Six4 and Six5 showed the same ventral body wall closure defects as those seen in human omphalocele. SIX4 and SIX5 were localized in surface ectodermal cells and somatic mesoderm-derived mesenchymal and coelomic epithelial cells (CECs) in the PAW. −/−;Six5−/−Six4 fetuses exhibited a large omphalocele with protrusion of both the liver and intestine, or a small omphalocele with protrusion of the intestine, with complete penetrance. The umbilical ring of −/−;Six5−/−Six4 embryos was shifted anteriorly and its lateral size was larger than that of normal embryos at the E11.5 stage, before the onset of myoblast migration into the PAW. The proliferation rates of surface ectodermal cells in the left and right PAW and somatic mesoderm-derived cells in the right PAW were lower in −/−;Six5−/−Six4 embryos than those of wild-type embryos at E10.5. The transition from CECs of the PAW to rounded mesothelial progenitor cells was impaired and the inner coelomic surface of the PAW was relatively smooth in −/−;Six5−/−Six4 embryos at E11.25. Furthermore, Six4 overexpression in CECs of the PAW promoted ingression of CECs. Taken together, our results suggest that Six4 and Six5 are required for growth and morphological change of the PAW, and the impairment of these processes is linked to the abnormal positioning and expansion of the umbilical ring, which results in omphalocele.Omphalocele is a human congenital anomaly in ventral body wall closure and may be caused by impaired formation of the primary abdominal wall (PAW) and/or defects in abdominal muscle development. Here, we report that mice doubly deficient in homeobox genes  Summary: The homeobox genes Six4 and Six5 are involved in the regulation of cell proliferation and mesothelium formation in the primary body wall, and −/−;Six5−/−Six4 mice are a suitable animal model for human middle-type omphalocele. IGF2 by abnormal imprinting causes Beckwith-Wiedemann syndrome, which is characterized by a large body, large internal organs and severe omphalocele  defect theory and the secondary abdominal wall defect theory (see reviews by Tfap2α (previously known as AP-2α) and Pitx2 mutants show severe ventral body wall closure defects such as thoracoabdominoschisis and omphalocele  and cell repulsion . HoweverSix4/AREC3 and Six5, which are categorized as Six4/Six5 subfamily members  mice exhibited phenotypes that resembled human middle-type omphalocele, including both small and large omphaloceles. We analyzed the morphological defects, cell proliferation and apoptosis in the PAW, and properties of coelomic epithelial cells (CECs) in −/−;Six5−/−Six4 embryos. From the results, we propose that the regulation of cell proliferation and morphological change in the PAW at an early stage is a basis for omphalocele phenotype, and that −/−;Six5−/−Six4 mice are a suitable animal model for reproducing human middle-type omphalocele.In this study, we focus on  members ,b of Six members . We foun−/−Six4, −/−Six5 and +/−;Six5−/−Six4 mice are viable, and developmental abnormalities in these mice have not been reported 18.5 in the expected Mendelian ratio (Table S1). As we found various types of ventral body wall closure defects in fetuses analyzed at E18.5, we classified them into three types: large omphalocele with the protrusion of the liver and intestine, small omphalocele with the protrusion of the intestine, and enlarged umbilical ring with anteriorly elongated shape of the umbilical ring compared with normal fetuses (Table S2). A small number of −/−Six4 fetuses exhibited an enlarged umbilical ring at E18.5 (Table S2), but others showed no abnormalities in the ventral body wall . All of the −/−;Six5−/−Six4 fetuses exhibited small or large omphaloceles at E18.5  with complete penetrance. These results indicate that Six4 and Six5 together are indispensable for ventral body wall closure in mice and that Six4 is the primary player between the two.reported . SIX4 hareported ,b, suggeand Six5 . Howeverody wall E,E′. Altral body I,I′, a shalocele J,J′. Remat E18.5 F-H′ . −/−Six5 and +/−;Six5−/−Six4 fetuses had normal skeletal structure . In −/−;Six5−/−Six4 fetuses, the numbers of ribs and sternums were normal, but formation of the ensiform cartilage was impaired . These results indicate that the defects in −/−;Six5−/−Six4 fetuses are mostly confined to the abdominal body wall.Patients with omphalocele with severe protrusion of the liver and intestine often show other malformations at the thoracic and lumbar levels . Therefo−/−;Six5−/−Six4 mice during abdominal wall formation.Previous studies have shown that abnormal formation of the secondary abdominal wall, including abdominal muscle differentiation, may be the cause of the large omphalocele phenotype in humans  and in m−/−;Six5−/−Six4 fetuses with large omphalocele and −/−Six5 or +/−;Six5+/−Six4 fetuses, which have no apparent defects in ventral body wall closure, at E15.5 and E18.5 . The abdominal wall was closed at the umbilical cord level and it completely covered internal organs at the level anterior to the umbilical cord in −/−Six5 fetuses . The herniated livers and intestines in −/−;Six5−/−Six4 fetuses were covered with the peritoneum and amnion at the umbilical cord level . The anterior abdominal wall was completely lost in −/−;Six5−/−Six4 fetuses .We compared ventral body wall formation in −/−;Six5−/−Six4 fetuses in a pattern similar to that observed in −/−Six5 fetuses at E18.5 . The epaxial extensor muscle  and panniculus carnosus muscle (PC) were normally positioned in −/−;Six5−/−Six4 fetuses and −/−Six5 fetuses . All abdominal muscles were separated and positioned in the ventral wall of −/−;Six5−/−Six4 embryos in a pattern similar to that in control +/−;Six5+/−Six4 embryos at E15.5 . Previous studies have shown that SIX family proteins are involved in skeletal muscle differentiation by activating the expression of muscle regulatory factors MYOG and MYOD1 . In addition, MYOG and MYOD1 were expressed in abdominal muscle cells in +/−;Six5+/−Six4 and −/−;Six5−/−Six4 embryos with similar patterns, except the downregulation in the PC at E13.5 . These results suggest that there are no severe defects in abdominal muscle differentiation in −/−;Six5−/−Six4 embryos. Our results indicate that omphalocele formation in −/−;Six5−/−Six4 mice is not attributable to impaired abdominal muscle differentiation and clearly suggest the other causes of large omphalocele.Myoblasts start to migrate into the PAW at E10.5 and differentiate into four types of hypaxial muscles, specifically the transversus abdominis muscle (TA), internal oblique muscle (IO), external oblique muscle (EO) and rectus abdominis muscle (RA), by E13.5 . These mnd MYOD1 . MYOG wa−/−;Six5−/−Six4 embryos, defects in PAW formation have been considered as the cause of omphalocele. Therefore, we next examined the effects of Six4 deficiency in a Six5-null background on PAW formation along the anterior-posterior axis in comparison with the phenotypes of control (−/−Six5) and −/−;Six5−/−Six4 embryos at E13.0  was formed in the region posterior to the thoracic diaphragm    but thicker than that of −/−Six5 embryos at the 50% position of the AnPAW   but thicker than that of −/−Six5 embryos at the 50% position of the LaPAW    but similar at the 50% position of the PoPAW  to that in −/−Six5 embryos  B-D in cost, n=3) A,C,D. Thst, n=3) E-F″,K,L.st, n=3) G-H″,M,N.s n=3/3, G,H. The st, n=3) G,H,M,O.  embryos I-J″,P,Q.