± 1.010.0 ± 1.023.0 ± 0.41
10. Mount Hartigan−76.8196−126.02632811600n.d.naBoudette Peak caldera3.58.50 ± 0.666.02 ± 0.505
Tusing Peak−76.8717−126.07422652450n.d.Tusing Peak caldera3.58.36 ± 0.827.57 ± 0.603
11. Mount Sidley−77.0570−126.13464181220026002000250Summit caldera4.54.43 ± 0.0614.37 ± 0.064.24 ± 0.083
Weiss Peak−77.0352−126.0428Weiss Peak caldera2.55.77 ± 0.124.87 ± 0.0654.66 ± 0.104.51 ± 0.025
12. Mount Waesche−77.1686−126.89383292120024002000160Summit caldera1.51.0 ± 0.2<0.1 ± 05
Chang Peak−77.1201−126.76482920700Chang Peak caldera102.01 ± 0.101.09 ± 0.103
McCuddin Mountains VF
13. Mount Flint−75.7223−129.059026959002000200052Summit caldera2 × 320.46 ± 0.079.67 ± 0.203.75 ± 0.063
Ames Range VF252
14. Mount Andrus−75.8083−132.34742978140018001600115Summit caldera4.512.71 ± 0.0611.18 ± 0.19511.60 ± 0.80<0.13
15. Mount Kosciusko−75.7161−132.20002909130018001600107Summit caldera3.5 × 59.2 ± 1.31
16. Mount Kauffman−75.6308−132.376623648001800160030Summit caldera?indet.7.1 ± 1.51
Flood Range VF
17. Mount Bursey−76.0260−132.51542787100022002000290Summit caldera56.04 ± 0.4810.49 ± 0.121
Koerner Bluff−76.0078−132.9713Koerner Bluff caldera6 × 8.510.12 ± 0.349.31 ± 0.7438.56 ± 0.060.25 ± 0.032
18. Mount Moulton−76.0373−135.126230781700240017000325Summit caldera5.54.03 ± 0.1411.04 ± 0.041
Prahl Crag−76.0513−134.6589Prahl Crag caldera4.5 × 75.95 ± 0.051
Kohler Dome−76.0459−134.2753Kohler Dome caldera?5n.d.
19. Mount Berlin−76.0547−135.86553478210019001400200Summit caldera1.50.0279 ± 0.00640.01 ± 0.00537
Merrem Peak−76.0378−135.9588Merrem Peak Caldera2.5 × 10.578 ± 0.0090.143 ± 0.006160.214 ± 0.0186

*Elevations from USGS 1:250 000 Anatarctic Topographic Map Series.

Estimated volume exposed above the ice sheet, from LeMasurier and Thomson (1990) 

n refers to the number of samples included in the edifice-building age range.

§n refers to the number of samples included in the post-edifice-building age range.

Boyd Ridge is included in this table because it is likely to be an additional central volcano, although it lacks a caldera and adequate exposure to identify it as such.

Other notes: the ‘?’ in the Caldera column indicates uncertainty about the idenfication of a caldera; indet., indeterminate.

Bracketing ages: italicized ages are conventional K–Ar ages, all other ages are 40Ar/39Ar ages; ages between bracketing ages are mostly 40Ar/39Ar ages but also include some K–Ar ages.

Central volcanoes (numbered 1–19) are either single centres or are coalesced centres, often identified by more than one caldera, indicated by inset.

LeMasurier (1972) proposed a generalized pattern of stratigraphic relationships among most central volcanoes that is characterized by a mafic ‘basal succession’ overlain by a felsic shield volcano, followed by late-stage post-edifice-building parasitic scoria cones and tuff cones situated on the shield flanks. In this model, the basal succession consists of subglacial to subaerial basaltic lavas and volcaniclastic rocks that form a platform that overlies a flat pre-Cenozoic basement. The basal succession includes many reported hyaloclastite deposits that were interpreted as erupted in subglacial environments (LeMasurier 1972; LeMasurier and Rex 1982, 1983; LeMasurier 1990h). The well-exposed Mount Murphy was used as a representative example, and was described as a basaltic shield volcano surmounted by a felsic shield volcano. Although this simple tripartite model of volcanoes built on a flat erosion surface may be accurate in places, detailed studies of the stratigraphy and evolution of some of the more dissected and better-exposed volcanoes indicate more complicated and varied histories, and a less flat erosion surface at many volcanoes (Panter et al. 1994; Wilch et al. 1999; Wilch and McIntosh 2000, 2002; Smellie 2001). These cases will be addressed in the individual descriptions of volcanoes.

As mentioned, post-edifice-building volcanoes are common in MBL and are superimposed on or are proximal to most of the central volcanoes. The number of preserved and exposed parasitic cones ranges from about two to 10 per central volcano. In a few cases, such as Mount Sidley, these cones are close in age to the main edifice-building interval (<0.4 myr younger); in other cases, such as Mount Cumming and Mount Bursey, the parasitic cones post-date main edifice construction by millions of years (Table 2). The parasitic volcanoes are located on the flanks of the central volcanoes and are typically erosional remnants of mafic scoria cones, characterized by partial crater rims or vent deposits and/or mixtures of welded and non-welded pyroclastic deposits with subordinate clastogenic or massive lavas. Some of the parasitic volcano lavas and tuffs contain large quantities of crustal and mantle xenoliths. Parasitic tuff cones and glaciovolcanic tuyas are less common and are identified by diagnostic lithofacies (Russell et al. 2014). Post-edifice-building parasitic intermediate and felsic domes, tuff cones, tuyas, and lavas occur at some locations. Post-edifice-building satellite volcanoes (not on the flanks of, but proximal to, major central volcanoes) and monogenetic volcanoes not proximal to the central volcanoes (e.g. in the Hobbs Coast and Jones and Hudson mountains volcanic fields) range from mostly intact to highly-eroded volcanoes, dominated by scoria cones, with fewer tuyas, tuff cones and lava domes.

A wide range of volcanic lithofacies has been documented in the MBLVG. Terminology used to describe the character and interpreted origin of the MBLVG rocks has changed over time, from more general in early studies (pre-1990) to more specific and detailed in later studies. In more recent studies there has been a trend toward more detailed non-genetic descriptions of the rocks (often following suggestions of McPhie et al. 1993) separate from the interpretations of the rock origins and palaeoenvironments (e.g. see Smellie et al. 1993; Wilch and McIntosh 2000, 2002, 2007; Smellie 2001). The more detailed lithofacies approach has led to reinterpretations of some sequences and a refinement in interpretations of others. An emphasis of the MBLVG physical volcanology research has been on differentiating rocks erupted in contact with the atmosphere with no influence of external water from rocks erupted into or in contact with ice or meltwater. Wilch and McIntosh (2000, 2002, 2007) used the general terms ‘wet’ and ‘dry’ to differentiate rocks erupted in contact with ice from those erupted subaerially. The ice-contact rocks are very important because they offer the potential to make inferences about the history of the WAIS at the time of eruptions. Smellie and Edwards (2016) discuss the lithofacies approach to describe and interpret ancient volcanic sequences in glacial environments in more detail. This method is particularly powerful because the alkaline rocks are readily datable using K–Ar or 40Ar/39Ar techniques.

Here, volcanic lithofacies are subdivided into three classes: primary coherent lithofacies; primary fragmental lithofacies; and secondary epiclastic lithofacies. The coherent lithofacies include lava and dykes (after McPhie et al. 1993). Coherent lava is the most common lithofacies found at the MBLVG outcrops. For this study, coherent lava lithofacies are subdivided into dominant subaerial (also referred to as ‘dry’) and subordinate subaqueous (or subglacial) (also referred to as ‘wet’) end members, based on the presence or absence of features (e.g. hyaloclastite breccia) characteristic of interactions with external water.

The subaerial lava lithofacies includes lavas with reddened brecciated bases and pāhoehoe tops or reddened brecciated tops that are interpreted to result from subaerial lava effusion and emplacement, without recognizable water or ice interaction. Many lavas, particularly felsic lavas, lack diagnostic features of subaerial or subaqueous environments. Unless these lavas are associated with other subaqueous deposits, they have been tentatively interpreted as subaerial lavas. Likewise, a common descriptor of MBL lavas in the LeMasurier and Thomson (1990) volume is ‘flow rock’, which we infer to be subaerial lavas, unless reinterpreted otherwise in subsequent studies. Other textural or morphological descriptors of subaerial lavas (such as compound, sheet, dome, clastogenic) are used when possible. Clastogenic lava, derived from agglutinated pyroclastic spatter, is a common subaerial lithofacies in both mafic and felsic end members, and shows pyroclastic and flowage textures. Intrusive bodies, including dykes, sills and irregular intrusions, are common at more deeply-eroded volcanoes.

The subglacial (or subaqueous) lava lithofacies show evidence of quenching, and are interpreted as water-cooled and generally associated with subaqueous eruptive or depositional environments (see Smellie and Edwards 2016). Two lava lithofacies, pillow lava and blocky or curvi-columnar-jointed sheet lava (see Smellie and Edwards (2016)), are common in several MBL sequences. Two variations of pillow lava lithofacies are recognized: compound nested pillows with minor interpillow hyaloclastite breccia; and lobe hyaloclastites composed of irregular pillow lobes with abundant (>10%) interpillow hyaloclastite breccia. Pillow lava lithofacies are common in MBL at Mount Rees, Mount Steere, Mount Murphy and Mount Takahe but are rare elsewhere. The blocky or curvi-columnar-jointed sheet lava includes slightly glassy, compound and simple lavas and intrusive bodies, with irregular to hackly jointing and rare crude pillow structures. In rare cases, lava apophyses locally deform bedding and exhibit hackly jointing. These apophyses are interpreted as dykes or lava that intruded wet volcaniclastic sediments.

For this study, primary fragmental lithofacies include ‘dry’ and ‘wet’ end members to differentiate deposits erupted and emplaced with no interaction between magma or lava and external water from those deposits that show evidence of external water interactions. The principal ‘dry’ end member is autobreccia. Subaerial autobreccias are ubiquitous in MBL, and are recognized by welding textures, reddening caused by deuteric oxidation and a lack of thick glassy ‘quenched’ margins. The principal ‘wet’ end member is hyaloclastite. Use of the term hyaloclastite requires some additional explanation because the term has been confusing in the literature when describing the MBLVG deposits. Early work on MBL volcanoes in LeMasurier and Thomson (1990 and references therein) followed a convention proposed by Fisher and Schmincke (1984) that hyaloclastites include glassy volcaniclastic deposits without regard to the origin of the glassy clasts. In other words, hyaloclastite included deposits formed by cooling contraction granulation, as well as by phreatomagmatic explosions. Hyaloclastites were typically characterized as palagonitized glassy breccia. In general, any palagonitized volcaniclastic deposits were referred to as hyaloclastites and were interpreted as evidence of past ice-sheet expansions. By contrast, following Smellie and Skilling (1994) and Wilch and McIntosh (2000, 2002), differentiated hyaloclastite formed from cooling contraction granulation in contact with water from hyalotuff (Honnorez and Kirst 1975) associated with explosive phreatomagmatic eruptions. In their usage, Wilch and McIntosh (2000, 2002) referred to locally reworked glassy material associated with pillow-lava sequences as reworked hyaloclastite. At several sites (e.g. Mount Takahe and Mount Rees), hyaloclastites are characterized as well stratified, suggesting that they were locally redeposited on underwater slopes by sediment gravity-flow processes. The definition of hyaloclastite proposed by White and Houghton (2006) only includes in situ deposits, whereas reworked deposits, common in lava-fed deltas associated with tuyas, are called breccia or lapilli tuff depending on grain size. In this chapter, we follow Wilch and McIntosh (2000, 2002) with the slightly less-constrained usage of the term hyaloclastite to include locally reworked hyaloclastite. The MBLVG hyaloclastites described here are likely to include peperite, a glassy breccia formed by lava or intrusions quenching and granulating in contact with wet sediment (Skilling et al. 2002). The distinction between hyaloclastite and peperite was not made in the field. In a few instances, it appears that the hyaloclastites and parent pillow lavas were moving downslope together and pillow lobes or fingers were injected up into the hyaloclastite breccia.

Other fragmental lithofacies (breccia, tuff breccia, lapilli tuff and tuff) differentiate deposits based on grain size and do not imply specific emplacement or eruption conditions. These lithofacies are dominated by two types: pyroclastic fall and dilute pyroclastic density current deposits. Pyroclastic fall deposits are products of magmatic or phreatomagmatic explosions and typically form crudely- to well-stratified deposits that mantle topography. A magmatic origin is inferred where pyroclasts are well sorted and exhibit uniform, moderate to high vesicularity and angular to fluidal shapes. Reddening by deuteric oxidation and welding are common features, and are interpreted as indications of close proximity to a subaerial vent (Walker and Croasdale 1971). The most common pyroclastic fall deposits are basaltic welded lapilli and bomb-rich outcrops at intact and eroded scoria cones, associated with mildly-explosive Strombolian-style eruptions. We make these interpretations using an ensemble of criteria, understanding the caveat that criteria to distinguish magmatic from phreatomagmatic fragmentation processes are not entirely diagnostic (White and Valentine 2016). A phreatomagmatic origin is inferred by relatively poor sorting, fine grain sizes, variable vesicularity, ash-coated or accretionary lapilli (in basalts), a predominance of sideromelane glass (in basalts) and blocky clast morphology (Wohletz 1983; Fisher and Schmincke 1984). In MBL, dilute turbulent pyroclastic density current deposits are identified by comparison to similar deposits described elsewhere (e.g. Wohletz and Sheridan 1983; Chough and Sohn 1990; Dellino and La Volpe 2000; Douillet et al. 2015) and typically form moderately- to poorly-sorted, planar, cross-stratified or massive beds that thicken and thin laterally. Individual clasts in these dilute pyroclastic density current (surge) deposits resemble those in phreatomagmatic fall deposits, except that they are often much more rounded and abraded due to turbulent lateral transport. In the past, some phreatomagmatic surge and fall deposits have been variably referred to as hyaloclastite (LeMasurier 1990g), hydroclastic tuff (LeMasurier and Rex 1990b) and hyalotuffs (Wilch and McIntosh 2000, 2002). These deposits commonly include signs of wet phreatomagmatic origin followed by subaerial deposition, such as ash-coated lapilli and bedding-plane sags caused by ballistic impacts. Pyroclastic density current and fall deposits are commonly interlayered. In western and central MBL, many fall and surge deposits contain clasts that exhibit a continuum of characteristics from phreatomagmatic explosivity (blocky shapes, a wide range of vesicularity) to magmatic explosivity (fluidal to cuspate shapes, high vesicularity).

Explosively-erupted felsic lithofacies are relatively rare compared to mafic counterparts but are, nonetheless, important in the eruptive histories of polygenetic central volcanoes of the MBLVP. Felsic welded pyroclastic fall deposits with variably-flattened fiamme were preserved in many caldera walls of the central volcanoes. At Mount Berlin, these caldera-wall pyroclastic fall deposits are associated with highly-explosive Plinian eruptions; and a similar eruption style is inferred at other calderas. Less commonly, felsic welded ignimbrites have been recognized at Mount Sidley and Mount Berlin.

Secondary epiclastic deposits are less common in MBL and include deposits formed by glacial, mass-flow and fluvial processes. Heterogeneous rock types, subrounded clast shapes, sedimentary structures and stratigraphic context characterize these sedimentary deposits. Despite the intraglacial setting of the MBL volcanic province, interbedded glacial tills/tillites and glacial–erosional unconformities are uncommon, limited to Mount Murphy, the Hobbs Coast Volcanic Field and Mount Aldaz.

Geochronology is an integral part of past studies of the MBLVG. This chapter summarizes published K–Ar and 40Ar/39Ar data, and presents unpublished 40Ar/39Ar data. Early work on dating MBLVG rocks relied on conventional K–Ar ages (summarized in LeMasurier and Thomson 1990), as well as on a few fission-track ages (Seward et al. 1980; Palais et al. 1988). Conventional K–Ar ages of MBLVG rocks (e.g. in the LeMasurier and Thomson 1990 volume and publications referenced therein) reported uncertainties at the 1σ level (66% confidence level). More recent studies use 40Ar/39Ar dating, and several studies presenting 40Ar/39Ar data report ages at the 2σ level (e.g. Wilch et al. 1999; Smellie 2001; Wilch and McIntosh 2002). In this paper, analytical 2σ level (95% confidence level) uncertainties are applied to all preferred ages.

In addition to correcting all uncertainties to 2σ, age corrections were applied to accommodate changes in standards and decay constants since the original calculation of ages. For example, the conventional K–Ar results prior to 1989 used a different K–Ar decay constant and the ages have been corrected to the now-accepted constant (Steiger and Jaeger 1977), following the method of Dalrymple (1979). Most of the 40Ar/39Ar results used the Fish Canyon sanidine (FCs) standard to monitor the neutron flux during irradiation but used now-outdated FCs ages in the calculation of sample ages. These samples were corrected to the accepted FCs age of Kuiper et al. (2008). Corrections applied to each sample are included in the Supplementary material Table S1.

In this study, 40Ar/39Ar ages are preferred over conventional K–Ar ages. Only 40Ar/39Ar ages are included in data tables for sites where both methods were applied; K–Ar ages are included for sites that have only been dated by the conventional K–Ar method. The reasons for the preference for 40Ar/39Ar ages over K–Ar ages are documented by Wilch (1997) and Wilch and McIntosh (2000, 2002, 2007), and are based on higher precision results, the ability to assess the reliability of results, the ability to date smaller samples including single crystals and the ability to date very young events.

Summaries of MBLVG age data are presented in a series of tables embedded in the volcano summaries. A compilation of more complete age data is available as Supplementary material. In total, results from 330 age analyses (273 40Ar/39Ar, 52 K–Ar and five fission track) are presented in this chapter, including 52 previously unpublished 40Ar/39Ar ages mostly from Mount Murphy and Mount Takahe. Complete 40Ar/39Ar datasets of new data from Mount Murphy and Mount Takahe are included in Supplementary material Tables S2 and S3.

At some localities, multiple samples are correlated to the same eruption event and a total of 198 eruption events are recognized: 80% of the eruption events are associated with the central volcanoes (Fig. 6a) and 20% are associated with minor isolated volcanoes (Fig. 6b). The number of samples and level of detail of the geochronology are quite variable at the MBLVG volcanoes. Well-dated, deeply-dissected volcanoes, including Mount Rees, Mount Steere, Mount Sidley and Mount Murphy, exhibit long-lived, >1 myr, eruption histories.

Fig. 6.

Geochronology of the MBLVG. (a) Central volcano ages. The upper plot shows individual ages of central volcanoes with 2σ (SD) uncertainties. Ages are colour coded by volcano and listed in the legend from youngest to oldest. The lower plot shows relative probability distribution of ages. (b) Minor volcano ages. Plots organized similar to that in (a).

Fig. 6.

Geochronology of the MBLVG. (a) Central volcano ages. The upper plot shows individual ages of central volcanoes with 2σ (SD) uncertainties. Ages are colour coded by volcano and listed in the legend from youngest to oldest. The lower plot shows relative probability distribution of ages. (b) Minor volcano ages. Plots organized similar to that in (a).

The physical volcanology and geochronology of central and satellite volcanoes of the MBLVG is summarized here. The descriptions of the MBLVG volcanoes are organized by volcanic province and volcanic field, as listed in Figure 3. Table 2 summarizes key physical characteristics of the central volcanoes of the MBLVP, as well as age ranges of edifice- and post-edifice-building eruptions. Rock compositional types for each volcano and their petrology are provided by Panter et al. (2021).

The Eastern MBL Volcanic Field consists of three isolated central volcanoes (Mount Murphy, Mount Takahe and Toney Mountain), satellite centres near Mount Murphy and isolated centres in the Kohler Range.

Mount Murphy is a deeply-dissected polygenetic coastal volcano with a summit elevation of 2703 m asl (Fig. 7). Mount Murphy covers an area of about 35 × 45 km, and has an irregular shape with multiple peaks and elongate resistant spurs that are likely to have resulted from, or are enhanced by, glacial erosion. Large cirques are eroded into the volcano and multiple alpine glaciers flow down the volcano flanks (Andrews and LeMasurier 1973). Mount Murphy forms an obstruction to regional ice flow, with inland ice levels at c. 800 m asl and coastal ice levels at c. 200 m asl. The WAIS abuts the south side of Mount Murphy, and drains around the west and east sides of Mount Murphy through Pope Glacier and Haynes Glacier, respectively. These glaciers flow into the Crosson Ice Shelf located just north of Mount Murphy. Mount Murphy is one of the few MBLVP central volcanoes that lacks a recognizable summit caldera, although some of the tops of cirque headwalls resemble caldera rims.

Fig. 7.

Geology of Mount Murphy volcano in the Eastern MBL Volcanic Field, with volcanic rock outcrops, lithofacies and ages. The thick dashed line shows the outline of the volcano at the ice-sheet surface. Note the elevation difference between the upstream (800 m asl) and downstream (200 m asl) sides of the volcano. Google Earth Pro image accessed June 2019. Image sources: Landsat from United States Geological Survey and 2019 Digital Globe. The variable resolution image was downloaded from Google Earth Pro and processed in Adobe Photoshop. Processing included conversion to black and white, and adjustment of brightness and contrast to enhance uniformity of surface tones.

Fig. 7.

Geology of Mount Murphy volcano in the Eastern MBL Volcanic Field, with volcanic rock outcrops, lithofacies and ages. The thick dashed line shows the outline of the volcano at the ice-sheet surface. Note the elevation difference between the upstream (800 m asl) and downstream (200 m asl) sides of the volcano. Google Earth Pro image accessed June 2019. Image sources: Landsat from United States Geological Survey and 2019 Digital Globe. The variable resolution image was downloaded from Google Earth Pro and processed in Adobe Photoshop. Processing included conversion to black and white, and adjustment of brightness and contrast to enhance uniformity of surface tones.

Most of the work at Mount Murphy has concentrated on the well-exposed SW ridge and on satellite nunataks west of the main edifice. There are many outcrops at Mount Murphy that have not been visited, especially on the eroded spurs, slopes and cliffs away from the west flank. Volcanic rocks on the western flank of Mount Murphy crop out between <400 and 2446 m asl. Several satellite nunataks are exposed as interfluves in Pope Glacier, just west of the main edifice. The lowest rock outcrops at Mount Murphy between c. 200 and 400 m asl expose the bedrock underlying the volcano (LeMasurier et al. 1990c).

LeMasurier et al. (1990c) characterized Mount Murphy as a basaltic shield surmounted by a much smaller felsic shield, based on 5° dip angles, on slope-forming basanitic lavas as high as 1900 m asl, and a transition to felsic rocks above that elevation. The total original volume of Mount Murphy is estimated at 580 km3 (LeMasurier et al. 1990c). The published descriptions of Mount Murphy have largely focused on volcanic sequences along the lower SW flank of Mount Murphy, at Sechrist Peak, and at three satellite nunataks, west and SW of the central volcano (McIntosh et al. 1985; LeMasurier et al. 1994; Smellie 2001; Wilch and McIntosh 2002) (Fig. 7). Stratigraphic sequences at these localities are composed of intercalated volcanic and glacial deposits that provide evidence of higher palaeo-ice levels during Miocene–Pleistocene times (LeMasurier et al. 1994; Wilch and McIntosh 2002). Here, we expand upon these interpretations with new field and 40Ar/39Ar data from additional Mount Murphy localities. In total, 26 new 40Ar/39Ar ages are presented along with descriptions of the rock units (Table 3). A summary of the volcanic geology and geochronology of the main edifice is presented first, followed by a description of the satellite nunataks. The overall evolution was not discussed.

Table 3.

Mount Murphy and Mount Kohler range summary age table

Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Mount Murphy
Murphy Shield Building, Bucher Peak
W85-060A5.18 ± 0.18North side, lava, 60 m below summit ice cap, c. 2390 m21
90-72A7.14 ± 0.05West cliff mesa, light-grey lava, 2% plag21
90-001A7.64 ± 0.14SW ridge, bomb interior, highest cone, trachyte–benmoreite?, 10% xtal, cpx, plag, amph?, porphy lava, spatter/flow, 5% pyx + ol + plag21
90-005A8.49 ± 0.46SW ridge, lava underlying tillite21
90-34A8.33 ± 0.04Main cliff, 50 m-thick felsic dome 725–775 m asl18
90-37A8.91 ± 0.20Main cliff, thin (subaerial?) lava at 685 m18
90-37A8.97 ± 0.15Main cliff, thin (subaerial?) lava at 685 m18
90-37A8.95 ± 0.13Mean age (n = 2)
90-39A9.38 ± 0.14Main cliff, lava assoc. w/hyaloclastite 555 m18
90-39A9.15 ± 0.12Main cliff, lava assoc. w/hyaloclastite 555m18
90-39A9.25 ± 0.24Mean age (n = 2)
90-33A9.19 ± 0.09Main cliff, 1.5 m lava (striated), 527 m18
90-33A9.51 ± 0.12Main cliff, 1.5 m lava (striated), 527 m18
90-33A9.31 ± 0.32Mean age (n = 2)
90-50A9.47 ± 0.13West cliff, pillow 3 m above tillite, 510 m18
90-50A9.44 ± 0.13West cliff, pillow 3 m above tillite, 510 m18
90-509.46 ± 0.10Mean age (n = 2)18
Sechrist Peak
90-69A0.607 ± 0.086West flank21
90-139A0.609 ± 0.028Dense intrusive interior21
0.609 ± 0.027Mean age (n = 2)
Murphy post-shield building, Bucher Peak
90-007A1.86 ± 0.09Basaltic bomb, lowest cinder cone21
W85-045A2.72 ± 0.08Basaltic in situ flow21
90-002A3.65 ± 0.16Basaltic flow, middle xeno-bearing cone21
Murphy shield building, Bucher south face
90-145A5.96 ± 0.30Dense black bomb interior, high in spatter-fed flow21
90-143A7.28 ± 0.12Dense spatter-fed lava, 15% pyx + plag + ol21
90-144A7.04 ± 0.10Green trachtye fragment21
90-146A8.71 ± 0.16Trachyte, west end of Hawkins–Murphy saddle21
Kay Peak
90-123A8.94 ± 0.03Aphanitic vesicular lava21
90-112A8.90 ± 0.11Second lowest unit, trachytic? lava, porphy w/4% ol + cpx + plag21
90-115A6.93 ± 0.49Crystal-rich basaltic lava, 15% ol + cpx + plag21
Grew Peak
90-127A4.31 ± 0.55Aphyric dyke in epiclast/hyaloclastite21
90-130A5.07 ± 0.16Aphyric pillow basalt21
90-131A5.81 ± 0.44Porphy lava, ol? plag, pyx, within hyalo/pillow unit21
Eisberg Head
90-133A5.51 ± 0.40J.L. Smellie sample MB58.3B21
Callendar Ridge
90-153A6.22 ± 0.35Lower lava, aphyric, coarse plag matrix21
90-149A5.93 ± 0.35Lowest exposed lava, aphyric21
Hawkins Peak
90-148A5.28 ± 0.06Bomb interior, 1% plag21
Satellite nunataks
Icefall Nunatak
90-47A6.60 ± 0.13High subaerial lava18
90-48A6.89 ± 0.20Lobe in low hyaloclastite18
Turtle Peak
90-94A4.76 ± 0.15Very vesicular pāhoehoe top18
90-92A5.72 ± 0.23Upper pillow18
90-87A5.95 ± 0.60Lower-flow foot breccia, angular clasts18
Hedin Nunatak
90-99A6.28 ± 0.24Subaerial lava, upper tuya18
90-110A6.58 ± 0.12Lowest tindar18
Dorrel Rock
90-105A35.50 ± 0.12Gabbro intrusion21
90-105A34.83 ± 0.45Gabbro intrusion21
90-105A34.92 ± 0.12Gabbro intrusion21
A35.20 ± 0.42Mean age (n = 3)
60AA>34.46 ± 0.22Gabbro intrusion13
60AA>37.05 ± 0.26Gabbro intrusion13
60CA33.93 ± 0.24Dyke13
60DA35.81 ± 0.32Dyke13
84K10.1 ± 3.4Thin lava on basement, proximal to tuff breccia1
Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Mount Murphy
Murphy Shield Building, Bucher Peak
W85-060A5.18 ± 0.18North side, lava, 60 m below summit ice cap, c. 2390 m21
90-72A7.14 ± 0.05West cliff mesa, light-grey lava, 2% plag21
90-001A7.64 ± 0.14SW ridge, bomb interior, highest cone, trachyte–benmoreite?, 10% xtal, cpx, plag, amph?, porphy lava, spatter/flow, 5% pyx + ol + plag21
90-005A8.49 ± 0.46SW ridge, lava underlying tillite21
90-34A8.33 ± 0.04Main cliff, 50 m-thick felsic dome 725–775 m asl18
90-37A8.91 ± 0.20Main cliff, thin (subaerial?) lava at 685 m18
90-37A8.97 ± 0.15Main cliff, thin (subaerial?) lava at 685 m18
90-37A8.95 ± 0.13Mean age (n = 2)
90-39A9.38 ± 0.14Main cliff, lava assoc. w/hyaloclastite 555 m18
90-39A9.15 ± 0.12Main cliff, lava assoc. w/hyaloclastite 555m18
90-39A9.25 ± 0.24Mean age (n = 2)
90-33A9.19 ± 0.09Main cliff, 1.5 m lava (striated), 527 m18
90-33A9.51 ± 0.12Main cliff, 1.5 m lava (striated), 527 m18
90-33A9.31 ± 0.32Mean age (n = 2)
90-50A9.47 ± 0.13West cliff, pillow 3 m above tillite, 510 m18
90-50A9.44 ± 0.13West cliff, pillow 3 m above tillite, 510 m18
90-509.46 ± 0.10Mean age (n = 2)18
Sechrist Peak
90-69A0.607 ± 0.086West flank21
90-139A0.609 ± 0.028Dense intrusive interior21
0.609 ± 0.027Mean age (n = 2)
Murphy post-shield building, Bucher Peak
90-007A1.86 ± 0.09Basaltic bomb, lowest cinder cone21
W85-045A2.72 ± 0.08Basaltic in situ flow21
90-002A3.65 ± 0.16Basaltic flow, middle xeno-bearing cone21
Murphy shield building, Bucher south face
90-145A5.96 ± 0.30Dense black bomb interior, high in spatter-fed flow21
90-143A7.28 ± 0.12Dense spatter-fed lava, 15% pyx + plag + ol21
90-144A7.04 ± 0.10Green trachtye fragment21
90-146A8.71 ± 0.16Trachyte, west end of Hawkins–Murphy saddle21
Kay Peak
90-123A8.94 ± 0.03Aphanitic vesicular lava21
90-112A8.90 ± 0.11Second lowest unit, trachytic? lava, porphy w/4% ol + cpx + plag21
90-115A6.93 ± 0.49Crystal-rich basaltic lava, 15% ol + cpx + plag21
Grew Peak
90-127A4.31 ± 0.55Aphyric dyke in epiclast/hyaloclastite21
90-130A5.07 ± 0.16Aphyric pillow basalt21
90-131A5.81 ± 0.44Porphy lava, ol? plag, pyx, within hyalo/pillow unit21
Eisberg Head
90-133A5.51 ± 0.40J.L. Smellie sample MB58.3B21
Callendar Ridge
90-153A6.22 ± 0.35Lower lava, aphyric, coarse plag matrix21
90-149A5.93 ± 0.35Lowest exposed lava, aphyric21
Hawkins Peak
90-148A5.28 ± 0.06Bomb interior, 1% plag21
Satellite nunataks
Icefall Nunatak
90-47A6.60 ± 0.13High subaerial lava18
90-48A6.89 ± 0.20Lobe in low hyaloclastite18
Turtle Peak
90-94A4.76 ± 0.15Very vesicular pāhoehoe top18
90-92A5.72 ± 0.23Upper pillow18
90-87A5.95 ± 0.60Lower-flow foot breccia, angular clasts18
Hedin Nunatak
90-99A6.28 ± 0.24Subaerial lava, upper tuya18
90-110A6.58 ± 0.12Lowest tindar18
Dorrel Rock
90-105A35.50 ± 0.12Gabbro intrusion21
90-105A34.83 ± 0.45Gabbro intrusion21
90-105A34.92 ± 0.12Gabbro intrusion21
A35.20 ± 0.42Mean age (n = 3)
60AA>34.46 ± 0.22Gabbro intrusion13
60AA>37.05 ± 0.26Gabbro intrusion13
60CA33.93 ± 0.24Dyke13
60DA35.81 ± 0.32Dyke13
84K10.1 ± 3.4Thin lava on basement, proximal to tuff breccia1

Notes: method A is 40Ar/39Ar; method K is K–Ar. Abbreviations: amph, amphibole; cpx, clinopyroxene; hyalo, hyaloclastite; ol, olivine; plag, plagioclase; porphy, porphyritic; pyx, pyroxene.

The main edifice that forms the highest and most voluminous part of Mount Murphy is deeply dissected by glacial erosion on the west and south sides, offering exposures of the interior of the volcano. The most complete sequence is exposed along the ridge that extends SW from Bucher Peak (2446 m asl), over Sechrist Peak (1350 m asl) to the base of volcano near 400 m asl (Figs 5 & 6). Compositions within this stratigraphic interval are almost all basanitic, with the exception of one 50 m-thick trachyte lava between 725 and 775 m asl, and several other trachytic lavas near the top of the sequence (1980–2390 m asl). 40Ar/39Ar ages for this sequence range from 9.46 ± 0.10 Ma for a pillow lava near the base to 5.18 ± 0.18 Ma for a trachytic lava just below Bucher Peak summit. A post-edifice-building Pleistocene tuff cone at Sechrist Peak is superimposed on the sequence; the late-stage tuff cone is considered after the description of the main stratigraphic sequence.

The geology exposed along this SW ridge is summarized in a composite stratigraphic section (Fig. 8: modified and expanded from Wilch and McIntosh 2002). Between 400 and 725 m asl, compositions are basanitic, and the sequence is dominated by pillow lavas, hyaloclastite breccias, bedded hyalotuffs reworked by traction currents, and complexly-jointed subaqueous lavas (Kubbaberg type) (Fig. 9a, b). These lithofacies are interlayered with subordinate amounts of subaerially-erupted hydrothermally-altered or palagonitized Strombolian tuff and welded bomb sequences (Fig. 9c), and columnar-jointed subaerial lavas with pāhoehoe-type fluidal top surfaces and oxidized basal breccias. At least three glacial erosion surfaces occur in the lower sequence, each associated with a polished and striated underlying surface, tillite and/or striated clasts, (Fig. 9d, e) (McIntosh et al. 1985; LeMasurier et al. 1994; Smellie 2001, 2008; LeMasurier 2002; Wilch and McIntosh 2002). Large irregular basanitic water-cooled lavas and intrusive or incursive bodies occur within the section, locally deforming and shearing stratified reworked hyalotuffs (Fig. 9f). Six 40Ar/39Ar ages from this interval range from 9.46 to 8.96 Ma (Wilch and McIntosh 2002). All workers agree that this sequence records complex interactions between glacial ice and a growing polygenetic volcano but there is incomplete agreement about the thickness of the interacting ice. McIntosh et al. (1985) and Wilch and McIntosh (2002) interpreted the interlayered subglacial and subaerial lithofacies as a record of the fluctuations of a regional ice sheet. Smellie (2001, p. 10; 2008) interpreted the lower Murphy sequence as a ‘subglacial “sheet-flow type” formed when the slopes of the volcano were mantled by relatively thin “ice” (probably mainly firn and/or snow <100 m thick)’. LeMasurier (2002) described the sequences as a series of complex lava-fed deltas, which have since been interpreted as ‘a‘ā-lava-fed deltas (Smellie and Edwards 2016). The lower Murphy sequence apparently records eruptive environments that varied from thin, temperate, mantling ice (yielding glacial unconformities, tillites, subaqueous lavas, and fluvially-reworked hyalotuffs and hyaloclastites), to regional ice (yielding lava-fed deltas) and subaerial emergence (yielding welded and non-welded Strombolian deposits and columnar-jointed lavas with oxidized basal breccias). Interpreting regional ice-sheet levels from the lower Sechrist Peak–Bucher Peak Ridge sequence is not straightforward, in part because of uncertainties regarding ice thickness but also because of potential feedback effects on the palaeo-ice-sheet level by the growing Mount Murphy edifice, as further discussed in a subsequent section (‘Synthesis of volcanic records of glaciation in the WAIS’).

Fig. 8.

Composite stratigraphic section of the Bucher–Sechrist Ridge, Mount Murphy, Eastern MBL Volcanic Field. The composite section is based on a traverse shown in Figure 7. The figure is an expanded and updated version of a figure in Wilch and McIntosh (2002).

Fig. 8.

Composite stratigraphic section of the Bucher–Sechrist Ridge, Mount Murphy, Eastern MBL Volcanic Field. The composite section is based on a traverse shown in Figure 7. The figure is an expanded and updated version of a figure in Wilch and McIntosh (2002).

Fig. 9.

Photographs of outcrops and lithofacies at Mount Murphy, Eastern MBL Volcanic Field. (a) Sequence near the base of the SW ridge, Mount Murphy, showing diamictite and bedded hyalotuffs, between two striated, polished glacial unconformities. The upper unconformity is overlain by a thick complexly-jointed, water-cooled, Kubbaberg-type lava. The bedded hyalotuffs were reworked and deposited by traction currents, probably water moving downslope in cavities or meltwater channels beneath warm-based, relatively thin, slope-mantling ice. Annotated photograph from Smellie (2008), with permission from Elsevier © 2008. (b) Tightly nested pillow lavas near the base the section along the lower SW ridge, Mount Murphy, formed by lava emplaced within meltwater-filled subglacial chambers. (c) Hydrothermally altered or palagonitized Strombolian welded bomb sequence, lower SW ridge, Mount Murphy. The flattened and mutually molded forms of the vesicular bombs indicate deposition of molten pyroclasts in a subaerial environment. (d) Polished, striated top of subaerially-erupted basanitic lava overlain by younger subaerially-erupted basanitic lava, lower SW ridge, Mount Murphy. This sequence attests to alternating subaerial and subglacial environments. (e) Polished, striated basanitic clast in glacial till, evidence for wet-based glaciation. A 15 cm pencil is for scale. Till overlies a striated unconformity near the base of the lower SW ridge of Mount Murphy. (f) Bedded hyalotuffs deformed and sheared by overlying or incursive basanitic water-cooled lavas, lower SW ridge of Mount Murphy. (g) Vertical and overturned beds of hyalotuff at Sechrist Peak, Mount Murphy, evidence of slope failure causing soft-sediment deformation during accumulation in a sub-ice-sheet meltwater chamber. (h) Glacially-eroded surface at the south end of the upper surface of Turtle Peak. This polished and striated surface with overlying granitic erratic boulders records post-volcanic overriding and erosion by the WAIS.

Fig. 9.

Photographs of outcrops and lithofacies at Mount Murphy, Eastern MBL Volcanic Field. (a) Sequence near the base of the SW ridge, Mount Murphy, showing diamictite and bedded hyalotuffs, between two striated, polished glacial unconformities. The upper unconformity is overlain by a thick complexly-jointed, water-cooled, Kubbaberg-type lava. The bedded hyalotuffs were reworked and deposited by traction currents, probably water moving downslope in cavities or meltwater channels beneath warm-based, relatively thin, slope-mantling ice. Annotated photograph from Smellie (2008), with permission from Elsevier © 2008. (b) Tightly nested pillow lavas near the base the section along the lower SW ridge, Mount Murphy, formed by lava emplaced within meltwater-filled subglacial chambers. (c) Hydrothermally altered or palagonitized Strombolian welded bomb sequence, lower SW ridge, Mount Murphy. The flattened and mutually molded forms of the vesicular bombs indicate deposition of molten pyroclasts in a subaerial environment. (d) Polished, striated top of subaerially-erupted basanitic lava overlain by younger subaerially-erupted basanitic lava, lower SW ridge, Mount Murphy. This sequence attests to alternating subaerial and subglacial environments. (e) Polished, striated basanitic clast in glacial till, evidence for wet-based glaciation. A 15 cm pencil is for scale. Till overlies a striated unconformity near the base of the lower SW ridge of Mount Murphy. (f) Bedded hyalotuffs deformed and sheared by overlying or incursive basanitic water-cooled lavas, lower SW ridge of Mount Murphy. (g) Vertical and overturned beds of hyalotuff at Sechrist Peak, Mount Murphy, evidence of slope failure causing soft-sediment deformation during accumulation in a sub-ice-sheet meltwater chamber. (h) Glacially-eroded surface at the south end of the upper surface of Turtle Peak. This polished and striated surface with overlying granitic erratic boulders records post-volcanic overriding and erosion by the WAIS.

From 725 to 850 m asl, the Sechrist Peak–Bucher Peak Ridge includes subaerial basalt and trachyte–benmoreite lavas (LeMasurier et al. 1990c). An approximately 50 m-thick columnar-jointed trachyte lava between 725 and 775 m asl is dated to 8.33 ± 0.04 Ma (Wilch and McIntosh 2002) (Table 3). A basanitic lava near 850 m asl is dated at 7.14 ± 0.05 Ma. From 850 to 1750 m asl, outcrops are sparse and sporadic, and consist of pyroclastic deposits and a few subordinate lavas. Poor exposures and variable dips within this interval hinder determinations of the stratigraphic relationships between these units and the well-exposed, stratigraphically-coherent, lower part of the composite section, described above. Some of the topographically higher units are older than topographically lower units, and other units are significantly younger, apparently overlying eroded older units. Some of these younger units are post-edifice-building basanites ranging in age from 3.65 ± 0.16 to 0.61 ± 0.03 Ma. The highest pyroclastic deposit in this interval is a 7.64 ± 0.14 Ma trachytic Strombolian pyroclastic deposit near 1750 m asl, possibly associated with the main edifice-building phase of Mount Murphy.

The uppermost part of the Mount Murphy SW ridge sequence, exposed between 1980 and 2390 m asl, consists of a series six or more basanitic–trachytic subaerial lavas. The lowest lava near 1980 m asl, poorly dated at 8.49 ± 0.46 Ma, has a polished and striated upper surface, overlain by a 1 m-thick sedimentary unit, possibly of glacial origin, composed of dipping (25° downslope) laminated clays containing 0.5–10 cm pebbles; in some cases themselves polished and striated. This laminated deposit contains recycled Neogene microfossils; it is not clear whether these were emplaced during glaciation or later by aeolian processes (LeMasurier et al. 1994). These glacial features suggest erosion and deposition by wet-based local ice on the slopes of Mount Murphy. The uppermost lava, exposed near 2390 m asl and covered by a small summit ice cap at the top of Bucher Peak, is basanitic and dated at 5.18 ± 0.18 Ma.

The main edifice of Mount Murphy has additional very limited outcrops south and east of the summit area between 1600 and 2400 m asl. Basanitic pyroclastic deposits and trachytic lavas exposed in these two areas range in age from 8.71 ± 0.16 to 5.96 ± 0.30 Ma.

Between 1050 and 1350 m asl, along Sechrist Peak–Bucher Peak Ridge, a sequence of deformed tephra-dominated tuya deposits is inset against Murphy edifice-building lavas. Two 40Ar/39Ar ages of 0.61 ± 0.09 and 0.61 ± 0.03 Ma indicate that the Sechrist Peak sequence formed several million years after Mount Murphy's main edifice-building period. Much of the sequence consists of thin beds of fine vesicular hyalotuffs, in many cases graded, which may represent volcanically-generated turbidites that accumulated within a subglacial meltwater cavity. These beds are intruded by dykes and irregular intrusive or incursive lava bodies that are locally pillowed at the margins or are brecciated and intermixed with the hyalotuffs. Some of the bedding is strongly deformed, in some cases folded and tilted to vertical or overturned orientations (Fig. 9g). This soft-sediment deformation is probably related to intrusions and slope failures resulting from the changing geometry of the water-filled chamber during and after emplacement. Towards the top of Sechrist Peak, the bedded hyalotuffs contain accretionary lapilli, bomb fragments, reddened Strombolian scoria and local subaerial lavas that suggest the eruptive vent became emergent. The deposits at Sechrist Peak probably erupted below and near the surface of a formerly higher WAIS. The passage zone of the Sechrist emergent sequence is not well defined but is located about 500 m above the local current 800 m elevation of the WAIS, just upstream of Sechrist Peak. The currently exposed surfaces of some of the more competent Sechrist lavas are striated and littered with exotic granitic erratics, suggesting glacial overriding subsequent to 0.61 Ma.

Kay Peak tops a spur that extends to the NNW from the summit of Mount Murphy (Fig. 7). Several outcrops on the eroded SE face of Kay Peak expose a volcanic sequence overlying a smooth, undulating and glacially-striated unconformity eroded into basement metagabbro. Lavas in the sequence at Kay Peak range in age from 8.90 ± 0.11 Ma near the base of the sequence to 6.93 ±0.49 Ma near the top. The basement contact, exposed at elevations as low as 640 m asl, is discontinuously overlain by 0.5 m of crudely-bedded volcanic-rich till containing rounded and striated clasts. The till is in turn overlain by sequence of pillow lavas, massive lavas and pillow hyaloclastites. The lowest lava has well-developed pillows throughout its vertical extent. The upper surface of this pillow lava and most of the overlying massive lavas are eroded, with an undulatory striated top surface. Eroded top surfaces are locally overlain by thin layers of volcanic-rich tillite or crudely-bedded hyaloclastite or reworked hyaloclastic sediments composed of glassy clasts probably derived from the quenching of the overlying lava. Clasts of pre-volcanic basement were not observed in the tillite or bedded sediments. A basanitic lava higher in the sequence has well-developed columnar jointing, and very limited development of pillows and hyaloclastite at the base of the columns. The top of the columnar lava is dramatically glacially sculpted into polished striated knobs with more than 20 m of relief. This well-developed glacial-erosion surface is in turn overlain by sequence of volcaniclastic rocks containing intact volcanic bombs, in turn capped by a deposit of Strombolian bombs and lapilli grading upwards into a black vesicular lava with a locally reddened, oxidized breccia on the upper surface. Minor localized pillows and hyaloclastite at the base of this otherwise subaerial lava is interpreted as evidence of limited interaction with snow or moisture in hollows along the underlying glacial surface. Such ‘wet’ to ‘dry’ upward transitions are seen in many lavas at Mount Rees and Mount Steere in the Crary Mountains.

The sequence at Kay Peak is interpreted to have been emplaced in a glacial valley on a bedrock surface adjacent to the slope of developing Mount Murphy. The valley experienced alternating periods of glacial erosion, emplacement of subglacial hyaloclastites and both subglacial and subaerial lavas, and fluvial reworking of hyalotuff and hyaloclastite debris. Similar sequences have been observed in deposits formed beneath valley glaciers in Iceland and elsewhere (Walker and Blake 1966; Smellie et al. 1993).

Grew Peak is located on a spur that extends NNE from the Mount Murphy summit (Fig. 7). Grew Peak, like Kay Peak, exposes a volcanic sequence overlying a pre-volcanic glacial-erosion surface cut into basement metasediments that extends as low as 840 m asl. The glacial unconformity consists of a polished striated surface at the base of a U-shaped valley cut into basement rock, subsequently filled by basanitic lavas and hyaloclastites. The glacial unconformity is directly overlain by 5 m of crudely-bedded, fines-poor basanitic lapilii tuff, containing angular fragments of basanitic lava. The sediment is in turn overlain locally by a massive basanitic lava dated at 5.81 ± 0.44 Ma. This is overlain by a more extensive hyaloclastite sequence containing pillows and lobes of basanitic lava as large as 6 m in length with quenched glassy margins. A pillow from this sequence yielded an age of 5.07 ± 0.16 Ma. Like the sequence at Kay Peak, the sequence at Grew Peak is likely to have formed by lavas flowing down an ice-filled valley on the slope of Mount Murphy. The sequence at Grew Peak is cut by basanitic dykes dated at 4.31 ± 0.55 Ma.

Callender Peak is located along the mostly snow- and ice-covered spur that extends ENE from the summit of Mount Murphy (Fig. 7). Outcrops near Callender Peak and the adjacent ridge extending to the north expose two lavas: a lower aphanitic basanite dated at 6.22 ± 0.35 Ma; and an upper porphyritic basanite dated at 5.93 ± 0.35 Ma. The upper lava contains 15% phenocrysts including plagioclase as large as 2 cm in diameter. Both lavas are vesicular, with brecciated, reddened flow margins indicating subaerial emplacement. The upper surfaces of exposures at Callender Peak are mantled with granitic and gneissic erratics, which indicate overriding by the regional ice sheet.

Hawkins Peak located SE of the summit of Mount Murphy has very limited exposures due to extensive ice cover. Two outcrops were visited. The lower outcrop near 1400 m asl is an agglutinated basaltic pyroclastic deposit containing red-oxidized lapilli and bombs to 1 m in diameter, and abundant ultramafic peridotite and pyroxenite xenoliths. A higher outcrop near 1600 m is a similar deposit of strongly-welded reddened bombs and lapilli with sparse xenoliths. A date from this unit is 5.28 ± 0.06 Ma, suggesting that it represents late edifice building rather than post-edifice-building activity.

There are three complex volcanic nunataks and one hypabyssal plutonic nunatak near Mount Murphy. The three volcanic nunataks are Icefall Nunatak, Hedin Nunatak and Turtle Peak, located 2–5 km west of the base of Mount Murphy on a north–south-orientated glacial interfluve between Pope Glacier to the west and an unnamed glacier to the east (Fig. 7). The plutonic nunatak is Dorrel Rock, located 5 km south of the volcanic nunataks. Of the volcanic nunataks, only Icefall Nunatak has been described in detail (Smellie 2001). General descriptions of Hedin Nunatak and Turtle Peak have been presented in several papers (McIntosh et al. 1985; LeMasurier et al. 1994; Smellie 2001; Wilch and McIntosh 2002; Smellie and Edwards 2016). The three volcanic nunataks are classified complex tuyas (Russell et al. 2014), each consisting of stacked sequences of multiple lava-fed deltas and associated source vents that accumulated in englacial meltwater lakes within a formerly thicker regional ice sheet.

Smellie (2001) provided a detailed lithofacies analysis and reconstruction of the volcanic history of Icefall Nunatak that includes consideration of the syneruptive ice thickness and hydrological conditions. The volcanic reconstruction is based on detailed mapping, identification and interpretation of 12 volcanic lithofacies and three distinct unconformities exposed in a 200 m-high by 900 m-long bluff section. Smellie (2001) interpreted Icefall Nunatak as a polygenetic basaltic volcanic centre that erupted between 6.89 ± 0.10 and 6.61 ± 0.08 Ma (Smellie 2001). At least five source vents were suggested on the basis of variable phenocryst content or stratigraphic context. Three distinct eruptive stages are differentiated by erosion surfaces and lithofacies assemblages. Deposits associated with Stage I have a maximum thickness of 60 m. Stage I is dominated by a monomict, clast-supported, lithic (juvenile) breccia composed mostly of holocrystalline lapilli to block-size clasts, with less abundant non-vesicular sideromelane clasts. Lenses of blocky-jointed and pillow lava are exposed lower in the section. The lavas and breccia are overlain by two thin (1–6 m) variably scoriaceous lavas separated by 4 m of a massive tuff breccia. The lavas transform laterally into scoriaceous breccia. The uppermost (and youngest) part of the Stage I sequence includes stratified lapilli tuff deposits which contain some highly vesicular clasts. Stage I deposits are inferred to have been emplaced in an ice-contact englacial lake that formed in response to heat from magma and an active eruption. The breccia and blocky-jointed pillow lava are all consistent with subaqueous eruption and deposition. The vesiculation observed in the uppermost volcaniclastic sediments suggests a reduction of hydraulic pressure over the vent and lower water levels. Smellie (2001) interpreted these uppermost deposits as turbidite deposits forming on the slopes of the growing edifice. He further inferred that the ice sheet re-established itself over the Stage I volcano before the onset of Stage II volcanism, although there is no evidence for a significant time break between the two stages.

Deposits associated with Stage II comprise most of the exposed rocks at Icefall Nunatak and have maximum thickness of >150 m. Stage II is interpreted as a subaqueous to emergent tuff cone sequence formed in an englacial lake. The sequence is similar to the middle–late stages of a glaciovolcanic tuya sequence (e.g. Jones 1969, 1970). Stage II is dominated by a wide variety of subaqueous tuff cone lithofacies that include abundant angular, sideromelane-rich, stratified volcaniclastic tuff, lapilli tuff, tuff breccia and breccia. The sequence transitions from a subaqueous tuff cone through a passage zone to subaerial deposition of lava that extends into subaqueous deposition of a lava-fed pillow/hyaloclastite delta. There is evidence of at least two slope-failure events during the construction of the emergent tuff cone sequence. The palaeowater level (a proxy for ice level) is at c. 700 m asl, about 100 m above the current local ice level.

Stage III deposits include thin columnar basalt lavas and thicker entablature overlain by a scoria cone remnant. Stage III is interpreted as emergent volcanism that occurred over thin ice/and or snow.

Turtle Peak and Hedin Nunatak are both stacked sequences of three or more basanitic lava-fed deltas deposits with passage zones that record varying ice levels during their eruption and deposition. Turtle Peak outcrops are larger and better exposed than those at Hedin Nunatak.

Turtle Peak is a flat-topped elongate edifice and measures approximately 2.5 km north–south and 1.5 km east–west. Cliffs as high as 200 m wrap around the east, north and west sides of Turtle Peak, and the flat top is largely free of snow and ice. Together, these outcrops provide excellent three-dimensional exposures of three subhorizontally stacked lava-fed delta sequences and the vent areas for the lower and upper delta sequences. The three lava-fed delta sequences conformably overlie each other without evidence of significant erosion between them (Fig. 10). The lowest sequence consists of gently (7°–12°) south-dipping pillow lava, pillow lava breccia, bedded hyaloclastite and reworked volcaniclastic sediment that formed foreset-bedded breccias as they accumulated in a meltwater chamber below the local ice-sheet level. Dykes and associated faulting and soft-sediment deformation near the north end of Turtle Peak suggest proximity to the eruptive source. Strombolian bombs and lapilli tuff intermixed with pillow breccias in this area suggest that the eruptive vent periodically emerged above water level. The gently-dipping breccia sequence is capped by a near-horizontal tabular, columnar-jointed lava, prominently exposed in both the east and west faces of Turtle Peak, which records a transition from subglacial to subaerial conditions either as the pile built above water level or following draining of the meltwater chamber. A 40Ar/39Ar age of 5.95 ± 0.60 Ma was obtained from a breccia clast in the lower lava-fed delta. The subaerial lava capping the lower delta sequence is overlain by a similar middle lava-fed delta sequence of pillows, hyaloclastites and reworked volcaniclastic sediments with variable but generally gentle northward dips. This middle pillow hyaloclastite delta sequence closely resembles the lower sequence but lacks a capping subaerial lava. A pillow from the middle sequence yielded a 40Ar/39Ar age of 5.72 ± 0.23 Ma, which overlaps within uncertainty with the age of the lower sequence. The middle sequence is overlain by a third lava delta sequence, which again consists of a sequence of pillows, hyaloclastite and reworked volcaniclastic sediments, gently south-dipping and capped at the north end of Turtle Peak by a subaerially-erupted Strombolian scoria cone and pāhoehoe lava that record emergent conditions in the vent area. The cliffs below the vent area expose dykes that appear to have fed the scoria cone and are spatially associated with, and probably caused, extensive deformation of adjacent foreset-bedded pillow breccia and hyaloclastite. The pāhoehoe lava yielded a 40Ar/39Ar age of 4.76 ± 0.15 Ma. The prominent subaerial lava capping the lower lava-fed delta at Turtle Peak is exposed as a flat erosional surface at the south end of the upper surface of Turtle Peak. This surface is highly polished, striated and littered with granitic glacial erratics, indicating overriding and erosion by the WAIS some time after activity ceased near 4.8 Ma (Fig. 9h).

Fig. 10.

Field sketch map showing the volcanic stratigraphy of the Turtle Peak satellite nunatak, located just east of Mount Murphy.

Fig. 10.

Field sketch map showing the volcanic stratigraphy of the Turtle Peak satellite nunatak, located just east of Mount Murphy.

Hedin Nunatak, located 4 km north of Turtle Peak, is smaller in area (1 × 1 km) but taller (300 m) than Turtle Peak. The top is ice-covered and outcrops are most extensive on the north side. The sequence at Hedin Nunatak is quite similar to that at Turtle Peak, consisting of at least one source tuya and three lava-fed delta sequences containing foreset-bedded pillow/hyaloclastite breccias, some capped by compound subaerial lavas. 40Ar/39Ar ages near the base and top are 6.58 ± 0.12 and 6.28 ± 0.24 Ma, respectively, slightly older than Turtle Peak but similar to the age of Icefall Nunatak, which may have been the source area for some of Hedin Nunatak's lava-fed delta deposits.

Taken together, the sequences of stacked lava-fed deltas and associated source tuyas at Icefall Nunatak, Turtle Peak and Hedin Nunatak provide a record of variably higher ice-sheet levels during their eruptive activity between 6.5 and 4.7 Ma. Local ice levels were periodically at least 200 m higher than the level of the current ice sheet but determining the regional level of the WAIS from these data is complicated by the position of these nunataks on a glacial interfluve and by their near coastal location, as discussed further below.

Dorrel Rock is an isolated nunatak with a peak elevation of 790 m asl situated about 8 km SW of the base of Mount Murphy and 7 km SSW of Turtle Peak. Dorrel Rock forms an interfluve in the heavily-crevassed and rapidly-descending adjacent north-flowing Pope Glacier. Dorrel Rock, along with Turtle Peak, appears to extend the buttressing effect of Mount Murphy on the WAIS, with upstream ice elevations between 600 m asl and downstream elevations at c. 200 m asl.

Dorrel Rock is characterized as an intrusive igneous complex composed of a well-exposed pegmatitic gabbro cut by benmoreite and trachytic dykes (LeMasurier 1990b; Rocchi et al. 2006). The compositions at Dorrel Rock are alkaline and similar to the MBLVG rocks (Rocchi et al. 2006), and the ages are close to the earliest MBLVG volcanism at Mount Petras in the McCuddin Mountains. Biotite in the gabbro was dated by 40Ar/39Ar to be >34.2 Ma and a dyke to c. 33.5 Ma (Rocchi et al. 2006). New 40Ar/39Ar ages on three different mineral phases in the gabbro yielded ages of 35.51± 0.12 Ma (hornblende), 34.83 ± 0.45 Ma (biotite) and 34.92 ± 0.12 Ma (potassium feldspar), suggesting an emplacement age of c. 35 Ma (Table 3).

Rocchi et al. (2006) estimated that the intrusion was emplaced at a depth of at least 3 km. They argued that most of the exhumation of Dorrel Rock occurred early between 34 and 27 Ma, based on the inference of broad-scale uplift of an MBL structural dome centred at Mount Petras in central MBL. It is interesting to note that the timing of intrusion and uplift proposed at Dorrel Rock coincides with a period of seafloor spreading that formed the Adare Basin (43–26 Ma: Cande et al. 2000; Cande and Stock 2006), as well as the intrusion of similar alkaline magmas in North Victoria Land (Meander Intrusive Group, 48–23 Ma: Rocchi et al. 2002), which altogether signify a widespread tectonomagmatic phase.

The Kohler Range is situated near the Walgreen Coast, about 90 km west of Mount Murphy and 100 km NE of Toney Mountain. Two volcanic rock patches rest on a bedrock unconformity (LeMasurier 1990b). The rocks are described as a yellow tuff at Morrison Bluff (observed from the air but not visited) and a ‘thin veneer’ of basalt lava at Leister Peak. Whole-rock geochemical data (XRF analyses) of the basalt sample are provided by LeMasurier (1990b) and there is a K–Ar age of 10.1 ± 3.4 Ma (LeMasurier 1972) (Table 3).

Mount Takahe (3460 m asl) is an isolated undissected late Quaternary polygenetic central volcano in eastern Marie Byrd Land, located 80 km SE of Toney Mountain and 90 km SW of Mount Murphy (Figs 5 & 11). The volcano has an 8 km-diameter, snow-filled, circular caldera that is c. 2100 m above ice level and a nearly symmetrical circular base, 30 km in diameter. Flank slopes range from 7° to 10°.

Fig. 11.

Satellite image map of Mount Takahe of the Eastern MBL Volcanic field, showing outcrops, lithofacies and ages. Outlines of volcano (gold) and caldera (blue) are approximate. Image source: Google Earth Pro image accessed June 2019; for more information on source and an explanation of image processing see the caption to Figure 7.

Fig. 11.

Satellite image map of Mount Takahe of the Eastern MBL Volcanic field, showing outcrops, lithofacies and ages. Outlines of volcano (gold) and caldera (blue) are approximate. Image source: Google Earth Pro image accessed June 2019; for more information on source and an explanation of image processing see the caption to Figure 7.

Mount Takahe has been studied in detail (McIntosh et al. 1985; Palais et al. 1988; LeMasurier and Rex 1990b; Wilch et al. 1999; LeMasurier 2002). Most outcrops on the volcano have been visited. Three groups of outcrops are described here: summit caldera outcrops; lower flank subglacial–subaerial sequences; and lower flank dominantly subaerial sequences. Reconstructions of the volcanic history of Mount Takahe provide key data for WAIS palaeo-ice-level history and for ice-core tephrochronology (McIntosh et al. 1985; Palais et al. 1988; Wilch et al. 1999; LeMasurier 2002). Dating of the young volcanic rocks at Mount Takahe has been challenging. Most K–Ar ages were reported as <0.1 Ma (LeMasurier and Rex 1990b). Fission-track analyses of two caldera-rim samples yielded young ages with large uncertainties (Palais et al. 1988) (Table 4). 40Ar/39Ar dating shows improvement with ages ranging from 194.5 ± 6.3 to 8.3 ± 5.4 ka (Wilch 1997; Wilch et al. 1999). New previously unpublished data presented here provide additional improvements in dating precision for selected outcrops (Fig. 11; Table 4).

Table 4.

Mount Takahe summary age table

Sample IDMethodCorrected preferred age ± 2 SD (ka)DescriptionRef.
Summit caldera
W85009A8.3 ± 5.4Bucher rim (south) pumice and obsidian bombs20
W85013A94.5 ± 7.9Bucher rim (north) lava beneath welded fall20
W85015A103.3 ± 7.4Bucher rim (north), welded fall obsidian20
W85015F135 ± 156Bucher rim (north), welded fall obsidian15
MT85006A156.1 ± 8.22835 m outcrop lava overlain by welded fall20
W85022A169.4 ± 13.8Bucher rim (south), lava20
W85019F175 ± 176Bucher rim, obsidian in moraine15
W85011A194.5 ± 6.3Bucher rim (south), 10 m lava over W85-0920
Post-shield flank outcrops
Oeschger Bluff
MT10A7.1 ± 13.0Pāhoehoe lava16
Roper Point
MT85-3A105.3 ± 28.0Subaerial lava16
98-83A113.1 ± 2.1Tephra in moraine21
98-86A137.7 ± 3.0Tephra in moraine21
98-87A193.9 ± 9.5Tephra in moraine21
Cadenazzi Rock
98-90A89.1 ± 4.1Lava
Steuri Glacier
MT85-9A45.6 ± 7.0Subaerial lava16
Gill Bluff
98-039A26.0 ± 5.1Lithic clast from debris flow21
98-017A21.5 ± 3.3Pillow hyaloclastite21
98-024A22.0 ± 3.7Pillow hyaloclastite21
98-027A20.2 ± 3.7Debris mixed/pillow hyaloclastite21
A21.2 ± 4.1Mean age (n = 3 samples)
Stauffer Bluff
98-055A69.5 ± 9.6Pillow hyaloclastite below the passage zone21
98-058A61.3 ± 10.5Juvenile bomb from accretionary lapilli hyalotuff21
98-064A56.7 ± 19.4Sample from pillow hyaloclastite just below the passage zone21
A64.7 ± 13.3Mean age (n = 3 samples)
Moll Spur (upper)
98-100-01A19.7 ± 12.1Pillow hyaloclastite, pillow interior21
98-100-05A22.8 ± 3.1Pillow hyaloclastite, pillow interior21
98-100-06A17.7 ± 3.9Pillow hyaloclastite, pillow interior21
98-100-07A20.0 ± 14.7Pillow hyaloclastite, pillow interior21
98-100A20.8 ± 4.7Mean age (n = 4 analyses)21
98-103A16.4 ± 2.8Matrix-rich pillow hyaloclastite21
A17.5 ± 4.8Mean age (n = 2 samples)21
Moll Spur (lower)
98-113-1A34.3 ± 13.7Pillow hyaloclastite from the eastern side21
98-113-2A36.3 ± 9.2Pillow hyaloclastite from the eastern side21
98-113-4A37.6 ± 4.0Pillow hyaloclastite from the eastern side21
98-113A37.2 ± 7.0Mean age (n = 3 analyses)21
98-110A40.7 ± 10.7Pillow hyaloclastite from the western side21
98-115A28.4 ± 7.5Pillow hyaloclastite from the eastern side21
A34.5 ± 9.2Mean age (n = 3 samples)
Sample IDMethodCorrected preferred age ± 2 SD (ka)DescriptionRef.
Summit caldera
W85009A8.3 ± 5.4Bucher rim (south) pumice and obsidian bombs20
W85013A94.5 ± 7.9Bucher rim (north) lava beneath welded fall20
W85015A103.3 ± 7.4Bucher rim (north), welded fall obsidian20
W85015F135 ± 156Bucher rim (north), welded fall obsidian15
MT85006A156.1 ± 8.22835 m outcrop lava overlain by welded fall20
W85022A169.4 ± 13.8Bucher rim (south), lava20
W85019F175 ± 176Bucher rim, obsidian in moraine15
W85011A194.5 ± 6.3Bucher rim (south), 10 m lava over W85-0920
Post-shield flank outcrops
Oeschger Bluff
MT10A7.1 ± 13.0Pāhoehoe lava16
Roper Point
MT85-3A105.3 ± 28.0Subaerial lava16
98-83A113.1 ± 2.1Tephra in moraine21
98-86A137.7 ± 3.0Tephra in moraine21
98-87A193.9 ± 9.5Tephra in moraine21
Cadenazzi Rock
98-90A89.1 ± 4.1Lava
Steuri Glacier
MT85-9A45.6 ± 7.0Subaerial lava16
Gill Bluff
98-039A26.0 ± 5.1Lithic clast from debris flow21
98-017A21.5 ± 3.3Pillow hyaloclastite21
98-024A22.0 ± 3.7Pillow hyaloclastite21
98-027A20.2 ± 3.7Debris mixed/pillow hyaloclastite21
A21.2 ± 4.1Mean age (n = 3 samples)
Stauffer Bluff
98-055A69.5 ± 9.6Pillow hyaloclastite below the passage zone21
98-058A61.3 ± 10.5Juvenile bomb from accretionary lapilli hyalotuff21
98-064A56.7 ± 19.4Sample from pillow hyaloclastite just below the passage zone21
A64.7 ± 13.3Mean age (n = 3 samples)
Moll Spur (upper)
98-100-01A19.7 ± 12.1Pillow hyaloclastite, pillow interior21
98-100-05A22.8 ± 3.1Pillow hyaloclastite, pillow interior21
98-100-06A17.7 ± 3.9Pillow hyaloclastite, pillow interior21
98-100-07A20.0 ± 14.7Pillow hyaloclastite, pillow interior21
98-100A20.8 ± 4.7Mean age (n = 4 analyses)21
98-103A16.4 ± 2.8Matrix-rich pillow hyaloclastite21
A17.5 ± 4.8Mean age (n = 2 samples)21
Moll Spur (lower)
98-113-1A34.3 ± 13.7Pillow hyaloclastite from the eastern side21
98-113-2A36.3 ± 9.2Pillow hyaloclastite from the eastern side21
98-113-4A37.6 ± 4.0Pillow hyaloclastite from the eastern side21
98-113A37.2 ± 7.0Mean age (n = 3 analyses)21
98-110A40.7 ± 10.7Pillow hyaloclastite from the western side21
98-115A28.4 ± 7.5Pillow hyaloclastite from the eastern side21
A34.5 ± 9.2Mean age (n = 3 samples)

Notes: method A is 40Ar/39Ar; method F is fission track.

Ref. code: 15, Seward et al. (1980); 16, Wilch (1997); 20, Wilch et al. (1999); 21, this study. See also Supplementary Material Table S1.

Outcrops at and near the caldera rim include a 60 m section of welded and non-welded pyroclastic lapilli-tuff deposits, obsidian-bearing bomb-and-block layers, hydrovolcanic tuffs, and lavas (McIntosh et al. 1985; Palais et al. 1988; Wilch et al. 1999). Anorthoclase from a sample of a non-welded obsidian and pumice bomb-and-block layer near the top of this sequence has a 40Ar/39Ar age of 8.3 ± 5.4 ka (Wilch et al. 1999). This young unit, currently being rapidly eroded, overlies a sequence of lavas and densely-welded pyroclastic deposits with ages ranging from 194.5 ± 6.3 to 94.5 ± 7.9 ka (Wilch et al. 1999). These older units indicate that Mount Takahe volcano reached its present elevation by 194.5 ka.

The c. 8.3 ka tephra on the Mount Takahe caldera rim has been correlated with 8.2 kyr old tephra layers found in ice cores across West Antarctica, including the Byrd, WAIS Divide and Siple Dome ice cores (Wilch et al. 1999; Dunbar et al. 2021). The c. 8.3 ka eruption was highly dispersive and was likely to have been a Plinian eruption. The Takahe tephra provides an important time stratigraphic horizon in these climate archives (Dunbar et al. 2021). Palais et al. (1988) interpreted Mount Takahe as the source of many tephra layers in the Byrd ice core. Two Mount Takahe eruption intervals were inferred from the tephra record: alternating subaerial and phreatomagmatic from 30 to 20 ka; and sustained phreatomagmatic eruptions from 20 to 14 ka. These younger eruptions do not appear to be represented in the caldera-rim sequences but are coeval with some late-stage flank eruptions described below.

Volcanic deposits on the lower flanks are younger than the oldest caldera-rim deposits, suggesting that they resulted from post-edifice-building flank eruptions. Three subglacial–subaerial passage-zone sequences associated with these late-stage flank eruptions are preserved near the base of the volcano at Gill Bluff, Möll Spur and Stauffer Bluff (McIntosh et al. 1985; Palais et al. 1988; LeMasurier 2002; unpublished data of Wilch and McIntosh).

The Gill Bluff promontory extends out from the NW side of Mount Takahe, where it rises to more than 500 m above the ice-sheet surface. The bluff is composed of trachyte lava and volcaniclastic deposits. Gill Bluff has extensive exposures on the NE- and SW-facing sides of the promontory, and both sides exhibit complex multi-level passage-zone transitions from subglacial to subaerial lithofacies and environments. The passage zone is best exposed on the NE-facing bluff, where it exhibits variable water levels (McIntosh et al. 1985; LeMasurier 2002). Gill Bluff is interpreted as a prograding lava delta that formed when lava flowing down the side of Mount Takahe encountered the WAIS and formed a lava-fed delta in an ice-marginal meltwater lake (McIntosh et al. 1985). Individual horizontally orientated subaerial lava units can be traced to dipping units of pillow hyaloclastite in delta foresets. LeMasurier (2002) described a rising vertical passage zone at Gill Bluff across which gently-dipping subaerial felsic lavas pass into 20°–25° dipping pillow hyaloclastite foreset beds. LeMasurier (2002) offered two alternative interpretations of the rising passage zone: either it represents a stable ice-sheet level with a fluctuating ice-marginal lake level; or it represents changing ice levels at different times of the eruption.

We offer an alternative description and interpretation of the stratigraphic sequence (Fig. 12). We observed that nearest to Mount Takahe, the passage zone forms a stable horizontal surface at a measured 413 m above the present ice-sheet surface (elevation determined by differential GPS) (Fig. 12). The delta prograded into an ice-marginal meltwater lake for a period until the passage zone dropped steeply by about 200 m in elevation, where it re-established a horizontal surface. This passage-zone fall is attributed to partial draining of the meltwater lake. After another period of stability and delta progradation at this lower lake level, the passage zone exhibits a steep incremental rise, depositing hyaloclastite foresets associated with the rising and prograding topset beds. The rising passage reached about 400 m above present ice level, close to the original lake level. We observed no stratigraphic evidence for any time gaps such as unconformities or changes in weathering, composition or age in the Gill Bluff sequences. Therefore, we infer that the entire bluff sequence was erupted over a short interval. A similar sequence of events in a glaciovolcanic pāhoehoe-lava-fed delta on James Ross Island (Antarctic Peninsula) was described by Smellie (2006) and similarly ascribed to dynamic fluctuations in the surface elevation of a meltwater lake caused by variable subglacial meltwater drainage. The Gill Bluff sequence was known to be young (K–Ar age of <0.1 Ma). Three anorthoclase separate samples with the most 40Ar/39Ar precise ages yielded concordant age spectra with a weighted mean age of 21.2 ± 4.1 ka (Table 4).