−/−;Six5−/−Six4 embryos (n=3) with patterns similar to wild-type and −/−Six5 embryos at E12.5 . These results suggest that umbilical ring formation is impaired in −/−;Six5−/−Six4 embryos before the protrusion of the liver.To analyze the detailed three-dimensional structure of the umbilical region, we performed micro-CT scanning analysis, which is a powerful method to identify structural phenotypes in mutant mice even at the early stages  Fig. 3)Fig. 3). −/−;Six5−/−Six4 embryos at the stage before E12.5. Physiological herniation and looping of the midgut were observed in wild-type and −/−Six5 embryos , and −/−;Six5−/−Six4 embryos showed similar patterns to wild-type and −/−Six5 embryos at E11.5 . Remarkably, bending of the PAW was missing at the distal levels in −/−;Six5−/−Six4 embryos at E11.5 . The overlay images of SIX4 and SIX5 immunofluorescence indicated that individual PAW cells showed the differential expression levels of SIX4 and SIX5 . As SIX1, another member of the Six family protein, is known to compensate for the roles of SIX4 in other tissues, we examined possible contribution of SIX1 in PAW development. However, SIX1 was not detected in the PAW at E10.75 . Consistent with protein localization, Six4 mRNA was completely absent in the PAW  and Six5 mRNA was expressed in the PAW at E11.5 and E12.5 . These results suggest that the expressions of Six4 and Six5 are differentially regulated in the PAW and they cooperatively function in the PAW at a stage before E11.5.The expression of reported , but thetibodies  at E10.0tibodies . SIX4 anand SIX5 C,F. In cand SIX5 G-I. ExprAs formation of the umbilical ring was impaired before the migration of muscle progenitor cells into the PAW, we next hypothesized that impaired cell proliferation in the PAW may cause abnormal PAW morphogenesis. To examine the cell proliferation rate in the PAW at the anterior to middle levels, we labeled proliferating cells with 5-ethynyl-2′-deoxyuridine (EdU) at E10.5 and counted the numbers of EdU-positive cells after 30 min within a 250 μm area from the lateral somitic frontier of the PAW M.Fig. 5P=0.22, Student's t-test, n=3)      C,D,M,N, st, n=3) C-F,M,N. st, n=3) C,D,M,N. st, n=3) C,E,M,N. st, n=3) D,F,M,N, P=0.17, Student's t-test, n=3)  than those in wild-type embryos  , resulting in the disappearance of the left-right asymmetrical ratio   A,C,D,O.  embryos B,E,F,O. st, n=3) A,C,D,O. st, n=3) B,E,F,O. Six4 and Six5 enhances cell death frequency in the PAW      G,I,J,P. st, n=3) H,K,L,P. st, n=3) G,I,J,P. st, n=3) H,K,L,P c−/−;Six5−/−Six4 embryos at E11.5 Six4−/−;SFig. S5H,I) but not at E9.25 and E9.5 . Laminin was accumulated with a fragmented distribution at the right and left sides of the PAW in wild-type embryos at E10.5  and mesenchyme was observed at E10.25  or Six4-expression vector (pCAGIG-flag-Six4). In the control vector introduction, 64.9±3.1%, 17.9±2.1% and 17.2±2.4% of EGFP-positive cells were detected in the CE, in the mesenchymal region adjacent to the laminin-positive basement membranes and in the mesenchymal region, respectively  but increased the ratio in the mesenchymal region adjacent to the laminin-positive basement membranes  and mesenchymal region   Q. These Table S3). However, in most of these mutant mice, anomalies other than omphalocele, such as neural tube defects, cleft palate and limb malformation, were observed but defects in the PAW have not been examined (summarized in Table S3). In the present study, we found that −/−;Six5−/−Six4 mice show severe ventral body wall closure defects with few other anomalies, they have defects in PAW formation, and that omphalocele formation is in complete penetrance (Table S2). Therefore, we emphasize that −/−;Six5−/−Six4 mice are a suitable animal model for large and small middle-type omphalocele reproducing typical human omphalocele.Various gene mutations cause body wall closure defects in mice  and, to −/−;Six5−/−Six4 embryos than in −/−Six5 embryos at the 0% position at E13.0 when the liver was partially protruded , the thinner abdominal wall in −/−;Six5−/−Six4 embryos was thought to be attributable to defects in the PAW before E11.5. Indeed, we detected the early occurrence of structural anomalies in −/−;Six5−/−Six4 embryos through high-resolution micro-CT analysis. The umbilical ring of −/−;Six5−/−Six4 embryos was larger than that of wild-type embryos before myoblast migration, suggesting that the initial change in ventral body wall closure had already occurred at E11.5 , analyses of −/−;Six5−/−Six4 mice first revealed that the impaired morphogenesis of the PAW may lead to middle-type omphalocele.Notably, the phenotypes of previously identified mouse mutants with ventral body wall closure defects have not been fully compared with omphalocele in newborn humans. Our analyses revealed that the AnPAW, LaPAW and PoPAW were thinner in  the PAW , Fig. S7at E11.5 . The PAW−/−;Six5−/−Six4 embryos at E10.5 before the onset of morphological change of the PAW. We observed the differential cell proliferation rates between the left and right somatic mesoderm-derived cells in the PAW of wild-type embryos at E10.5 , genes which are involved in the regulation of EMT in cultured cell lines . PCR-based genotyping was performed as previously described . escribed . Six4+/−escribed . Midday For Hematoxylin staining, E18.5 fetuses were embedded in paraffin and cut into 8-μm sections using a microtome . For Hematoxylin or Hematoxylin & Eosin staining for fetuses, frozen sections were prepared using cryostats . For skeletal staining, fetuses were fixed in 4% paraformaldehyde (PFA)/phosphate buffered saline (PBS) at 4°C for a week, and then the skin, mesenchyme and internal organs were removed. Fetuses were dehydrated in ethanol and incubated in acetone for 1 day. After incubation in ethanol, fetuses were incubated in Alcian Blue solution  at 37°C overnight and hydrated. Fetuses were incubated in 0.5% trypsin diluted with sodium borate solution for 6 h at 37°C and subsequently incubated in Alizarin Red S (Sigma) diluted with 0.5% KOH at room temperature overnight. Stained fetuses were washed with 0.5% KOH and treated with glycerol.−/−Six5: AnPAW level, 13 sections; LaPAW level, 14-17 sections; PoPAW level, nine sections. −/−;Six5−/Six4−: AnPAW level, 12 sections; LaPAW level, 20 sections; PoPAW level, six sections) were stained with anti-laminin antibody and DAPI. The distance of the PAW between the midline or distal position of the PAW and the RA was measured using cellSens software (Olympus), then the midline , half and proximal positions were defined as the 0%, 50% and 100% positions, respectively.The immunostaining of frozen sections was performed as described previously . Mouse ewww.osirix-viewer.com) and 3D Slicer (www.slicer.org) software.Imaging procedures were performed as previously reported . EmbryosIn situ hybridization was performed on 12 μm frozen sections as described previously  were used for hybridization.eviously . DIG-laband Six5  and Pitx−/−;Six5−/−Six4 embryos were used to obtain counts of EdU-positive cells and TUNEL-positive cells in the left and right PAW regions.A cell proliferation assay using EdU was performed as previously reported . To labepCAGIG-flag-Six4 was generated by inserting Flag-tagged mouse Six4 cDNA into multi-cloning sites of pCAGIG, which contain the CAG promoter and IRES-EGFP  was injected into the coelomic cavity of the right side in E11.0 mouse embryos using a fine glass needle . For quantification, EGFP-positive cells were counted in the total of 20 or 21 sections obtained from three embryos in control and Six4 experiments, respectively. The embryos were electroporated in a 20 mm×20 mm electro-chamber using an electroporator   and further cultured for 16 h for analyses.Whole-embryo culture and electroporation were performed according to a protocol described previously for rat embryos . Embryo CAX-EGFP  or pCAGIImages were taken using microscopes  equipped with a CCD-camera , a confocal laser microscope , a virtual slide microscope  and a microscope  equipped with an HD-color camera  and a recorder .


Ice‐wedge polygons are widespread periglacial features and influence landscape hydrology and carbon storage. The influence of climate and topography on polygon development is not entirely clear, however, giving high uncertainties to projections of permafrost development. We studied the mid‐ to late Holocene development of modern ice‐wedge polygon sites to explore drivers of change and reasons for long‐term stability. We analyzed organic carbon, total nitrogen, stable carbon isotopes, grain size composition and plant macrofossils in six cores from three polygons. We found that all sites developed from aquatic to wetland conditions. In the mid‐Holocene, shallow lakes and partly submerged ice‐wedge polygons existed at the studied sites. An erosional hiatus of ca 5000 years followed, and ice‐wedge polygons re‐initiated within the last millennium. Ice‐wedge melt and surface drying during the last century were linked to climatic warming. The influence of climate on ice‐wedge polygon development was outweighed by geomorphology during most of the late Holocene. Recent warming, however, caused ice‐wedge degradation at all sites. Our study showed that where waterlogged ground was maintained, low‐centered polygons persisted for millennia. Ice‐wedge melt and increased drainage through geomorphic disturbance, however, triggered conversion into high‐centered polygons and may lead to self‐enhancing degradation under continued warming. During times of peat accumulation, ice‐wedge polygons act as considerable sinks in the global carbon cycle.Large‐scale climate trends may trigger growth or degradation of ice‐wedge polygons at the regional scale. Widespread permafrost degradation was, for example, recorded during the early Holocene thermal maximum in the western Canadian Arctic.Reconstruction of ice‐wedge polygon development on the Yukon Coastal Plain during the mid‐ to late Holocene.Identification of drivers triggering (i) initiation of ice‐wedge polygon development and (ii) conversion of low‐centered polygons into high‐centered polygons.Determination of factors promoting stability in ice‐wedge polygons.We addressed the spatial heterogeneity2Ice‐wedge polygons are most widespread in regions underlain by continuous permafrost.One way to classify different morphological types of ice‐wedge polygons is to distinguish low‐centered polygons from high‐centered polygons.Ice‐wedge formation may be related to large‐scale climate trends. Thermal contraction cracking requires severe ground frost in winter,Ice‐wedge polygons also experience drastic geomorphological changes, most recently induced by permafrost thaw. Increased thaw has been observed to produce thicker active layers and degrading ice wedges,Studies of long‐term ice‐wedge polygon development have shown that ice‐wedge polygons may exist in a relatively stable state over millennia.3The study area is situated on the terrestrial part of the Canadian Beaufort Sea shelf. It is characterized by a subarctic, maritime climate, a flat to slightly undulating topography, and ice‐rich unconsolidated sediments shaped by periglacial processes in the western part and by Pleistocene glaciations superimposed by periglacial processes in the eastern part.http://climate. weather. gc.ca). About half of the scarce precipitation falls as snow, resulting in a thin snow cover (mean 25 cm), which is locally variable due to strong wind redistribution and prevails for 250 days per year on average. The topography of the plain is characterized by a flat coastal zone and rolling hills toward the Mountain range. This study focused on the flat coastal reaches, which were shaped by (i) Late Pleistocene advances of the Laurentide Ice Sheet, which reached its furthest extent about 16.2 ka BPCarex sp) dominating sites with impeded drainage, and tussock cottongrass (Eriophorum vaginatum) dominating better drained, elevated surfaces.Salix spp., Betula glandulosa and Rubus chamaemorus, while in river valleys sheltered conditions promote taller growth of the shrubby taxa Salix spp., Alnus crispa and Betula glandulosa.The Yukon Coastal Plain stretches across 240 km of coastline from the Mackenzie Delta in the east to the Alaskan border in the west and is bordered by the Beaufort Sea in the north and by the British Mountains in the south, leaving it 10–40 km wide  and elevated surfaces  within the polygon.Ptarmigan Polygon Figure b was sitBetula glandulosa, Salix pulchra, Rubus chamaemorus, Ledum decumbens, Vaccinium vitis‐idaea, Hierochloë alpina, Eriophorum vaginatum and Luzula confusa.Roland Polygon Figure c was loc44.1Field work at Komakuk Polygon (YC12‐KP) and Roland Polygon (YC12‐RP) was conducted in August 2012, and field work at Ptarmigan Polygon (YC13‐PP) was conducted in July 2013. The sampling approach was identical for all three polygons. The results from a detailed survey of microtopography and vegetation at the sites have been published in Wolter et al.4.214C and as conventional radiocarbon ages following the conventions of Stuiver, PolachThe six active‐layer cores were photographed and described in the laboratory, before being subsampled in 1 cm increments. In three cores , the lowermost samples could not be reasonably divided further, so the lowermost 1.5 or 2 cm was taken as one sample. In total, 24 radiocarbon dates were obtained from terrestrial plant macrofossils (Table 13C) was conducted on freeze‐dried, ground, carbonate‐free subsamples at Helmholtz Centre, Potsdam GFZ German Research Centre for Geosciences, Potsdam, Germany, using a Thermo Fisher Scientific DELTAplusXL mass spectrometer. Stable carbon isotope analyses on subsamples of core YC12‐RP‐Mr were measured at Alfred Wegener Institute Helmholtz Centre for Polar and Marine Research, Potsdam, Germany, using a Finnigan MAT DELTA‐S mass spectrometer. All δ13C values are expressed as per mil relative to the Vienna PeeDee Belemnite standard (‰ vs. VPDB). Grain size analyses were carried out on carbonate‐free and organic‐free subsamples using a Beckman Coulter LS 200 laser diffraction particle sizer. In the upper centimeters of the cores, grain size analyses were precluded by very low contents of minerogenic material in the peat. Grain sizes are given as volume percent (vol.%). Plant macrofossil analyses were conducted on selected subsamples . For each subsample, 50 mL of sediment was wet sieved through 1 mm and 250 μm mesh sizes. Due to the large amounts of coarse organic material in the samples, we limited the analyses to picking and identifying vascular plant remains in the >1 mm fraction. This approach provided an overview of vascular plant taxa that were present in the cores, while a full plant macrofossil analysis would have included smaller seeds.We measured total organic carbon (TOC) and total nitrogen (TN) on freeze‐dried, ground subsamples using an Elementar Vario Max C analyzer (TOC) and an Elementar Vario EL III analyzer (TN). Element contents are expressed as weight percent (wt.%). The analysis of stable carbon isotopes . Sediment data were range‐transformed to ensure the data are on the same scale. PCAs were conducted using the function “princomp” in R. Hellinger and range transformations were executed using the function “tran” in the package analogueThe zonation presented for the cores was delineated using the Constrained Incremental Sum of Squares (CONISS) algorithm55.1c1 and KPc2 for the core, which corresponded with the facies break and were mainly distinguished by a sharp upward increase in TOC  had a median basal age of 1597 cal years BP. The core showed a distinct sedimentary facies break at 14 cm depth, which was accompanied by an age inversion , TOC was low to intermediate, while TOC/TN showed relatively low values and δ13C was relatively high , wet terrestrial  and aquatic  plant macrofossils was preserved in this zone (Table In zone KPc2 (0–13 cm depth) uniformly showed very high TOC contents, increasing TOC/TN toward the top of the core and decreasing δ13C  and wet taxa (Carex sp.) were found. Remains of aquatic plant taxa were absent.Zone KPc2 0–1 cm depthr1 and KPr2 corresponded roughly with the age hiatus.The active layer core from the rim of Komakuk Polygon YC12‐KP‐Mr) showed a hiatus of about 5000 years between 16 and 17 cm depth (Table 2‐KP‐Mr sr1 (16–31 cm depth) TOC values were intermediate to very high, with a high variability and accompanied by stable TOC/TN values and decreasing δ13C values upcore  showed very high and uniform TOC contents, while TOC/TN ratios increased strongly toward the top of the core, and δ13C values were stable  dominated in this zone, while remains of wet terrestrial taxa (Carex sp.) were scarce and aquatic taxa (Potamogeton sp.) disappeared above 15 cm core depth (Table r2A to KPr2B.Zone KP5.2c1 and PPc2 in the active layer core (YC13‐PP‐Mc) and subjacent permafrost core (PG2161)  had post‐bomb radiocarbon dates (Table c1 (24–86 cm depth) organic matter was characterized by low to intermediate TOC contents, uniformly low TOC/TN ratios and relatively high δ13C values , organic matter was characterized by high TOC contents, while TOC/TN ratios were similar to those found in zone PPc1, and δ13C was slightly lower than in PPc1  had a median basal age of 1127 cal years BP. CONISS ordination validated by broken stick modeling supported no zonation in the core Figure b. TOC vae Figure b. The te‐PP‐Mr ha5.313C, we established two stratigraphic zones in the core  showed a hiatus but no age inversions  was characterized by low TOC values, low TOC/TN ratios and relatively high δ13C values  alongside seeds of wet terrestrial Carex and emergent aquatic Hippuris vulgaris, Menyanthes trifoliata, and Potentilla palustris  were frequent in this zone.RPs Figure a. The seis Table . Remainsc2 (0–21 cm depth) showed high TOC values and a strong increase in TOC/TN, accompanied by a marked decrease in δ13C  became particularly abundant in RPc2, while Carex sp. seed occurrence declined gradually and aquatic remains disappeared entirely.Zone RPC Figure a. The str1 and RPr2 were delineated based on CONISS ordination of parameters characterizing organic matter , and the lower zone was divided into subzones RPr1A and RPr1B  showed a similar hiatus and had a median basal age of 7085 cal years BP that was nearly identical to the one in the center core YC12‐RP‐Mc from the same polygon (Table 2‐RP‐Mr sr1 (12–27 cm depth) had high TOC contents, while TOC/TN ratios were intermediate and δ13C decreased slightly upcore   Table . Remainsr2 (0–11 cm depth), TOC contents were high and TOC/TN ratios increased strongly upcore, while δ13C was stable  and generally on wet ground.Potamogeton is a true aquatic genus indicating water depths on the scale of decimeters to meters.Eriophorum vaginatum, Betula glandulosa, Ledum decumbens and Vaccinium vitis‐idaea are all typical of mesic conditions, as found on elevated, better drained surfaces in ice‐wedge polygons on the Yukon Coastal Plain.13C illustrates that the main variance pattern in the core data is indeed associated with a hydrological gradient from submerged to well‐drained conditions  than those of submerged polygons, but they overlap largely. Overall, organic‐rich lake sediment from shallow lakes is clearly separated from peat of submerged, wet and mesic ice‐wedge polygon microsites.This development is supported by PCA on plant macrofossil and sediment data Figure . PCA of 6.16.1.1c1 , fine‐grained sediments with TOC values around 10 wt.