Fig. 12.

Field sketch of a NE-facing outcrop at Gill Bluff, Mount Takahe. Volcanic lithofacies abbreviations: L, subaerial lava; PH, pillow hyaloclastite; DF, debris flow; talus, talus colluvium. Numbers 1–8 designate the sequence of subaerial lavas; numbers 1–7B designated the correlated equivalent sequence of subaqueous deposits in a flow-foot delta. The passage-zone surface is highlighted in red and shows an initial high level at 413 m above current ice level, a drop to low level at c. 200 m above ice level and a rise to c. 400 m above ice level. Composite photograph by T.I. Wilch.

Fig. 12.

Field sketch of a NE-facing outcrop at Gill Bluff, Mount Takahe. Volcanic lithofacies abbreviations: L, subaerial lava; PH, pillow hyaloclastite; DF, debris flow; talus, talus colluvium. Numbers 1–8 designate the sequence of subaerial lavas; numbers 1–7B designated the correlated equivalent sequence of subaqueous deposits in a flow-foot delta. The passage-zone surface is highlighted in red and shows an initial high level at 413 m above current ice level, a drop to low level at c. 200 m above ice level and a rise to c. 400 m above ice level. Composite photograph by T.I. Wilch.

The highest horizontal passage zone at +413 m above modern ice level is interpreted as the minimum thickening of the ice-sheet surface at the time of the eruption, c. 21 ka, coincident with the Last Glacial Maximum. The subsequent descending, horizontal and ascending passage zones are interpreted as indicators of partial draining, stability and refilling of an englacial lake during the eruption phase. Deposits at the passage-zone boundary show mixtures of subaqueous and subaerial lithofacies (Fig. 13a). Highly-vesicular welded breccias and massive holocrystalline lavas indicative of subaerial conditions crop out above the passage zone. Pillow lava, bedded hyaloclastite breccia and interbedded debris-flow deposits crop out below and proximal to the passage zone (Fig. 13b, c). These subaqueously emplaced deposits are locally deformed, as evidenced by disrupted bedding, shear surfaces and normal faults. Towards the toe of the delta, glass-rich clastic deposits are finer grained and more uniformly bedded. In places, these distal delta-toe deposits exhibit reverse faulting that may have resulted from pressure due to ice readvance (Fig. 13d).

Fig. 13.

Photographs of outcrops and lithofacies at Mount Takahe, Eastern MBL Volcanic Field. (a) View of palagonitized Strombolian breccia just above a falling passage zone at Gill Bluff, Mount Takahe. Unit is underlain by pillow hyaloclastite. Located near the passage zone at +413 m above current ice level. (b) View of debris flow and colluvium deposits below the low passage zone (+200 m above ice level) at Gill Bluff, Mount Takahe. Note the steep dips of debris units and sharp shear boundaries between units. (c) View of the complex stratigraphy in a SW-facing Gill Bluff sequence at Mount Takahe. Basal tuff breccia is interpreted as remobilized hyaloclastite debris flow. In the middle of the photograph, fine-grained stratified and deformed lapilli tuff and tuff (yellow) are interpreted as remobilized sediment. This is overlain by a ‘lobe hyaloclastite’ consisting of large irregular-shaped pillow lobes with coarse hyaloclastite breccia. The glassy margins are up to 15 cm thick. (d) View of Z-folded, planar-bedded trachytic tuff and lapilli tuff at the distal toe of the Gill Bluff flow-foot delta. (e) View of trachyte pillows and hyaloclastite breccia in the upper part of Möll Spur flow-foot delta about 60 m below the passage zone (at +575 m above current ice level). Sample 98-100 was collected from a pillow interior at this location. Four separate anorthoclase mineral aliquots from this sample were 40Ar/39Ar dated, yielding a mean age of 20.8 ± 2.8 ka. (f) View of a massive shear zone in the central part of the Möll Spur subglacial sequence. (g) View of incursive pillows in the complex sediment–hyaloclastite–pillow sequences deposited in an englacial lake during the growth of the parasitic Stauffer Bluff tuya. The location of the photograph is shown in Figure 15. (h) View of planar-stratified lapilli tuff with subangular lithic and scoria clasts on the summit of the Stauffer Bluff tuya, Mount Takahe, Eastern MBL Volcanic Field. Note the lithic block sitting in the bedding-plane sag below the snow in the centre of the photograph. Other bedding-plane sags and pyroclastic bombs were observed. The deposit is interpreted as a combination of pyroclastic fall and base surge deposits associated with an emergent phreatomagmatic eruption.

Fig. 13.

Photographs of outcrops and lithofacies at Mount Takahe, Eastern MBL Volcanic Field. (a) View of palagonitized Strombolian breccia just above a falling passage zone at Gill Bluff, Mount Takahe. Unit is underlain by pillow hyaloclastite. Located near the passage zone at +413 m above current ice level. (b) View of debris flow and colluvium deposits below the low passage zone (+200 m above ice level) at Gill Bluff, Mount Takahe. Note the steep dips of debris units and sharp shear boundaries between units. (c) View of the complex stratigraphy in a SW-facing Gill Bluff sequence at Mount Takahe. Basal tuff breccia is interpreted as remobilized hyaloclastite debris flow. In the middle of the photograph, fine-grained stratified and deformed lapilli tuff and tuff (yellow) are interpreted as remobilized sediment. This is overlain by a ‘lobe hyaloclastite’ consisting of large irregular-shaped pillow lobes with coarse hyaloclastite breccia. The glassy margins are up to 15 cm thick. (d) View of Z-folded, planar-bedded trachytic tuff and lapilli tuff at the distal toe of the Gill Bluff flow-foot delta. (e) View of trachyte pillows and hyaloclastite breccia in the upper part of Möll Spur flow-foot delta about 60 m below the passage zone (at +575 m above current ice level). Sample 98-100 was collected from a pillow interior at this location. Four separate anorthoclase mineral aliquots from this sample were 40Ar/39Ar dated, yielding a mean age of 20.8 ± 2.8 ka. (f) View of a massive shear zone in the central part of the Möll Spur subglacial sequence. (g) View of incursive pillows in the complex sediment–hyaloclastite–pillow sequences deposited in an englacial lake during the growth of the parasitic Stauffer Bluff tuya. The location of the photograph is shown in Figure 15. (h) View of planar-stratified lapilli tuff with subangular lithic and scoria clasts on the summit of the Stauffer Bluff tuya, Mount Takahe, Eastern MBL Volcanic Field. Note the lithic block sitting in the bedding-plane sag below the snow in the centre of the photograph. Other bedding-plane sags and pyroclastic bombs were observed. The deposit is interpreted as a combination of pyroclastic fall and base surge deposits associated with an emergent phreatomagmatic eruption.

Möll Spur is a prominent steep ridge composed of trachyte lava and volcaniclastic deposits on the south side of Mount Takahe (Figs 11 & 14). Möll Spur rises up to 800 m above the level of the ice sheet and, similar to Gill Bluff, was formed by late-stage lava eruption on the flank of Mount Takahe. A passage zone from subglacial to subaerial lithofacies occurs at 575 m above the present ice surface at Möll Spur. Below the passage zone, outcrops consist of trachytic pillow lava and palagonitized, matrix- to clast-supported, hyaloclastite breccia deposits (Fig. 13e) (LeMasurier 2002). New40Ar/39Ar analyses suggest that the Möll Spur sequence was erupted at two different times. Five analyses of three samples from the lowermost pillow and hyaloclastite deposits yield a weighted mean age of 34.5 ± 9.2 ka (Table 4). These subglacially-erupted pillow and lobe hyaloclastite deposits occur from the base of the slope up to about 300 m above ice level. Many of the clastic deposits, especially in the basal sequences, show evidence of syn- or post-depositional deformation, including faulting and slickensided shear surfaces (Fig. 13f). The deformation is likely to have resulted from slumping and collapse during or shortly after deposition into an unstable englacial lake formed between Mount Takahe and the ice sheet. Five 40Ar/39Ar analyses of anorthoclase separated from three samples in the uppermost pillow and hyaloclastite deposits yield a weighted mean age of 17.5 ± 4.8 ka. It is possible that these Möll Spur deposits were part of a more explosive Mount Takahe eruption that has been documented in the WAIS Divide ice core as the ‘17.7 ka Mount Takahe Event’ (McConnell et al. 2017). The dated Möll Spur samples were collected between 300 and 525 m above ice level. The passage zone is at +575 m and is overlain by a massive subaerial lava that forms a prominent cliff section (McIntosh et al. 1985; LeMasurier 2002). The Möll Spur sequences place minimum WAIS ice level constraints of >+300 m at c. 35 ka and +575 m at c. 18 ka. The inferred ice thickening at c. 18 ka is coincident with the Last Glacial Maximum.

Fig. 14.

View of Möll Spur, Mount Takahe, Eastern MBL Volcanic Field, showing subaerial lava at top of the spur with a flow-foot pillow hyaloclastite delta below the passage zone.

Fig. 14.

View of Möll Spur, Mount Takahe, Eastern MBL Volcanic Field, showing subaerial lava at top of the spur with a flow-foot pillow hyaloclastite delta below the passage zone.

Stauffer Bluff, situated on the NE flank of Mount Takahe, is composed of hawaiite lava and volcaniclastic deposits; 40Ar/39Ar ages of three hawaiite groundmass samples from Stauffer Bluff provide a mean age of 64.7 ± 13.3 ka (Table 4). Stauffer Bluff is part of a 1 km-diameter relatively flat-topped edifice that rises 525 m above the ice-sheet surface (Fig. 15). The Stauffer Bluff edifice is interpreted as a late-stage parasitic tuya with a vent likely to have been centred on the edifice. The bluff face comprises an exposed flank of the tuya, and is composed of stratigraphically complex subaqueously deposited primary and reworked lithofacies (Fig. 16). The dominant subaqueous lithofacies are thick sequences of hawaiite pillow lava dipping outward from the edifice, palagonitized hyaloclastite breccia, and redeposited fine-grained hyalotuff and hyaloclastite sediments (Figs 10 & 13g). Lithofacies also include sills intruded into glassy lapilli tuff and beds of reworked scoria lapilli forming peperite. The sills are composed of nested pillows with no hyaloclastite matrix that were presumably intruded into water-saturated sediment (Fig. 13g). The bluff-face strata are interpreted as a lava-fed delta formed in an englacial lake. The passage zone from subaqueous to subaerial conditions is covered by snow and ice but is estimated to be located near the break in slope between the bluff face and the bluff top, 400 m above the present ice level. Outcrops on the top of the bluff consist of horizontally to steeply-bedded, palagonitized lapilli tuff and tuff breccia that are interpreted as phreatomagmatic hyalotuff deposits emplaced in an emergent environment (Fig. 13h). Ash-coated lapilli are common in the lapilli-tuff deposits. Large reddened pyroclastic bombs (up to 1 m in diameter) form bedding-plane sags in the tuff deposits. The Stauffer Bluff sequence places minimum WAIS ice level constraints of >+400 m at 65 ka.

Fig. 15.

View of Stauffer Bluff tuya at the NE base of Mount Takahe volcano. The edge of the summit caldera is visible at the top of the photograph. The photograph was taken from an LC-130 airplane by T.I. Wilch.

Fig. 15.

View of Stauffer Bluff tuya at the NE base of Mount Takahe volcano. The edge of the summit caldera is visible at the top of the photograph. The photograph was taken from an LC-130 airplane by T.I. Wilch.

Fig. 16.

Field interpretation of outcrop at the base of Stauffer Bluff tuya, Mount Takahe, Eastern MBL Volcanic Field (see Fig. 14). The outcrop is situated below the passage zone. Volcanic lithofacies abbreviations: Ph, pillow with some hyaloclastite; Phb, hyaloclastite with some pillow lava; S, reworked fine-grained hyalotuff sediment (mostly lapilli tuff); SP, pillow lavas intruded into sediment (S). Photograph by T.I. Wilch.

Fig. 16.

Field interpretation of outcrop at the base of Stauffer Bluff tuya, Mount Takahe, Eastern MBL Volcanic Field (see Fig. 14). The outcrop is situated below the passage zone. Volcanic lithofacies abbreviations: Ph, pillow with some hyaloclastite; Phb, hyaloclastite with some pillow lava; S, reworked fine-grained hyalotuff sediment (mostly lapilli tuff); SP, pillow lavas intruded into sediment (S). Photograph by T.I. Wilch.

All other flank outcrops at Mount Takahe are interpreted as resulting from subaerial eruptions above the level of the ice sheet and without significant interaction with meltwater. Cadenazzi Rock, located on the west flank of Mount Takahe at 350 m above the ice sheet surface, is a 50 m-high bluff composed of 89.1 ± 4.1 ka mugearite pyroclastic rocks (Table 4). Palagonitized lapilli tuff with rare large (up to 1 m) lithic blocks and pumiceous bombs (up to 15 cm) are exposed in the lower part of the bluff. The upper section of the bluff includes lapilli tuff, rich in ash-coated lapilli with beds of well-sorted achnelith-rich lapillistone deposits. Cadenazzi Rock is interpreted as a pyroclastic deposit that resulted from a mix of magmatic and phreatomagmatic eruptions that formed in a subaerial environment where meltwater had intermittent access to the vent.

Downslope from Cadenazzi Rock, Roper Point consists of eroded outcrops of scoriaceous hawaiite lava, 40Ar/39Ar dated to 105 ± 28 ka (Wilch 1997), overlain by glacial till. Roper Point is the oldest dated flank outcrop at Mount Takahe. The till at Roper Point is a poorly-consolidated heterolithic silty gravelly diamict and abundant volcanic erratics. Erratic lithologies are dominated by subaerial lithofacies, including a variety of lava fragments, as well as pumice-rich moderately- to densely-welded pyroclastic rocks. The pyroclastic erratics appear to be derived from welded fall and ignimbrite deposits. The till also includes trachytic pumice lapilli and ash concentrations up to 10 cm thick, with subrounded pumice clasts as large as 2 cm in diameter. The tephra deposits are presumed to be locally derived and are interpreted as reworked. Three 40Ar/39Ar dates of anorthoclase separates from different tephra deposits range from 193.9 ± 9.5 ka to 113.1 ± 2.1 ka (Table 4). The till, situated about 270 m above the local level of the ice sheet, is inferred to represent deposition since the Last Glacial Maximum and after the eruption of the Möll Spur and Gill Bluff glaciovolcanic sequences.

On the SW side of Mount Takahe, two outcrops are present near Steuri Glacier (Fig. 11). The eroded rim of a monogenetic basaltic scoria cone is preserved on slopes NW of the glacier, and subaerial trachytic lava and breccia are preserved SE of the glacier. The Steuri Glacier trachyte is about 600 m above present ice level and is 40Ar/39Ar dated to 45 ± 7 ka (Wilch 1997). On the SE side of Mount Takahe, a stacked sequence of pāhoehoe basanite lavas is exposed at Oeschger Bluff, about 250 m above the level of the ice sheet. The lava yielded an imprecise 40Ar/39Ar age of 7 ± 13 ka (Wilch 1997). On the north side of Mount Takahe, undated trachyte lava is exposed at Knezevitch Rock, less than 200 m above the level of the ice sheet.

In summary, Mount Takahe is a Late Quaternary volcano with a large summit caldera that was constructed by 194.5 ka and the last known eruption was about 8 ka. Because of the Holocene activity, Mount Takahe is still considered to be active. Mount Takahe provides critical data for reconstructing late Pleistocene WAIS ice levels, which are discussed later in this chapter.

Toney Mountain is a massive, elongate, east–west-orientated central volcano with basal dimensions of 55 × 15 km (Figs 1, 2 & 17). The volcano rises c. 2000 m above the WAIS and has a summit peak elevation of 3595 m. Toney Mountain produces a significant damming effect on the north-flowing WAIS. Ice levels on the upstream side of the volcano are 500 m higher than on the downstream side. The summit area of the volcano has a circular caldera, about 3 km in diameter. The edifice appears to be mostly undissected, except at its eastern end. Relatively steep constructional slopes (13°–21°) occur below the caldera on the north and south flanks.

Fig. 17.

Satellite image map of studied outcrops of Toney Mountain in the Eastern MBL Volcanic Field. Outlines of volcano (gold) and caldera (blue) are approximate. There are only three analysed outcrops. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Fig. 17.

Satellite image map of studied outcrops of Toney Mountain in the Eastern MBL Volcanic Field. Outlines of volcano (gold) and caldera (blue) are approximate. There are only three analysed outcrops. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Only five samples from three outcrop localities have been analysed for age (K–Ar, fission track) and geochemistry. Cox Bluff, located c. 15 km east of the summit caldera between 1600 and 1800 m above sea level, exposes a 200 m-thick sequence of Late Miocene (c. 9–10 Ma) subaerial hawaiite lava (LeMasurier et al. 1990d). Zurn Peak, located 7 km NNE of the caldera between 1200 and 1400 m above sea level, exposes trachyte and comendite rhyolite lava. The trachyte yielded a K–Ar age of 1.0 ± 0.4 Ma and the comendite yielded a fission track age of 0.29 ± 0.4 Ma (Seward et al. 1980; LeMasurier et al. 1990d) (Table 5). A third (unnamed) outcrop situated 2–3 km SW of the caldera at about 2900 m asl exposes Pleistocene (K–Ar 0.5 ± 0.2 Ma) benmoreite lava (LeMasurier et al. 1990d). On satellite images, this site appears to be part of a cone rim that is c. 0.8 km in diameter (Fig. 17) and is likely to be a parasitic cone. Several other parasitic cones are situated along the long axis of the volcano; some of these cones have been sampled but the outcrops were too altered for analysis (LeMasurier et al. 1990d). Paulsen and Wilson (2010) noted the alignment of some these cones and included them in their stress-pattern analysis.

Table 5.

Toney Mountain and Crary Mountains summary age table

Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Toney Mountain
80K9.1 ± 1.2Cox Bluff- horizontal lava flows9
80K10.1 ± 1.2Cox Bluff- horizontal lava flows9
80K9.6 ± 1.0Mean age (n = 2)
76AK1.0 ± 0.4Zurn Peak, trachyte flow rock9
75K0.5 ± 0.2Unnamed outcrop at 2900 m asl; flow rock9
76DF0.29 ± 0.10Zurn Peak, comendite15
Crary Mountains: Mount Rees
Trabucco Cliff (listed in stratigraphic order from top to bottom)
92-174A9.06 ± 0.06Lava flow top of section18
92-175A9.06 ± 0.06Lava flow top of section18
92-6A9.09 ± 0.07Massive flow, above 92-1518
92-15A9.12 ± 0.04Lava flow middle of section18
92-1A9.25 ± 0.53Lobe hyaloclastite18
Tasch Peak Ridge (listed in stratigraphic order from top to bottom)
92-36A7.62 ± 0.06Dyke18
92-38A8.32 ± 0.132230 m pillow hyaloclastite18
92-34A8.34 ± 0.092269 m intrusive hyaloclastite18
92-41A8.61 ± 0.16Hyaloclastite lobes18
92-31A9.10 ± 0.281844 m pillow lobe interior18
92-28A9.19 ± 0.061817 m spatter-fed lava18
92-23A9.03 ± 0.121739 m pillow lobe18
92-59A9.46 ± 0.241643 m lava flow18
Isolated outcrops
92-117A6.91 ± 0.26SW of summit; eroded parasitic cone18
92-114A7.01 ± 0.21Eroded parasitic cone18
92-112A9.14 ± 0.19North end of Mount Rees, lava-flow, striae (240 m)18
92-109A8.81 ± 0.08lava-flow, striae (328 m)18
Crary Mountains: Mount Steere
Outcrops on west, north and NE sides
92-118A7.64 ± 0.05Flow-banded lava, west of summit18
92-95A5.81 ± 0.04Flow-banded lava clast from moraine, base of north side18
92-64A8.16 ± 0.08Dyke, NE outcrop18
92-53A8.34 ± 0.08Lava-flow dome, NE outcrop18
92-93A8.35 ± 0.08Dyke, NE outcrop18
92-63A8.38 ± 0.08Dyke, NE outcrop18
92-51A8.43 ± 0.06Lava flow, exposed plug, NE outcrop18
92-107A8.45 ± 0.06Dyke, 2413 m18
92-104A8.46 ± 0.08Flow-banded lava, 2278 m18
92-108A8.48 ± 0.06Flow-banded lava, 2413 m18
92-91A8.57 ± 0.09Flow-banded lava, east side18
92-181A8.01 ± 0.20Dyke, fine grained, east side18
92-182A8.63 ± 0.06Flow-banded lava, east side18
92-183A8.63 ± 0.06Flow-banded lava, east side18
92-178A8.66 ± 0.04Flow-banded lava, east side18
Ridge SE of Lie Cliff (samples listed in stratigraphic order)
92-169A6.49 ± 0.43SE ridge: 1814 m lava flow18
92-165A7.47 ± 0.07SE ridge: 1736 m lava, near 92-16218
92-162A6.78 ± 0.05SE ridge: 1736 m lava, near 92-16518
Lie Cliff (samples listed in stratigraphic order)
92-80A7.92 ± 0.06Dyke, 1.5 m wide18
92-85A8.49 ± 0.331631 m lava flow18
92-82A8.39 ± 0.211600 m hyaloclastite lobe18
92-79A8.54 ± 0.06Lowest subaerial lava, thin flows18
92-86A8.63 ± 0.231558 m lava flow18
Ridge NW of Lie Cliff (samples listed in stratigraphic order)
92-89A7.78 ± 0.06Dyke, intrudes entire section18
92-193A8.30 ± 0.181798 m feeder dyke, lava18
92-192A8.33 ± 0.07Feeder dyke, lava18
92-194A8.30 ± 0.221753 m glassy breccia18
92-189A8.56 ± 0.461747 m glassy lava18
92-190A8.38 ± 0.641743 m pillow lava18
92-186A8.51 ± 0.111646 m subaerial lava18
Crary Mountains: Mount Frakes
Morrison Rocks
92-145A1.83 ± 0.10Subaerial lava18
92-142A1.84 ± 0.05Subaerial lava18
92-128A2.55 ± 0.06Subaerial lava18
92-125A2.57 ± 0.09Subaerial lava18
92-130A3.93 ± 0.03Subaerial lava18
92-122A4.22 ± 0.05Lava flow18
92-121A4.24 ± 0.04Lava flow18
92-127A4.31 ± 0.03Lava flow18
4.26 ± 0.05Mean age (n = 3)
English Rock
92-151A0.032 ± 0.010Parasitic cone18
92-151A0.035 ± 0.010Parasitic cone18
92-151A0.034 ± 0.014Mean age (n = 2)
92-157A0.837 ± 0.079Parasitic cone18
92-157A0.862 ± 0.036Parasitic cone18
92-157A0.858 ± 0.066Mean age (n = 2)
92-159A1.62 ± 0.02Parasitic cone18
Crary Mountains: Boyd Ridge
92-135aA2.24 ± 0.19Runyon Rock: juvenile clast from hyaloclastite16
92-135bA2.00 ± 0.3016
92-135A2.17 ± 0.32Mean age (n = 2)
92-134A2.05 ± 0.05Runyon Rock, molded and polished clasts from debris flow16
92-139A1.29 ± 0.03Subaerial lava from north top side16
Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Toney Mountain
80K9.1 ± 1.2Cox Bluff- horizontal lava flows9
80K10.1 ± 1.2Cox Bluff- horizontal lava flows9
80K9.6 ± 1.0Mean age (n = 2)
76AK1.0 ± 0.4Zurn Peak, trachyte flow rock9
75K0.5 ± 0.2Unnamed outcrop at 2900 m asl; flow rock9
76DF0.29 ± 0.10Zurn Peak, comendite15
Crary Mountains: Mount Rees
Trabucco Cliff (listed in stratigraphic order from top to bottom)
92-174A9.06 ± 0.06Lava flow top of section18
92-175A9.06 ± 0.06Lava flow top of section18
92-6A9.09 ± 0.07Massive flow, above 92-1518
92-15A9.12 ± 0.04Lava flow middle of section18
92-1A9.25 ± 0.53Lobe hyaloclastite18
Tasch Peak Ridge (listed in stratigraphic order from top to bottom)
92-36A7.62 ± 0.06Dyke18
92-38A8.32 ± 0.132230 m pillow hyaloclastite18
92-34A8.34 ± 0.092269 m intrusive hyaloclastite18
92-41A8.61 ± 0.16Hyaloclastite lobes18
92-31A9.10 ± 0.281844 m pillow lobe interior18
92-28A9.19 ± 0.061817 m spatter-fed lava18
92-23A9.03 ± 0.121739 m pillow lobe18
92-59A9.46 ± 0.241643 m lava flow18
Isolated outcrops
92-117A6.91 ± 0.26SW of summit; eroded parasitic cone18
92-114A7.01 ± 0.21Eroded parasitic cone18
92-112A9.14 ± 0.19North end of Mount Rees, lava-flow, striae (240 m)18
92-109A8.81 ± 0.08lava-flow, striae (328 m)18
Crary Mountains: Mount Steere
Outcrops on west, north and NE sides
92-118A7.64 ± 0.05Flow-banded lava, west of summit18
92-95A5.81 ± 0.04Flow-banded lava clast from moraine, base of north side18
92-64A8.16 ± 0.08Dyke, NE outcrop18
92-53A8.34 ± 0.08Lava-flow dome, NE outcrop18
92-93A8.35 ± 0.08Dyke, NE outcrop18
92-63A8.38 ± 0.08Dyke, NE outcrop18
92-51A8.43 ± 0.06Lava flow, exposed plug, NE outcrop18
92-107A8.45 ± 0.06Dyke, 2413 m18
92-104A8.46 ± 0.08Flow-banded lava, 2278 m18
92-108A8.48 ± 0.06Flow-banded lava, 2413 m18
92-91A8.57 ± 0.09Flow-banded lava, east side18
92-181A8.01 ± 0.20Dyke, fine grained, east side18
92-182A8.63 ± 0.06Flow-banded lava, east side18
92-183A8.63 ± 0.06Flow-banded lava, east side18
92-178A8.66 ± 0.04Flow-banded lava, east side18
Ridge SE of Lie Cliff (samples listed in stratigraphic order)
92-169A6.49 ± 0.43SE ridge: 1814 m lava flow18
92-165A7.47 ± 0.07SE ridge: 1736 m lava, near 92-16218
92-162A6.78 ± 0.05SE ridge: 1736 m lava, near 92-16518
Lie Cliff (samples listed in stratigraphic order)
92-80A7.92 ± 0.06Dyke, 1.5 m wide18
92-85A8.49 ± 0.331631 m lava flow18
92-82A8.39 ± 0.211600 m hyaloclastite lobe18
92-79A8.54 ± 0.06Lowest subaerial lava, thin flows18
92-86A8.63 ± 0.231558 m lava flow18
Ridge NW of Lie Cliff (samples listed in stratigraphic order)
92-89A7.78 ± 0.06Dyke, intrudes entire section18
92-193A8.30 ± 0.181798 m feeder dyke, lava18
92-192A8.33 ± 0.07Feeder dyke, lava18
92-194A8.30 ± 0.221753 m glassy breccia18
92-189A8.56 ± 0.461747 m glassy lava18
92-190A8.38 ± 0.641743 m pillow lava18
92-186A8.51 ± 0.111646 m subaerial lava18
Crary Mountains: Mount Frakes
Morrison Rocks
92-145A1.83 ± 0.10Subaerial lava18
92-142A1.84 ± 0.05Subaerial lava18
92-128A2.55 ± 0.06Subaerial lava18
92-125A2.57 ± 0.09Subaerial lava18
92-130A3.93 ± 0.03Subaerial lava18
92-122A4.22 ± 0.05Lava flow18
92-121A4.24 ± 0.04Lava flow18
92-127A4.31 ± 0.03Lava flow18
4.26 ± 0.05Mean age (n = 3)
English Rock
92-151A0.032 ± 0.010Parasitic cone18
92-151A0.035 ± 0.010Parasitic cone18
92-151A0.034 ± 0.014Mean age (n = 2)
92-157A0.837 ± 0.079Parasitic cone18
92-157A0.862 ± 0.036Parasitic cone18
92-157A0.858 ± 0.066Mean age (n = 2)
92-159A1.62 ± 0.02Parasitic cone18
Crary Mountains: Boyd Ridge
92-135aA2.24 ± 0.19Runyon Rock: juvenile clast from hyaloclastite16
92-135bA2.00 ± 0.3016
92-135A2.17 ± 0.32Mean age (n = 2)
92-134A2.05 ± 0.05Runyon Rock, molded and polished clasts from debris flow16
92-139A1.29 ± 0.03Subaerial lava from north top side16

Notes: method A is 40Ar/39Ar; method K is K/Ar; method F is fission track.

Toney Mountain is the only volcano in MBL for which there is a depth estimate of the sub-ice volcano–basement contact. Based on a 1959–60 seismic traverse across the west end of the edifice, the contact occurs at a depth of 3000 m below sea level (Bentley and Clough 1972), suggesting a total relief of the volcano of c. 6600 m, and a possible volume ranging from 2800 to 3613 km3 (LeMasurier et al. 1990d; LeMasurier 2013).

In summary, Toney Mountain is a major central volcano with a well-defined caldera that appears to have been active in Late Miocene and Pleistocene times. Toney Mountain is the least-studied central polygenetic volcano of the MBLVP and, although it is mostly snow- and ice-covered, there are many outcrops that could be studied to better understand its eruptive history.

The Crary Mountains Volcanic Field consist of three large, coalesced central volcanoes, Mount Rees, Mount Steere and Mount Frakes, and the much smaller Boyd Ridge, which are aligned roughly NW–SE and are less dissected and progressively younger toward the SE (Fig. 18). The Late Miocene Mount Rees and Mount Steere, and Pliocene Mount Frakes, together cover an area of about 33 × 15 km. Boyd Ridge, located about 13 km SE of these central volcanoes, is an east–west-orientated, lower-relief ridge that covers an area of about 13 × 7 km. Mount Steere and Mount Frakes each have >2 km-diameter summit calderas. The Crary Mountains produce a significant damming effect on the NE-flowing WAIS. Ice levels on the upstream SW side of the volcanoes are 200–400 m higher than on the downstream NE side. Deep cirques cut into the east and NE sides of Mount Steere and Mount Rees expose thick stratigraphic sequences and intrusive rocks, making these two volcanoes among the best exposed in Marie Byrd Land.

Fig. 18.

Satellite image map showing ages and outcrops of the Crary Mountain Volcanic Field. There is a general progression from NW to SE for the four major volcanoes in the Crary Mountains. Both Mount Rees and Mount Steere are deeply dissected and many dykes are exposed. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Fig. 18.

Satellite image map showing ages and outcrops of the Crary Mountain Volcanic Field. There is a general progression from NW to SE for the four major volcanoes in the Crary Mountains. Both Mount Rees and Mount Steere are deeply dissected and many dykes are exposed. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

On the basis of initial observations and four K–Ar ages, the Crary Mountains were characterized as two felsic shield volcanoes (Mount Steere and Mount Frakes) on a platform of basaltic lava and pyroclastic rocks (LeMasurier et al. 1990a). Subsequent detailed fieldwork, geochemical analysis of 70 samples and 40Ar/39Ar dating of 77 samples (Table 5) provide a more comprehensive record of volcanism in the Crary Mountains (Wilch 1997; Panter et al. 2000; Wilch and McIntosh 2002; Chakraborty 2007). This summary of the Crary Mountains volcanoes is drawn mostly from Wilch (1997) and Wilch and McIntosh (2002), with a minor amount of unpublished data.

Mount Rees and Mount Steere are overlapping Late Miocene polygenetic central volcanoes (Fig. 18). Given their similar ages, lithofacies and histories, Mount Rees and Mount Steere are presented together in this subsection. Mount Rees is more deeply dissected, lower in elevation (2709 m compared to 3558 m), and slightly older (9.5–6.9 Ma) than Mount Steere (8.7–5.8 Ma). Mount Rees (c. 10 × 7 km) has a smaller footprint than Mount Steere (c. 9 × 12.5 km). The eroded Mount Rees lacks a summit caldera, whereas Mount Steere has a mostly intact c. 2 km-diameter summit caldera. Extensive outcrops of the dissected volcanoes are located on the east flank of Mount Rees and on the north flank of Mount Steere. Outcrops up to 800 m above current ice level have been examined and sampled.

Mount Rees has thick stratigraphic sequences at Trabucco Cliff and Tasch Peak Ridge consisting mostly of mafic–intermediate volcanic rocks with subordinate interlayered felsic lavas (Wilch and McIntosh 2002). The mafic–intermediate rock outcrops (Fig. 19) are characterized by two alternating lithofacies: (1) unbrecciated lavas with oxidized bases (subaerial lithofacies); and (2) palagonitized glassy hyaloclastite breccias and pillow lavas (subaqueous or water-contact lithofacies). In places, oxidized clastogenic lava makes up the subaerial lithofacies. 40Ar/39Ar analyses of five samples from different levels in the Trabucco Cliff sequence agree with the stratigraphic order (from 9.25 ± 0.53 to 9.06 ± 0.06 Ma) but are analytically indistinguishable from one another (Table 5). No glacial unconformities or tillites were observed in either the subaqueous or subaerial lithofacies sequences. A short interval of eruptions is consistent with the lack of unconformities in the sections. At Tasch Peak Ridge (Fig. 20), seven 40Ar/39Ar ages are mostly in stratigraphic order and range from 9.46 ± 0.24 to 8.32 ± 0.13 Ma. The mostly intermediate–mafic, alternating lavas and hyaloclastite breccias are cut by a trachytic dyke, 40Ar/39Ar dated to 7.62 ± 0.06 Ma. At both Trabucco Cliff and Tasch Peak Ridge, the contacts between the alternating subaerial and subaqueous lithofacies are dipping conformably with the constructional slopes of the volcano.

Fig. 19.

View of the alternating subaqueous and subaerial lithofacies at Trabucco Cliff, Mount Rees, Crary Mountains Volcanic Field. The layered sequences and passage zones conform to the constructional slopes of the volcanoes. The location of Trabucco Cliff is shown in Figure 18. Adapted and updated from Wilch and McIntosh (2002).

Fig. 19.

View of the alternating subaqueous and subaerial lithofacies at Trabucco Cliff, Mount Rees, Crary Mountains Volcanic Field. The layered sequences and passage zones conform to the constructional slopes of the volcanoes. The location of Trabucco Cliff is shown in Figure 18. Adapted and updated from Wilch and McIntosh (2002).

Fig. 20.

Stratigraphic section from Tasch Peak Ridge at Mount Rees, Crary Mountains Volcanic Field. Volcanic lithofacies are interpreted as either ‘dry’ (i.e. subaerial) or ‘wet’ (i.e. subglacial/ice-contact). Figure adapted from Wilch and McIntosh (2002). The location of Tasch Peak Ridge is shown in Figure 18.

Fig. 20.

Stratigraphic section from Tasch Peak Ridge at Mount Rees, Crary Mountains Volcanic Field. Volcanic lithofacies are interpreted as either ‘dry’ (i.e. subaerial) or ‘wet’ (i.e. subglacial/ice-contact). Figure adapted from Wilch and McIntosh (2002). The location of Tasch Peak Ridge is shown in Figure 18.

The stacked slope-parallel alternating subaerial and subaqueous (meltwater-contact) lithofacies at the base of Mount Rees (and also Mount Steere, described below) comprise slope-forming constructional lithofacies that lack interbedded fluvial deposits, glacial deposits and glacial unconformities. Wilch and McIntosh (2002) interpreted the sloping passage-zone sequence from subaqueous to subaerial lithofacies as being formed by lavas interacting with slope ice and snow to form pillows and hyaloclastites until they built above the level of slope ice and formed subaerial lavas. Furthermore, they interpreted the lack of glacial tills and unconformities as evidence of volcanic interactions with thin or cold-based ice on the slopes of the growing Mount Steere and Mount Rees volcanoes, above the level of the WAIS. The Mount Rees and Mount Steere sequences appear to resemble glaciovolcanic sequences in northern Victoria Land, Antarctica, which have been interpreted similarly by Smellie et al. (2011b) as lavas that interacted with a thin cold-based ice field or topography-draping ice. Smellie et al. (2011b, p. 135) considered the ‘non-horizontal lava-fed delta passage zone(s) orientated parallel to underlying pre-volcanic bedrock slope’ to be diagnostic of eruption in a glacial setting and interpreted the northern Victoria Land sequences to be slope parallel ‘a‘ā-lava-fed hyaloclastite deltas formed in sloping, open-top lava-melted channels in the ice (see also Smellie et al. 2013). Although generally similar, the Crary Mountain sequences include both wet to dry and dry to wet transitions, leading Wilch and McIntosh (2002) to suggest that lava flowing through open channels or tunnels in the ice were in some cases resubmerged after becoming emergent. Wilch and McIntosh (2002) concluded that the alternating wet and dry lithofacies, coupled with the sloping passage zones, precludes formation in a water-filled melt-water chamber within a formerly higher WAIS. Smellie and Edwards (2016), using the field descriptions of Wilch and McIntosh (2002), described the Mount Rees and Mount Steere sequences as ‘a‘ā-lava-fed deltas.