% indicated the presence of a lake environment rather than peat from an ice‐wedge polygon. While lake sediment in deeper parts of lakes in the region exhibits slightly lower TOC, lower C/N and higher δ13C,c1.The center core exhibited two zones that we interpreted as lake sediments overlain by peat. In zone KPc2  peat established. Strong parallels in composition, plant macrofossil assemblage and thickness of this peat layer and the dated peat in KPr1 from the same ice‐wedge polygon suggested a similar age range for KPc2 r2B, and their increasing abundance caused increased C/N ratios in the peat. We suggest that this vegetation shift accompanied the conversion of a low‐centered polygon into an intermediate‐centered polygon within the last 300 years.The peat in KPng Table . Radiocang Table , limitinThe cores taken from Komakuk Polygon were spaced only 5 m apart, and reconstruction of polygon development at the site indicated the presence of a mid‐Holocene peatland, followed or interrupted by a phase of aquatic conditions in a lake margin environment, during which sediment was lost from the polygon rim by erosion and/or decomposition of organic material. Regular peat growth reinitiated during the last 300 years. Both cores showed shifts from aquatic vegetation to mesic ice‐wedge polygon vegetation, which is characterized by different habitats existing in close proximity. The conversion of a low‐centered polygon to an intermediate‐centered polygon occurred within the last 300 years, probably very recently.6.26.2.1c1, only small amounts of plant material of mixed origin were present, and low and stable TOC and C/N values suggested a lake sedimentation environment. This interpretation is supported by the PCA of sediment parameters , sedge peat established, as evident from stable high TOC contents, consistently low C/N ratios , a lake environment existed, as indicated by abundant occurrences of Charophyte oogonia, Potamogeton seeds and Daphnia ephippiae along with low TOC, low C/N, relatively high δ13C and a fine‐grained sediment texture , to the initiation of a low‐centered polygon , and subsequent gradual conversion to a high‐centered polygon (last century). During the period corresponding to zone RPe Figure , Table 4y Figure c, in whic2  peat established and aquatic taxa disappeared . In accordance with the available dates we suggest that the transition to drier conditions occurred within the last 100 years. Macroremains of plants were entirely from mesic taxa that were identified at the site during a vegetation survey in 2012  were found  matched the center core. RPd Figure , Table 4eg, Bliss. Daphniar1B drier conditions established, indicated by decreasing δ13C, rising C/N, decreasing amounts of Carex seeds, absence of aquatics and increasing dominance of mesic terrestrials  are determined by climate and geomorphology. Ice‐wedge polygon initiation and conversion of low‐centered into high‐centered polygons is therefore strongly related to the dynamics of and the interplay between the two.ad 1600–1850; seeInvestigations into radiocarbon dates have revealed broad climate‐induced simultaneous patterns of peatland initiation.Topographic evidence favors geomorphic causes for ice‐wedge polygon initiation on the Yukon Coastal Plain, where most polygon fields, including the ones we studied, are situated in drained thaw lake basins. The initiation of peat accumulation in Ptarmigan Polygon was probably linked to sea level rise in addition to lake drainage. The lake was probably already drained when Workboat Passage Figure  was flooThe conversion of low‐centered polygons to high‐centered polygons is thought to be linked to internal self‐organization,The conversion of one polygon type to another may result from internal self‐organization through two main processes: lateral movement of material adjacent to ice wedges may widen ice‐wedge troughs and displace material toward the polygon center, where a mound establishes.Improved drainage may result from a change in topographic gradient and thus in surface flow patterns, or from ice‐wedge degradation promoting drainage of polygon centers into the surrounding ice‐wedge troughs. The modern positions of Komakuk Polygon and Roland Polygon on elevated surfaces above lakeshore bluffs of several meters height Figure a, c indiIn the two studied ice‐wedge polygons that experienced conversion from low‐centered to intermediate‐centered (Komakuk Polygon) or high‐centered (Roland Polygon), both rim cores and one center core show a hiatus of at least 5000 years Figure  caused bThe changes we observed in peatland initiation and change from low‐centered to high‐centered were mostly caused by geomorphological change . In permafrost‐affected landscapes, climatic change may trigger widespread geomorphological change, especially where unconsolidated ice‐rich sediments dominate. Such climate‐induced geomorphological change may have locally variable impacts, but its frequency is likely to increase under climatic change. Regionally synchronized ice‐wedge polygon development requires a higher amplitude and seasonality of temperature and precipitation change than evident for the mid‐ to late Holocene. Our findings indicate that modern warming, however, may have triggered regional‐scale conversion from low‐centered polygons to high‐centered polygons. This process may rapidly initiate irreversible self‐enhancing erosion of ice‐wedge polygons.Roland Polygon experienced stability for at least 2000 years during the mid‐Holocene  ice‐wedge polygon initiation was lake drainage. The main driver triggering (ii) conversion of low‐centered polygons to high‐centered polygons was improved drainage through ice‐wedge degradation and changes in the local topographic gradient. By contrast, we found that stable conditions prevailed for millennia during the late Holocene in ice‐wedge polygons under low‐amplitude climatic change as long as a non‐negative water balance was maintained in the polygon field. Hence, geomorphic disturbance was the main driver of locally variable ice‐wedge polygon dynamics in periods of low climatic forcing. Extreme climatic change triggered simultaneous developments, such as widespread peat initiation during the Holocene Thermal Maximum. Modeling of the development of polygon fields through time must thus focus on temperature constraints as well as landscape water balance and flow paths.


Sophora flavescens have a variety of biological benefits for humans; however, cancer immunotherapy effect has not been fully investigated. The flavonoids (1–6) isolated from S. flavescens showed IDO1 inhibitory activities (IC50 4.3–31.4 µM). The representative flavonoids (4–6) of S. flavescens were determined to be non-competitive inhibitors of IDO1 by kinetic analyses. Their binding affinity to IDO1 was confirmed using thermal stability and surface plasmon resonance (SPR) assays. The molecular docking analysis and mutagenesis assay revealed the structural details of the interactions between the flavonoids (1–6) and IDO1. These results suggest that the flavonoids (1–6) of S. flavescens, especially kushenol E (6), as IDO1 inhibitors might be useful in the development of immunotherapeutic agents against cancers. Indoleamine 2,3-dioxygenase 1 (IDO1), a tryptophan catabolising enzyme, is known as a tumour cell survival factor that causes immune escape in several types of cancer. Flavonoids of   Sophora flavescens has been used traditionally to treat skin diseases, viral hepatitis, and cancer owing to the presence of medicinal components such as alkaloids and flavonoids,L-Tryptophan (L-Trp) is an essential amino acid and can be converted to kynurenine (Kyn), a natural ligand of the aryl hydrocarbon receptor, by Trp-catabolising enzymes such as indoleamine 2,3-dioxygenase 1 (IDO1), indoleamine 2,3-dioxygenase 2 (IDO2), and tryptophan 2,3-dioxygenase (TDO),,L-Trp to Kyn and inducing amino acid starvation in cancer environment. In addition, IDO1 induces differentiation of regulatory T cells and avoids the immune responsesThe S. flavescens have been reported as tyrosinase inhibitors, anti-inflammatory and anti-oxidant substances, but have not been reported as IDO1 inhibitors yet,1–6) from S. flavescens as a novel class of IDO inhibitor.In fact, some flavonoids of S. flavescens were purified by the method of Lee et al.1–6) were identified from spectroscopic data and chemical evidence through comparisons with previous reports,The flavonoids of Escherihia coli DH5α and T7 strain were used for gene cloning and protein production, respectively. The human IDO1 clone was provided from Korea Human Gene Bank, Medical Genomics Research Center, KRIBB, Korea. The full-length of IDO1 gene was amplified by PCR using pfu DNA polymerase . The PCR product from IDO1 mRNA was cloned into expression vector pET28a to obtain pET28a-IDO1 and confirmed using DNA sequencing . The E. coli T7 competent cells were transformed with this plasmid and expression of IDO1 protein was induced with 0.2 mM isopropyl-β-D-1-thiogalactopyranoside  for 4 h at 37 °C. The expressed IDO1 protein was purified using Ni-NTA agarose  at 4 °C. The eluted IDO1 protein was dialysed using storage buffer (20 mM Tris-HCl buffer (pH 8.0 and 200 mM NaCl)). The purified IDO1 protein was analysed with SDS-PAGE.The L-Trp (200 µM). The metabolites were serially diluted in dimethyl sulfoxide  and transferred to a 96-well plates. The reaction was conducted at 37 °C for 1 h and stopped by adding 40 µL of trichloroacetic acid (30%); the solution was then heated at 65 °C for 15 min. The plates were centrifuged at 4000 rpm for 15 min to remove debris. Next, the supernatant (125 µL) from each well was transferred to a 96-well plates. p-dimethylaminobenzaldehyde  in acetic acid was then added to each well, and the absorbance value of the plate was read at a wavelength of 480 nm using a SpectraMAX-190 ELISA reader .The enzyme assays were performed with previous descripted in Takikawa et al.2.HeLa  cells were purchased from the American Type Culture Collection (ATCC) and cultured in DMEM supplemented with fetal bovine serum  and penicillin–streptomycin (1%). The cells were maintained at 37 °C with 5% CO4 cells/well). Next, the cells were stimulated with human recombinant IFN-γ  in complete DMEM for 24 h. After replacing the medium with DMEM containing IFN-γ, the cells were treated with the flavonoids (1–6) or DMSO  for 6 h, and then the supernatant (125 µL) from each well was transferred to a 96-well plate. The cell-based assay was similarly performed to the IDO1 enzyme assay.HeLa cells were seeded into 48-well plates (2 × 103 cells/well) were cultured in 96-well plates and pretreated with human recombinant IFN-γ for 24 h. After following an overnight incubation, each of the cells was treated with serially diluted flavonoids (4–6) for 6 h. EZ-Cytox (10 µL) was added to each of the wells of the plate and incubated for 1 h. The absorbance was determined at 450 nm with SpectraMAX-190 ELISA reader . The cell viability was normalised to the control group.The cell viability was tested using the EZ-Cytox  reagent according to the manufacturer’s recommendations. Briefly, HeLa cells (5 × 10L-Trp as substrate of IDO1 were added to the reaction mixture with four different concentrations of flavonoids (0–45 µM). After incubation at 37 °C for 1 h, the dissociation constant (iκ) of the three flavonoids were determined by measuring the activity at five concentrations of L-Trp (0 to 50 µM).The kinetics assay was performed by adding purified protein to the reaction mixture (200 µL) which contained a final concentration of potassium phosphate buffer , ascorbic acid (20 mM), catalase (10 mM), methylene blue (10 µM), and distilled water up to 200 µL. Various concentrations of 4), (2S)-2'-methoxy kurarinone , and kushenol E , or serially diluted KE (6) in 20 µL reaction buffer containing potassium phosphate buffer (50 mM) and sodium chloride (100 mM), and then incubated for 1 h. The measurement of protein denaturation via an increase in temperature was determined using real-time PCR device  using the FRET channel. In the presence of IDO1 enzyme, SYPRO-Orange has a maximum excitation near 490 nm and a maximum emission near 600 nm. The heating rate and fluorescence reading were 0.5 °C/10 s depending on the temperature range 25–95 °C.The thermal stability of proteins was measured by SYPRO-Orange  as previously described in Soderhole et al.6) in PBS and 5% DMSO at 25 °C using a Reichert instrument  using a sensor chip CM5  . Amine coupling of His-IDO1 (WT and P314A) was performed with 0.2 M 1-ethyl-3-(3-dimethylaminopropyl)-carbodiimide, 0.05 M N-hydroxysuccinimide, and 1 M ethanolamine . The results were analysed using Scrubber software .The SPR assay was performed with 10 mM stock solution and five serially diluted concentrations of KE (1–6) were built using the Maestro build panel and minimised using the Impact module of Maestro in the Schrödinger suite program1–6) were docked into the prepared receptor grid by LibDock in Discovery Studio 2018 suite 6).The crystal structure of IDO1 from the Protein Data Bank (PDB code 6AZU (holoenzyme form)) was used for docking simulationsTM mutagenesis kit . Oligonucleotides for generation of Pro182 (P182A), Phe185 (F185A), double mutation (P182A and F185A), and Pro 314 (P314A) to alanine mutants (mutated residues are underlined) were as follows:For P182A: 5′-CAATCAAAGTAATTGCTACTGTATTCAAGG-3′, 5′-CCTTGAATACAGTAGCAATTACTTTGATTG-3′.For F185A: 5′-AATTCCTACTGTAGCCAAGGCAATGCAAAT-3′, 5′-ATTTGCATTGCCTTGGCTACAGTAGGAATT-3′.For double mutation (P182A and F185A): 5′-AATCAAAGTAATTGCTACTGTAGCCAAGGCAATGCAAATG-3′, 5′-CATTTGCATTGCCTTGGCTACAGTAGCAATTACTTTGATT-3′.For P314A: 5′- CATTAGAGTCAAATGCCTCAGTCCGTGAGT-3′, 5′- ACTCACGGACTGAGGCATTTGACTCTAATG-3′.Point mutations were introduced using the Muta-DirectAll construct sequences were confirmed before use.S. flavescens indicated the possibility for cancer treatments and anti-inflammatory agents,,S. flavescens. For this reason, all of the flavonoids (1–6) isolated from S. flavescens were tested for IDO1 inhibitory activity (1), sophoraflavanone B (2), (−)-kurarinone (3), kushenol F , (2S)-2'-methoxy kurarinone , and kushenol E  from S. flavescens exhibited IDO1 inhibitory activities (50 = 7.7 µM) is three times more effective than (−)-kurarinone , KF  and MK , respectively. The presence of a hydroxyl group at B and C ring is also important for inhibitory potency: i.e. compare (−)-kurarinone (3) versus MK  and sophoraflavanone B versus noranhydroicaritin (IC50 = 19.3 versus 23.5 µM). We also tested 1-methyl-tryptophan (1-MT) as a positive control and steppogenin, a non-prenyl and non-lavandulyl flavonoid backbone, as negative control.The diverse therapeutic effects of flavonoids from activity . The nortivities . Our resL-Trp degradation by IDO1 (4), MK (5), and KE (6) are more potent than other flavonoids (1–3) in cells  showed the strongest inhibitory effect on Kyn production (IC50 = 4.3 ± 0.3 µM).To investigate IDO1 activity in tumour cell, HeLa cells were exposed to recombinant IFN-γ (50 ng/ml) to induce endogenous IDO1 expression and analysed the Kyn production level from  by IDO1 . The cel by IDO1 . Three fin cells  and reduectively . As the ectively . KE (6) 4–6) were quantitatively analysed by the Lineweaver–Burk plots. The experiments were performed using L-Trp, as the substrate at a concentration of 0–50 µM and with the flavonoids at a concentration of 0–45 µM. This analysis . The iκ values of KF (4), MK (5), and KE (6) were 20.2 ± 3.1, 13.7 ± 0.03, and 9.5 ± 0.6 µM, respectively (4–6) showed a typical tendency of non-competitive inhibition behaviour.To investigate the inhibition mode, the kinetics of representative IDO1 inhibitory active flavonoids  to IDO1 was determined by the temperature stability analysis using SYPRO-Orange. The SYPRO-Orange binds to the hydrophobic residues in the unfolded regions of proteins, and is denatured by temperature4–6) to IDO1 with SYPRO-Orange in order to confirm whether they change unfolding temperature of IDO1 enzyme without L-Trp. The flavonoids (4–6) were increased the mean melting temperature compared to the DMSO-treated group by +1.5 ± 0.9, +0.4 ± 0.7, and +5.7 ± 0.8 °C, respectively (6) exhibited the strongest IDO1 inhibitory activity, we tested experiments with various concentrations of KE (4–6), especially KE (6) increase the structural stability of the IDO1 enzyme by directly binding.The thermal stability assay has been widely used to confirm the interactions between compounds and their target proteins. The interaction of compound is expected to enhance the thermal stability of the proteinectively . Since Kns of KE . These b1–6) to the IDO1, enzymes were applied to the molecular docking analysis. The molecular docking analysis using LibDock program revealed the mode of binding of flavonoids (1–6) at the allosteric site of IDO1 (1–6) were calculated to be docked in the hydrophobic valley region (arrow) of the enzyme. For this docking analysis, the ligand binding sites in IDO1 were searched and there were five candidate sites (data not shown). Among them, however, we selected the sites where six kinds of flavonoids were commonly combined and where a site had the highest stable energy. The Ser309, Asn313, and Ile178 residues in the binding pocket played critical roles in the high-affinity binding of the flavonoids (1–6) through hydrogen-bond formation. Furthermore, pi–pi stacking interactions were observed between Phe306 and the flavonoids (1–6). However, we anticipated that Pro182, Phe185, and Pro314 were more important for binding of the flavonoids (1–6). This was because prenyl and lavanduyl residues of the flavonoids interacted with the site of the IDO1 enzyme. In 6) was shown to bind to Pro182 and Phe185, and the bottom prenyl residue was shown to bind to Pro314 through alkyl formation with A ring on the prenyl and lavanduyl groups, but other flavonoids (1–5) interacted only with Pro182 and Phe185 (1–6), especially KE (6), have the most inhibitory activity against IDO1.To visualise the binding of flavonoids  bi-prenyl residues. We generated four types of mutations, P182A, F185A, P182A/F185A double mutation, and P314A. Next, we confirmed that Pro182 and Phe185 were the crucial binding sites of the flavonoids, because their mutagenesis  eliminated the IDO1 activity  binding to IDO1.To confirm which binding sites were crucial for IDO1 inhibition by these flavonoids, we performed mutagenesis in the His-IDO1 structure. According to the molecular docking results, Proactivity . Indeed,± 1.8 µM . These r6), we utilised a surface plasmon resonance (SPR) assay. Specifically, His-IDO1 was linked to the sensor chip and then incubated with successive injections of the KE (6) solution at a flow rate of 30 µL/min in PBS buffer-containing 5% DMSO. The SPR results showed that KE (6) was strongly bound to WT IDO1 with a DK value of 6.5 ± 0.1 µM (DK of 8.9 ± 0.2 µM (DK) was calculated from the