We contend that these alternating lithofacies associated with sloping passage zones are likely to be common features developed during the growth of ice-mantled Antarctic volcanoes. These sequences have not been recognized at other MBL volcanoes, possibly in part due to the lack of dissection and extensive snow- and ice-cover of most MBL volcanoes. Sequences of this type are diagnostic of eruption in an environment mantled with thin (<100 m), cold-based ice and provide only a maximum elevation for syneruptive regional ice sheets.

Two other outcrop areas were sampled at Mount Rees (Wilch and McIntosh 2002). On the NW side of Mount Rees, subaerially-erupted trachyte lava and bombs are exposed in multiple outcrops situated between 200 and 500 m above modern ice levels. Two samples are 40Ar/39Ar dated to 9.14 ± 0.19 and 8.81 ± 0.08 Ma (Table 5). West of Tasch Peak, basanite lapilli and bombs are exposed at three sites situated about 750 m above current ice levels. Bomb interiors from two locations yielded 40Ar/39Ar ages of 7.01 ± 0.21 and 6.91 ± 0.26 Ma. These basanite outcrops are interpreted as remnants of parasitic scoria cones.

The north face of Mount Steere is deeply eroded, with multiple separate outcrops exposing abundant felsic flow-banded lava and breccia, cut by numerous felsic–mafic dykes. The oldest dated rocks at Mount Steere are a series of hydrothermally-altered and brecciated, strongly flow-banded trachyte and rhyolite lava, situated between 1830 and 2030 m asl on the lower NE flank of the volcano. Potassium feldspar from three samples date from 8.66 ± 0.04 to 8.63 ± 0.06 Ma (Table 5). The lava sequence is cut by a hydrothermally altered and brecciated phonolite dyke that is dated at 8.01 ± 0.20 Ma. Separate outcrops of trachyte lava yield 40Ar/39Ar ages ranging from 8.57 ± 0.09 to 8.43 ± 0.08 Ma. These outcrops are cut by multiple dykes, with ages ranging from 8.45 ± 0.06 to 8.16 ± 0.08 Ma.

On the lower east flank of Mount Steere at Lie Cliff and a ridge to its north, outcrops include a 300 m-thick section of alternating subaqueous and subaerial lithofacies dominated by hawaiite lava and breccia, with subordinate trachyte lava. The sections resemble the alternating slope-forming stratigraphic sequences at Mount Rees and are interpreted in the same way. These sequences were erupted over a short interval that is not differentiated by 40Ar/39Ar dating of eight samples, with overlapping ages ranging from 8.63 ± 0.23 to 8.30 ± 0.22 Ma (Table 5). These subaqueous and subaerial sequences are intruded by four trachyte and phonolite dykes with ages ranging from 8.33 ± 0.07 to 7.78 ± 0.06 Ma.

Post-8 Ma eruptions at Mount Steere are represented at three localities. A flow-banded phonolite lava from the west side of Mount Steere is dated to 7.64 ± 0.05 Ma (Table 5). A discontinuous rock ridge SE of Lie Cliff exposes a slightly younger (7.47 ± 0.07–6.49 ± 0.43 Ma) series of dominantly basanite subaerial lava and pyroclastic deposits over a 400 m elevation range. The pyroclastic rocks are mostly subaerial lapilli tuff and bomb-rich agglutinate deposits with some interbedded palagonitized, fine-grained, laminated tuff deposits. Minor trachyte lava occurs in the sequence. This sequence is interpreted as resulting from a series of dry magmatic eruptions with minor intermittent interactions with external water producing phreatomagmatic eruptive phases. The uppermost hydrovolcanic tuff unit at c. 1810 m asl is cut by a glacial unconformity, as evidenced by a polished planar contact with truncated clasts. The units beneath the unconformity are dated to 6.49 ± 0.43 Ma and provide a maximum age of the erosion event. Finally, the youngest rock at Mount Steere is a flow-banded mugearite lava on the lower north flank of Mount Steere, downslope from older lavas and dykes. This subaerially-erupted mugearite lava is dated to 5.81 ± 0.04 Ma and suggests that the eroded Mount Steere slope developed prior to c. 5.8 Ma.

Mount Frakes is the least-dissected and highest of three central volcanoes in the Crary Mountains Volcanic Field, with a nearly circular 3 × 2.5 km-diameter summit caldera at 3654 m asl and flank slopes of 11°–15° (Fig. 18). The nearly symmetrical base of Mount Frakes at the level of the ice sheet is 13.5 × 14.5 km. Numerous isolated outcrops are exposed at Morrison Rocks on the south flank of the volcano and upslope from English Rocks on its west flank.

Morrison Rocks consists of a series of outcrops on the south slope of Mount Frakes between 2225 and 2990 m asl. The oldest 40Ar/39Ar dated rocks (4.26 ± 0.05 Ma: Table 5) are crystal-rich (15–30%) phonolite lavas, with anorthoclase phenocrysts up to 4 cm in length. In situ phonolite lava crops out as high as 2550 m asl; phonolite and other felsic erratics occur among more mafic outcrops as high as 2990 m asl. The phonolite outcrops include large lava-flow levees running down the volcano flank. Several outcrops of mafic (basanite and hawaiite) welded Strombolian deposits are also exposed at Morrison Rocks. Spatter ramparts associated with some of these deposits are interpreted as evidence of fissure vents. The mafic rocks have 40Ar/39Ar dates of 3.93 ± 0.03, 2.55 ± 0.06 and 1.84 ± 0.05 Ma, and are interpreted as representing post-edifice-building subaerial eruptions. One of the basanite outcrops includes a 6 m-thick section of well-bedded, well-sorted lapilli tuff. Most of the bedding is planar with some cross-bedded layers. This section lacks bedding-plane sags and ash-coated/accretionary lapilli. It is interpreted as a ‘dry’ turbulent low-density pyroclastic density current deposit that resulted from water–magma interaction during a late-stage low-volume eruption.

Remnants of late-stage basanitic scoria cone deposits also crop out on the western side of Mount Frakes at English Rock. Three outcrops were sampled and dated to 1.62 ± 0.02 Ma, and 858 ± 66 and 34 ± 14 ka (Table 5). The outcrops include clastogenic lava and bombs up to 1.5 m in length. The youngest deposits, situated c. 150 m above the level of the ice sheet, limit syneruptive ice-sheet expansion to <150 m above the present ice level at c. 33.9 ka.

Other than evidence for minor intermittent water–magma interaction in one outcrop at Morrison Rocks, there is no evidence for glaciovolcanic interactions at Mount Frakes. The overall absence of glaciovolcanic sequences at Mount Frakes may simply reflect the lack of dissection and limited exposure.

Boyd Ridge is a low-relief east–west elongate ridge located 13 km SE of Mount Frakes (Fig. 18). At the east end and lee side of Boyd Ridge, Runyon Rock exposes a 130 m-thick stratigraphic section. The base of the section is a 2 m-thick hydrothermally altered phonolite lava breccia. This is overlain by 5 m of massive and poorly-sorted tuff breccia that is dominated by a matrix of angular finely-vesicular hawaiite lapilli and larger (up to 2 m in diameter) phonolite lava clasts that range from angular to very well rounded. Some of the phonolite clasts are polished. The deposit also contains abundant vesicular hawaiite clasts, and some of the clasts and other xenoliths are coated by hawaiite lava. This is interpreted as a debris-flow deposit or, possibly, an explosion breccia emplaced at the time of hawaiite volcanism. The basal tuff breccia grades into a uniform massive to slightly-bedded, matrix-supported, palagonitized hyaloclastite lapilli tuff, with rare larger hawaiite juvenile blocks. The lapilli are dominated by hawaiite clasts. There are no pillow lavas in the hawaiite hyaloclastite sequence. The 130 m-thick hyaloclastite deposit reaches an elevation of 1820 m asl, about 200 m below the upstream ice levels; therefore, in the absence of Boyd Ridge, a hyaloclastite sequence like that at Runyon Rock would be expected to form at the same location today.

Two samples have been dated from the Runyon Rock sequence. A phonolite clast from the debris-flow deposit is 40Ar/39Ar dated to 2.05 ± 0.05 Ma, and provides a maximum age for the debris-flow deposit and overlying hyaloclastite. A juvenile hawaiite clast from the hyaloclastite deposit was dated to 2.17 ± 0.32 Ma (Table 5). The maximum age of the phonolite clast and the imprecise eruption age suggest that the hyaloclastite was emplaced between 2.10 and 1.85 Ma. A second outcrop at Boyd Ridge, not identified on USGS topographical maps and located 1 km SW of Runyon Rock, is composed of subaerially erupted phonotephrite lava and lapilli tuff, 40Ar/39Ar dated to 1.29 ± 0.03 Ma.

A NW–SE alignment and systematic age progression of the three Crary Mountains polygenetic volcanoes (Mount Rees (9.5–6.9 Ma), Mount Steere (8.7–5.8 Ma) and Mount Frakes (4.3 Ma)) was noted in the structural analysis of MBL volcanoes by Paulsen and Wilson (2010). Younger volcanism occurred at the east–west-aligned Boyd Ridge (c. 2–1.3 Ma) and at post-edifice-building volcanoes on the flank of Mount Frakes (3.9–0.03 Ma). The age progression is consistent with the degree of dissection: Mount Rees is much more eroded than Mount Steere, and both of these two Late Miocene volcanoes are more eroded than the Pliocene Mount Frakes. It is likely that the erosion of Mount Rees and Mount Steere occurred early in their histories in the Late Miocene. The fact that the Mount Frakes is undissected suggests that only limited erosion has occurred since early Pliocene time. Numerous erratics are scattered around the flanks of the Crary Mountains. These erratics are mostly alkaline volcanic lithologies and are likely to result from a combination of alpine/local and ice-sheet/regional glaciation. A few granite erratics occur at Morrison Rocks on the south lower flank of Mount Frakes. An altimeter reading on the highest granite erratic was 2440 m asl, about 600 m above the elevation of the adjacent ice-sheet surface. This granite erratic is interpreted as evidence of former ice-sheet expansion(s) since 1.6 Ma, the age of the youngest volcanic rocks at Morrison Rocks. The lack of glacial till and striated clasts suggests that the ice sheet was likely to have been cold-based.

Mount Rees and Mount Steere are among the more deeply-eroded volcanoes in Marie Byrd Land, and thus offer more complete records of volcanism than do other less-dissected volcanoes. The deep dissection is expressed geomorphologically, and also by the abundance of exposed dykes and the diversity of locally-derived volcanic glacial erratics on the surface. The overlapping ages and bimodal lithology of these volcanoes suggest chemically diverse volcanism since the early stages of the growth of both volcanoes. The alternating subaqueous and subaerial lithofacies at Mount Rees and Mount Steere suggests that during their construction in the Late Miocene (9.5–8.3 Ma) both volcanoes were mantled with snow and ice as they are today. Although extensive work has been completed in the Crary Mountains, many outcrops remain unvisited, especially in the upper elevations at Mount Steere and upper west flank at Mount Frakes.

The Mount Siple Volcanic Field consists of one volcano, Mount Siple, an undissected composite volcano on the MBL coast (Fig. 1). The massive volcano rises from sea level to 3100 m asl, with a 4.5 km-diameter ice-filled summit caldera and a base diameter at sea level of 45 × 35 km (Fig. 21). Because the WAIS terminates at the coast, Mount Siple is not buried by the ice sheet and thus has the largest exposed volume of volcanoes in the MBLVP (c. 1800 km3) (LeMasurier and Rex 1990a). The volcano is almost completely ice-covered with a few low outcrops along the coast, and limited outcrop near and in the summit caldera wall.

Fig. 21.

Satellite image map of Mount Siple volcano showing the approximate volcano outline (orange), caldera rim (dashed blue), studied outcrops (yellow), lithofacies and ages. Mount Siple is an island in the Southern Ocean, separated from the Bakutis Coast by the Getz Ice Shelf. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Fig. 21.

Satellite image map of Mount Siple volcano showing the approximate volcano outline (orange), caldera rim (dashed blue), studied outcrops (yellow), lithofacies and ages. Mount Siple is an island in the Southern Ocean, separated from the Bakutis Coast by the Getz Ice Shelf. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

A total of four coastal or lower flank outcrops have been visited (Table 6) (LeMasurier and Rex 1990a; Wilch 1997; Wilch et al. 1999). LeMasurier and Rex (1990a) identified two basaltic tuff cone remnants at Lovill Bluff and at a site 1 km to the south, and a subhorizontal basaltic lava 150 m asl about 4 km north of Lovill Bluff. The sites near Lovill Bluff yielded K–Ar ages with large uncertainties (2.0 ± 1.4 and 1.1 ± 1.0 Ma) and the ages are considered unreliable. The Lovill Bluff outcrop has not been dated reliably: LeMasurier and Rex (1990a) reported a <0.1 Ma K–Ar age; and Wilch (1997),40Ar/39Ar dated a hawaiite clast to 8 ± 88 ka. The Lovill Bluff tuff cone deposits include well-sorted glass-rich lapilli tuff with aphyric hawaiite lava and altered scoria clasts. On the NE side of Mount Siple, at about 320 m asl, a hawaiite tuff cone deposit was dated to 746 ± 36 ka (Wilch 1997). This deposit includes glass-rich, planar-bedded lapilli tuff, significant soft-sediment deformation, ash-coated lapilli and bedding-plane sags.

Table 6.

Mount Siple Volcanic Field summary age table

Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Mount Siple
Summit caldera shield building
93-278A0.1711 ± 0.0054200 m below summit lava20
93-277A0.2329 ± 0.0067Caldera wall, densely-welded fall, 20 m-thick exposure20
93-277A0.2366 ± 0.0173Caldera wall, densely-welded fall, 20 m-thick exposure20
93-277A0.2250 ± 0.0069Caldera wall, densely-welded fall, 20 m-thick exposure20
93-277A0.2296 ± 0.0076Mean age (n = 3)20
Flank deposits post-shield building
93-270A0.746 ± 0.036Lithic clasts in lapilli tuff16
93-275A0.008 ± 0.088Lithic clasts in lapilli tuff16
K1.1 ± 1.0Tuff cone deposit8
W83-5K2.0 ± 1.430 m-thick subhorizontal basanite lava8
Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Mount Siple
Summit caldera shield building
93-278A0.1711 ± 0.0054200 m below summit lava20
93-277A0.2329 ± 0.0067Caldera wall, densely-welded fall, 20 m-thick exposure20
93-277A0.2366 ± 0.0173Caldera wall, densely-welded fall, 20 m-thick exposure20
93-277A0.2250 ± 0.0069Caldera wall, densely-welded fall, 20 m-thick exposure20
93-277A0.2296 ± 0.0076Mean age (n = 3)20
Flank deposits post-shield building
93-270A0.746 ± 0.036Lithic clasts in lapilli tuff16
93-275A0.008 ± 0.088Lithic clasts in lapilli tuff16
K1.1 ± 1.0Tuff cone deposit8
W83-5K2.0 ± 1.430 m-thick subhorizontal basanite lava8

Notes: method A is 40Ar/39Ar; method K is K/Ar.

The summit of Mount Siple was first visited in January 1994. Wilch et al. (1999) described a 20+ m-thick densely-welded, fiamme-rich, pyroclastic fall deposit exposed at the highest point of the caldera rim. Three analyses of anorthoclase derived from the trachyte deposit yielded a mean40Ar/39Ar age of 229.6 ± 7.6 ka. A trachyte lava associated with a subsidiary vent 210 m lower in elevation than the summit crater yielded an age of 171.1 ± 5.4 ka. Although there has been speculation about recent eruptions at Mount Siple (Global Volcanism Program 1988), there is no evidence of this at the summit caldera.

In summary, Mount Siple is a Pleistocene polygenetic central volcano with a 227 ka summit caldera that has been active since at least 746 ka. Parasitic volcanism may be active in the Holocene but existing age data are unreliable. The limited outcrops and observations preclude making more detailed interpretations of the volcanic history. At this stage there is no indication that Mount Siple should be considered ‘active’.

The Executive Committee Range (ECR) Volcanic Field in central Marie Byrd Land consists of five central volcanoes that are aligned north–south and progressively young toward the south (Fig. 22). The elevation of the ice-sheet surface is higher at the ECR volcanoes (2000–2600 m asl) than at other MBLVP volcanoes. The diameters of the volcanoes at the level of the ice sheet range from 3 to 15 km. All five of the ECR Volcanic Field volcanoes have well-defined summit calderas, and three of the five have two calderas; caldera diameters range from 2 to 10 km. The volcanoes appear to be deeply buried by the WAIS, in some cases exposing little more than the summit calderas.

Fig. 22.

Satellite image map of the Executive Committee Range (ECR) Volcanic Field. Note the age progression from north to south. NASA Earth Observatory image by Jesse Allen, using Landsat data from the United States Geological Survey. Accessed 21 June 2019 from https://earthobservatory.nasa.gov/images/85238/antarcticas-tallest-volcano

Fig. 22.

Satellite image map of the Executive Committee Range (ECR) Volcanic Field. Note the age progression from north to south. NASA Earth Observatory image by Jesse Allen, using Landsat data from the United States Geological Survey. Accessed 21 June 2019 from https://earthobservatory.nasa.gov/images/85238/antarcticas-tallest-volcano

Mount Hampton, located at the north end of the ECR Volcanic Field, contains an intact slightly NW-elongated 6.5 × 5.5 km summit caldera that rises to 3223 m asl and is 400–700 m above the ice sheet (Fig. 23). The flanks slope from 10° to 20° and the base only extends c. 2.5 km beyond the caldera. The remnant of a second, smaller (c. 3.5 km-diameter) and older caldera is situated NW of the summit caldera. The remnant caldera has an intact NW rim that includes Whitney Peak. The two calderas are composed of felsic lavas (trachyte and low-silica rhyolite) with superimposed parasitic basanite cones. The basanite deposits include both crustal and mantle xenoliths.

Fig. 23.

Simplified geological map of Mount Hampton in the ECR Volcanic Field. Image extracted from NASA Earth Observatory image by Jesse Allen, using Landsat data from the United States Geological Survey. Accessed 21 June 2019 from https://earthobservatory.nasa.gov/images/85238/antarcticas-tallest-volcano

Fig. 23.

Simplified geological map of Mount Hampton in the ECR Volcanic Field. Image extracted from NASA Earth Observatory image by Jesse Allen, using Landsat data from the United States Geological Survey. Accessed 21 June 2019 from https://earthobservatory.nasa.gov/images/85238/antarcticas-tallest-volcano

Three rock types are exposed on the older caldera rim and slopes near Whitney Peak (observations of the authors) (Fig. 23). The caldera rim is dominated by a plagioclase-rich (5–25%) trachyte lava, and includes a benmoreite lava (LeMasurier and Kawachi 1990c; observations of the authors). In places, the well-exposed trachyte lava exhibits strong and contorted flow foliations, defined by alternating variably glassy layers. An associated lithic-rich trachyte welded fall, with 5:1 flattening of pumice clasts, crops out locally. The trachyte lava is 40Ar/39Ar dated to 13.36 ± 0.05 Ma (Table 7). The third lithology is a late-stage xenolith-bearing basanite that overlies the trachyte lava. Dip direction variations among the basanite outcrops suggest that they constitute multiple parasitic vents. Outcrops include variably-welded lapilli and bombs up to 40 cm, and pyroclastic deposits are locally transitional to spatter-fed, flow-foliated lavas. This unit was not described by LeMasurier and Kawachi (1990c) and has not been dated.

Table 7.

Executive Committee Range Volcanic Field summary age table

Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Mount Hampton
Hampton caldera
20DK8.6 ± 1.0Near summit caldera7
22BK10.1 ± 0.8West flank7
22DK11.4 ± 1.2West flank7
25K10.7 ± 0.8SE flank7
90-201A11.09 ± 0.04Crystal-rich phonolite21
MB-74.2A11.43 ± 0.04Crystal-rich phonolite21
19CK11.7 ± 1.0Flow rock7
Whitney Peak caldera
90-174A13.36 ± 0.05Phonolite21
23AK13.4 ± 1.0Flow rock7
24AK13.7 ± 1.0Flow rock7
23CK13.7 ± 1.0Flow rock7
Mount Cumming
28K10.4 ± 1.0South flank, lava7
27K10.0 ± 1.0Near LaVaud Peak, lava4
26AK3.0 ± 0.4Near Annexstad Peak, parasitic cone4
Mount Hartigan
43AK6.02 ± 0.50North caldera, lava7
45CK7.57 ± 0.60South caldera, lava7
46BK7.86 ± 1.00South caldera near Mintz, lava7
48K8.36 ± 0.82South caldera, near Tusing, lava7
42BK8.50 ± 0.66North caldera, lava7
Mount Sidley
Stage IV – post-shield activity
K168A4.24 ± 0.08Strombolian tephra and lava, parasitiic cone12
Stage III – formation of the breached Sidley caldera
K85A4.37 ± 0.06Tuff cone deposits and lava, flank vent12
K51A4.43 ± 0.06Unwelded ignimbrite, flank vent12
K137A4.31 ± 0.06Welded fall with fiamme, flank vent12
Stage II – flank activity
MB29.4A4.59 ± 0.04Welded fall with abundant lithic fragments12
K105A4.61 ± 0.08Porphyritic lava and breccia from endogenous dome and commingled lava, flank vent12
K55A4.66 ± 0.10Porphyritic lava and breccia from endogenous dome and commingled lava, flank vent12
MB33.3A4.64 ± 0.04Porphyritic lava and breccia from endogenous dome, flank vent12
MB35.5A4.51 ± 0.02Poorly-welded pyroclastic fall, flank vent12
Stage I – Weiss caldera formed
K149A4.87 ± 0.06Vitric and porphyritic lava and basal breccia12
MB42.3A5.15 ± 0.14Vitric and porphyritic lava and basal breccia12
K106A5.43 ± 0.04Lava and basal breccia12
K108A5.60 ± 0.14Porphyritic lava and breccia12
K68A5.77 ± 0.12Porphyritic lava and breccia12
Mount Waesche
Flank deposits
41AK<0.1 ± 0.00SW flank7
35A*K0.170 ± 0.60SW flank, cinder cone7
39AK0.200 ± 0.40SW flank7
33CK1.000 ± 0.20SW flank7
A0.49 ± 0.02SW flank, trachyte dome
Chang Peak caldera
32AK1.6 ± 0.4Chang Peak, caldera wall7
32AF1.48 ± 0.33Chang Peak, caldera wall15
A1.09 ± 0.10Chang Peak, NW flank lava11
A2.01 ± 0.10Chang Peak, NW flank lava11
Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Mount Hampton
Hampton caldera
20DK8.6 ± 1.0Near summit caldera7
22BK10.1 ± 0.8West flank7
22DK11.4 ± 1.2West flank7
25K10.7 ± 0.8SE flank7
90-201A11.09 ± 0.04Crystal-rich phonolite21
MB-74.2A11.43 ± 0.04Crystal-rich phonolite21
19CK11.7 ± 1.0Flow rock7
Whitney Peak caldera
90-174A13.36 ± 0.05Phonolite21
23AK13.4 ± 1.0Flow rock7
24AK13.7 ± 1.0Flow rock7
23CK13.7 ± 1.0Flow rock7
Mount Cumming
28K10.4 ± 1.0South flank, lava7
27K10.0 ± 1.0Near LaVaud Peak, lava4
26AK3.0 ± 0.4Near Annexstad Peak, parasitic cone4
Mount Hartigan
43AK6.02 ± 0.50North caldera, lava7
45CK7.57 ± 0.60South caldera, lava7
46BK7.86 ± 1.00South caldera near Mintz, lava7
48K8.36 ± 0.82South caldera, near Tusing, lava7
42BK8.50 ± 0.66North caldera, lava7
Mount Sidley
Stage IV – post-shield activity
K168A4.24 ± 0.08Strombolian tephra and lava, parasitiic cone12
Stage III – formation of the breached Sidley caldera
K85A4.37 ± 0.06Tuff cone deposits and lava, flank vent12
K51A4.43 ± 0.06Unwelded ignimbrite, flank vent12
K137A4.31 ± 0.06Welded fall with fiamme, flank vent12
Stage II – flank activity
MB29.4A4.59 ± 0.04Welded fall with abundant lithic fragments12
K105A4.61 ± 0.08Porphyritic lava and breccia from endogenous dome and commingled lava, flank vent12
K55A4.66 ± 0.10Porphyritic lava and breccia from endogenous dome and commingled lava, flank vent12
MB33.3A4.64 ± 0.04Porphyritic lava and breccia from endogenous dome, flank vent12
MB35.5A4.51 ± 0.02Poorly-welded pyroclastic fall, flank vent12
Stage I – Weiss caldera formed
K149A4.87 ± 0.06Vitric and porphyritic lava and basal breccia12
MB42.3A5.15 ± 0.14Vitric and porphyritic lava and basal breccia12
K106A5.43 ± 0.04Lava and basal breccia12
K108A5.60 ± 0.14Porphyritic lava and breccia12
K68A5.77 ± 0.12Porphyritic lava and breccia12
Mount Waesche
Flank deposits
41AK<0.1 ± 0.00SW flank7
35A*K0.170 ± 0.60SW flank, cinder cone7
39AK0.200 ± 0.40SW flank7
33CK1.000 ± 0.20SW flank7
A0.49 ± 0.02SW flank, trachyte dome
Chang Peak caldera
32AK1.6 ± 0.4Chang Peak, caldera wall7
32AF1.48 ± 0.33Chang Peak, caldera wall15
A1.09 ± 0.10Chang Peak, NW flank lava11
A2.01 ± 0.10Chang Peak, NW flank lava11

Notes: method A is 40Ar/39Ar; method K is K/Ar; method F is fission track.

The dominant edifice-building lithology at Mount Hampton is crystal-rich anorthoclase phonolite (kenyte) (LeMasurier and Kawachi 1990c). The crystal-rich phonolite includes both well-exposed lava and welded fall lithofacies. A subordinate but similar lithology is crystal-poor phonolite lava. The stratigraphic relationship of the two phonolite lavas is uncertain. Previously unpublished 40Ar/39Ar ages of the crystal-rich phonolite are 11.43 ± 0.04 and 11.09 ± 0.04 Ma; the crystal-poor phonolite is K–Ar dated to 8.6 ± 1.0 Ma (Table 7). The crystal-rich phonolite lava is overlain by a third lithology, a variably-welded xenolith-bearing basanite. Two K–Ar ages for the same basanite sample are 11.4 ± 1.2 and 10.1 ± 0.8 Ma (LeMasurier and Rex 1989). The texture of the basanite varies from non-welded vesicular lapilli tuff and bombs to densely-welded spatter-fed lavas. Bombs are 0.5–1.5 m in diameter. The basanite appears to post-date the erosion of underlying phonolites. However, some spatter-fed basanitic lavas also appear to be cut by the caldera wall, suggesting that this basanite was erupted before caldera collapse.

Glacial moraines are found near both Mount Hampton calderas and contain locally-derived volcanic clasts, including phonolite, trachyte and basanite.

LeMasurier and Wade (1968) and LeMasurier and Kawachi (1990c) noted that the caldera rim has large conical snow and ice mounds, and interpreted these as dormant fumarolic ice towers and indications of recent activity. This interpretation is incorrect and the features are probably snow (rime) mushrooms resulting from wind sculpting (Whiteman and Garibotti 2013). There is no indication of recent volcanic activity or geothermal features.

Mount Cumming, located 12 km south of Mount Hampton, is a 3.5 km-diameter circular caldera that protrudes only 200 m above the ice sheet surface (Figs 22 & 24). The exposed base of the volcano extends <1 km beyond the caldera rim. The caldera rim is intact and includes limited outcrops composed of variably foliated pantelleritic trachyte lava, locally underlain by a non-welded to densely-welded fall deposit with flattened glassy fiamme and resorbed cognate xenoliths. Remnants of parasitic basanite scoria cones on the lower slopes are composed of xenolith-rich (e.g. granulite and ultramafic) welded fall deposits transitional to spatter-fed lavas. Conventional K–Ar ages of the trachyte are 10.4 ± 1.0 and 10.0 ± 1.0 Ma, and one basanite cone is dated to 3.0 ± 0.1 Ma (Table 7) (LeMasurier and Rex 1989; LeMasurier and Kawachi 1990b).

Fig. 24.

Simplified geological map of Mount Cumming in the ECR Volcanic Field. Image extracted from NASA Earth Observatory image by Jesse Allen, using Landsat data from the United States Geological Survey. Accessed 21 June 2019 from https://earthobservatory.nasa.gov/images/85238/antarcticas-tallest-volcano

Fig. 24.

Simplified geological map of Mount Cumming in the ECR Volcanic Field. Image extracted from NASA Earth Observatory image by Jesse Allen, using Landsat data from the United States Geological Survey. Accessed 21 June 2019 from https://earthobservatory.nasa.gov/images/85238/antarcticas-tallest-volcano

Mount Hartigan, located 12 km SSW of Mount Cumming at the centre of the ECR, is made up of two overlapping north–south-aligned volcanoes, each defined by a 3.5 km-diameter circular caldera rim and snow-/ice-filled caldera (Fig. 22). The calderas are less well defined than at Mount Hampton and Mount Cumming. The northern volcano is composed of trachyte, rhyolite and mugearite lava; the southern volcano is composed of hawaiite and mugearite lava (LeMasurier 1990f). The volcanoes overlap in age, with two samples from the northern volcano K–Ar dated to 8.50 ± 0.66 and 6.02 ± 0.50 Ma; and three from the southern volcano ranging in age from 8.36 ± 0.82 to 7.57 ± 0.60 Ma (LeMasurier and Rex 1989) (Table 7).

Mount Sidley, located 12 km south of Mount Hartigan near the southern end of the ECR Volcanic Field, is one of the most well-exposed and well-studied volcanoes in the MBLVG (Fig. 25). This summary is derived mostly from Panter et al. (1994), who published a detailed reconstruction of the history of the volcanic complex, including 40Ar/39Ar dating and geochemistry.

Fig. 25.

Simplified geological map on satellite image base map Mount Sidley of the ECR Volcanic Field. Geology adapted from figure 4 in Panter et al. (1994); base map extracted from NASA Earth Observatory image by Jesse Allen, using Landsat data from the United States Geological Survey. Accessed June 21, 2019 from https://earthobservatory.nasa.gov/images/85238/antarcticas-tallest-volcano

Fig. 25.

Simplified geological map on satellite image base map Mount Sidley of the ECR Volcanic Field. Geology adapted from figure 4 in Panter et al. (1994); base map extracted from NASA Earth Observatory image by Jesse Allen, using Landsat data from the United States Geological Survey. Accessed June 21, 2019 from https://earthobservatory.nasa.gov/images/85238/antarcticas-tallest-volcano

Mount Sidley (4285 m asl) is the tallest volcano in Antarctica. The volcano rises 2200 m above the WAIS level and has basal dimensions of 14 × 19 km. Early workers (González-Ferrán and González-Bonorino 1972; LeMasurier 1972) noted two calderas: the older Weiss caldera, a c. 2.5 km-diameter, snow- and ice-filled caldera situated north of the summit peak that is truncated by the younger Sidley caldera, a 4.5 km-diameter and >1200 m-deep breached caldera that opens to the south. Parks Glacier occupies the floor of the breached Sidley Caldera and flows from the headwall at the north end of the caldera into the ice sheet SE of the volcano.

Panter et al. (1994) described a complex four-stage geological history spanning c. 1.5 myr (5.8–4.2 Ma); each stage is defined by one or more geographical shifts of eruptive centres and by geochemical changes of the magma (Fig. 25). Stage I, Sidley Activity, comprised the largest eruptive volume (90% or c. 180 km3) of the Sidley massif and resulted in the construction of three successive volcanic edifices, the Byrd (informal name by Panter et al. 1994), Weiss and Sidley volcanoes, culminating in the formation of a c. 3.5 km-diameter summit caldera. Weiss caldera is the only Stage I caldera preserved mostly intact and is the protocaldera of the later-stage breached caldera. The reconstructed calderas delineate a migration of eruptive centres towards the SW. Phonolite–tephriphonolite lavas dominate these sequences, although minor amounts of pumice-bearing pyroclastic material are interbedded within and between the lava sequences. Minor pillow lava and hyaloclastite breccia deposits occur locally and are attributed to limited glaciovolcanic interactions. Erosional unconformities separate the volcano sequences. The stratigraphy of Stage I is well dated by 40Ar/39Ar geochronology, with ages from 5.77 ± 0.12 Ma at the base of the sequence to 4.87 ± 0.06 Ma at the top of the Sidley volcano sequence.

Stage II, Pirrit Activity, is characterized by a petrological shift to trachyte, as well as a shift in eruption style from the central vent to multiple small monogenetic eruptive centres distributed on the flanks of a largely undissected Mount Sidley. Rocks include thin (5–10 m) lavas associated with small vents, and thick (> 50 m) lavas and carapace breccias associated with endogenous domes. One of the large eruptions is evidenced by a single vent with a thick (70 m) stratigraphic sequence of deposits that includes a basal trachyte/phonolite pyroclastic fall deposit that transitions up into a unit of alternating foliated trachyte lavas and carapace breccias. Stage II lasted from 4.66 ± 0.10 to 4.51 ± 0.02 Ma, and produced c. 7% (18 km3) of the erupted volume of Mount Sidley (Panter et al. 1994).

Stage III, Doumani Activity, was a short-lived (4.43–4.37 Ma) trachyte pyroclastic eruptive phase that began with emplacement of a lithic-rich welded fall deposit followed closely by non-welded ignimbrite and associated surge deposits and co-ignimbrite breccias (Panter et al. 1994). Panter et al. (1994) speculated that this explosive phase was coincident with the initial collapse of the present-day breached caldera, which was created by a combination of syneruptive explosive landslide and post-eruptive erosional processes. South of the main Sidley edifice at Doumani Peak, an ignimbrite sequence is dated to 4.43 ± 0.06 Ma and is overlain conformably by a mugearite/benmoreite tuff cone sequence, dated to 4.37 ± 0.06 Ma. This stage comprises an estimated 1% of the erupted volume (c. 2.5 km3).

Stage IV is characterized by multiple basanite parasitic scoria cones on the flanks of Mount Sidley. The presence of mantle and lower-crustal xenoliths was interpreted as evidence of rapid ascent rates. A sample from one cone was dated to 4.24 ± 0.08 Ma. This stage comprises <<1% of the erupted volume (c. 0.25 km3).

Panter et al. (1994) note that magmatic activity at Mount Sidley migrated southwestwards at a rate of 6 cm a−1 and follows the same pattern as the ECR Volcanic Field overall. The migration has been interpreted as being due to fracture propagation caused by regional tectonic stresses (LeMasurier and Rex 1989) or related to plate-boundary forces (Paulsen and Wilson 2010). Magma injection into a complex system of conduits and chambers may have aided periodic dilation of pre-existing structures, resulting in the eruption of evolved magmas in discrete pulses (Panter et al. 1994).

The Mount Waesche massif, located at the south end of the ECR, consists of two coalesced polygenetic volcanoes: the older Chang Peak volcano defined by a NNE-elongated 10 × 6 km-diameter caldera; and the younger Mount Waesche volcano defined by the prominent peak and c. 1.5 km-diameter caldera superimposed and centred on the southern rim of Chang Peak caldera (Fig. 22). Outcrops are limited at Chang Peak and include comendite (peralkaline rhyolite) pumice and vitrophyre that is K–Ar dated to 1.6 ± 0.4 Ma (LeMasurier and Rex 1989). Two additional comendite outcrops on the western slope of the caldera have been 40Ar/39Ar dated. A flow-banded vitrophyre and lava with abundant spherulites and lithophysae is dated to 2.01 ± 0.10 Ma, and a similar but more phenocryst-rich vitrophyic lava is dated to 1.09 ± 0.10 Ma (Table 7). Thus, the Chang Peak caldera and volcano appear to be early Pleistocene in age.

The younger Mount Waesche volcano has excellent exposures on the SW flank that include a wide range of compositions, from basalts to intermediate types (mugearite, phonotephrite and tephriphonolite) to phonolite and trachyte (Panter et al. 2021). Conventional K–Ar ages of hawaiite samples range from 1.0 ± 0.2 to <0.1 Ma (LeMasurier and Rex 1989) (Table 7). Numerous englacial tephra layers are exposed in a blue-ice ablation zone of the ice sheet at the base of the south flank (Dunbar et al. 2021).