ratio of the dissociation rate constant divided by the association rate constant (DK = dk/ak). Based on this, we confirmed the direct binding of KE (6) to IDO1.To determine the interaction between IDO1 and KE  from S. flavescens showed inhibitory effects on IDO1, in both enzyme-based and cell-based assays. The representative flavonoids (4–6) were identified as non-competitive IDO1 inhibitors through kinetic analyses, and the direct interactions with IDO1 were confirmed by thermal stability assay and SPR assay. We checked the amount of active His-IDO1 through by measuring absorbance at 280/404 nm absorbance (data not shown), because IDO1 can exist in equilibrium in its forms of haeme-bound (holo-IDO1) and haeme-free (apo-IDO1) form. These results indicate that the half of the measured His-IDO1 amount was in fact the apo-form. However, this did not affect the results of the IDO1 enzyme kinetics assay, but could explain the high DK values obtained in the SPR assay.In this study, the flavonoids (1–6) are important to IDO1 inhibitory activity, because steppogenin, a non-prenyl and non-lavandulyl flavonoid, does not affect IDO1 inhibitory activity. Molecular docking analysis supported the results of IDO1 enzyme assay. In the molecular docking results, the prenyl residues of active flavonoids have binding affinity to Pro182 and Phe185. The flavonoids (1–6) were found to be embedded in the inner side of IDO1 completely (6) were accomplished by hydrogen bonds and alkyl interactions (1–6) were bound to the Pro182 and Phe185 residues through alkyl interactions. In particular, Pro314 of KE (6) imparted a binding site on another prenyl residue (6) had the strongest activity as compared to other flavonoids of S. flavescens. Therefore, flavonoid prenyl or lavandulyl residues exhibited IDO1 inhibitory activity, as a difference to steppogenin. Thus, we considered that the alkyl interactions by prenyl were important in IDO1 inhibition.Initially, we determined that the residues of the flavonoids . Therefore, these molecules should be further investigated due to their IDO1 inhibitor ability and their potential use in cancer immunotherapy.According to previous reports, substrate


Recent studies suggest that Antarctica has the potential to contribute up to ~15 m of sea-level rise over the next few centuries. The evolution of the Antarctic Ice Sheet is driven by a combination of climate forcing and non-climatic feedbacks. In this review we focus on feedbacks between the Antarctic Ice Sheet and the solid Earth, and the role of these feedbacks in shaping the response of the ice sheet to past and future climate changes. The growth and decay of the Antarctic Ice Sheet reshapes the solid Earth via isostasy and erosion. In turn, the shape of the bed exerts a fundamental control on ice dynamics as well as the position of the grounding line—the location where ice starts to float. A complicating issue is the fact that Antarctica is situated on a region of the Earth that displays large spatial variations in rheological properties. These properties affect the timescale and strength of feedbacks between ice-sheet change and solid Earth deformation, and hence must be accounted for when considering the future evolution of the ice sheet. The evolution of the Antarctic Ice Sheet is driven by a combination of climate forcing and non-climatic feedbacks. In this review, the authors focus on feedbacks between the Antarctic Ice Sheet and the solid Earth, and the role of these feedbacks in shaping the response of the ice sheet to past and future climate changes. In the intervening period the bed of the AIS has been continuously reshaped by erosion and sedimentation, periodically flexed by glacial isostasy, and abruptly altered by tectonic and volcanic activity, with the latter two processes also playing a role in determining the thermal conditions at the bed . The pre-glacial topography of the Antarctic continent determined the location and style of glacial inception ~34 Ma ago7 but they have received renewed attention over the last decade, reflecting the important role they play in the interpretation of satellite measurements of contemporary ice-sheet change8. A number of recent studies have also sought to understand the strength of feedbacks acting in the opposite direction, that is, the impact of solid Earth deformation on ice dynamics and the potential for this deformation to delay or prevent unstable ice-sheet retreat10. Such feedbacks were first considered in the 1980s11 but have only recently begun to be implemented into fully-coupled models12, where the evolving shape of the solid Earth and the depth of the ocean adjacent to an ice sheet grounded below sea level act as fundamental boundary conditions on the dynamics of the ice sheet.Interactions between ice sheets and the solid Earth have long been studied within the field of Glacial Isostatic Adjustment (GIA)—defined here as the response of the solid Earth and the global gravity field to changes in the distribution of ice and water on Earth’s surface. The first numerical models of GIA were developed in the 1970s13 indicates that there are significant spatial variations in mantle properties beneath Antarctica, which suggests mantle viscosities, and hence relaxation timescales, may vary by up to several orders of magnitude from the global mean. Indeed, in the northern Antarctic Peninsula, geodetic evidence14 suggests that contemporary ice loss is triggering a viscous response orders of magnitude more rapidly than was previously assumed possible in Antarctica – on decadal rather than millennial timescales. Feedbacks on ice dynamics are likely to be enhanced in such regions15, and the ongoing GIA signal is likely to be dominated by recent ice-sheet change (few millennia or less), as opposed to the ice loss that followed the Last Glacial Maximum (LGM)16. Here, we define the GIA signal to be the ongoing response of the solid Earth, the gravity field, and/or relative sea level to past ice-sheet change.The strength of any feedbacks between glacial isostasy and ice dynamics depends on the rate at which the solid Earth responds to ice-sheet change, which, in turn, depends on the rheological properties of the mantle. Seismic evidence18 suggests that the response to surface load change can alter bed slopes across West Antarctica by 0.25–0.4 m/km  and that water-depth change around the margin of the ice sheet can deviate from eustatic by >100 m. Coupled ice sheet-sea level models can now be used to quantify the impact of such changes on ice sheet evolution19. Over timescales of millions of years, changes to the shape of the ice sheet bed will additionally reflect processes associated with erosion and deposition, the isostatic response to sediment redistribution, and mantle convection. Feedbacks between ice sheet evolution and long-term landscape evolution have been hypothesised20, but they have not been modelled within a coupled framework. The impacts of long-wavelength isostasy-driven changes to ice sheet boundary conditions are reviewed here, but smaller-scale subglacial controls, such as the material properties of the bed and variations in subglacial hydrology and geomorphology, are discussed elsewhere2.The growth and decay of the AIS is driven by a combination of climate forcing and non-climatic feedbacks, but modelling studies that seek to understand the controls on AIS change often neglect to consider how the ice sheet alters its own boundary conditions. Over the timescale of the last deglaciation, GIA model outputIn this review we proceed by briefly summarising the current state of knowledge on ice-sheet change and glacial isostasy across Antarctica, before discussing the impact of spatial variations in Earth rheology on glacial isostasy and the impact of GIA on ice dynamics. Feedbacks between ice sheet and landscape evolution are discussed, and we conclude by identifying a number of future research priorities that link ice history, Earth rheology, and the overarching issue of global sea-level change.22. While the third component is well defined by physical theory, Antarctic ice-sheet reconstructions and solid Earth rheology are subject to substantial uncertainty and debate.Forward model predictions of GIA-related solid Earth deformation, gravity-field change, and polar motion rely on: a reconstruction of the spatial and temporal pattern of past ice loading, a viscoelastic Earth model that describes the time-dependent response of the solid Earth to surface loading changes, and the iterative consideration of physical feedbacks associated with polar motion and coastline position as they are altered by the deformation of the solid Earth and the gravity field23. On a global scale, such data-model comparisons have been used to estimate the timing of ice-sheet growth and retreat over glacial cycles25, but the combined dependence of model output on Earth rheology and ice history leads to non-uniqueness26, and it remains challenging to determine how ice was partitioned between the different ice sheets27.GIA modelling can be used to infer past large-scale ice-sheet change via comparison of model output with a range of constraint data relating to sea-level change and solid Earth deformation28. This range reflects differences in methodology and the interpretation of data constraints. The most common reconstruction approaches consider either numerical models of ice-sheet dynamics29 or local geological, geomorphological, and geodetic data constraints on past ice extent30, or a hybrid32. Data constraints are sparse as they rely on sampling of spatially limited bedrock exposures34 or interpretation of ice core records35, while numerical ice-sheet models are restricted by a limited representation of reality. Continent-wide ice-sheet reconstructions can be produced by simulating the response of dynamic models to past climate changes38, but at present there is not enough data to tightly constrain such models. Consequently, there is significant variation in the predicted magnitude and spatial pattern of isostatic deformation across Antarctica due to past ice-sheet change39 . However, the time-varying