A geological map and detailed description of the geology are presented by Dunbar et al. (2021) and only a summary is given here based on that account. Exposed along the rim of the ice-filled summit caldera of Mount Waesche are deposits of basanite lava, agglutinated lava and welded Strombolian tephra containing xenolith-rich bombs. A heterolithic debris deposit is also found around the rim of the summit caldera and on the upper slopes of Mount Waesche. The debris is composed of fragments of lapilli tuff, green felsic (trachytic?) lavas, gabbro and hypabyssal plutonic lithologies. One large block (c. 2 m in diameter) of phonolite contains xenoliths and basalt clasts that are flattened and aligned with reaction rims, all of which suggest that the block is a remnant of a welded pyroclastic fall or flow deposit. The origin of the debris deposit is considered either as a result of glacial deposition or explosive volcanic or a combination of both, and has since been modified by downslope movement and solifluction processes (Smellie et al. 1990; Dunbar et al. 2021). Ongoing work (2018–19 fieldwork) includes extensive 40Ar/39Ar geochronology and detailed mapping.

The McCuddin Mountains Volcanic Field is located in central Marie Byrd Land, north of the ECR and east of the Ames Range (Figs 1 & 2). The field includes Mount Flint, a low-relief central volcano, and monogenetic volcanoes at Mount Petras and the USAS Escarpment. Together these represent the oldest known phase of MBLVG volcanism, mostly from 36.58 to 20.46 Ma, followed by Late Miocene and Pliocene (9.67–3.75 Ma) parasitic volcanism at the Mount Flint central volcano (Table 8).

Table 8.

McCuddin Mountains and Ames Range volcanic fields summary age table

Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
USAS Escarpment
Near Mount Galla
MB70.1A26.40 ± 0.21Hyalotuff (tuff cone)16
Mount Aldaz
90-181A19.0 ± 1.4Mount Aldaz, spatter lava over glacial unconformity and bedrock16
90-180A19.68 ± 0.31Mount Aldaz, basanitic lava16
Mount Flint
93-279A9.67 ± 0.20SW parasitic cone, agglutinate lava16
93-281A3.75 ± 0.06West upper parasitic cone, agglutinate lava16
93-284A8.66 ± 0.24West lower parasitic cone agglutinate lava16
93-288A20.45 ± 0.10Reynolds Ridge, hypabbyssal lava16
93-291A20.48 ± 0.12Reynolds Ridge, lava, 20–25% feldspar16
20.46 ± 0.07Mean age (n = 2)
Mount Petras
93-323A36.71 ± 0.51SW flank, mugearite lava xenolith17
93-329A27.53 ± 0.23SW flank, dense haw SW saddlebomb interior17
93-343A28.96 ± 0.22Aphyric haw lava, 2 m thick17
93-337A28.26 ± 0.38Near summit, aphyric glassy haw lava, rare xenoliths17
93-332A36.58 ± 0.22Near summit,massive mugearite lava17
93-333A28.22 ± 0.52Near summit,dense haw bomb interior17
Ames Range
Mount Andrus
60K<0.1Upper Lind Ridge, parasitic cone6
41K11.3 ± 0.8Near Rosenburg Glacier7
44K11.6 ± 0.8Near Rosenburg Glacier parasitic cone6
93-308A9.28 ± 0.06Lava, parasitic cone16
93-309A12.71 ± 0.06Pumiceous lava, massive16
93-311A11.19 ± 0.08Pumiceous lava16
93-312A11.18 ± 0.19Pumiceous lava16
Mount Kosciusko
40BK10.00 ± 0.80Parasitic cone2
40BK8.66 ± 0.70Parasitic cone2
K9.20 ± 1.30Mean age (n = 2)
Mount Kauffman
67B-8K5.9 ± 1.0Lava6
AR39BK7.6 ± 0.6Lava7
K7.1 ± 1.5Mean age (n = 2)
Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
USAS Escarpment
Near Mount Galla
MB70.1A26.40 ± 0.21Hyalotuff (tuff cone)16
Mount Aldaz
90-181A19.0 ± 1.4Mount Aldaz, spatter lava over glacial unconformity and bedrock16
90-180A19.68 ± 0.31Mount Aldaz, basanitic lava16
Mount Flint
93-279A9.67 ± 0.20SW parasitic cone, agglutinate lava16
93-281A3.75 ± 0.06West upper parasitic cone, agglutinate lava16
93-284A8.66 ± 0.24West lower parasitic cone agglutinate lava16
93-288A20.45 ± 0.10Reynolds Ridge, hypabbyssal lava16
93-291A20.48 ± 0.12Reynolds Ridge, lava, 20–25% feldspar16
20.46 ± 0.07Mean age (n = 2)
Mount Petras
93-323A36.71 ± 0.51SW flank, mugearite lava xenolith17
93-329A27.53 ± 0.23SW flank, dense haw SW saddlebomb interior17
93-343A28.96 ± 0.22Aphyric haw lava, 2 m thick17
93-337A28.26 ± 0.38Near summit, aphyric glassy haw lava, rare xenoliths17
93-332A36.58 ± 0.22Near summit,massive mugearite lava17
93-333A28.22 ± 0.52Near summit,dense haw bomb interior17
Ames Range
Mount Andrus
60K<0.1Upper Lind Ridge, parasitic cone6
41K11.3 ± 0.8Near Rosenburg Glacier7
44K11.6 ± 0.8Near Rosenburg Glacier parasitic cone6
93-308A9.28 ± 0.06Lava, parasitic cone16
93-309A12.71 ± 0.06Pumiceous lava, massive16
93-311A11.19 ± 0.08Pumiceous lava16
93-312A11.18 ± 0.19Pumiceous lava16
Mount Kosciusko
40BK10.00 ± 0.80Parasitic cone2
40BK8.66 ± 0.70Parasitic cone2
K9.20 ± 1.30Mean age (n = 2)
Mount Kauffman
67B-8K5.9 ± 1.0Lava6
AR39BK7.6 ± 0.6Lava7
K7.1 ± 1.5Mean age (n = 2)

Notes: method A is 40Ar/39Ar method K is K/Ar.

Mount Flint is a low-relief central volcano that rises to 2695 m, about 900 m above the ice-sheet level (Fig. 26). The edifice is elongate in an east–west direction and is about 10 × 6 km in diameter. The volcano is almost completely covered with snow and ice, with limited outcrops exposed on the west flank. A flat area near the summit edifice suggests a possible c. 2 × 3 km-diameter oval caldera. Three outcrops were visited and are interpreted as erosional remnants of parasitic basanite scoria cones (LeMasurier et al. 1990b; Wilch 1997). The rock includes reddened welded bombs and lapilli, and spatter-fed lava (Wilch 1997). One scoria cone contains abundant mantle xenoliths, whereas two others contain crustal xenoliths. The 40Ar/39Ar plateau ages of the three remnant cones are 9.67 ± 0.20, 8.66 ± 0.24 and 3.75 ± 0.06 Ma (Table 8). The post-edifice-building cones provide minimum ages for the Mount Flint edifice. No outcrops have been identified that are part of the underlying (pre-parasitic cone) low-relief edifice.

Fig. 26.

Satellite image map of Mount Flint volcano of the McCuddin Mountains Volcanic Field, showing the approximate volcano outline (orange), caldera rim (dashed blue), studied outcrops (yellow), lithofacies and ages. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on source and an explanation of image processing.

Fig. 26.

Satellite image map of Mount Flint volcano of the McCuddin Mountains Volcanic Field, showing the approximate volcano outline (orange), caldera rim (dashed blue), studied outcrops (yellow), lithofacies and ages. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on source and an explanation of image processing.

Reynolds Ridge is a north–south-orientated, 1 km-long, linear, glacially-eroded outcrop ridge situated 5 km NW of the base of Mount Flint. The ridge consists of a 20 m-thick trachyte lava (LeMasurier et al. 1990b) underlain by a 30 m-thick potassium-feldspar-rich, coarse-grained (crystals up to 4 cm) syenite (i.e. intrusive equivalent of trachyte) (Wilch 1997). Our field observations show that the hypabyssal syenite intrudes into the lava. 40Ar/39Ar ages of anorthoclase separates from the syenite hypabyssal intrusion and trachyte lava are 20.45 ± 0.10 and 20.48 ± 0.12 Ma, respectively. Reynolds Ridge records the earliest felsic (trachyte) MBLVG volcanism in the Early Miocene at 20.5 Ma.

The relationship between early Miocene Reynolds Ridge and Mount Flint are uncertain. LeMasurier's (1990b) interpretations that Reynolds Ridge is either a dome at the base of Mount Flint or a separate remnant of a trachyte shield volcano are both viable possibilities in the absence of more data. We tentatively assign Reynolds Ridge to the edifice-building phase of Mount Flint (see Table 2).

Mount Petras is a glacially dissected nunatak (2867 m asl), with about 900 m of relief exposed above the level of the WAIS. Mount Petras is located about 10 km SE of Mount Flint. The eroded spurs that comprise the main nunatak cover an area of about 5 × 8 km.

Mount Petras is an unusual and important locality because it exposes the oldest known volcanic rocks in the MBLVP and these rocks overlie a currently high-standing non-conformity eroded into Cretaceous rhyodacite basement rocks, K–Ar dated as 82.9 ± 5.7 Ma (LeMasurier and Wade 1976; age adjusted for the decay constant of Steiger and Jaeger 1977). Wilch and McIntosh (2000) significantly revised the original interpretations of Mount Petras summarized in LeMasurier (1990c). The differences in interpretations are important as they have implications for the timing of the onset of MBLVP volcanism, the earliest volcanic records of glaciation, and the nature and significance of pre-volcanic erosion surfaces.

LeMasurier (1990c) interpreted volcanic rocks at Mount Petras as remnants of an Oligocene–early Miocene (25–22 ka) subglacially erupted table mountain overlying a flat, uplifted unconformity surface. The helicopter-supported reconnaissance fieldwork focused on the largest volcanic outcrop at Mount Petras, located on the SW flank. Volcanic rocks there were interpreted as 200 m of subhorizontally-stratified basaltic hyaloclastite, composed of weakly vesicular clasts and lacking any significant subaerial component (LeMasurier 1990c). These interpretations of the deposit characteristics, together with the observation of interbedded rounded basement clasts, led LeMasurier (1990c) to conclude that the volcanic rocks were the remnants of subglacially-erupted sequences and evidence for a Late Oligocene–Early Miocene thick ice sheet in West Antarctica. The original estimates of c. 400 m relief of the erosion surface (LeMasurier and Wade 1976) was revised to <100 m (LeMasurier et al. 1981).

Wilch and McIntosh (2000) used detailed volcanic lithofacies and age data to reach very different conclusions; they identified five eruptive episodes based on 40Ar/39Ar ages, geochemistry, lithofacies analysis and field relationships (Fig. 27). The first stage of volcanism occurred at 36.58 ± 0.22 Ma with an apparently subaerial extrusion of massive mugearite lava (Fig. 27; Table 8). This is the oldest known eruption in the MBLVP. The second stage of eruptions included four pyroclastic hawaiite events from three different vents dated between 28.96 ± 0.22 and 27.53 ± 0.23 Ma. Lithofacies associated with the second stage include welded tuff breccia attributed to a subaerial Strombolian eruption, and stratified lapilli tuff, massive lapilli tuff and ash-coated lapilli tuff attributed to subaerial to shallow-water Surtseyan eruptions. Two of the lapilli tuff units contained rare (<5%) lithic clasts. The lithic clasts are sub-angular to subrounded basement and mugearitic blocks, which are up to 10 cm in diameter and, in some cases, coated with hawaiite lava. No signs of glacial moulding or polish were observed on any of the lithic clasts. Intact and disintegrated pyroclastic bombs and blocks occur as large clasts up to 30 cm in length in some deposits.

Fig. 27.

Topographical profile of the summit area of Mount Petras of the McCuddin Mountains Volcanic Field, showing a schematic cross-section of volcanic outcrops, interpretations and ages. Figure adapted from figure 2 in Wilch and McIntosh (2000). The inset map is derived from the United States Geological Survey McCuddin Mountains 1:250 000 topographical map.

Fig. 27.

Topographical profile of the summit area of Mount Petras of the McCuddin Mountains Volcanic Field, showing a schematic cross-section of volcanic outcrops, interpretations and ages. Figure adapted from figure 2 in Wilch and McIntosh (2000). The inset map is derived from the United States Geological Survey McCuddin Mountains 1:250 000 topographical map.

Wilch and McIntosh (2000) observed that the outcrops and inferred vents were in contact with basement rocks and that topographical relief on the basement unconformity was >400 m, consistent with the early estimate by LeMasurier and Wade (1976). Wilch and McIntosh (2000) concluded that the 29–27.5 Ma Mount Petras eruptions involved intermittent interaction with water derived from a thin, local ice cap or from snow and ice on the slopes of a relatively high-relief (> 400 m) bedrock nunatak. We believe there is no evidence of extensive ice-sheet glacial conditions, as suggested by LeMasurier (1990c).

The 29–27.5 Ma pyroclastic deposits at Mount Petras provide the oldest terrestrial evidence for glacial ice in MBL but offer no evidence for a thick, continental, ice sheet at that time. The mixed Surtseyan and Strombolian eruptions imply local or intermittent contact with external water, which Wilch and McIntosh (2000) infer resulted from melting of a thin, local ice cap or ice and snow on slopes. The 29–27.5 Ma tuff cone deposits overlie an erosional unconformity, with >400 m of topographical relief. The relatively high-relief pre-volcanic environment is suggestive of ongoing erosion and is inconsistent with interpretations of a regional, low-relief, early Cenozoic West Antarctic Erosion Surface (e.g. Rocchi et al. 2006).

The USAS Escarpment is an east–west-orientated north-facing escarpment located at about 76° S (Fig. 1). It is mostly snow- and ice-covered with five nunataks of basement and/or volcanic rocks exposed at the break in slope along 50 km of the escarpment. Ice on the north side of the escarpment is c. 100–300 lower than ice on the south side. LeMasurier (1990c) noted that the USAS structural escarpment is aligned with Mount Petras. Two volcanic nunataks along the escarpment have been described by LeMasurier (1990c) and unpublished observations of the authors.

At Mount Aldaz, basaltic pyroclastic fall deposits and spatter-fed lavas overlie a felsic hypabyssal intrusive basement outcrop. The top surface of basement rock is smoothed, undulating and striated, and interpreted as a glacial unconformity. Bedrock and volcanic sequence are separated by a 1 m-thickness of strongly deformed laminated sandstone, interpreted as tillite, overlain by a well-sorted, lithic-rich, Strombolian scoria unit grading upwards into a welded bomb deposit, which is overlain by aphyric lava breccias and multiple coherent lavas. The Strombolian deposits are cut by a dyke that appears to have caused local palagonitization in adjacent scoria beds. The volcanic sequence is 40Ar/39Ar dated to 19.68 ± 0.31 Ma (Table 8). The basement rock, currently at about 2300 m asl, was eroded by a wet-based glacier prior to 19.7 Ma. The Mount Aldaz glacial unconformity is more indicative of ice-sheet glaciation than the Mount Petras outcrops, and records the earliest terrestrial evidence for a significant WAIS in the MBLVG. The >19.7 Ma unconformity at Mount Aldaz provides the oldest terrestrial evidence for an ice sheet in MBL.

At an unnamed outcrop c. 4 km east of Mount Galla, a palagonitized thinly-bedded hawaiite lapilli tuff is dominated by vesicular lapilli with subordinate fracture-bounded non-glassy lithic clasts and rare basement clasts. The bedded lapilli tuff exhibits local channelling and dune bedding. The bedded, palagonitized lapilli tuff is locally underlain by basanitic monomict breccia, possibly a vent-clearing explosion breccia. The sequence is interpreted as a tuff cone remnant and 40Ar/39Ar dated to 26.40 ± 0.21 Ma (Table 8).

The north–south-trending Ames Range Volcanic Field covers an area of about 30 × 15 km and consists of three coalesced central volcanoes: Mount Andrus, Mount Kosciuscko and Mount Kauffman (Fig. 28). The volcanoes are moderately dissected and glaciated with snow- and ice-covered slopes. Based on the limited field observations and sampling, it appears that trachyte lava dominates the Ames Range volcanoes, with a general trend of younger volcanism towards the north. LeMasurier and Rex (1989) noted that the Ames Range is aligned with the Quaternary Shepard Island volcano, 135 km north of Mount Kauffman, and may represent renewed felsic activity along the Ames Range lineament. Paulsen and Wilson (2010) used the ages and alignment of volcanoes in the Ames Range and Koerner Bluff at Mount Bursey in the Flood Range, located 10 km to the south, to establish the timing of a change in regional stress orientation after 6 Ma.

Fig. 28.

Satellite image map of the Ames Range Volcanic Field: Mount Andrus, Mount Kosciusko and Mount Kauffman. The map shows studied outcrops (yellow), scoria cone remnants (red), ages, lithofacies and approximate outlines of volcano extent (orange) and caldera rim (dashed blue). Abbreviations: BV, Brown Valley; CG, Coleman Glacier; RG, Rosenberg Glacier. Data suggest an age progression from south to north. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Fig. 28.

Satellite image map of the Ames Range Volcanic Field: Mount Andrus, Mount Kosciusko and Mount Kauffman. The map shows studied outcrops (yellow), scoria cone remnants (red), ages, lithofacies and approximate outlines of volcano extent (orange) and caldera rim (dashed blue). Abbreviations: BV, Brown Valley; CG, Coleman Glacier; RG, Rosenberg Glacier. Data suggest an age progression from south to north. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Mount Andrus, at the south end of the Ames Range, is about 13 × 13 km in map view and has an estimated volume of 115 km3 (Fig. 28). The volcano has a partially preserved summit caldera that reaches 2978 m asl and is breached on its west side. Coleman Glacier originates in the caldera, and flows west through the breach and down into the WAIS at about 1600 m asl. There are extensive rock exposures on the south and north sides of Coleman Glacier.

Lind Ridge on the south side of Coleman Glacier has almost continuous outcrop from 2000 to 2800 m asl where it joins the caldera rim. Early reconnaissance reports provided no specific details on exposed rock at Lind Ridge, except that outcrops are dominated by trachyte lava, with minor hydroclastic deposits and no magmatic tuff (LeMasurier 1990a). Wilch (1997) provided additional age and geochemical data on four samples from the ridge (Table 8). A trachyte lava near the base of the ridge was dated to 12.71 ± 0.06 Ma; this lava had a massive base and a frothy top that included welded bombs (Wilch 1997). The main ridge outcrops include a welded pyroclastic fall deposit, vertically foliated clastogenic pumiceous trachyte lava and massive trachyte lava. The pumiceous lava and massive lava were 40Ar/39Ar dated to 11.19 ± 0.08 and 11.18 ± 0.19 Ma, respectively (Table 8). We interpret the welded trachyte fall and clastogenic lava deposits as products of Strombolian-style eruptions. Two trachyte lava samples from the north flank of Mount Andrus yielded identical K–Ar ages of 11.3 ± 0.8 Ma (LeMasurier 1990a).

González-Ferrán and González-Bonorino (1972) mentioned four parasitic scoria cones on the lower west flank of Mount Andrus, noted that three of them appeared aligned north–south and suggested their vent locations are fault-controlled. Analysis of satellite imagery suggests that there are remnants of numerous parasitic scoria cones on the south and west flanks of Mount Andrus (Fig. 28). The cones have partially preserved crater rims that appear to be breached on one side. Narrow linear ridges extending down the slope from a few of the breached rims may be lava-flow levees. Only two parasitic cones have been analysed geochemically and both are basanite (LeMasurier 1990a; Wilch 1997). The lowest elevation outcrop at Mount Andrus is a basanite lava from a glacially eroded scoria cone situated SW of Lind Ridge, with a 40Ar/39Ar age of 9.28 ± 0.06 Ma (Table 8). A parasitic scoria cone located midway up Mount Andrus near Lind Ridge was K–Ar dated but the analysis yielded no radiogenic argon and the reported age was <0.1 Ma (LeMasurier 1990a).

Mount Kosciusko overlaps with Mount Andrus in the area of Rosenberg Glacier (Fig. 28). It has a summit elevation of 2909 m asl, a c. 3 km-diameter summit caldera and covers an area of about 16 × 16 km. The Mount Kosciusko caldera is about 10 km NNE of the Mount Andrus caldera. Outcrops appear to be limited to one locality north of Rosenberg Glacier, and include older phonolite lav,a K–Ar dated to 10.0 ± 0.8 Ma, and younger parasitic Strombolian scoria cones, K–Ar dated to 10.3 ± 0.8 and 8.66 ± 0.70 Ma, with a mean age of 9.2 ± 1.3 Ma (Table 8) (LeMasurier and Rex 1983; LeMasurier 1990a).

Mount Kauffman, situated at the north end of the Ames Range, is the smallest central volcano and the least studied in the Ames Range (Fig. 28). It is connected to Mount Kosciusko by Gardiner Ridge and has a summit elevation of 2364 m asl. Only one locality has been visited and the rock was described as trachyte lava (LeMasurier 1990a). Two conventional K–Ar dating analyses of one exposure yielded non-overlapping ages of 7.6 ± 0.6 and 5.9 ± 1.0 Ma (Table 8), suggesting that the volcano is Late Miocene in age and the youngest of the Ames Range. Based on helicopter reconnaissance, LeMasurier (1990a) described inaccessible cliff exposures on the SE side as interbedded lavas and fragmental debris, and possibly representing a caldera or an explosion crater.

The Flood Range Volcanic Field in western Marie Byrd Land (Figs 1 & 5) consists of three east–west-aligned trachytic central shield volcanoes: Mount Bursey (10–6 Ma), Mount Moulton (6–4 Ma) and Mount Berlin (2.8 Ma–active) (Table 9). Mount Bursey, although part of the Flood Range Volcanic Field, forms a solitary edifice located about 20 km east of Mount Moulton and about 10 km south of the north–south-oriented Ames Range Volcanic Field. Mount Moulton and Mount Berlin are separated by a high-elevation (2100 m asl), c. 10 km-wide, saddle. Mount Moulton and Mount Berlin are significant obstacles to the north-flowing WAIS, with surface ice elevations 600–800 m higher on the upstream (south) sides than on the downstream sides of these volcanoes. All three of the Flood Range central volcanoes are covered with snow and ice, and appear to be relatively undissected. Each of the three central volcanoes is a compound polygenetic volcano with two east–west-aligned summit calderas. The preservation of calderas indicates that minimal erosion has occurred, although past overriding by the WAIS at Mount Bursey and Mount Moulton cannot be precluded. Very limited outcrops on Mount Bursey and Mount Moulton provide a glimpse of their histories. Mount Berlin is the best-exposed volcano in the volcanic field and provides the most detailed geological history in the range.

Table 9.

Flood Range Volcanic Field summary age table

Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Mount Bursey
93-156A8.56 ± 0.06Starbuck crater lava16
93-158A0.25 ± 0.03Syrstad Rock, bomb, not in situ16
24K10.40 ± 0.80Syrstad Rock, lava2
93-160A10.08 ± 0.06Koerner Bluff, lava16
93-161A10.12 ± 0.34Koerner Bluff, bomb16
29K9.31 ± 0.74Felsic cone on caldera rim3
28AK0.49 ± 0.12Hutt Peak parasitic cone3
27K6.04 ± 0.48Heaps Rock, lava3
Mount Moulton
93-146A1.12 ± 0.15Gawne Nunatak, bomb16
93-146A1.00 ± 0.10Gawne Nunatak, bomb16
A1.04 ± 0.04Gawne Nunatak, bomb16
93-318A5.95 ± 0.05Prahl Crag, obsidian welded fall16
2AF4.80 ± 0.61Prahl Crag15
93-345A4.03 ± 0.14Edward's Spur, platy lava16
Mount Berlin
Summit caldera trachyte deposits
93-16A0.0104 ± 0.0053Fumarolic ice-cave floor trachyte lava20
93-25A0.0184 ± 0.0058Non-welded pumice fall20
93-22A0.0375 ± 0.0120Laminated pumice, top of upper welded unit20
93-22A0.0274 ± 0.0031Laminated pumice, top of upper welded unit20
93-22A0.0255 ± 0.0031Laminated pumice, top of upper welded unit20
93-22A0.0268 ± 0.0037Mean age (n = 3 analyses)20
93-17A0.0253 ± 0.0059Pumiceous rheomorphic tuff, upper welded unit20
93-21A0.0280 ± 0.0037Welded fall, below WCM93-2220
93-21A0.0246 ± 0.0029Welded fall, below WCM93-2220
93-21A0.0259 ± 0.0039Mean age (n = 2 analyses)20
93-23A0.0344 ± 0.0130East side, lower welded trachyte unit20
93-23A0.0259 ± 0.0049East side, lower welded trachyte unit20
93-23A0.0330 ± 0.0108East side, lower welded trachyte unit20
93-23A0.0279 ± 0.0064Mean age (n = 3 analyses)20
93-15A0.0284 ± 0.0110Spatter lava with cognate xenoliths20
93-15A0.0241 ± 0.0029Spatter lava with cognate xenoliths20
93-15A0.0261 ± 0.0042Spatter lava with cognate xenoliths20
93-15A0.0249 ± 0.0029Mean age (n = 3 analyses)20
A0.0259 ± 0.0020Welded fall deposits (n = 5 samples)20
Merrem Peak caldera trachyte
93-130A0.1432 ± 0.00571 m-thick pumiceous phonolite fall (non-welded)20
93-123A0.1631 ± 0.0260Black/yellow fall (SE)20
93-128A0.1862 ± 0.0046bomb (WNW)20
93-129A0.1763 ± 0.0170Pumiceous trachyte fall (NE)20
93-129A0.1877 ± 0.0046Pumiceous fall (NE)20
93-133A0.1843 ± 0.0048Welded trachyte fall (NE)20
93-133A0.1934 ± 0.0075Welded trachyte fall (NE)20
93-133A0.1869 ± 0.0091Mean age (n = 2 analyses)20
93-139A0.1844 ± 0.0046Welded clastogenic flow20
93-140A0.182 ± 0.014Welded clastogenic flow20
A0.1860 ± 0.0029Welded trachyte fall deposit (n = 5 samples)20
Flank trachyte deposits
93-011A0.234 ± 0.011Near-vent trachyte lava (NE flank)20
93-152A0.231 ± 0.012Wedemeyer Rock, welded ignmibrite (SE)20
A0.2328 ± 0.0083Flank trachyte deposits mean age (n = 2 samples)20
Mefford Knoll mafic deposits
93-001A0.2142 ± 0.0346Basanite cinder cone20
93-001A0.1994 ± 0.0487Basanite cinder cone20
93-001A0.210 ± 0.031Mean age (n = 2 analyses)20
93-004A0.215 ± 0.028Basanite cinder cone20
93-004A0.216 ± 0.036Basanite cinder cone20
93-004A0.216 ± 0.022Mean age (n = 2 analyses)20
93-008A0.210 ± 0.080Hawaiite flow levee20
A0.214 ± 0.018Mefford Knoll mafic mean age (n = 3 samples)20
Merrem Peak (SW) trachyte lava
93-134A0.456 ± 0.175Lava20
93-127A0.460 ± 0.041Foliated lava20
A0.460 ± 0.040SW flank trachyte lava 2 (n = 2 samples)20
93-125A0.580 ± 0.015Foliated lava20
93-126A0.573 ± 0.012Clastogenic trachyte lava20
93-126A0.571 ± 0.013Clastogenic trachyte lava20
93-126A0.572 ± 0.009Mean age (n = 2 analyses)20
93-137A0.583 ± 0.013Lava20
93-138A0.588 ± 0.040Clastogenic lava20
93-135A0.594 ± 0.018Clastogenic phonolite20
93-135A0.588 ± 0.048Clastogenic phonolite20
93-135A0.560 ± 0.020Clastogenic phonolite20
93-135A0.587 ± 0.036Mean age (n = 3 analyses)20
93-009A0.597 ± 0.028Trachyte flow levee (NW)20
A0.578 ± 0.009SW flank trachyte lava 1 (n = 6 samples)20
Brandenberger Bluff trachyte/phonolite tuya
93-014A2.71 ± 0.07Phonotephrite cinder cone, upslope from bluff20
93-010A2.80 ± 0.11Phonolite dome lava20
93-037A2.86 ± 0.24Fractured phonolite lava at bluff base20
93-053A2.76 ± 0.09Trachyte clast in hyalotuff20
93-121A2.79 ± 0.12Phonolite clast in hyalotuff20
93-250A2.72 ± 0.13Trachyte clast in hyalotuff20
A2.77 ± 0.06Mean age (n = 5 samples)20
Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Mount Bursey
93-156A8.56 ± 0.06Starbuck crater lava16
93-158A0.25 ± 0.03Syrstad Rock, bomb, not in situ16
24K10.40 ± 0.80Syrstad Rock, lava2
93-160A10.08 ± 0.06Koerner Bluff, lava16
93-161A10.12 ± 0.34Koerner Bluff, bomb16
29K9.31 ± 0.74Felsic cone on caldera rim3
28AK0.49 ± 0.12Hutt Peak parasitic cone3
27K6.04 ± 0.48Heaps Rock, lava3
Mount Moulton
93-146A1.12 ± 0.15Gawne Nunatak, bomb16
93-146A1.00 ± 0.10Gawne Nunatak, bomb16
A1.04 ± 0.04Gawne Nunatak, bomb16
93-318A5.95 ± 0.05Prahl Crag, obsidian welded fall16
2AF4.80 ± 0.61Prahl Crag15
93-345A4.03 ± 0.14Edward's Spur, platy lava16
Mount Berlin
Summit caldera trachyte deposits
93-16A0.0104 ± 0.0053Fumarolic ice-cave floor trachyte lava20
93-25A0.0184 ± 0.0058Non-welded pumice fall20
93-22A0.0375 ± 0.0120Laminated pumice, top of upper welded unit20
93-22A0.0274 ± 0.0031Laminated pumice, top of upper welded unit20
93-22A0.0255 ± 0.0031Laminated pumice, top of upper welded unit20
93-22A0.0268 ± 0.0037Mean age (n = 3 analyses)20
93-17A0.0253 ± 0.0059Pumiceous rheomorphic tuff, upper welded unit20
93-21A0.0280 ± 0.0037Welded fall, below WCM93-2220
93-21A0.0246 ± 0.0029Welded fall, below WCM93-2220
93-21A0.0259 ± 0.0039Mean age (n = 2 analyses)20
93-23A0.0344 ± 0.0130East side, lower welded trachyte unit20
93-23A0.0259 ± 0.0049East side, lower welded trachyte unit20
93-23A0.0330 ± 0.0108East side, lower welded trachyte unit20
93-23A0.0279 ± 0.0064Mean age (n = 3 analyses)20
93-15A0.0284 ± 0.0110Spatter lava with cognate xenoliths20
93-15A0.0241 ± 0.0029Spatter lava with cognate xenoliths20
93-15A0.0261 ± 0.0042Spatter lava with cognate xenoliths20
93-15A0.0249 ± 0.0029Mean age (n = 3 analyses)20
A0.0259 ± 0.0020Welded fall deposits (n = 5 samples)20
Merrem Peak caldera trachyte
93-130A0.1432 ± 0.00571 m-thick pumiceous phonolite fall (non-welded)20
93-123A0.1631 ± 0.0260Black/yellow fall (SE)20
93-128A0.1862 ± 0.0046bomb (WNW)20
93-129A0.1763 ± 0.0170Pumiceous trachyte fall (NE)20
93-129A0.1877 ± 0.0046Pumiceous fall (NE)20
93-133A0.1843 ± 0.0048Welded trachyte fall (NE)20
93-133A0.1934 ± 0.0075Welded trachyte fall (NE)20
93-133A0.1869 ± 0.0091Mean age (n = 2 analyses)20
93-139A0.1844 ± 0.0046Welded clastogenic flow20
93-140A0.182 ± 0.014Welded clastogenic flow20
A0.1860 ± 0.0029Welded trachyte fall deposit (n = 5 samples)20
Flank trachyte deposits
93-011A0.234 ± 0.011Near-vent trachyte lava (NE flank)20
93-152A0.231 ± 0.012Wedemeyer Rock, welded ignmibrite (SE)20
A0.2328 ± 0.0083Flank trachyte deposits mean age (n = 2 samples)20
Mefford Knoll mafic deposits
93-001A0.2142 ± 0.0346Basanite cinder cone20
93-001A0.1994 ± 0.0487Basanite cinder cone20
93-001A0.210 ± 0.031Mean age (n = 2 analyses)20
93-004A0.215 ± 0.028Basanite cinder cone20
93-004A0.216 ± 0.036Basanite cinder cone20
93-004A0.216 ± 0.022Mean age (n = 2 analyses)20
93-008A0.210 ± 0.080Hawaiite flow levee20
A0.214 ± 0.018Mefford Knoll mafic mean age (n = 3 samples)20
Merrem Peak (SW) trachyte lava
93-134A0.456 ± 0.175Lava20
93-127A0.460 ± 0.041Foliated lava20
A0.460 ± 0.040SW flank trachyte lava 2 (n = 2 samples)20
93-125A0.580 ± 0.015Foliated lava20
93-126A0.573 ± 0.012Clastogenic trachyte lava20
93-126A0.571 ± 0.013Clastogenic trachyte lava20
93-126A0.572 ± 0.009Mean age (n = 2 analyses)20
93-137A0.583 ± 0.013Lava20
93-138A0.588 ± 0.040Clastogenic lava20
93-135A0.594 ± 0.018Clastogenic phonolite20
93-135A0.588 ± 0.048Clastogenic phonolite20
93-135A0.560 ± 0.020Clastogenic phonolite20
93-135A0.587 ± 0.036Mean age (n = 3 analyses)20
93-009A0.597 ± 0.028Trachyte flow levee (NW)20
A0.578 ± 0.009SW flank trachyte lava 1 (n = 6 samples)20
Brandenberger Bluff trachyte/phonolite tuya
93-014A2.71 ± 0.07Phonotephrite cinder cone, upslope from bluff20
93-010A2.80 ± 0.11Phonolite dome lava20
93-037A2.86 ± 0.24Fractured phonolite lava at bluff base20
93-053A2.76 ± 0.09Trachyte clast in hyalotuff20
93-121A2.79 ± 0.12Phonolite clast in hyalotuff20
93-250A2.72 ± 0.13Trachyte clast in hyalotuff20
A2.77 ± 0.06Mean age (n = 5 samples)20

Notes: method A is 40Ar/39Ar; method K is K/Ar; method F is fission track.

Mount Bursey (Fig. 29), located at the east end of the Flood Range Volcanic Field, is a 20 × 30 km ice-covered massif with two east–west-aligned calderas at Hutt Peak and adjacent to Koerner Bluff (LeMasurier 1990e). Mount Bursey is interpreted as two coalesced trachytic shield volcanoes (LeMasurier 1990e). The Koerner Bluff caldera is situated 300 m lower than the Hutt Peak caldera and is interpreted to be the oldest caldera, 40Ar/39Ar dated to 10.08 ± 0.06 Ma (Table 9). Outcrops associated with the Koerner Bluff caldera include a phonolite lava dome at Koerner Bluff, a felsic crater rim called Starbuck Crater, basaltic rocks at Syrstad Rock, a parasitic felsic cone on the west side of the caldera rim and a parasitic hawaiite cone also on the caldera rim (LeMasurier 1990e; Wilch 1997). The phonolite lava dome is flat topped, and is composed entirely of flow-foliated lava with slickensided flow surfaces. A bomb from the hawaiite scoria cone on the caldera rim was 40Ar/39Ar dated to 10.12 ± 0.34 Ma, consistent with the 10.08 ± 0.06 Ma age for the lava dome. A felsic cone on the north side of the caldera was K–Ar dated to 9.31 ± 0.74, also consistent with the dome age. Starbuck Crater, downslope of Koerner Bluff, is an eroded crater rim with a 40Ar/39Ar age of 8.56 ± 0.06 Ma. The eroded crater is composed of subaerially-erupted trachyte that exhibits welded spatter transitional to clastogenic lava textures. A basaltic sample from Syrstad Rock was dated by conventional K–Ar to 10.4 ± 0.8 Ma. Mugearite pyroclastic rocks also crop out at Syrstad Rock; a mugearite sample has a 40Ar/39Ar date of 0.25 ± 0.03 Ma. The summit caldera at Hutt Peak was dated to 6.04 ± 0.48 Ma, based on a conventional K–Ar date for a sample from Heaps Rock just below the caldera. A mafic cone on the caldera rim of Hutt Peak was K–Ar dated to 0.49 ± 0.12 Ma and interpreted as a late-stage parasitic vent (LeMasurier 1990e).

Fig. 29.

Satellite image map of the Mount Bursey volcano at the east end of the Flood Range Volcanic Field. The map shows studied outcrops, ages, lithofacies and approximate outlines of volcano extent (orange) and caldera rim (dashed blue). Asterisks indicate conventional 40Ar/39Ar ages. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Fig. 29.

Satellite image map of the Mount Bursey volcano at the east end of the Flood Range Volcanic Field. The map shows studied outcrops, ages, lithofacies and approximate outlines of volcano extent (orange) and caldera rim (dashed blue). Asterisks indicate conventional 40Ar/39Ar ages. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

In summary, edifice-building activity associated with Koerner Bluff caldera occurred from 10.08 to 8.56 Ma, based on high-precision 40Ar/39Ar dating (Table 9). The higher and younger Mount Bursey caldera formed at 6.04 Ma, based on K–Ar dating. Mount Bursey late-stage Pleistocene parasitic eruptions occurred at 0.49 and 0.25 Ma. All of Mount Bursey outcrops appear to have resulted from subaerial eruptions and there is no evidence of ice–magma interactions, although the Starbuck Crater appears to have been overridden by ice at some point after the rocks were erupted. No erratics were observed at Starbuck Crater.

Mount Moulton is an almost entirely snow- and ice-covered, 15 × 40 km-long, east–west-orientated massif, with an estimated volume of 325 km3 (Fig. 30) (LeMasurier and Kawachi 1990d). There are two distinct c. 5–7 km-diameter summit calderas: the eastern caldera associated with the older Prahl Crags, here named the Prahl Crags caldera; and the younger western caldera at Britt Peak, here named the Mount Moulton caldera. There may be a third smaller caldera near Kohler Dome, about 9 km east of the Prahl Crags caldera (González-Ferrán and González-Bonorino 1972), although its morphology is less distinct than the other two. The Mount Moulton calderas are aligned east–west, and the massif and calderas are aligned with Mount Bursey and Mount Berlin, located just to the east and west, respectively. The caldera morphologies become increasingly distinct from east to west, suggesting progressive younging towards the west. Mount Moulton has three known outcrops, Prahl Crags, Edwards Spur at the NW end of the massif, and Gawne Nunatak at the western flank near Wells Saddle which separates Mount Moulton and Mount Berlin.

Fig. 30.

Satellite image map of the Mount Moulton volcano, located in the centre of the Flood Range Volcanic Field. The map shows studied outcrops, ages, lithofacies and approximate outlines of volcano extent (orange) and caldera rim (dashed blue). The eastern caldera at Kohler Dome is speculative. Possible caldera and degree of erosion suggest that Mount Moulton is composed of three coalesced shield volcanoes that are progressively younger to the west. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Fig. 30.

Satellite image map of the Mount Moulton volcano, located in the centre of the Flood Range Volcanic Field. The map shows studied outcrops, ages, lithofacies and approximate outlines of volcano extent (orange) and caldera rim (dashed blue). The eastern caldera at Kohler Dome is speculative. Possible caldera and degree of erosion suggest that Mount Moulton is composed of three coalesced shield volcanoes that are progressively younger to the west. Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Prahl Crags consists of a peralkaline rhyolite lava and an associated welded fall deposit, with obsidian fiamme, and is 40Ar/39Ar dated to 5.95 ± 0.05 Ma (Table 9). Prahl Crags is interpreted as a fragment of the eastern caldera rim and represents the early phase of Mount Moulton volcanism. Edwards Spur consists of a rheomorphic welded trachyte fall deposit, with a slightly ropy texture and interlayered with a platy lava (Wilch 1997). Although located near the base of the volcano, Edwards Spur is the only felsic rock proximal to the western caldera. The 40Ar/39Ar age of Edwards Spur is 4.03 ± 0.14 Ma (Table 9) and is inferred to date the growth of the younger caldera. Gawne Nunatak is a remnant of a parasitic hawaiite scoria cone, consisting of weakly-welded pyroclastic lapilli and bomb deposits. The 40Ar/39Ar age of Gawne Nunatak at 1.04 ± 0.04 Ma (Table 9) provides the only constraint on post-edifice-building volcanism. The sampled outcrops at Mount Moulton all exhibit evidence of subaerial eruptions and deposition, with no evidence of glaciovolcanic interactions.

In 1993–94 a very significant finding at Mount Moulton was a sequence of dipping englacial tephra layers exposed in blue ice in the eastern caldera, just north of Prahl Crags (Wilch et al. 1999). The englacial tephra layers were sampled in reconnaissance and 40Ar/39Ar dated by Wilch et al. (1999), and revisited and studied in detail in 1999–2000 and 2003–4 (Dunbar et al. 2008). A total of 48 tephra layers were observed. Most are trachytic in composition with some mafic layers (Dunbar et al. 2008). 40Ar/39Ar dating of K-feldspar phenocrysts from coarse-grained pumice-rich trachytic tephra layers yielded eight stratigraphically consistent ages, ranging from 500 to 10 ka. The trachytic tephra layers were erupted from explosive eruptions of Mount Berlin, located 30 km west of the site; the tephra and their chronology are described further by Dunbar et al. (2021). A few of the tephra layers are mafic in composition and may have been derived from parasitic vents on the flanks of Mount Berlin or Mount Moulton (Dunbar et al. 2008).

Mount Berlin is an active volcano with steaming fumaroles and fumarolic ice caves. It is a predominantly trachytic polygenetic composite volcano with a prominent summit caldera and an older subsidiary caldera at Merrem Peak (LeMasurier and Wade 1968; LeMasurier and Kawachi 1990a; Wilch et al. 1999) (Figs 31, 32 & 33). Mount Berlin, at the west end of the Flood Range Volcanic Field, has an estimated volume of 125 km3 (LeMasurier and Kawachi 1990a) and basal dimensions of 18 × 18 km at the level of the WAIS. Mount Berlin is largely undissected, except for the older north-facing Brandenberger Bluff. The volcanic history is recorded in caldera-rim and flank deposits on the volcano, and in distal englacial tephra deposits at Mount Moulton and in ice cores (see Dunbar et al. 2021). Wilch et al. (1999) provided a detailed analysis of the volcanic geology and documented three stages in the growth of Mount Berlin, summarized below.

Fig. 31.

Satellite image map of Mount Berlin, in the western Flood Range Volcanic Field. The map shows outcrop age ranges and rock types, and approximate outlines of volcano extent (orange) and caldera rim (dashed blue). The three stages of the growth of Mount Berlin are discussed in the text; the number (n) of age determinations included in the age ranges of each stage are listed after the ages. In 1993, there were multiple steaming ice towers on the caldera rim; the ice tower visited is marked IT. Letter m designates local moraine. The summits of Merrem Peak (3000 m asl) and Mount Berlin (3478 m asl) are shown with stars. Image source: Google Earth Pro image accessed June 2019; the caption to Figure 7 for more information on the source and an explanation of image processing. Elevations from the Mount Berlin (1973) quadrangle, scale 1:250 000, USGS Reconnaissance Series, Antarctica, United States Geological Survey.

Fig. 31.

Satellite image map of Mount Berlin, in the western Flood Range Volcanic Field. The map shows outcrop age ranges and rock types, and approximate outlines of volcano extent (orange) and caldera rim (dashed blue). The three stages of the growth of Mount Berlin are discussed in the text; the number (n) of age determinations included in the age ranges of each stage are listed after the ages. In 1993, there were multiple steaming ice towers on the caldera rim; the ice tower visited is marked IT. Letter m designates local moraine. The summits of Merrem Peak (3000 m asl) and Mount Berlin (3478 m asl) are shown with stars. Image source: Google Earth Pro image accessed June 2019; the caption to Figure 7 for more information on the source and an explanation of image processing. Elevations from the Mount Berlin (1973) quadrangle, scale 1:250 000, USGS Reconnaissance Series, Antarctica, United States Geological Survey.

Fig. 32.

Satellite image map of the Mount Berlin summit caldera at the west end of the Flood Range Volcanic Field. Circles locate fumarolic ice towers and steaming vents identified in 1993. The double circle locates the ice tower entrance to the ice cave that was entered, and from which lava on the caldera floor was sampled and dated to 10.3 Ma (Wilch et al. 1999). A c. 150 m-thick exposure of caldera-wall deposits are highlighted in red. The satellite image was derived from Google Earth in June 2019; image sources: 2019 Digital Globe.

Fig. 32.

Satellite image map of the Mount Berlin summit caldera at the west end of the Flood Range Volcanic Field. Circles locate fumarolic ice towers and steaming vents identified in 1993. The double circle locates the ice tower entrance to the ice cave that was entered, and from which lava on the caldera floor was sampled and dated to 10.3 Ma (Wilch et al. 1999). A c. 150 m-thick exposure of caldera-wall deposits are highlighted in red. The satellite image was derived from Google Earth in June 2019; image sources: 2019 Digital Globe.

Fig. 33.

Photographs of lithofacies and features of Mount Berlin, Flood Range Volcanic Field. (a) Lapilli tuff with ash-coated lapilli clasts from near the top of Brandenberger Bluff at Mount Berlin. (b) The 143.2 ka non-welded pyroclastic fall breccia and lapilli tuff near the rim of the Merrem Peark caldera. (c) The 25.9 ka welded fall deposit in the wall of the Berlin summit caldera. The person is pointing to the 1 m-long flattened pumice bomb. Tephra from this and many other Mount Berlin eruptions was found 30 km away in the summit caldera at Prahl Crags on Mount Moulton, as well as in multiple West Antarctic ice cores. (d) A steaming fumarolic ice tower in the Mount Berlin summit caldera. A lava sample from the floor of the ice cave was dated to 10.4 ± 5.4 ka.

Fig. 33.

Photographs of lithofacies and features of Mount Berlin, Flood Range Volcanic Field. (a) Lapilli tuff with ash-coated lapilli clasts from near the top of Brandenberger Bluff at Mount Berlin. (b) The 143.2 ka non-welded pyroclastic fall breccia and lapilli tuff near the rim of the Merrem Peark caldera. (c) The 25.9 ka welded fall deposit in the wall of the Berlin summit caldera. The person is pointing to the 1 m-long flattened pumice bomb. Tephra from this and many other Mount Berlin eruptions was found 30 km away in the summit caldera at Prahl Crags on Mount Moulton, as well as in multiple West Antarctic ice cores. (d) A steaming fumarolic ice tower in the Mount Berlin summit caldera. A lava sample from the floor of the ice cave was dated to 10.4 ± 5.4 ka.

Brandenberger Bluff, located on the north side of Mount Berlin, is a c. 350 m-high Pliocene (2.77 ± 0.06 Ma (n = 5): Fig. 31; Table 9) lava and volcaniclastic edifice that exhibits evidence for subglacial, emergent and subaerial glaciovolcanic palaeoenvironments (Wilch et al. 1999). The north-facing stratigraphic sequence is about 250 m thick and 1 km wide, and is composed of steeply-dipping (20°–30°), well-stratified, fine-grained trachytic/phonolitic vitric tuff that overlies a highly jointed and brecciated, aphyric, glassy trachyte/phonolite lava (Wilch 1997). The deposits include thick sections of cyclic packages of normally-graded vitric lapilli tuff and vitric tuff and massive fine vitric tuff that are interpreted as volcaniclastic turbidites that formed fan deposits in an ice-marginal lake. The dominance of a fine-grained glassy matrix in the deposits is attributed to intense phreatomagmatic fragmentation during the eruption. The volcaniclastic sediment sequences are interrupted by areas of large (tens of metres in diameter) slide blocks of similar material that are interpreted as collapse breccias formed by gravitational failure of fan deposits during growth of the subaqueous fan. Throughout the stratigraphic sequence both pumice lapilli and blocky, aphyric, glassy lapilli are common. Stratified vitric lapilli tuff at the top of the bluff exhibits shallow-dipping planar beds and cross-beds, and contains abundant ash-coated lapilli, interpreted as subaerial phreatomagmatic pyroclastic density current deposits (Fig. 34a). The margins of the bluff top surface include areas of intense soft-sediment deformation. The Brandenberger Bluff sequence and edifice are interpreted as a felsic variation of a tephra-dominated tuya (Smellie and Edwards 2016), composed of subaqueously-emplaced lava and reworked tuff deposits at its base and on its flanks, and subaerially-emplaced phreatomagmatic tuff deposits at the top of the bluff. We infer that this sequence was deposited in an intraglacial lake setting. The abundant deformation evident in the volcaniclastic deposits may have developed as supporting ice walls collapsed. The basal lava and several clasts from within the tuff cone sequences were dated by the 40Ar/39Ar method, with a mean age of 2.77 ± 0.06 Ma (n = 5) (Table 9). Phonotephritic scoria cone deposits, located upslope and about 100 m in elevation above the top of the bluff, yield an overlapping age of 2.71 ± 0.07 Ma (Table 9) that suggests a single eruptive phase. These subaerially-erupted phonotephrite tuff deposits indicate complete emergence above ice level.

Fig. 34.

Satellite image map of the Hobbs Coast nunataks in western MBL. The map shows studied outcrops, ages, lithofacies and inferred eruptive environments. Adapted from Wilch and McIntosh (2007). Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Fig. 34.

Satellite image map of the Hobbs Coast nunataks in western MBL. The map shows studied outcrops, ages, lithofacies and inferred eruptive environments. Adapted from Wilch and McIntosh (2007). Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

The passage zone from subaqueous to subaerial depositional environments is situated near the top of Brandenberger Bluff, approximately 250 m above the level of the WAIS on the north side of Mount Berlin. Currently, ice-sheet flow is obstructed by Mount Berlin and the elevation of the WAIS on the upstream (south) side of Mount Berlin is at about 1800–2000 m asl, about 400–600 m higher than on the downstream side (Fig. 31). In middle Pliocene time, when the Brandenberger Bluff tuya emerged above ice level, it is likely that the Mount Berlin edifice did not exist and there was probably no local obstruction to ice flow. In the absence of Mount Berlin, local ice elevations at Brandenberger Bluff would be at c. 1800 m asl. Remarkably, the exposed passage zone at Brandenberger Bluff may actually record a lower late Pliocene WAIS level, because of changes in ice-flow patterns after its eruption, caused by growth of the Mount Berlin edifice.

The second and most voluminous phase of activity at Mount Berlin is characterized by growth of the volcano at Merrem Peak to 3000 m asl and by eruptions from the 2.5 × 1 km-diameter Merrem Peak caldera. The constructional slopes of Merrem Peak volcano average about 13°. The earliest activity associated with this stage of activity is recorded in pumiceous foliated and clastogenic trachyte lava located just west of Merrem Peak, and dated to 578 ± 9 ka (n = 6: Table 9). Three subsequent eruptive episodes at Merrem Peak caldera are recorded by proximal deposits west of the caldera, including a trachytic lava dated to 460 ± 40 ka, an 18+ m-thick, densely- to incipiently-welded, trachytic pumiceous pyroclastic breccia dated to 186.0 ± 2.9 ka, and a 1 m-thick, phonolitic, non-welded pumice lapilli layer dated to 143.2 ± 5.7 ka (Fig. 33b). Both of the 186.0 and 143.2 ka fall deposits mantle topography. The youngest (143.2 ka) fall deposit is situated on the rim of the Merrem Peak caldera.

Compositionally-diverse eruptions also occurred during this Merrem Peak caldera interval of volcanism. A pyroclastic vent breccia and a near-vent trachyte lava (234 ± 11 ka) are located on the NE flank of Merrem volcano. On the south flank, a welded trachytic ignimbrite (231 ± 12 ka) contains c. 30% fiamme, with typical aspect ratios of 1:10 (height:length). The ignimbrite may have originated from the caldera. The NE flank trachyte lava and south flank ignimbrite are geochemically identical, with a mean age of 232.8 ± 8.3 ka. A benmoreite lava overlies the ignimbrite and signifies a c. 230 ka or younger effusive eruption. On the NW flank, basanite–hawaiite scoria cone remnants yield a mean age of 214 ± 18 ka.

The final and still-active phase of Mount Berlin volcanism was marked by growth of the volcano by more than 400 m to 3478 m asl and a southeastward shift of the vent area to the 2 km-diameter summit caldera (Figs 31 & 32). The constructional slopes of the volcano above Merrem Peak caldera range from 19° to 30°. Two prominent welded trachytic pyroclastic fall units, totalling more than 150 m in thickness, are exposed in the eastern wall of the summit caldera. These pyroclastic fall deposits are composed of slightly flattened pumiceous bombs and abundant cognate xenoliths. The welding is interpreted as agglutination with minor load-pressure compaction. Recumbent folds seen in one unit suggest that part of the pyroclastic deposit flowed rheomorphically. Ages of five samples from the two lower welded units are analytically indistinguishable and range from 27.9 ± 6.4 to 24.9 ± 2.9 ka, with a mean age of 25.9 ± 2.0 ka (Table 9) (Fig. 33c). These are locally overlain by a >10 m-thick sequence of welded and non-welded pyroclastic fall beds. The fall deposits include a 6 m-thick lithic- and ash-rich explosion breccia containing bomb and lithic clasts up to 50 cm in diameter, and a 2 m-thick densely welded obsidian fall. Anorthoclase phenocrysts from the obsidian yielded a maximum age of 18.4 ± 5.8 ka. These caldera-wall deposits at Mount Berlin were previously interpreted as lava on the basis of reconnaissance investigations (LeMasurier and Kawachi 1990a).

Several fumarolic ice towers and steaming vents along the summit caldera rim attest to ongoing geothermal activity (Figs 32 & 33d). One ice tower opens into an underlying ice-cave system more than 70 m long. Lava exposed on the cave floor is dated to 10.4 ± 5.3 ka (Table 9). In 1993, surface temperatures of the cave-floor lava were as high as 12°C. Distal englacial tephra from Mount Berlin have been documented in the ice-filled Mount Moulton caldera, located 30 km east of the summit caldera (Wilch et al. 1999; Dunbar et al. 2008). On the basis of the size and density of Mount Moulton pumice clasts, and distance from their Mount Berlin source, Wilch et al. (1999) calculated eruption column heights ranging from 28 to 40 km, indicative of highly-explosive Plinian eruptions.

Mount Berlin is a mostly trachytic polygenetic central volcano that was constructed in three stages (Wilch et al. 1999). Stage I was the growth of the c. 500 m-high, 2.77 myr-old, trachytic Brandenberger Bluff tephra-dominated tuya, in which successive subglacial, phreatomagmatic and subaerial lithofacies record the emergence of the volcano above a palaeo-ice-sheet surface. This volcano formed in an englacial lake, with a passage zone about 250 m above current ice level to the north of the bluff. Stage II is the growth of the Merrem Peak shield volcano with eruptions from the 3000 m asl Merrem Peak caldera from 578 to 143 ka. Most of the estimated volume of Mount Berlin was erupted during construction of the Merrem Peak volcano. This stage was mostly trachytic in composition and was dominated by variably-welded pumiceous pyroclastic rocks. The third and ongoing stage of Mount Berlin volcanism is recorded in explosively-erupted pyroclastic deposits dating back to 25.9 ka and preserved in the summit caldera that is up to 3478 m asl. Englacial tephra found at Mount Moulton and in ice cores records highly-explosive Mount Berlin volcanism during the Merrem Peak and Berlin summit caldera stages (Dunbar et al. 2021).

Several monogenetic, mostly basaltic, volcanic centres located at inland and island sites near the Hobbs Coast in western Marie Byrd Land comprise the Hobbs Coast Volcanic Field (Figs 1 & 34). The inland nunataks include volcanic centres mostly located on the north–south-orientated Demas Range, just east of Berry Glacier; Bowyer Butte is the exception, located 20 km to the west and just west of Venzke Glacier (Fig. 34). Granite bedrock is exposed in contact with or beneath the volcanic rocks at several locations. Basement and volcanic rocks of the inland nunataks exhibit abundant evidence of past glaciations. Wilch and McIntosh (2007) described the lithofacies characteristics, stratigraphic relationships and 40Ar/39Ar geochronology of all sites except Bowyer Butte; the summary here is based on this study unless otherwise noted. Ages of volcanism are mostly Late Miocene but range from 11.45 to 2.57 Ma (Table 10). The discussion of the inland volcanic centres is organized according to interpretations of dominant eruptive palaeoenvironments.

Table 10.

Hobbs Coast Volcanic Field summary age table

Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Shepard Island and nearby islands
Shepard Island
93-255A0.481 ± 0.049Lava-flow interior, ponded flows in Mathewson Point tuff cone16
93-258A0.56 ± 0.14Lava clast from within tuff beds16
93-265A0.41 ± 0.09Mount Petinos, lithic clast in tuff beds16
S9FK0.43 ± 1.20Worley Point platy lava6
Grant Island
GI11EK0.7 ± 0.2Hyaloclastite from tuff cone6
Cruzen Island
CI51CK2.75 ± 0.26Subaerial basalt top of tuya6
Hobbs Coast Nunataks
Coleman Nunatak
93-199A2.57 ± 0.06Dense lava, below upper surge deposit19
93-185A2.64 ± 0.28Lava, SW moat19
93-184A2.70 ± 0.14Lava, SW moat19
93-192A2.61 ± 0.20Lava, SW moat19
93-188A2.49 ± 0.12Basal lava flow SW moat19
93-208A2.59 ± 0.12North end, clast within scoriaceous dyke19
93-207A2.62 ± 0.09North end, dyke19
93-205A2.79 ± 0.27North end, dyke19
93-202A2.84 ± 0.23North end, dyke19
A2.60 ± 0.08Mean (n = 9)
Cousins Rock
93-218A4.92 ± 0.10Bomb19
93-219A4.96 ± 0.11Lava19
A4.94 ± 0.14Mean (n = 2)
Shibuya Peak
93-221A5.07 ± 0.14Bomb19
93-223A5.35 ± 0.20Lava crusts19
93-229A4.80 ± 0.11Dyke19
A4.98 ± 0.29Mean (n = 2)
Patton Bluff
93-307A11.45 ± 0.23Lava, 15 m thick19
Kouperov Peak
93-303A9.24 ± 0.11South end, lava over unconformity (n = 2)19
93-304A9.07 ± 0.57Lava (n = 2)19
A9.27 ± 0.22Mean (n = 2)
Kennel Peak
93-297A8.01 ± 0.87Pillow lobe interior19
Holmes Bluff
93-299A6.36 ± 0.07Lava filling valley19
50K8.39 ± 0.66Subaerial lava at south end, near Kennel Peak6
Bowyer Butte
57DK9.82 ± 1.805 m basalt lava on hyaloclastite over glacially-striated basement6
Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Shepard Island and nearby islands
Shepard Island
93-255A0.481 ± 0.049Lava-flow interior, ponded flows in Mathewson Point tuff cone16
93-258A0.56 ± 0.14Lava clast from within tuff beds16
93-265A0.41 ± 0.09Mount Petinos, lithic clast in tuff beds16
S9FK0.43 ± 1.20Worley Point platy lava6
Grant Island
GI11EK0.7 ± 0.2Hyaloclastite from tuff cone6
Cruzen Island
CI51CK2.75 ± 0.26Subaerial basalt top of tuya6
Hobbs Coast Nunataks
Coleman Nunatak
93-199A2.57 ± 0.06Dense lava, below upper surge deposit19
93-185A2.64 ± 0.28Lava, SW moat19
93-184A2.70 ± 0.14Lava, SW moat19
93-192A2.61 ± 0.20Lava, SW moat19
93-188A2.49 ± 0.12Basal lava flow SW moat19
93-208A2.59 ± 0.12North end, clast within scoriaceous dyke19
93-207A2.62 ± 0.09North end, dyke19
93-205A2.79 ± 0.27North end, dyke19
93-202A2.84 ± 0.23North end, dyke19
A2.60 ± 0.08Mean (n = 9)
Cousins Rock
93-218A4.92 ± 0.10Bomb19
93-219A4.96 ± 0.11Lava19
A4.94 ± 0.14Mean (n = 2)
Shibuya Peak
93-221A5.07 ± 0.14Bomb19
93-223A5.35 ± 0.20Lava crusts19
93-229A4.80 ± 0.11Dyke19
A4.98 ± 0.29Mean (n = 2)
Patton Bluff
93-307A11.45 ± 0.23Lava, 15 m thick19
Kouperov Peak
93-303A9.24 ± 0.11South end, lava over unconformity (n = 2)19
93-304A9.07 ± 0.57Lava (n = 2)19
A9.27 ± 0.22Mean (n = 2)
Kennel Peak
93-297A8.01 ± 0.87Pillow lobe interior19
Holmes Bluff
93-299A6.36 ± 0.07Lava filling valley19
50K8.39 ± 0.66Subaerial lava at south end, near Kennel Peak6
Bowyer Butte
57DK9.82 ± 1.805 m basalt lava on hyaloclastite over glacially-striated basement6

Notes: method A is 40Ar/39Ar; method K is K/Ar.

Subglacial volcanic palaeoenvironments are inferred at two localities: Kennel Peak (Wilch and McIntosh 2007) and Bowyer Butte (LeMasurier 1990d). At Kennel Peak, a c. 125 m-thick section of hawaiitic pillow lavas and interpillow hyaloclastite breccia (8.01 ± 0.87 Ma) unconformably overlies a smooth moulded outcrop of granitic basement rocks (Wilch and McIntosh 2007). A <1 m-thick glacial till occurs locally along the unconformity between the volcanic and basement rocks. The top of the thick pillow hyaloclastite sequence is eroded and overlain by crudely-graded, planar-stratified hyaloclastite lapilli tuff. The pillow lavas and hyaloclastites at Kennel Peak are inferred to have been deposited in a subglacial chamber onto thin basal till and glacially-moulded bedrock. The vesicularity and steep dips of the interbedded pillow lava and hyaloclastite, and the presence of similar age (K–Ar 8.39 ± 0.66 Ma) subaerial lavas nearby at an outcrop just north of Kennel Peak (LeMasurier 1990d), suggest that the pillow hyaloclastite deposits form part of a subaqueous lava-fed delta sequence, possibly associated with subaerial lava effusion. The pillow lava and hyaloclastite sequence extends only c. 25 m above today's ice surface, suggesting that the syneruptive local palaeo-ice level was at least slightly higher than today's ice level (Wilch and McIntosh 2007). At Bowyer Butte, a volcanic sequence of subaerial lava over thin hyaloclastite beds (1–5 m) resting on a striated granite surface has a low precision K–Ar age of 9.82 ± 1.8 Ma (LeMasurier 1990d).

Other volcanic centres along the Berry Glacier show evidence of emergent phreatomagmatic and subaerial eruptions that, in places, deposited volcanic rocks on a previously glaciated bedrock surface. At Patton Bluff, the oldest known volcanic outcrop in the area consists of 11.45 ± 0.23 Ma subaerially-erupted basaltic lava that is situated near, but not in direct contact with, local basement rocks. At Kouperov Peak, two samples with duplicate 40Ar/39Ar ages that average 9.27 ± 0.22 Ma consist of subaerially-erupted basaltic lava which unconformably overlies basement rocks. The unbrecciated lava is about 30 m above the present ice surface and the lava–basement contact is poorly exposed but extends over >60 m of topographical relief. The basement rocks are locally striated to within 2 m of the lava, although no striations were observed beneath the lava. The Kouperov Peak locality is tentatively interpreted as a glacial unconformity (older than 9.2 Ma) that is overlain by a subaerial lava.

At Holmes Bluff, a c. 55 m-thick, columnar-jointed, valley-filling 6.36 ± 0.07 Ma (Table 10) lava with a 1.3 m-thick welded and ropy basal breccia overlies a 1.5 m-thick achnelith-rich lapilli tuff and a partially-moulded, slightly-weathered, basement unconformity. The sequence is tentatively interpreted as a basement glacial unconformity (older than 6.4 Ma) that was subsequently exposed and locally weathered before being covered by a thin, subaerially-erupted, pyroclastic fall deposit and a thick, valley-filling, subaerial lava.

Three samples from Shibuya Peak with a mean age of 4.98 ± 0.29 Ma are from an unusual volcaniclastic deposit with abundant exotic granitoid clasts exposed in a >100 m-thick section (Table 10). LeMasurier (1990d) inferred a subglacial eruptive environment based on the presence of bedded basaltic hyaloclastite and interbedded tillite. Wilch and McIntosh (2007) inferred that the volcanic rocks resulted from subaerial Strombolian and phreatomagmatic eruptions, noting the bedding and cross-stratification, the variably vesicular lapilli tuff and tuff breccia, and the mix of sideromelane and tachylite glass clasts. They noted that some of the granitoid cobbles and boulders were encrusted with a thin (1 cm) coat of basaltic lava, and argued that this is consistent with eruption through a tillite but not consistent with the clasts being part of an interbedded glacial tillite. This interpretation provides evidence for a pre-volcanic (>c. 5 Ma) glaciation at this locality. An outcrop of subaerially-erupted phonotephrite and welded agglutinate and slightly reworked phreatomagmatic vitric tuff at Cousins Rock, a small (150 × 150 m) nunatak located near Shibuya Peak, yielded an age of 4.94 ± 0.14 Ma, similar in age to the Shibuya Peak sample.

The 1 × 3 km Coleman Nunatak is one of the largest and the youngest (2.60 ± 0.08 Ma) monogenetic volcanic centres exposed in the Hobbs Coast Volcanic Field (Table 10). Coleman Nunatak has a 75 m-thick, 1 km-wide sequence, well exposed in the wind moat on the south end of the nunatak, which suggests that the volcano is an emergent tuff cone sequence. The base of the sequence includes minor pillow and glassy lava overlain by compound lava with subaerial breccia, and interbedded vitric lapilli tuff and vitric lapillistone units. The vitric lapilli tuff units include bread-crust bombs with bedding-plane sags and ash-coated lapilli, indicative of magmatic and phreatomagmatic fragmentation processes. Discontinuous, erosive, massive, planar and cross-stratified beds are interpreted as turbulent pyroclastic density current deposits (Fisher and Schmincke 1984). The combination of turbulent pyroclastic density current deposits and minor subaqueous lavas is indicative of an emergent intraglacial environment, where interaction with meltwater controlled the explosive phreatomagmatic eruption style and modified the lavas slightly. Cross-cutting exposures of steeply-dipping lapilli tuff and breccia are interpreted as evidence of a vent funnel unconformity overlain by vent slurry lapilli tuff (Sohn and Park 2005; White and Ross 2011).

The inland nunataks of the Hobbs Coast Volcanic Field provide compelling evidence for Miocene and Pliocene ice-sheet glaciation near coastal West Antarctica. The syneruptive thicknesses of the palaeo-ice sheets are difficult to determine because the volcanic rocks and glacial deposits and surfaces are positioned on glacial interfluves. The subglacial lithofacies at two sites and emergent–subaerial lithofacies at the other five sites are generally consistent with regional ice levels similar to the current WAIS at multiple times between 11.5 and 2.6 Ma.

Three small volcanic islands are located offshore and are connected to the WAIS by the Getz Ice Shelf (Fig. 1). Shepard Island and Grant Island are situated 30 km offshore and located 50–70 km NE of Holmes Bluff, the northeasternmost inland nunatak described above. Cruzen Island is situated 30 km offshore along the Rupert Coast, and is located about 240 km west of Shepard Island and 160 km west of Bowyer Butte, the northwesternmost inland nunatak described above.

Shepard Island is a mostly snow- and ice-covered, low-relief, 12 km-diameter volcanic island, with outcrops concentrated along the coast. The island has four main outcrop areas: Mathewson Point, Mount Petinos, Worley Point and Mount Colburn. The rocks are dominated by phreatomagmatic tuff cone sequences erupted between 500 and 400 ka. Our observations confirm and amplify those by LeMasurier (1990g), and show that Mathewson Point is a remnant of a monogenetic basanite tuff cone composed of well-bedded and cross-bedded lapilli tuff. Other features include abundant ash-coated lapilli and small bombs (up to 12 cm) and with bedding-plane sags. The 0.8 km-diameter tuff cone rim at Mathewson Point is well preserved with inward- and outward-dipping beds. A late-stage subaerial lava is ponded in the centre of the tuff cone crater. A lava clast from within the lapilli tuff beds is 40Ar/39Ar dated to 0.56 ± 0.14 Ma and the ponded lava is dated to 0.481 ± 0.049 Ma (Table 10). The lapilli tuff and ponded lava are likely to have the same age of about 481 ka.

A c. 500 m-thick sequence of palagonitized hawaiite volcaniclastic deposits and lava extends from sea level to the summit of Mount Petinos. The lapilli tuff and tuff breccia exhibit many of the same textures as the Mathewson Point tuff cone, including planar bedding, accretionary lapilli and bedding-plane sags, suggesting that the sequence resulted from phreatomagmatic tuff cone eruptions. A lithic clast in the lapilli tuff was 40Ar/39Ar dated to 0.41 ± 0.09 Ma (Table 10). Undated trachyte lava is exposed at Mount Colburn, the highest point on Shepard Island. Worley Point consists of trachyte lava, K–Ar dated to 0.43 ± 1.2 Ma (LeMasurier 1990g). The Mount Colburn trachyte may be coeval with a trachyte lava collected at Worley Point.

Grant Island is a mostly snow- and ice-covered, large (35 × 20 km), low-relief island that has a single outcrop situated on Mount Obiglio, the high point (510 m asl) of the island. Mount Obiglio is a small isolated cone composed of palagonitized Strombolian tephra and K–Ar dated to 0.7 ± 0.2 Ma (LeMasurier 1990g).

Cruzen Island is a snow- and ice-capped volcanic island, 1.7 × 2.2 km in diameter. The island is interpreted as a 200 m-high, flat-topped, basalt tuya K–Ar dated to 2.75 ± 0.26 Ma (LeMasurier 1990g). Cliff outcrops expose 50–75 m of hyaloclastite overlain by horizontal subaerial lavas. A passage zone from subglacial to subaerial palaeoenvironment is at least 75 m asl and suggests expanded ice cover in the late Pliocene.

The Fosdick Mountains Volcanic Field is located in the Ford Range in far western MBL (Fig. 1). The volcanic field consists of at least 18 separate Quaternary (?) small volcanic centres, many of which are exposed on top of Paleozoic and Mesozoic bedrock (Fenner 1938; LeMasurier and Wade 1990; Gaffney and Siddoway 2007) (Fig. 35). During the course of extensive bedrock and structural geological studies of the Ford Range (Siddoway et al. 2004), many volcanic outcrops were visited and two geochemical studies (Gaffney and Siddoway 2007; Chatzaras et al. 2016) provide information on some of the volcanic geology.

Fig. 35.

Satellite image map of volcanoes of the Fosdick Mountains in western Marie Byrd Land. Volcanic outcrop locations are from Gaffney and Siddoway (2007). Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Fig. 35.

Satellite image map of volcanoes of the Fosdick Mountains in western Marie Byrd Land. Volcanic outcrop locations are from Gaffney and Siddoway (2007). Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Gaffney and Siddoway (2007) described several of the volcanic outcrops and noted that many volcanic sequences overlie periglacial features and glacial deposits, post-dating glacial event(s). They also noted that many of the rocks preserve delicate features, suggesting they have not been overridden by wet-based glacial ice (Gaffney and Siddoway 2007). In the Ochs Glacier area, volcanic necks and irregular conduits intrude bedrock (Gaffney and Siddoway 2007), including an eroded diatreme near Marujupu Peak (Chatzaras et al. 2016). At Mount Avers, thin vesicular pāhoehoe basanite lavas unconformably overlie the glaciated bedrock summit area (871 m asl) and descend to 610 m asl (C. Siddoway pers. comm.). An array of subvertical, 0.3–1.5 m-wide, dykes exposed over 500 m of vertical elevation cuts through the Mount Avers bedrock and have similar compositions to the summit lavas; the dykes are interpreted as feeder dykes for the summit lava (Gaffney and Siddoway 2007). At Mount Perkins, a c. 200 m sequence of 2–4 m-thick, massive (subaerial?) basanite lavas is capped by lava breccia. The Mount Perkins lava sequence is glacially incised and unconformably overlain by golden-tan-coloured (palagonitized?) tephra layers that dip inwards to a summit depression. The Mount Perkins sequence is tentatively interpreted here as a glacially-eroded subaerial lava sequence overlain by a tuff cone. Recess Nunatak consists of black, vesiculated, vertically-orientated flow-banded, glassy tephra interpreted as a glacially-eroded remnant of a vent complex (Gaffney and Siddoway 2007). Recess Nunatak is the only known site in the Fosdick Mountain Volcanic Field with crustal xenoliths. Many of the nunataks have ultramafic mantle xenoliths, which have been analysed for geochemistry and strain fabrics (Gaffney and Siddoway 2007; Chatzaras et al. 2016).