gravity field also contains a gravity-change signal associated with past ice-sheet change; GIA modelling is used to determine the ‘GIA correction’ that should be applied to remove the signal associated with past change8. Uncertainties associated with this process remain the dominant source of error in gravimetry-based estimates of the Antarctic contribution to contemporary sea-level rise41. However, newer GIA model predictions32 yield improved inter-model agreement, as do inverse GIA solutions derived by combining ice elevation and GRACE data44. Recent estimates for the magnitude of the net Antarctic GIA signal vary over the range ~40–80 Gt/yr45, but basin-level differences remain substantial and in some cases different studies do not even agree on the sign of the mass change in each basin43, which hampers the advance of glaciological insights into the processes governing present change.The effect of GIA on geodetic and gravimetric measurements, and the resulting impact on estimates of ice-mass change, has driven much of the recent interest in Antarctic GIA. Contemporary ice-sheet change can be determined from analysis of Earth’s time-varying gravity field, as measured by the Gravity Recovery and Climate Experiment (GRACE) satellites between 2002 and 201747. For the same reasons, forward models underestimate the signal in the northern Antarctic Peninsula, but then so do inverse solutions because the GIA signal in this region has a shorter spatial wavelength than inverse solutions can resolve (~300 km). Comparison of the new generation of GIA models with GPS-derived uplift rates—which must be corrected for elastic deformation associated with contemporaneous ice-mass change49—demonstrates that important differences remain, notably in West Antarctica and especially in the Amundsen Sea and northern Antarctic Peninsula regions39.Differences in estimates of the present-day GIA signal are particularly acute in the region of the Amundsen Sea embayment, where forward models have substantially less GIA signal than inverse solutions. This may be due to the fact that the GIA signal in this low mantle viscosity region predominantly reflects significant decadal-to-centennial ice load change that is not accounted for in most forward modelsThe pattern and magnitude of the solid Earth response to ice-sheet growth and decay is strongly dependent on the rheology of the interior of the solid Earth, and the choice of rheological model has a large effect on the modelled GIA response see Box . For exa53, with greater heterogeneity than observed in areas of Northern Hemisphere continental glaciation54. East Antarctica shows high upper mantle seismic velocities characteristic of cold cratonic regions worldwide, whereas West Antarctica shows upper mantle structure consistent with much warmer tectonically active zones , which are similar to estimates of upper mantle viscosity in Fennoscandia61. The emerging picture suggests cratonic East Antarctica is characterised by higher upper mantle viscosity than West Antarctica, with exceptionally low upper mantle viscosity, on the order of 1018 to 1019 Pa s, beneath some regions of West Antarctica.Upper mantle viscosity variations of several orders of magnitude beneath West Antarctica have been quantified by modelling observed GPS uplift in specific settings. Nield et al.63. Ivins and Sammis64 formulated an approach for converting shear velocity anomalies to viscosity anomalies relative to a global reference 1-D viscosity model, using the observed scaling of shear velocity and density anomalies to infer the temperature scaling, and then using olivine diffusion creep (linear) flow laws to constrain the viscosity perturbation. Wu et al.65 proposed a similar method that uses experimentally determined temperature derivatives of shear velocity, including the effect of anelasticity. An alternative method is to estimate the temperature structure of the mantle from the shear velocity structure, and then use a composite rheology that computes the strain associated with both diffusion and dislocation creep mechanisms56. Since the dislocation creep rheology is non-linear, the calculated effective viscosity of the mantle will depend on the stress field. One key assumption in any of these methods is that the observed seismic velocity anomalies result entirely from temperature anomalies, whereas it is well known that a portion of the velocity perturbations result from  compositional anomalies66. Wu et al.65 adjust their viscosity structure by the percentage of the velocity anomalies thought to be caused by thermal variations, which in their study was estimated to be between 65% and 100%. An alternative approach would involve correcting for (poorly known) spatial variations in mantle composition.Existing constraints on mantle viscosity across Antarctica draw on our ability to measure the solid Earth response to known surface-load change. In regions where this is not possible, the three-dimensional shear velocity structure of the upper mantle beneath Antarctica can be used to estimate the lateral and depth variation of viscosity. Although there is no direct physical correspondence between shear velocity and viscosity, their variation in the upper mantle is largely controlled by temperature65, the estimated mantle viscosity beneath Antarctica, at depths of 100 and 250 km, varies by 2–3 orders of magnitude across the continent . It is difficult to clearly delineate the regions of the mantle dominated by linear and power-law viscosities because the rheologies depend on stress and the poorly constrained parameters of water content, grain size and activation volume153. Power-law rheologies can be introduced into GIA models using a composite rheology, where low stress portions of the model adopt a linear Maxwell rheology and high stress portions adopt a power law rheology assuming some transition stress156. Alternatively, strain is assumed to be the sum of diffusion and dislocation creep, as calculated using laboratory-based flow laws assuming parameters such as grain size and temperature157.The response of the Earth to changing loads has generally been described using a linear Maxwell viscoelastic rheology, with an instantaneous elastic response superposed on a longer-term Newtonian viscous relaxation159. These observations have resulted in increased interest in the use of Burgers rheology in GIA studies160. However, similar to the power-law case, there are few data to constrain these more complex models and most recent GIA models continue to use a Maxwell rheology32.Recent studies show that the transient relaxation following major earthquakes is generally best fit by a Burgers (biviscous) rheology with two characteristic relaxation times or effective viscosities67, and that the stability and dynamics of marine ice sheets (ice sheets which rest on ground below sea level) are sensitive to the depth of water at their grounding lines, i.e., the point where they begin to float70. Marine-terminating ice sheets such as Antarctica lose most of their mass following the flow of grounded ice across their grounding lines into floating ice shelves .It has long been recognised that the evolution of ice sheets is influenced by the geometry and deformation of the underlying solid Earthves Fig. . The iceeat Fig. . Howevereat Fig. . The bat72. A local decrease in water depth can enhance grounding of the ice shelf at ice rises, stabilising the ice sheet, while an increase in water depth can lead to ungrounding at the ice rise, enhancing flow across the grounding line of the ice sheet. Modelling studies of the AIS have shown that accelerated retreat and thinning and eventually large-scale collapse of marine sectors of the ice sheet can occur when the surrounding ice shelves break up73, but that if ice shelves are able to re-ground, this can have a stabilising effect on the ice sheet74. Furthermore, changes in bathymetry may have implications for ocean circulation and heat transfer under the ice shelves. The role of these and other ice shelf processes in controlling the dynamics of the grounded portion of the ice sheet are discussed in more detail in a companion paper by Smith et al. (manuscript submitted).Water depth changes can also influence the degree to which ice shelves are able to stabilise the ice sheet. Ice shelves play an important role in the stability and evolution of marine ice sheets by providing resistance to the flow of grounded ice across the grounding line. The spreading of ice shelves is inhibited by friction along their sides and base, particularly where the ice shelf becomes locally grounded on bumps or pinning points in the bathymetry, forming ice rises71. Recently, a series of more realistic coupled models have been developed that capture Antarctic ice sheet and ice shelf dynamics, global sea level and solid Earth deformation, and the interactions between these systems36. It has been shown that GIA-related sea-level and solid Earth changes, including changes to the slope of the underlying bed, alter the stress field of the ice sheet in a way that acts to dampen and slow past75 and future76 ice-sheet growth and retreat in Antarctica  model, which treats the asthenosphere as a time-lagged relaxation towards equilibrium83, or a model with flow in a viscous half-space below an elastic plate84. The Earth rheology models adopted in the newly-developed fully-coupled models described above75 capture more realistically the full multi-normal-mode response of the Earth to both ice and water loading, thus enabling the computation of gravitationally self-consistent variations of the sea surface, and Earth rotational effects. By solving the sea-level equation7, these studies also account for migrating shorelines, including migration into regions previously occupied by marine-based ice. Ice model simulations15 over the last deglaciation incorporating an ELRA bed model (black line) and the full sea-level coupling (blue and red lines) are shown in Fig. The strength of the feedback between GIA processes and ice dynamics depends on the rheology of the solid Earth. Models representing the palaeo86 often do not