Two samples from Mount Perkins yielded indistinguishable 40Ar/39Ar ages of 1.41 ± 0.04 and 1.40 ± 0.03 Ma (Table 11); these new ages are considerably younger and more precise than previous K–Ar ages of 4.7 ± 1.0 and 3.5 ± 0.6 Ma (LeMasurier and Wade 1990). Otherwise, rocks of the Fosdick Mountain Volcanic Field are undated.

Table 11.

Fosdick Mountains summary age table

Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Fosdick Mountains
28K3.5 ± 0.6Mount Perkins5
30K4.7 ± 1.0Mount Perkins5
90207aA1.41 ± 0.04lava, Mount Perkins21
90207bA1.40 ± 0.03lava, Mount Perkins21
A1.41 ± 0.03Mean age (n = 2)21
Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Fosdick Mountains
28K3.5 ± 0.6Mount Perkins5
30K4.7 ± 1.0Mount Perkins5
90207aA1.41 ± 0.04lava, Mount Perkins21
90207bA1.40 ± 0.03lava, Mount Perkins21
A1.41 ± 0.03Mean age (n = 2)21

Notes: method A is 40Ar/39Ar; method K is K/Ar.

Ref. code: 5, LeMasurier and Rex (1982); 21, this study. See also Supplementary Material Table S1.

The Fosdick Mountains Volcanic Field offers the potential for detailed lithofacies and dating studies, similar to those undertaken by Wilch and McIntosh (2007) in the Hobbs Coast Volcanic Field. Such work would be likely to provide additional valuable information on the past extent of the WAIS at the time of the eruptions.

The Thurston Island Volcanic Province is located in western Ellsworth Land on the Thurston Island tectonic block, and includes the Hudson Mountains and Jones Mountains volcanic fields (Fig. 2). Many of the outcrops in the Hudson and Jones Mountains volcanic fields have been studied in reconnaissance, mostly in the 1960s (Rowley 1990; Rowley et al. 1990); detailed outcrop descriptions and geochronology data are not available.

The Hudson Mountains Volcanic Field consists of around 20, mostly snow- and ice-covered, volcanic nunataks at 200–750 m asl, located near the Walgreen Coast and just north of Pine Island Glacier in western Ellsworth Land (Fig. 36). No basement rocks are exposed (Rowley et al. 1990).

Fig. 36.

Satellite image map of volcanoes of the Hudson Mountains Volcanic Field of the Thurston Island Volcanic Province. Volcanoes are designated subaerial or subglacial/subaqueous; adapted from Rowley et al. (1990). Image Source: Google Earth Pro image accessed June 2019; see Figure 7 caption for more information on source and an explanation of image processing.

Fig. 36.

Satellite image map of volcanoes of the Hudson Mountains Volcanic Field of the Thurston Island Volcanic Province. Volcanoes are designated subaerial or subglacial/subaqueous; adapted from Rowley et al. (1990). Image Source: Google Earth Pro image accessed June 2019; see Figure 7 caption for more information on source and an explanation of image processing.

Previous workers suggested there are three, mostly ice-covered, major volcanoes at Teeters Nunatak, Mount Moses and Mount Manthe, which are overlain by smaller parasitic volcanoes (see Rowley et al. 1990). Analysed rocks range from tephrite to hawaiite in composition (Rowley et al. 1990). Three general lithofacies have been noted: a lower pillow lava and associated hyaloclastite tuff (as much as 60 m thick) observed at Mount Nickens; a middle subhorizontal to steeply-dipping stratified hyaloclastite tuff (as much as 350 m thick); and upper subhorizontal (including pāhoehoe) lavas (as much as 60 m thick) observed at Maish Nunatak, Mount Moses and Mount Manthe (see the map in Rowley et al. 1990). The ‘hyaloclastite’ tuffs are sideromelane rich and interpreted as subglacially erupted, and the upper lava is interpreted as suberially erupted and emplaced (Rowley et al. 1990). Samples from two sites, Mount Manthe and Velie Nunatak, have conventional K–Ar dates (Table 12). The most precise ages for Mount Manthe come from clasts within a hyaloclastite with a weighted mean age of 4.9 ± 0.6 Ma. The age of a lava from within a hyaloclastite at Velie Nunatak is 3.7 ± 0.4 Ma. Late Miocene ages (without errors) were reported by Lopatin and Polyakov (1974) for two other nunataks.

Table 12.

Thurston Island Volcanic Province summary age table

Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Hudson Mountains Volcanic Field
Velie Nunatak
28·3AK3.7 ± 0.40Lava within hyaloclastite tuff sequence5
Mount Manthe
42·6AK5.0 ± 0.6Clasts from near the top of a 200 m hyaloclastite section5
42·5AK4.7 ± 0.45
42·4AK5.1 ± 0.65
424.9 ± 0.6Mean age (n = 3)
H·6K5.6 ± 3.8Subaerial cap of a 200 m hyaloclastite section5
H-4K8.7 ± 2.0Clast within hyaloclastite roughly 50 m below the top5
H-2K5.0 ± 3.2Lava in the middle of a hyaloclastite tuff section5
Jones Mountains Volcanic Field
A7.68 ± 0.11Lava above basal pillows and unconformity from a peak west of Forbidden Rocks14
Sample IDMethodCorrected preferred age ± 2 SD (Ma)DescriptionRef.
Hudson Mountains Volcanic Field
Velie Nunatak
28·3AK3.7 ± 0.40Lava within hyaloclastite tuff sequence5
Mount Manthe
42·6AK5.0 ± 0.6Clasts from near the top of a 200 m hyaloclastite section5
42·5AK4.7 ± 0.45
42·4AK5.1 ± 0.65
424.9 ± 0.6Mean age (n = 3)
H·6K5.6 ± 3.8Subaerial cap of a 200 m hyaloclastite section5
H-4K8.7 ± 2.0Clast within hyaloclastite roughly 50 m below the top5
H-2K5.0 ± 3.2Lava in the middle of a hyaloclastite tuff section5
Jones Mountains Volcanic Field
A7.68 ± 0.11Lava above basal pillows and unconformity from a peak west of Forbidden Rocks14

Notes: method A is 40Ar/39Ar; method K is K/Ar.

There is great potential for new field and analytical work in the Hudson Mountains. Smellie and Edwards (2016) noted that the edifice shapes and described lithofacies are like those found in tuyas. Our Figure 36 is adapted from the Rowley et al. (1990) map and designates outcrops by lithofacies as either subaerial or subglacial/emergent. We agree that at least some of the subglacial/emergent lithofacies are associated with tuyas. It is likely that a well-documented and dated record of ice–volcanic interactions would provide valuable syneruptive records of past ice-sheet thicknesses.

The Jones Mountains Volcanic Field is located about 150 km NE of the Hudson Mountains Volcanic Field near the Eights Coast in western Ellsworth Land (Fig. 2). The mountains are mostly snow- and ice-covered, and include numerous nunataks that rise up to about 1500 m asl, up to 1000 m higher than the adjacent ice sheet (Fig. 37). The geology of the Jones Mountains has only been mapped in reconnaissance.

Fig. 37.

Satellite image map of volcanic outcrops (red) in the Jones Mountains Volcanic Field of the Thurston Island Volcanic Province. The dashed line separates the southern nunataks that have all volcanic rocks from northern nunataks that are basement rocks overlain by volcanic rocks on peaks. At Avalanche Ridge, the dashed line marks the glacial unconformity between Mesozoic basement and late Cenozoic subglacially emplaced volcanic rocks. Adapted from Rowley (1990), after Craddock et al. (1964). Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

Fig. 37.

Satellite image map of volcanic outcrops (red) in the Jones Mountains Volcanic Field of the Thurston Island Volcanic Province. The dashed line separates the southern nunataks that have all volcanic rocks from northern nunataks that are basement rocks overlain by volcanic rocks on peaks. At Avalanche Ridge, the dashed line marks the glacial unconformity between Mesozoic basement and late Cenozoic subglacially emplaced volcanic rocks. Adapted from Rowley (1990), after Craddock et al. (1964). Image source: Google Earth Pro image accessed June 2019; see the caption to Figure 7 for more information on the source and an explanation of image processing.

The nunataks in the north and east parts of the field consist of Mesozoic basement granitoids at lower elevations, overlain unconformably by alkali basalts on the peaks; in the south only, alkali basalts are exposed (Rowley 1990 after Craddock et al. 1964). There is widespread evidence of glaciation along the unconformity (Craddock et al. 1964; Rutford et al. 1968, 1972). Although thick sections (500–700 m) of basaltic rock are found in outcrop and numerous nunatak outcrops are shown on sketch maps (e.g. Rowley 1990), detailed lithofacies mapping and characterization have not been published.

Near Avalanche Ridge and Pillsbury Tower in the centre of the Jones Mountains, about 500–700 m of Miocene alkali basalt volcanic rocks overlie an unconformity cut into Mesozoic-aged basement rocks. The rocks were described as pillow lavas and nearly horizontal lapilli tuffs by Craddock et al. (1964). Hole et al. (1994) characterized the rocks as massive, highly-vesiculated pillow lava and palagonitized volcaniclastic rocks, subdivided into two units. Each unit has a c. 10 m-thick base composed of variably-stratified, cross-bedded and reworked volcaniclastic tuff and lapilli tuff. These basal units are interpreted as reworked mass flow deposits. The volcanic rocks are interpreted as subglacially erupted. Numerous attempts at K–Ar dating the basal pillow lavas yielded a wide range of ages ranging from 332 to 6 Ma (Table 12), and Rutford et al. (1972) suggested that the most reliable age for Jones Mountains volcanic rocks is between 12 and 7 Ma. The most recent and most reliable determination is a 40Ar/39Ar plateau age of 7.68 ± 0.11 Ma (Table 12) from a lava sample taken above the basal sequence from an unnamed nunatak 30 km west of Forbidden Rocks (Rutford and McIntosh 2007). The exact location is uncertain but the basanite had a higher K2O content than most samples from the Jones Mountains (Rutford et al. 1972).

Lenses of diamictite with striated and faceted erratic clasts are found along the basement unconformity (Rutford et al. 1972; Hole et al. 1994; Rutford and McIntosh 2007). The matrix of the diamictite is composed of volcanic glass and palagonite. Craddock et al. (1964) and subsequent workers have interpreted the diamictite as tillite. The nearly horizontal unconformity extends for 33 km and has a maximum estimated relief of c. 50 m (Rutford and McIntosh 2007). The unconformity is a polished and planed surface with striations and chattermarks (Craddock et al. 1964; Rutford et al. 1972). The unconformity is interpreted as a glacial unconformity that formed coeval with the eruption of the overlying glaciovolcanic sequence and deposition of volcanic-rich diamictite at c. 7.7 Ma (Table 12) (Rutford and McIntosh 2007).

Rutford and McIntosh (2007) noted that the interpretations of the Jones Mountains by Craddock et al. (1964) provided the first well-documented evidence for pre-Quaternary glaciations in Antarctica. There is a need of more detailed fieldwork in the Jones Mountains Volcanic Field, both at the unconformity locations and the several other volcanic nunataks. Establishing a high-resolution chronology will be challenging as the Jones Mountains rocks are largely basaltic with low K2O contents and glassy textures.

Volcanism and glaciation have been active geological processes in West Antarctica since middle Cenozoic time. As discussed in the individual volcano summaries, the MBLVG volcanic sequences provide snapshot records of eruption and emplacement conditions that can be used to provide syneruptive constraints on ice thickness and palaeoenvironmental conditions (Smellie 2018). A limitation of applying this approach in a currently glaciated area such as West Antarctica is that only records of ice levels that are thicker or equal to today's ice level can be observed. LeMasurier (1972) and LeMasurier and Rex (1982, 1983) reconstructed volcanic records of glaciation in West Antarctica, based on regional reconnaissance fieldwork and K–Ar geochronology. Other studies have provided updated records at selected MBL volcanoes (McIntosh et al. 1991; LeMasurier et al. 1994; Wilch et al. 1999; Smellie 2001; LeMasurier 2002; Wilch and McIntosh 2000, 2002, 2007; LeMasurier and Rocchi 2005). Here we develop a regional synthesis of glaciovolcanism in the MBLVG through time that merges palaeoenvironmental reconstructions and geochronology results.

In the introduction to this paper, we defined the common lithofacies at MBLVG volcanoes and interpreted them as subglacial (or ice-contact), emergent or subaerial to infer the eruptive and depositional palaeoenvironment and the influence of external water in the eruption dynamics. Smellie and Edwards (2016) provided a summary of the lithofacies approach. Detailed lithofacies descriptions and interpretations are a key first step in using the volcanic deposits to reconstruct past ice extent or conditions. In addition to lithofacies interpretations, there are other complexities of the West Antarctic glaciovolcanic environment that are critical to consider in reconstructing the volcanic record of glaciation in West Antarctica (Wilch 1997; Wilch and McIntosh 2000, 2002, 2007). These complexities are explained and addressed below in a conceptual model.

Here we review some general concepts and develop a conceptual model that addresses some of the complexities of the glaciovolcanic environment of West Antarctica. A major objective of the model is to differentiate records of local ice level from those of larger WAIS level. Aspects of the conceptual model are illustrated in Figure 38a–f and discussed below. Outcrop examples listed in Figure 38 were introduced in the volcano summaries and are discussed in further detail here.

Fig. 38.

Cartoons illustrating the key ideas of the conceptual model for interpreting the WAIS history from volcanic sequences. Details are described in the text.

Fig. 38.

Cartoons illustrating the key ideas of the conceptual model for interpreting the WAIS history from volcanic sequences. Details are described in the text.

Lithofacies analysis and interpretations of palaeoenvironments in relation to local ice conditions provide the foundation of these process-orientated studies. Early compilations of glaciovolcanism in MBL (e.g. LeMasurier and Rex 1982, 1983) relied heavily on general outcrop characteristics, typically interpreting rocks that exhibit any evidence of external water interactions (palagonitization, pillow lava, hyaloclastites, hydroclastic tuff) as evidence of formerly higher ice-sheet levels. If possible, direct evidence of glaciation, such as a tillite or a glacially-incised surface in a volcanic sequence, was used to add confidence to the interpretations. Hyaloclastite and lava-fed deltas and tuyas provide additional evidence of glaciovolcanic interactions (LeMasurier 2002; Smellie and Edwards 2016). Ideally, tuyas or lava-fed deltas preserve passage zones that can provide ‘dipstick’ measurements of syneruptive palaeo-ice levels (Fig. 38a). Lithofacies analyses and interpretations of glaciovolcanic sequences have become more detailed and nuanced over the past 30 years, with considerable advances in reconstructing the thickness and thermal properties of the glacial environment (Smellie and Edwards 2016). Although detailed lithofacies analysis is critical for understanding palaeoenvironments, an integrative reconstruction of volcanic records of ice-sheet glaciation also requires an understanding of the palaeoenvironments in a regional ice-sheet context.

One complication of the volcanic record of the WAIS is that all palaeo-ice-level elevations are relative to modern elevations and do not account for possible uplift or subsidence of volcanic centres (Fig. 38a). This uplift caveat is problematic in many terrestrial ice-sheet reconstructions and is a particular problem for reconstructions of older (pre-Quaternary) glaciations (e.g. Wilch et al. 1993). LeMasurier and Landis (1996) suggested that the progressively higher elevations of the pre-volcanic basement unconformity and older volcanoes towards the centre of the MBLVP are a surface expression of regional structural domal uplift that has been ongoing since inception of volcanism in middle Cenozoic times. Rocchi et al. (2006) suggested that the early Oligocene intrusion at Dorrel Rock near present-day Mount Murphy was rapidly exhumed (500 m Ma−1) and exposed at the surface between 34 and 27 Ma by fluvial and glacial erosion. According to their model, uplift beginning in the centre of the province near Mount Petras has proceeded since that time and has far outpaced erosion. Alternatively, the pre-volcanic unconformity may be a time-transgressive erosion surface that appears to be very flat in some places (e.g. Bowyer Butte) and quite rugged in other places (e.g. Mount Petras exhibits 400 m of relief and Mount Murphy exhibits 900 m of relief: McIntosh et al. 1991; Wilch and McIntosh 2000; Smellie 2001). Although the pattern of higher unconformities and older volcanoes towards the centre of the volcanic province is real, there are limited constraints on the timing and amount of (basement) surface uplift at any location. Wilson et al. (2013) conducted climate–ice sheet modelling experiments to account for calculated ice growth in Antarctica at the Eocene–Oligocene transition, and concluded that the surface topography of West Antarctica may have declined significantly due to glacial erosion and subsidence. Since the Oligocene the topography of West Antarctica has also responded glacioisostatically during glacial–deglacial cycles. In the absence of quantitative geological constraints on uplift or subsidence of MBLVG volcanoes, we propose tentative palaeo-ice-level elevations that assume no uplift or subsidence. The impact of this assumption is diminished by the fact that much of our WAIS reconstruction relies on the younger part of the volcanic history of the MBLVG (since 10 Ma) and the most definitive estimates of past ice level are in late Quaternary sequences.

The large polygenetic central volcanoes of the MBLVP are obstructions to regional WAIS ice flow, and produce higher ice levels upstream and lower ice levels downstream (Wilch et al. 1999; Wilch and McIntosh 2002) (Fig. 38b; Table 2). This ice-damming effect can give a false impression of palaeo-ice levels recorded in glaciovolcanic sequences. An example of the ice-damming effect is the coastal volcano, Mount Murphy, where upstream ice levels are 400–600 m higher than downstream ice levels. Elsewhere in MBLVP, many of the central volcanoes have coalesced to form linear ranges. Mount Berlin and Mount Moulton in the Flood Range form a continuous >60 km-long barrier to ice flow, with upstream ice levels 400–800 m higher than downstream ice levels. The tuya at Brandenberger Bluff on the north flank of Mount Berlin is >2 Ma older than the Mount Berlin polygenetic shield volcano and may have formed as an isolated nunatak. The passage zone at 250 m above modern local ice level may have formed when the regional ice-sheet level was similar to or lower than it is today. The more isolated MBLVP central volcanoes produce less significant ice-damming effects: for example, ice-sheet elevations surrounding the isolated inland Mount Takahe polygenetic volcano are <300 m higher on the upstream side.

Interpretations of early edifice-building sequences can be complicated by issues related to changing ice-flow patterns and ice-sheet levels during volcano construction. At many volcanoes, edifice-building outcrops are not exposed on the lower flanks of the large central volcanoes and problems relating to ice damming can be ignored. In other cases, some of the oldest rocks are exposed at the base of polygenetic volcanoes and palaeo-ice-level interpretations must account for the ice-damming effect (Fig. 38b).

Post-edifice-building flank eruptions and parasitic monogenetic eruptions that occur on pre-existing shields provide less-complicated palaeo-ice-level indicators (Fig. 38c). Presumably, ice-damming effects produced by the pre-existing shield had a similar effect on local ice-flow patterns and elevations as they do today. In flank deposits, lava-fed deltas with passage zones from subaqueous to subaerial environments (e.g. Mount Takahe) provide reliable ‘dipstick’ records of palaeo-ice levels that can used in glaciological models. Conversely, the lowest elevations of subaerially-erupted parasitic volcanoes (e.g. English Rock at Mount Frakes) provide limits on the maximum elevation of the syneruptive ice sheet at those locations. An important reminder is that these subaerial sequences indicate only maximum ice-sheet levels and may be recording times when there was little or no ice sheet developed in West Antarctica.

The volcanoes of the MBLVG are almost completely mantled in snow and ice of variable thickness ranging from no ice at the limited volcanic outcrops to probably >100 m in local glaciers and ice-filled calderas. This mantle of snow and ice provides ample opportunities for glaciovolcanic interactions above the level of the ice sheet. The possibility that glaciovolcanic sequences resulted from interactions with slope ice rather than the regional ice sheet must be evaluated at each site (Fig. 38d).

There are many non-Antarctic examples where historically erupted lavas have been observed overriding ice, cutting open channels into ice and tunnelling into ice on the flanks of volcanoes (see examples in Smellie and Edwards 2016). In these cases, the lavas were quenched and generally fragmented where they came into contact with melted ice. Lava–slope ice interactions offer conditions conducive to forming dipping ‘passage zones’ where conditions vary from subglacial to subaqueous to subaerial on both open volcanic slopes and within meltwater tunnels. In MBLVP volcanoes, glaciovolcanic interactions with slope ice above the level of the WAIS were dominant processes at Mount Rees and Mount Steere in the Late Miocene.

Models of eruptions beneath valley-confined glaciers provide possible analogues for the type of sequences produced on volcanoes with a moderately thin ice mantle (≤100 m). Eruptions beneath valley-confined glaciers have been inferred from volcanic sequences in Iceland (Walker and Blake 1966) and Alexander Island, Antarctica (Smellie et al. 1993). In both Icelandic and Alexander Island sequences, deposition in thermally-excavated subglacial tunnels resulted in a post-glacial landform in the form of a volcanogenic ‘esker’. Smellie et al. (2011a, b) describe volcanic sequences similar those at Mount Rees and Mount Steere from Hallett Coast volcanoes in North Victoria Land, and interpret them as ‘a‘ā-lava deltas formed beneath slope-mantling glacial ice. Explosive phreatomagmatic eruptions on volcano slopes may have an associated lava phase and produce dipping passage zones (Smellie et al. 1993) or they may simply produce monogenetic tuff cones or rings without an effusive pillow lava phase. Sohn (1996) suggested that phreatomagmatic pyroclastic density current eruptions are typically associated with deeper explosions in weak substrates, where water has only limited access to the vent. Such environments may have been common on the slopes of active Antarctic volcanoes.

Small volcanoes perched on interfluves between areas of faster flowing ice may exhibit records of changing local ice levels that give false impressions of the magnitude and elevations of regional ice-level variations (Fig. 38e). Examples of interfluve volcanoes include the satellite nunataks at Mount Murphy, and inland volcanic centres in the Hobbs Coast Volcanic Field and Fosdick Mountains Volcanic Field. Regional growth of the ice sheet may not cause higher ice levels at these nunataks if the local ice streams compensate by discharging ice at a greater rate. Records of higher local ice levels on the interfluves may indicate times of higher regional ice levels, if it can be shown that the passage zones lie above the level of the regional ice sheet. Records of unchanged ice levels are more difficult to interpret and may be associated with times of higher, similar or lower regional ice levels relative to today. Uncertainties about the timing of downcutting of the adjacent drainages weaken interpretations of regional palaeo-ice level using interfluve volcanoes.

A final point to consider is whether the site is recording inland or coastal ice-sheet thickening (Fig. 38f). Inland thickening by snow/ice accumulation will result in outward expansion and thickening in coastal regions (e.g. Cuffey and Paterson 2010). Lateral expansion of the grounding line will effectively raise ice-sheet levels more dramatically at coastal sites than at inland sites. Sea-level changes can further complicate ice flow at the margins of marine-based ice sheets, such as the WAIS. Sea-level rise can buoyantly lift and destabilize grounded ice margins, which may result in accelerated ice flow. Sea-level fall will result in expansion of grounded ice. Such sea-level changes may have affected local ice levels at coastal volcanoes (Mount Murphy) and at interfluve volcanoes adjacent to glaciers draining into the sea (the Hobbs Coast Volcanic Field).

Figure 39 provides a synthesis of palaeoenvironmental reconstructions and a proxy record of palaeo-ice levels of the WAIS since latest Eocene times. Supplementary material Table S4 summarizes data pertinent for palaeo-ice-level determinations at the time of the eruptions (syneruptive palaeo-ice levels), including the interpreted palaeoenvironment, the eruption age and the elevation relative to today's ice sheet. Subglacial deposits indicate that local syneruptive ice levels were at least as high as the outcrop elevation. Horizontal passage zones provide proxies for local syneruptive englacial lake levels and, by association, for minimum local palaeo-ice levels. Subaerial deposits suggest a ‘dry’ environment where no ice or meltwater was present and constrain the maximum local syneruptive elevation of the ice sheet relative to today's ice level. The regional ice levels (Supplementary material Table S4) account for complications discussed in the conceptual model above. A summary of the volcanic record of the WAIS through time is presented below.

Fig. 39.

Volcanic record of the WAIS in Marie Byrd Land with key events listed on the right. A plot of select eruption ages v. distance of the volcano from the coast. Subaerial volcanoes provide maximum syneruptive ice levels. Subaerial volcanoes listed as <100 m indicates that the outcrops are <100 m above the current ice-sheet level; those listed as <400 m include outcrops between 100 and 400 m above the current ice sheet level. Emergent eruptions are associated with Hobbs Coast Nunatak sites, near to present-day ice levels. Palaeo-ice-sheet ≥ today are sites with evidence of subglacial eruptions at or above today's ice level. At most sites, several caveats shown in Figure 37 apply, making it difficult to establish palaeo-ice levels with confidence. Palaeo-ice-level elevations are listed for late-stage eruption sites at Mount Takahe and Mount Murphy. Palaeo-ice sheet ≤ today refer to sites where there is evidence of sloping passage zones inferred to represent a thin mantle of ice on edifice side slopes.

Fig. 39.

Volcanic record of the WAIS in Marie Byrd Land with key events listed on the right. A plot of select eruption ages v. distance of the volcano from the coast. Subaerial volcanoes provide maximum syneruptive ice levels. Subaerial volcanoes listed as <100 m indicates that the outcrops are <100 m above the current ice-sheet level; those listed as <400 m include outcrops between 100 and 400 m above the current ice sheet level. Emergent eruptions are associated with Hobbs Coast Nunatak sites, near to present-day ice levels. Palaeo-ice-sheet ≥ today are sites with evidence of subglacial eruptions at or above today's ice level. At most sites, several caveats shown in Figure 37 apply, making it difficult to establish palaeo-ice levels with confidence. Palaeo-ice-level elevations are listed for late-stage eruption sites at Mount Takahe and Mount Murphy. Palaeo-ice sheet ≤ today refer to sites where there is evidence of sloping passage zones inferred to represent a thin mantle of ice on edifice side slopes.

The record of volcanism from 37 to 10 Ma is sparse but can be divided into three intervals: sporadic monogenetic volcanism with some evidence of glaciation (37–20 Ma); a period of apparent quiescence from 20 to 13.4 Ma; and then large-scale caldera-forming polygenetic volcanism from 13.4 to 10 Ma with no definitive evidence of glaciation. The earliest known volcanism in the MBLVG occurred in the late Eocene (36.58 ± 0.22 Ma), when massive mugearite lava was erupted at Mount Petras (Wilch and McIntosh 2000). Mid-Oligocene (29–27 Ma) Surtseyan and Strombolian deposits at Mount Petras provide the first terrestrial evidence for glacial ice in Marie Byrd Land (Wilch and McIntosh 2000). These eroded tuff cone deposits overlie bedrock on an unconformity with over 400 m of relief, suggesting a local hilly or mountainous terrain at 29–27 Ma. Wilch and McIntosh (2000) interpreted the deposits to be products of intermittent contact of magma with external water, probably derived from the melting of a thin, local ice cap or ice and snow on bedrock slopes. Three other MBL volcanic outcrops have been dated to the 36–20 Ma interval. A small hyalotuff outcrop situated at the level of the ice sheet near Mount Galla is dated to 26.40 ± 0.21 Ma and provides an additional suggestion of local late Oligocene glaciation, although by no means definitive. The volcanic record indicating Oligocene glaciers in West Antarctica is consistent with other proxy records from around Antarctica (e.g. Anderson 1999; Olivetti et al. 2015).

At Reynolds Ridge near the north end of Mount Flint, a 20.20 ± 0.08 Ma subaerial trachyte lava situated at the level of the ice sheet suggests that early Miocene syneruptive ice levels were lower than present levels. It is important to note that this and subsequent examples of subaerial deposits provide only maximum constraints and do not preclude the possibility that MBL was ice-free at the time of the eruptions. Palagonitized Strombolian tuff breccia deposits (19.68 ± 0.31) overlying striated bedrock at Mount Aldaz provide the first definitive indication for >19.7 Ma syneruptive ice levels similar to or lower than they are today.

Several polygenetic volcanoes in the Ames and Executive Committee ranges were apparently formed during the Middle–Late Miocene interval (13.4–10 Ma) interval, all of which are interpreted as consisting of subaerially-erupted, producing mostly felsic lavas and pyroclastic rocks (Fig. 39; see also Supplementary material Table S4). Many of the outcrops are situated near the level of the WAIS and suggest that maximum syneruptive WAIS levels were similar to or lower than today's level (Fig. 38: e.g. Mount Andrus, Mount Hampton and Mount Whitney). A small outcrop of palagonitized hyalotuff at Patton Bluff in the Hobbs Coast nunataks suggests limited hydrovolcanic interactions at 11.32 ± 0.25 Ma but does not provide definitive evidence of palaeo-ice.

The Late Miocene interval from 10 to 8 Ma marks an apparent acceleration in polygenetic volcanism in the MBLVG and provides the first substantial evidence for a widespread WAIS. Late Miocene (10–8 Ma) glaciovolcanic sequences are exposed at the inland Crary Mountains (Mount Steere and Mount Rees), at Kennel Peak and Bowyer Butte near the Hobbs Coast and at coastal Mount Murphy volcano. The Mount Murphy main shield sequence records fluctuating syneruptive (mostly 9.34 ± 0.10–8.84 ± 0.13 Ma) ice–volcano interactions with both thin-ice and regional ice-sheet conditions at different times. The lava-fed delta sequences suggest palaeo-ice levels up to 300 m higher than today's local ice level. Relatively low-elevation striated glacial unconformities and interbedded tillites record fluctuating ice flow across the growing volcano during this interval.

Interpretations of WAIS palaeo-ice levels from the main shield outcrops at Mount Murphy are complicated by three factors. First, the main shield sequences are located on the west side of Mount Murphy, where ice is currently descending from high upstream levels at c. 800 m asl to low downstream levels of c. 200 m asl. The elevations of these sequences (up to c. 700 m asl) are about the same as the elevation of the regional ice sheet on the upstream side. Therefore, these outcrops may record fluctuations of an ice sheet that was smaller than today's ice sheet, assuming that the much smaller Mount Murphy produced a less significant damming effect. Second, because Mount Murphy is at the coast where glaciers are feeding into the Getz Ice Shelf, the local ice configurations may have been very responsive to changes in sea level. Third, the coastal position of Mount Murphy may have facilitated draining of ice-marginal lakes formed by volcanism, so passage-zone sequences may be significantly lower than the palaeo-ice level. By this scenario, the main shield sequences do provide strong evidence for higher than present local ice levels in the Late Miocene. However, the complex setting of Mount Murphy precludes making interpretations about regional palaeo-ice levels based solely on the main shield sequence. The sequence does indicate the presence of an ice sheet but the ice sheet may have been similar in thickness to today's ice sheet.

Coeval glaciovolcanic sequences at Mount Rees and Mount Steere in the Crary Mountains (9.46 ± 0.24–8.30 ± 0.22 Ma) are interpreted as evidence of slope-ice interactions and imply that abundant local slope ice extended to near or below the level of the modern ice sheet. These outcrops are also on the lee or downstream lower-elevation side of the volcanoes. The Mount Rees and Mount Steere reconstructions are consistent with the interpretations of Mount Murphy that suggest lower or unchanged syneruptive WAIS levels in the Late Miocene.

Monogenetic volcanoes at Bowyer Butte (9.82 ± 0.90 Ma by K–Ar: from LeMasurier and Thomson 1990) and Kennel Peak (8.01 ± 0.87 Ma) provide indications of Late Miocene syneruptive palaeo-ice levels that were 50–200 m higher than today's local ice levels (LeMasurier 1990d; Wilch and McIntosh 2007). Interpretations of regional ice-level changes from these local ice levels is complicated by the fact that both sites are situated on interfluves between glaciers that are descending steeply from inland areas to the Getz Ice Shelf. A further complication is that both volcanoes are situated along prominent linear scarps, almost certainly related to faults of unknown age (LeMasurier and Landis 1996). The pillow lava, hyaloclastite breccia, tillite and striated basement rocks do provide additional evidence for widespread glaciation in West Antarctica between 10 and 8 Ma.

During the 10–8 Ma interval of abundant glaciovolcanism, late-stage scoria cones and subaerial lava were erupted at Mount Flint (9.67 ± 0.20 and 8.66 ± 0.24 Ma), Mount Bursey (8.45 ± 0.046 Ma), Mount Andrus (9.28 ± 0.06 Ma) and Kouperov Peak (9.24 ± 0.06 Ma). These subaerially-erupted deposits provide maximum elevation limits on Late Miocene palaeo-ice levels ranging from <50 to ≤300 m above present ice level (Supplementary material Table S4). Poorly-exposed, subaerially-erupted, edifice-building sequences in the Executive Committee Range Volcanic Field (Mount Cumming (10.0 ± 1.0 Ma), Mount Hartigan (8.50 ± 0.66 Ma)) and the Ames Range Volcanic Field (Mount Kosciusko (9.20 ± 1.3 Ma)) provide local limits on syneruptive palaeo-ice levels at <200–250 m above present local ice level (all ages by K–Ar method from LeMasurier and Thomson 1990). Finally, at Kay Peak on Mount Murphy (8.9–6.9 Ma), late-stage glaciovolcanic sequences record interactions between lavas and valley glaciers above the level of the regional ice sheet.

In summary, the record of 10–8 Ma volcanism in MBL provides substantial evidence for a widespread glaciation of West Antarctica by 9 Ma. Assuming no uplift/subsidence, the palaeo-ice levels appear to have been similar to today's level at several times during this interval. The possibility of much higher or lower ice levels during this interval cannot be verified or contradicted by the volcanic record. The volcanic record of a widespread ice sheet in West Antarctica in the Late Miocene is consistent with global cooling and ice-sheet growth recorded in Late Miocene benthic foraminifera records (e.g. Holbourn et al. 2013). ANDRILL 1B marine-sediment-core records from the Ross Sea show direct evidence of cyclic glaciation during this interval (Wilson et al. 2012).

Polygenetic volcanism, late-stage flank eruptions on older volcanoes and isolated monogenetic volcanism persisted through the latest Miocene, Pliocene and into the early Pleistocene. Monogenetic volcanism at Brandenberger Bluff north of Mount Berlin, at the satellite nunataks and on the flanks of Mount Murphy, at several nunataks of the Hobbs Coast, Hudson Mountains, and Jones Mountains volcanic fields provides abundant evidence for glaciovolcanism between 7.8 and 2.6 Ma. Passage zones in stacked lava-delta sequences at Turtle Peak, Hedin Nunatak, Icefall Nunatak near Mount Murphy and in the emergent Brandenberger Bluff tuya at Mount Berlin record local syneruptive palaeo-ice-level elevations as high as 350 m above today's local ice level. The nunataks at Mount Murphy are situated on interfluves at about the same elevation as the main shield outcrops described in the previous subsection. Extracting accurate WAIS syneruptive palaeo-ice-level elevations from these local ice levels is complicated because small changes in WAIS elevations or local erosion or fault displacement at the interfluve margins could have had dramatic effects on local ice level at the interfluve. In addition, ice flow in these localities may have been strongly affected by sea-level fluctuations.

The Brandenberger Bluff tuya (2.77 Ma) also has a passage zone at 250 m above local ice level. However, it is on the lee side of Mount Berlin, about 150 m lower than modern regional ice levels on the upstream side of Mount Berlin, implying a reduced WAIS at 2.77 Ma. This assumes that the regional ice configuration and elevations prior to growth of Mount Berlin were similar to today's configuration.

In the Hobbs Coast Volcanic Field, emergent hydrovolcanism was common during this time interval, with subaerial eruptions of fall and pyroclastic density current hyalotuffs at Cousins Rock (5.00 ± 0.08 Ma), Shibuya Peak (5.00 ± 0.29 Ma) and Coleman Nunatak (2.60 ± 0.01 Ma). The Shibuya Peak hyalotuff contains abundant moulded, well-rounded basement boulders, possibly resulting from eruption through a glacial debris layer. In the Hudson Mountains Volcanic Field and Jones Mountains Volcanic Field subglacial to subaerial transitions are documented at c. 5 and 7.7 Ma, respectively.

Taken together, the volcanic records at these sites imply that syneruptive palaeo-ice levels were similar to today's WAIS level at several locations during the 8–1 Ma interval. This may indicate that the WAIS configuration through this interval was relatively stable. Alternatively, the volcanic record may only document ice-sheet highstands, and records of lower palaeo-ice levels may be currently concealed beneath the ice.

The record of Quaternary volcanism offers strong evidence for three significant ice-sheet expansion events. A middle Pleistocene (609 ± 27 ka) ice-sheet highstand of +500 m is inferred at Mount Murphy, based on late-stage tuff cone eruptions at Sechrist Peak. The marine oxygen isotope record suggests that significant global-ice build-up occurred during marine isotope stage (MIS) 16 prior to termination VII at 621 ka (Lisiecki and Raymo 2005). The extreme thickening of the WAIS at Mount Murphy may have been driven, at least in part, by sea-level lowering and ice-sheet thickening as the ice sheet expanded onto the exposed continental shelf.

At about the same time (578 ± 9 ka), edifice-building subaerial lava and pyroclastic rocks were erupted from Merrem caldera, on the west flank of Mount Berlin. These rocks imply that syneruptive local ice was at a lower elevation than the lavas. The lowest elevation lava (c. 1850 m asl) is situated about 200 m higher than local ice levels but about 200 lower than the upstream level of the WAIS, in an area where present-day ice is descending from high upstream levels of about 2000 m asl to much lower downstream levels of about 1400 m asl. Although it is difficult to place an absolute elevation on the regional palaeo-ice level, it is probable that syneruptive ice-sheet thickening at below Merrem caldera was much less extreme than at Sechrist Peak. Alternatively, these two eruptions may have occurred in different parts of a glacial–interglacial cycle given the age and age uncertainties of the two events.

The Late Quaternary records from Mount Takahe are even more compelling because they show significant ice-sheet thickening at MIS 4 (71 ka) and MIS 2 (18 ka). The Stauffer Bluff tuya, dated to 65 ± 7 ka, forms a table-shaped edifice that rises 525 m above the ice-sheet level on the NE side of Mount Takahe. The highest-elevation, subglacially-emplaced, pillow hyaloclastites, just below the break in slope and inferred passage zone, are situated almost 400 m above the regional ice sheet. The Stauffer Bluff sequence suggests that the WAIS was at least c. 400 m higher than today at 65 ka, when global temperatures and sea level were lower and the global-ice volume was higher.

Passage-zone sequences on the lower flanks of Mount Takahe at Gill Bluff and Möll Spur record syneruptive ice-sheet highstands coincident with MIS 2. The post-edifice-building passage zones developed as subaerial lavas flowed into ice-marginal lakes forming lava-fed deltas on the lower flanks of Mount Takahe. The upper horizontal passage zone at Gill Bluff indicates a minimum former ice-sheet level at c. 413 m above the present ice-sheet surface at 21.9 ± 2.0 ka. The passage zone at Möll Spur indicates a minimum former ice-sheet level at c. 575 m above the present ice-sheet surface at 18.6 ± 4.4 ka. The lower and older pillow sequences at Möll Spur indicate that ice levels were at least 300 m above present ice level at 36.0 ± 5.0 ka. These post-edifice-building passage-zone sequences at Mount Takahe provide the first glaciovolcanic ice-sheet-level records correlated to the marine isotopic stage timescale.

Limits on the duration of a Last Glacial Maximum ice-sheet highstand at Mount Takahe are provided by lower-elevation (c. 150 m above ice level) subaerial deposits at Oeschger Bluff (7 ± 13 ka) and Steuri Glacier (45 ± 7 ka). Late-stage parasitic scoria cone deposits at English Rock, Mount Frakes (34 ± 14 ka) limit syneruptive ice-sheet expansion to <150 m in the Crary Mountains.

The MBLVG includes two volcanic provinces and 11 geographically distinct volcanic fields, with 19 large central polygenetic alkaline shield-like composite volcanoes and numerous smaller isolated volcanoes that rise above the WAIS. Our understanding of the extent and history of the MBLVG remains fragmentary, largely due to a lack of dissection, extensive local snow and ice cover, and partial or complete burial by the WAIS. We currently have no access the more than 90 hidden sub-ice-sheet volcanoes already documented by geophysical methods (e.g. Behrendt 2013; de Vries et al. 2018). Despite the incomplete record, many details of MBLVG volcanism are known. Key themes have emerged from this research, in some cases reinforcing and in other cases challenging past interpretations.

Documented MBLVG activity dates back to the latest Eocene (c. 37 Ma) and continues to today, with Mount Berlin and Mount Takahe considered active. Early volcanism (37–19 Ma) in the MBLVP was marked by infrequent monogenetic mafic eruptions, preserved in upland areas at Mount Petras in central MBL. The earliest felsic volcanism was the 20 Ma trachyte lava near Mount Flint. Following this, there appears to have been a c. 6 Ma hiatus before the onset of polygenetic volcanism in the middle Miocene (13.4 Ma) at Whitney Peak caldera in the northern ECR. Large-scale polygenetic volcanism has continued since the middle Miocene without major interruptions, producing 19 central volcanoes with 24 preserved large (2–10 km-diameter) summit calderas.

The ubiquity of large summit calderas (24 in the province) is a noteworthy aspect of the MBLVP (LeMasurier 1990h). An anomaly of these calderas is that, in the past, the dominant recognized volcanic outcrops were lavas, with lesser amounts of breccia, including hyaloclastite, and very rare pyroclastic rocks except in late-stage basaltic scoria cones. While lavas dominate, pyroclastic rocks are less rare than previously documented in LeMasurier and Thomson (1990). In this chapter, we presented published (Wilch et al. 1999) and new data documenting explosively-erupted pumiceous welded fall deposits at central volcano caldera-rim outcrops across the MBLVP. These outcrops have been preserved because the pyroclastic deposits are welded. Presumably, the eruptions that generated the caldera-rim welded deposits also generated non-welded tephra that was dispersed across the WAIS and into the Southern Ocean. The two youngest polygenetic central volcanoes, Mount Takahe and Mount Berlin, have Holocene-aged deposits in their caldera walls that have been correlated to tephra in ice cores and englacial tephra layers currently exposed in ablation areas (Dunbar et al. 2021). Although most of the central volcanoes of the MBLVP have shield-like morphologies, they are not simple lava shield volcanoes but have complex explosive and effusive histories.

The extensive application of 40Ar/39Ar geochronology has resulted in significant improvements in resolving the history of the volcanoes, especially where detailed analyses have been done. In many cases, the 40Ar/39Ar ages are similar to the K–Ar ages but in some cases the ages are significantly different (e.g. Wilch and McIntosh 2007). In almost every instance, 40Ar/39Ar ages are at least an order of magnitude more precise than conventional K–Ar ages. The high-precision geochronology can be powerful when tied to detailed fieldwork. For example, Panter et al. (1994) developed a well-dated history of the construction and caldera collapse and breach of Mount Sidley based on detailed field mapping combined with targeted 40Ar/39Ar dating. With detailed sampling and 40Ar/39Ar dating of multiple 200–800 m-thick stratigraphic sequences in the Crary Mountains and at Mount Murphy, Wilch and McIntosh (2002) compared glaciovolcanic interactions at coastal v. inland sites during the same time interval.

The biggest advance associated with the application of 40Ar/39Ar geochronology at MBL volcanoes has been the ability to date very young samples (to Holocene age) (Wilch et al. 1999; Dunbar et al. 2021). This has been critical in reconstructing the history of the youngest volcanoes in the province, especially Mount Berlin and Mount Takahe. Ongoing work at Mount Waesche will complete the first level of detailed work on the Late Pleistocene volcanoes of MBL. 40Ar/39Ar dating of young outcrops combined with tephrochronlogy analyses in blue-ice areas and ice cores has led to fingerprinting and correlating MBL tephra to climate proxy records. Lastly, the high resolution of young 40Ar/39Ar ages in glaciovolcanic sequences at Mount Takahe has permitted correlation of interpreted ice levels with Milankovitch-driven marine isotopic cycles.

Detailed lithofacies reconstructions have also advanced our understanding of palaeoenvironments and the eruptive histories of MBL volcanoes. A detailed field and geochronology study of Mount Petras led to major revisions of the earliest history of the province, challenging early interpretations of the glacial and the erosional history of Marie Byrd Land (Wilch and McIntosh 2000). A detailed lithofacies analysis of Icefall Nunatak near Mount Murphy provided an incremental but significant advance in understanding glaciovolcanic processes and how volcanic records can be used to understand local syneruptive ice conditions (Smellie 2001).

Finally, the palaeoenvironmental reconstructions and 40Ar/39Ar geochronology of the MBLVG provide snapshot records of the WAIS since the late Eocene, with the following major conclusions:

  • The earliest terrestrial indications for glacial ice in West Antarctica are middle Oligocene (29–27 Ma) tuff cone deposits at Mount Petras that suggest the presence of a thin local ice cap or ice and snow on bedrock slopes. Throughout the remainder of MBLVG volcanic record, interactions between volcanism and thin local ice continued to occur, suggesting that significant cover by local snow and ice was common in MBL since the Oligocene.

  • The first volcanic evidence for a widespread WAIS is Late Miocene c. 9 Ma glaciovolcanic sequences from across MBL. Two patterns emerge from the record of syneruptive palaeo-ice levels: former WAIS thickening was more extensive at coastal sites (such as Mount Murphy) than at inland sites (such as Mount Rees and Mount Steere); and, currently, the WAIS is in a near-maximum configuration that has existed at many times since 10 Ma but was rarely exceeded.

  • A middle Pleistocene ice-sheet highstand of +500 m is inferred at the coastal volcano, Mount Murphy, based on late-stage 609 ± 27 ka tuff cone deposits at Sechrist Peak. This ice-sheet highstand may correspond to global ice expansion (and eustatic lowering) at MIS 16 in the marine record.

  • The Stauffer Bluff tuya at Mount Takahe records minimum syneruptive ice levels at 400 m above today's ice at 64.7 ± 13.3 ka, coincident with a eustatic sea-level lowstand at MIS 4.

  • The Gill Bluff and Möll Spur and passage-zone sequences at Mount Takahe indicate minimum syneruptive ice levels at 413 m and c. 575 m above today's ice at 21.2 ± 4.1 and 17.5 ± 4.8 ka, respectively, which is consistent with a eustatic sea-level lowstand at MIS 2.

We are especially grateful to Phil Kyle, Ian Skilling and John Smellie for detailed and constructive reviews, which improved this chapter significantly. We thank US Navy VXE-6 squadron, Antarctic Support Associates, and Ken Borek Air Ltd for logistical support. We are grateful for field collaborators Nelia Dunbar, John Smellie and John Gamble; and field assistants Paul Rose, Chris Griffiths and Tony Teeling. Most of the 40Ar/39Ar geochronology analyses were completed at the New Mexico Geochronological Research Laboratory at New Mexico Tech and we wish to thank the NMGRL staff, particularly Lisa Peters and Matt Heizler, for their assistance. We are grateful to Sean McCuddy and Mike Wagg for additional assistance with lab work. Christine Merritt, Emily Ebaugh and Miriam Wilch provided invaluable technical assistance in preparation of maps and figures for this chapter. This work was supported by National Science Foundation grants DPP-8816342 (WAVE), DPP-9198806 (WAVE II), OPP-9419686, OPP-9725910, and OPP-9814782.

TIW: conceptualization (lead), data curation (equal), formal analysis (lead), funding acquisition (equal), investigation (equal), methodology (equal), project administration (lead), writing – original draft (lead), writing – review & editing (equal); WCM: conceptualization (supporting), data curation (equal), formal analysis (equal), investigation (equal), methodology (equal), visualization (supporting), writing – original draft (supporting), writing – review & editing (equal); KSP: conceptualization (supporting), data curation (equal), formal analysis (equal), investigation (equal), methodology (supporting), project administration (supporting), writing – original draft (supporting), writing – review & editing (equal).

Fieldwork was supported by the National Science Foundation with grants OPP-9725910 and DPP-918806 awarded to W.C. Mcintosh, and grant OPP-9814782 awarded to T.I. Wilch.

The datasets generated during the current study are available in Supplementary material Tables S1–S4.

1.
Anderson
J.B.
1999
.
Antarctic Marine Geology
 .
Cambridge University Press
,
Cambridge, UK
.
2.
Andrews
J.T.
and
LeMasurier
W.E.
1973
.
Rates of Quaternary glacial erosion and corrie formation, Marie Byrd Land, Antarctica
.
Geology
 ,
1
,
75
80
,
3.
Behrendt
J.C.
2013
.
The aeromagnetic method as a tool to identify Cenozoic magmatism in the West Antarctic Rift System beneath the West Antarctic Ice Sheet – A review; Thiel subglacial volcano as possible source of the ash layer in the WAISCORE
.
Tectonophysics
 ,
585
,
124
136
,
4.
Behrendt
J.C.
,
LeMasurier
W.E.
,
Cooper
A.K.
,
Tessensohn
F.
,
Trehu
A.
and
Damaske
D.
1991
.
Geophysical studies of the West Antarctic rift system
.
Tectonics
 ,
10
,
1257
1273
,
5.
Behrendt
J.C.
,
Saltus
R.
,
Damaske
D.
,
McCafferty
A.
,
Finn
C.A.
,
Blankenship
D.D.
and
Bell
R.E.
1996
.
Patterns of late Cenozoic volcanic and tectonic activity in the West Antarctic rift system revealed by aeromagnetic data
.
Tectonics
 ,
15
,
660
676
,
6.
Bentley
C.R.
and
Clough
J.W.
1972
. Antarctic subglacial structure from seismic refraction measurements. In:
Adie
R.J.
(ed.)
Antarctic Geology and Geophysics
 .
Universitetsforlaget
,
Oslo
,
683
691
.
7.
Blankenship
D.D.
,
Bell
R.E.
,
Hodge
S.M.
,
Brozena
J.M.
,
Behrendt
J.C.
and
Finn
C.A.
1993
.
Active volcanism beneath the West Antarctic ice sheet and implications for ice-sheet stability
.
Nature
 ,
361
,
526
529
,
8.
Cande
S.C.
and
Stock
J.M.
2006
. Constraints on the timing of extension in the Northern Basin, Ross Sea. In:
Futterer
D.K.
,
Damaske
D.
,
Kleinschmidt
G.
,
Miller
H.
and
Tessensohn
F.
(eds)
Antarctica: Contributions to Global Earth Sciences
 .
Springer
,
Berlin
,
319
326
.
9.
Cande
S.C.
,
Stock
J.M.
,
Müller
R.D.
and
Ishihara
T.
2000
.
Cenozoic motion between east and west Antarctica
.
Nature
 ,
404
,
145
150
,
10.
Chakraborty
S.
2007
.
The Geochemical Evolution of Alkaline magMas from the Crary Mountains, Marie Byrd Land, Antarctica
 . MS thesis,
Bowling Green State University
,
Bowling Green, Ohio, USA
.
11.
Chaput
J.
,
Aakaster
R.C.
et al
2014
.
The crustal thickness of West Antarctica
.
Journal of Geophysical Research: Solid Earth
 ,
119
,
378
395
,
12.
Chatzaras
V.
,
Kruckenberg
S.C.
,
Cohen
S.M.
,
Medaris
L.G.
Jr
,
Withers
A.C.
and
Bagley
B.
2016
.
Axial-type olivine crystallographic preferred orientations: the effect of strain geometry on mantle texture
.
Journal of Geophysical Research: Solid Earth
 ,
121
,
4895
4922
,
13.
Chough
S.K.
and
Sohn
Y.K.
1990
.
Depositional mechanics and sequences of base surges, Songaksan tuff ring, Cheju Island, Korea
.
Sedimentology
 ,
37
,
1115
1135
,
14.
Cooper
A.K.
and
Davey
F.J.
1985
.
Episodic rifting of phanerozoic rocks in the Victoria Land Basin, Western Ross Sea, Antarctica
.
Science
 ,
229
,
1085
1087
,
15.
Cooper
A.K.
,
Davey
F.J.
and
Hinz
K.
1991
. Crustal extension and origin of sedimentary basins beneath the Ross Sea and Ross Ice Shelf, Antarctica. In:
Crame
J.A.
,
Thomson
J.W.
and
Thomson
M.R.A.
(eds)
Geological Evolution of Antarctica
 .
Cambridge University Press
,
Cambridge, UK
,
299
304
.
16.
Courtillot
V.
,
Davaille
A.
,
Besse
J.
and
Stock
J.
2003
.
Three distinct types of hotspots in the Earth's mantle
.
Earth and Planetary Science Letters
 ,
205
,
295
308
,
17.
Craddock
C.
,
Bastien
T.W.
and
Rutford
R.H.
1964
. Geology of the Jones Mountains area. In:
Adie
R.J.
(ed.)
Antarctic Geology
 .
North-Holland
,
Amsterdam
,
171
187
.
18.
Cuffey
K.
and
Paterson
W.S.B.
2010
.
The Physics of Glaciers
 . 4th edn.
Academic Press
,
Amsterdam
.
19.
Dalrymple
G.B.
1979
.
Critical tables for conversion of K–Ar ages from old to new decay constants
.
Geology
 ,
7
,
558
560
,
20.
Dalziel
I.W.D.
and
Elliot
D.H.
1982
.
West Antarctica – Problem child of Gondwanaland
.
Tectonics
 ,
1
,
3
19
,
21.
Dellino
P.
and
La Volpe
L.
2000
.
Structures and grain size distribution in surge deposits as a tool for modelling the dynamics of dilute pyroclastic density currents at La Fossa di Vulcano (Aeolian Islands, Italy)
.
Journal of Volcanology and Geothermal Research
 ,
96
,
57
78
,
22.
de Vries
,
M.V.
,
Bingham
R.G.
and
Hein
A.S.
2018
.
A new volcanic province: an inventory of subglacial volcanoes in West Antarctica
.
Geological Society, London, Special Publications
 ,
461
,
231
248
,
23.
Douillet
G.A.
,
Taisne
B.
,
Tsang-Hin-Sun
E.
,
Muller
S.K.
,
Kueppers
U.
and
Dingwell
D.B.
2015
.
Syn-eruptive, soft-sediment deformation of deposits from dilute pyroclastic density current: triggers from granular shear, dynamic pore pressure, ballistic impacts and shock waves
.
Solid Earth
 ,
6
,
553
572
,
24.
Dunbar
N.W.
,
McIntosh
W.C.
and
Esser
R.P.
2008
.
Physical setting and tephrochronology of the summit caldera ice record at Mt. Moulton, West Antarctica
.
Geological Society of America Bulletin
 ,
120
,
796
812
,
25.
Dunbar
N.W.
,
Iverson
N.A.
,
Smellie
J.L.
,
McIntosh
W.C.
,
Zimmerer
M.J.
and
Kyle
P.R.
2021
. Active volcanoes in Marie Byrd Land.
Geological Society, London, Memoirs
 ,
55
,
26.
Emry
E.L.
,
Nyblade
A.A.
et al
2015
.
The mantle transition zone beneath West Antarctica: Seismic evidence for hydration and thermal upwellings
.
Geochemistry, Geophysics, Geosystems
 ,
16
,
40
58
,
27.
Fenner
C.N.
1938
.
Olivine fourchites from Raymond Fosdick Mountains, Antarctica
.
Bulletin of the Geological Society of America
 ,
49
,
367
400
,
28.
Fisher
R.V.
and
Schmincke
H.
1984
.
Pyroclastic Rocks
 .
Springer
,
Berlin
.
29.
Fretwell
P.
,
Pritchard
H.D.
et al
2013
.
Bedmap2: improved ice bed, surface and thickness datasets for Antarctica
.
Cryosphere
 ,
7
,
375
393
,
30.
Gaffney
A.M.
and
Siddoway
C.S.
2007
. Heterogeneous sources for Pleistocene lavas of Marie Byrd Land, Antarctica: new data from the SW Pacific diffuse alkaline magmatic province.
United States Geological Survey Open-File Report
,
2007-1047
,
31.
Global Volcanism Program
.
1988
. Report on Siple (Antarctica), Marie Byrd Land, Antarctica.
Scientific Event Alert Network Bulletin
 ,
13
:
12
.
Smithsonian Institution
,
32.
González-Ferrán
O.
and
González-Bonorino
F.
1972
. The volcanic ranges of Marie Byrd Land between long. 100° and 140°W. In:
Adie
R.J.
(ed.)
Antarctic Geology and Geophysics
 .
Universitetsforlaget
,
Oslo
,
261
276
.
33.
Granot
R.
and
Dyment
J.
2018
.
Late Cenozoic unification of East and West Antarctica
.
Nature Communications
 ,
9
,
3189
,
34.
Grosse
P.
and
Kervyn
M.
2018
.
Morphometry of terrestrial shield volcanoes
.
Geomorphology
 ,
304
,
1
14
,
35.
Grosse
P.
,
Euillades
P.A.
,
Euillades
L.D.
and
de Vries
B.V.
2014
.
A global database of composite volcano morphometry
.
Bulletin of Volcanology
 ,
76
,
16
,
36.
Hansen
S.E.
,
Graw
J.H.
et al
2014
.
Imaging the Antarctic mantle using adaptively parameterized P-wave tomography: evidence for heterogeneous structure beneath West Antarctica
.
Earth and Planetary Science Letters
 ,
408
,
66
78
,
37.
Heeszel
D.S.
,
Wiens
D.A.
et al
2016
.
Upper mantle structure of central and West Antarctica from array analysis of Rayleigh wave phase velocities
.
Journal of Geophysical Research: Solid Earth
 ,
121
,
1758
1775
,
38.
Holbourn
A.
,
Kuhnt
W.
,
Clemens
S.
,
Prell
W.
and
Andersen
N.
2013
.
Middle to late Miocene stepwise climate cooling: evidence from a high-resolution deep water isotope curve spanning 8 million years
.
Palaeoceanography
 ,
28
,
688
699
,
39.
Hole
M.J.
,
Storey
B.C.
and
LeMasurier
W.E.
1994
.
Tectonic setting and geochemistry of Mocene alkaline basalts from the Jones Mountains, West Antarctica
.
Antarctic Science
 ,
6
,
85
92
,
40.
Honnorez
J.
and
Kirst
P.
1975
.
Submarine basaltic volcanism: morphometric parameters for discriminating hyaloclastites from hyalotuffs
.
Bulletin of Volcanology
 ,
39
,
1
25
,
41.
Jones
J.G.
1969
.
Intraglacial volcanoes of the Laugarvatn region, south-west Iceland – I
.
Quarterly Journal of the Geological Society, London
 ,
124
,
197
211
,
42.
Jones
J.G.
1970
.
Intraglacial volcanoes of the Laugarvatn region, south-west Iceland – I
.
Journal of Geology
 ,
78
,
127
140
,
43.
Jordan
T.A.
,
Riley
T.R.
and
Siddoway
C.S.
2020
.
The geological history and evolution of West Antarctica
.
Nature Reviews
 ,
1
,
117
133
,
44.
Kuiper
K.F.
,
Deino
A.
,
Hilgen
F.J.
,
Krijgsman
W.
,
Renne
P.R.
and
Wijbrans
J.R.
2008
.
Synchronizing rock clocks of Earth history
.
Science
 ,
320
,
500
504
,
45.
LeMasurier
W.E.
1972
. Volcanic record of Cenozoic glacial history in Marie Byrd Land. In:
Adie
R.J.
(eds)
Antarctic Geology and Geophysics
 .
Universitetsforlaget
,
Oslo
,
251
260
.
46.
LeMasurier
W.E.
1990a
. Ames Range.
American Geophysical Union Antarctic Research Series
 ,
48
,
216
220
.
47.
LeMasurier
W.E.
1990b
. Miocene and older centers, Walgreen Coast.
American Geophysical Union Antarctic Research Series
 ,
48
,
203
207
.
48.
LeMasurier
W.E.
1990c
. Miocene–Oligocene centers, Mount Petras and USAS Escarpment.
American Geophysical Union Antarctic Research Series
 ,
48
,
239
243
.
49.
LeMasurier
W.E.
1990d
. Miocene–Pliocene centers, Hobbs Coast.
American Geophysical Union Antarctic Research Series
 ,
48
,
244
247
.
50.
LeMasurier
W.E.
1990e
. Mount Bursey.
American Geophysical Union Antarctic Research Series
 ,
48
,
221
224
.
51.
LeMasurier
W.E.
1990f
. Mount Hartigan.
American Geophysical Union Antarctic Research Series
 ,
48
,
199
202
.
52.
LeMasurier
W.E.
1990g
. Pliocene–Pleistocene centers, Hobbs Coast.
American Geophysical Union Antarctic Research Series
 ,
48
,
248
250
.
53.
LeMasurier
W.E.
1990h
. Summary.
American Geophysical Union Antarctic Research Series
 ,
48
,
146
163
.
54.
LeMasurier
W.E.
2002
. Architecture and evolution of hydrovolcanic deltas in Marie Byrd Land, Antarctica.
Geological Society, London, Special Publications
 ,
202
,
115
148
,
55.
LeMasurier
W.E.
2006
. What supports the Marie Byrd Land dome? An evaluation of potential uplift mechanisms in a continental rift system. In:
Fütterer
D.K.
,
Damaske
D.
,
Kleinschmidt
G.
,
Miller
H.
and
Tessensohn
F.
(eds)
Antarctica
 .
Springer
,
Berlin
,
299
302
.
56.
LeMasurier
W.E.
2008
.
Neogene extension and basin deepening in the West Antarctic rift inferred from comparisons with the East African rift and other analogs
.
Geology
 ,
36
:
247
250
,
57.
LeMasurier
W.E.
2013
.
Shield volcanoes of Marie Byrd Land, West Antarctic rift: oceanic island similarities, continental signature, and tectonic controls
.
Bulletin of Volcanology
 ,
75
,
58.
LeMasurier
W.E.
and
Kawachi
Y.
1990a
. Mount Berlin.
American Geophysical Union Antarctic Research Series
 ,
48
,
229
233
.
59.
LeMasurier
W.E.
and
Kawachi
Y.
1990b
. Mount Cumming.
American Geophysical Union Antarctic Research Series
 ,
48
,
195
198
.
60.
LeMasurier
W.E.
and
Kawachi
Y.
1990c
. Mount Hampton.
American Geophysical Union Antarctic Research Series
 ,
48
,
189
194
.
61.
LeMasurier
W.E.
and
Kawachi
Y.
1990d
. Mount Moulton.
American Geophysical Union Antarctic Research Series
 ,
48
,
225
228
.
62.
LeMasurier
W.E.
and
Landis
C.A.
1996
.
Mantle-plume activity recorded by low-relief erosion surfaces in West Antarctica and New Zealand
.
Geological Society of America Bulletin
 ,
108
,
1450
1466
,
63.
LeMasurier
W.E.
and
Rex
D.C.
1982
. Volcanic record of Cenozoic glacial history in Marie Byrd Land and western Ellsworth Land: Revised chronology and evalution of tectonic factors. In:
Craddock
C.
(ed.)
Antarctic Geoscience
 .
University of Wisconsin Press
,
Madison, WI
,
725
734
.
64.
LeMasurier
W.E.
and
Rex
D.C.
1983
. Rates of uplift and the scale of ice level instabilities recorded by volcanic rocks in Marie Byrd Land, West Antarctica. In:
Oliver
R.L.
,
James
P.R.
and
Jago
J.B.
(eds)
Antarctic Earth Sciences
 .
Australian Academy of Science
,
Canberra
,
660
673
.
65.
LeMasurier
W.E.
and
Rex
D.C.
1989
.
Evolution of linear volcanic ranges in Marie Byrd Land, West Antarctica
.
Journal of Geophysical Research
 ,
94
,
7223
7236
,
66.
LeMasurier
W.E.
and
Rex
D.C.
1990a
. Mount Siple.
American Geophysical Union Antarctic Research Series
 ,
48
,
185
188
.
67.
LeMasurier
W.E.
and
Rex
D.C.
1990b
. Mount Takahe.
American Geophysical Union Antarctic Research Series
 ,
48
,
169
174
.
68.
LeMasurier
W.E.
and
Rex
D.R.
1991
. The Marie Byrd Land Volcanic Province and its relation to the Cainozoic West Antarctic rift system. In:
Tingey
R.J.
(ed.)
The Geology of Antarctica
 .
Clarendon Press
,
Oxford
,
249
284
.
69.
LeMasurier
W.E.
and
Rocchi
S.
2005
.
Terrestrial record of post-Eocene climate history in Marie Byrd Land, West Antarctica
.
Geografiska Annaler Series A – Physical Geography
 ,
87A
,
51
66
,
70.
LeMasurier
W.E.
and
Thomson
J.E.
(eds).
1990
.
Volcanoes of the Antarctic Plate and Southern Oceans
 .
American Geophysical Union Antarctic Research Series
 ,
48
,
146
163
.
71.
LeMasurier
W.E.
and
Wade
F.A.
1968
.
Fumarolic activity in Marie Byrd Land Antarctica
.
Science
 ,
162
,
352
,
72.
LeMasurier
W.E.
and
Wade
F.A.
1976
. Volcanic history in Marie Byrd Land: implications with regard to southern hemisphere tectonic reconstructions. In:
González-Ferrán
O.
(ed.)
Proceedings of the International Symposium on Andean and Antarctic Volcanology Problems
 .
IAVCEI
,
Rome
,
398
424
.
73.
LeMasurier
W.E.
and
Wade
F.A.
1990
. Fosdick Mountains.
American Geophysical Union Antarctic Research Series
 ,
48
,
251
253
.
74.
LeMasurier
W.E.
,
McIntosh
W.C.
and
Rex
D.C.
1981
.
Mid-Tertiary glacial history recorded at Mt. Petras, Marie Byrd Land
.
Antarctic Journal of the United States
 ,
16
,
19
21
.
75.
LeMasurier
W.E.
,
Kawachi
Y.
and
Rex
D.C.
1990a
. Crary Mountains.
American Geophysical Union Antarctic Research Series
 ,
48
,
180
184
.
76.
LeMasurier
W.E.
,
Kawachi
Y.
and
Rex
D.C.
1990b
. Mount Flint–Reynolds Ridge.
American Geophysical Union Antarctic Research Series
 ,
48
,
212
215
.
77.
LeMasurier
W.E.
,
Kawachi
Y.
and
Rex
D.C.
1990c
. Mount Murphy.
American Geophysical Union Antarctic Research Series
 ,
48
,
164
168
.
78.
LeMasurier
W.E.
,
Kawachi
Y.
and
Rex
D.C.
1990d
. Toney Mountain.
American Geophysical Union Antarctic Research Series
 ,
48
,
175
179
.
79.
LeMasurier
W.E.
,
Harwood
D.M.
and
Rex
D.C.
1994
.
Geology of Mount Murphy Volcano: an 8-m.y. history of interaction between a rift volcano and the West Antarctic ice sheet
.
Geological Society of America Bulletin
 ,
106
,
265
280
,
80.
Lisiecki
L.E.
and
Raymo
M.E.
2005
.
A Pliocene–Pleistocene stack of 57 globally distributed benthic δ18O records
.
Palaeoceanography
 ,
20
,
17
,
81.
Lloyd
A.J.
,
Wiens
D.A.
et al
2015
.
A seismic transect across West Antarctica: evidence for mantle thermal anomalies beneath the Bentley Subglacial Trench and the Marie Byrd Land Dome
.
Journal of Geophysical Research: Solid Earth
 ,
120
,
8439
8460
,
82.
Loose
B.
,
Garabato
A.C.N.
,
Schlosser
P.
,
Jenkins
W.J.
,
Vaughan
D.
and
Heywood
K.J.
2018
.
Evidence of an active volcanic heat source beneath the Pine Island Glacier
.
Nature Communications
 ,
9
,
9
,
83.
Lopatin
B.G.
and
Polyakov
M.M.
1974
.
Geology of the volcanic Hudson Mountains, Walgreen Coast, West Antarctica
.
Antarktika
 ,
13
,
36
51
.
84.
Lough
A.C.
,
Wiens
D.A.
et al
2013
.
Seismic detection of an active subglacial magmatic complex in Marie Byrd Land, Antarctica
.
Nature Geoscience
 ,
6
,
1031
1035
,
85.
Martin
A.P.
,
Cooper
A.F.
,
Price
R.C.
,
Kyle
P.R.
and
Gamble
J.A.
2021
. Erebus Volcanic Province: petrology.
Geological Society, London, Memoirs
 ,
55
,
86.
McConnell
J.R.
,
Burke
A.
et al
2017
.
Synchronous volcanic eruptions and abrupt climate change similar to 17.7 ka plausibly linked by stratospheric ozone depletion
.
Proceedings of the National Academy of Sciences of the United States of America
 ,
114
,
10 035
10 040
,
87.
McIntosh
W.C.
,
LeMasurier
W.E.
,
Ellerman
P.J.
and
Dunbar
N.W.
1985
.
A reinterpretation of glaciovolcanic interaction at Mount Takake and Mount Murphy, Marie Byrd Land, Antarctica
.
Antarctic Journal of the United States
 ,
19
,
57
59
.
88.
McIntosh
W.C.
,
Smellie
J.L.
and
Panter
K.S.
1991
.
Glaciovolcanic interaction in eastern Marie Byrd Land, Antarctica
. In:
Proceedings of the Sixth International Symposium on Antarctic Earth Sciences, Saitama, Japan, 1991
. National Institute of Polar Research,
Tokyo
,
402
.
89.
McPhie
J.
,
Doyle
M.
and
Allen
R.
1993
.
Volcanic Textures: A Guide to the Interpretation of Textures in Volcanic Rocks
 .
Centre for Ore Deposit and Exploration Studies, University of Tasmania
,
Hobart, Australia
.
90.
Morgan
W.J.
1972
.
Deep mantle convection plumes and plate motions
.
AAPG Bulletin
 ,
56
,
203
213
.
91.
Olivetti
V.
,
Balestrieri
M.L.
,
Rossetti
F.
,
Thomson
S.N.
,
Talarico
F.M.
and
Zattin
M.
2015
.
Evidence of a full West Antarctic Ice Sheet back to the early Oligocene: insight from double dating of detrital apatites in Ross Sea sediments
.
Terra Nova
 ,
27
,
238
246
,
92.
Palais
J.M.
,
Kyle
P.R.
,
McIntosh
W.C.
and
Seward
D.
1988
.
Magmatic and phreatomagmatic volcanic activity at Mt. Takahe, West Antarctica, based on tephra layers in the Byrd ice core and field observations at Mt. Takahe
.
Journal of Volcanology and Geothermal Research
 ,
35
,
295
317
,
93.
Panter
K.S.
1995
.
Geology, Geochemistry and Petrogenesis of the Mt. Sidley Volcano, Marie Byrd Land, Antarctica
 . PhD thesis,
New Mexico Institute of Mining and Technology
,
Socorro, New Mexico, USA
.
94.
Panter
K.S.
,
McIntosh
W.C.
and
Smellie
J.L.
1994
.
Volcanic history of Mt. Sidley, a major alkaline volcano in Marie Byrd Land, Antarctica
.
Bulletin of Volcanology
 ,
56
,
361
376
,
95.
Panter
K.S.
,
Kyle
P.R.
and
Smellie
J.L.
1997
.
Petrogenesis of a phonolite-trachyte succession at Mt. Sidley, Marie Byrd Land, Antarctica
.
Journal of Petrology
 ,
38
,
1225
1253
,
96.
Panter
K.S.
,
Hart
S.R.
,
Kyle
P.
,
Blusztanjn
J.
and
Wilch
T.
2000
.
Geochemistry of Late Cenozoic basalts from the Crary Mountains: characterization of mantle sources in Marie Byrd Land, Antarctica
.
Chemical Geology
 ,
165
,
215
241
,
97.
Panter
K.S.
,
Wilch
T.I.
,
Smellie
J.L.
,
Kyle
P.R.
and
McIntosh
W.C.
2021
. Marie Byrd Land and Ellsworth Land: petrology.
Geological Society, London, Memoirs
 ,
55
,
98.
Paulsen
T.S.
and
Wilson
T.J.
2010
.
Evolution of Neogene volcanism and stress patterns in the glaciated West Antarctic Rift, Marie Byrd Land, Antarctica
.
Journal of the Geological Society, London
 ,
167
,
401
416
,
99.
Quartini
E.
,
Blankenship
D.D.
and
Young
D.A.
2021
. Active subglacial volcanism in West Antarctica.
Geological Society, London, Memoirs
 ,
55
,
100.
Ramirez
C.
,
Nyblade
A.
et al
2017
.
Crustal structure of the Transantarctic Mountains, Ellsworth Mountains and Marie Byrd Land, Antarctica: constraints on shear wave velocities, Poisson's ratios and Moho depths
.
Geophysical Journal International
 ,
211
,
1328
1340
,
101.
Reusch
A.M.
,
Nyblade
A.A.
,
Benoit
M.H.
,
Wiens
D.A.
,
Anandakrishan
S.
,
Voigt
D.
and
Shore
P.J.
2008
.
Mantle transition zone thickness beneath Ross Island, the Transantarctic Mountains, and Eas