include bedrock deformation and sea-level changes. Simulated Antarctic ice volume changes under moderate future climate warming with fixed bed and sea surface are compared with results from a coupled model9 in which these surfaces are allowed to vary  viscoelastic uplift occurring on centennial or even decadal timescales47, with consequent implications for ice sheet evolution.When considering short-term change, ice-sheet models designed to simulate decadal to centennial-scale transient ice dynamics in response to future warmingary Fig. . The imp15. Note that, in a coupled modelling context, altering the Earth structure influences GIA predictions both directly by altering the timing and geometry of the Earth’s response to surface loading, and indirectly by changing the ice loading itself. The blue and red lines in Fig. 15 over the last 40 ka, during a Pliocene warm period, and into the future under moderate and strong climate warming scenarios (see figure caption for details) adopting two different radially-varying models of Earth viscosity and lithospheric thickness within the coupled model. One is a relatively high viscosity Earth model with a thick lithosphere (HV), similar to models adopted in global GIA studies, and the other is representative of the Earth structure beneath much of West Antarctica, having a thinned lithosphere and a zone of low viscosity down to 200 km in the upper mantle (LVZ). For each of the time periods considered in Fig. Recent coupled modelling studies have begun to quantify the sensitivity of predicted ice dynamics, GIA and crustal deformation in Antarctica, as well as the AIS’s contribution to past and future global sea-level change, to the adopted radial viscosity structure of the solid Earth73), West Antarctica collapses early on regardless of the choice of Earth model, and differences in the ice sheet evolution between the LVZ and HV simulations occur mostly in East Antarctica  and changes in dynamic topography over the last 3 Ma are postulated to have altered the stability of some sectors of East Antarctica90. However, since the inception of the AIS, glacially-driven erosion and sedimentation, and the accompanying isostatic response94, have been the main drivers of topographic change across Antarctica. This topographic change has, in turn, played a role in controlling the sensitivity of the ice sheet to climate forcing. This section focuses on the extent to which feedbacks between the ice sheet and the solid Earth have shaped the long-term evolution of both.The previous section describes the impact of ice-load-driven changes to the elevation of the solid Earth surface and gravity field on ice dynamics. Many ice-sheet modelling studies account for ice-driven isostatic adjustment95 and higher mean topography96. Dated offshore sediment packages are testament to the volume of material that has been removed from Antarctica since 34 Ma (ref. 97). Backstripping techniques can be used to unpick the history of progradation and isostatic subsidence, and hence reconstruct the palaeotopography of the continental shelf98. It is more challenging to reconstruct onshore palaeotopography due to the difficulty of determining the volume of material that has been removed, and the source of that material. By using assumed drainage pathways89 or process-based erosion modelling tuned to match offshore sediment volumes92, and accounting for the isostatic response to ice and ocean loading, as well as sediment erosion and deposition, bounds can be placed on the pre-glaciation topography of Antarctica99. Current palaeotopography reconstructions have been used to suggest that the West Antarctic Ice Sheet could have formed much earlier than previously thought, during the Eocene-Oligocene transition100. However, when two palaeotopography end members are used to define boundary conditions for ice-sheet growth under ‘cold’ mid-Miocene conditions, the difference in modelled ice volume is equivalent to 20 m sea level101, with the shallower topography leading to the growth of a larger ice sheet. It is clear that the configuration of the ice sheet is highly sensitive to the underlying topography, but the question remains as to how Antarctic topography has evolved, and the degree to which this evolution was coupled to ice-sheet change via glacial erosion and the resulting isostatic response.Prior to 34 Ma, Antarctica was characterised by a fluvial landscape102, then at 14 Ma there was a switch to cold, polar conditions103, which resulted in the establishment of a cold-based, less erosive ice sheet104. A range of hypotheses have been proposed to explain this transition to the current ‘icehouse’ world106, but it remains to be tested whether landscape evolution played a role, either through a rapid change in topographic boundary conditions or by causing the system to pass some internal threshold as large portions of the ice sheet became marine grounded.To understand the role of internal, i.e., non-climatic, processes in controlling long-term ice-sheet change it is necessary to consider feedbacks between ice dynamics, erosion, deposition, isostasy, and water depth change. Extending the coupled ice sheet - sea level modelling approach described above to account for landscape evolution processes would allow the testing of hypotheses that seek to explain the dramatic changes that took place as Antarctica evolved from a terrestrial to a marine ice sheet against a background of long-term cooling. Specifically, between 34 and 14 Ma the volume of the AIS fluctuated significantly2. The resulting isostatic response of onshore uplift and offshore subsidence will have focused erosion on the inner shelf92, leading to the development of a reverse bed slope that is too deep to support ice-sheet advance107 and is susceptible to unstable grounding line retreat68. These processes are conceptually simple, but they will have been subject to control by spatially variable sea-level change. As the mean topography of Antarctica decreased through erosion, the ice sheet will have become more sensitive to sea-level change108 and ocean forcing109. Applying appropriate boundary conditions to understand long-term ice-sheet change is challenging due to uncertainties associated with palaeotopography and the growth of the Northern Hemisphere ice sheets, but initial studies that have explored the impact of spatially variable sea-level change on the pre-Pleistocene ice sheet highlight the importance of accounting for peripheral bulge growth when considering ice110 and ocean111 dynamics, and the damping effects of ice sheet - sea level feedbacks in regions of weak mantle viscosity15. The impact of spatial variability in Earth rheology on coupled ice sheet - landscape evolution has not yet been investigated.Ice sheet-driven processes that will have had an impact on ice-sheet dynamics include progradation of the continental shelf and long wavelength erosion and deposition86. The details of how the ice sheet and the underlying topography reached this state are not known, but it is clear that feedbacks will have played a role as the fluvially- and tectonically-controlled topography of 34 Ma was eroded and warped into its current configuration94, and it has been suggested that future landscape evolution will render the ice sheet even more unstable20. Models that seek to explore the past and future long-term evolution of the AIS should account for evolution of the underlying topography due to both sediment redistribution and isostasy, ideally within a coupled framework. The development of such models will improve understanding of sediment transport pathways, permitting stronger conclusions to be drawn from provenance studies112. Better quantification of bedrock erosion rates and offshore sediment packages are important targets for improving our understanding of feedbacks between ice sheet and landscape change.Regions of the AIS that are currently located on over-deepened reverse bed slopes have been suggested to be fundamentally unstableThe previous section highlights a number of advances that are required to further our understanding of the long-term evolution of the AIS and the role of changing topography. These include the development of more sophisticated numerical models that consider feedbacks between ice dynamics, glacial erosion, isostasy, and global sea-level change. Such processes will also play a role in controlling contemporary and future ice dynamics but, as highlighted above, quantifying present change is exacerbated by the need to interpret geodetic observations in terms of the response to both past and present ice-sheet change. In this section we discuss several areas that should be prioritised as we seek to better quantify the GIA signal across Antarctica and hence understand the processes responsible for contemporary ice-sheet change.87. Characteristic relaxation times of a Maxwell fluid may be approximated by dividing the viscosity by the shear modulus (~7 × 1010 Pa for the upper mantle). The relaxation time of mantle material with a viscosity of 1019 Pa s is therefore only ~5 years, or 1–2 orders of magnitude faster than global averages. Actual relaxation times will be somewhat larger due to the presence of higher-viscosity layers within the rheological profile, but this calculation illustrates that, for regions underlain by low viscosity mantle, detailed knowledge of glacial load changes over the past decades to centuries is needed to quantify the GIA signal in regions of low viscosity113. For viscosities of 1018 Pa s or lower the viscous relaxation almost occurs contemporaneously with the surface load change and, by implication, the elastic deformation, with the viscous component being around an order of magnitude greater than the elastic component47. Such a rapid response is at odds with the traditional idea that the present-day GIA signal is related to post-LGM ice loss, although, given the much longer relaxation time of the lower mantle, some component of the present-day GIA signal is likely still associated with large-scale, post-LGM ice-sheet change.Characterisation of absolute mantle viscosity is preliminary. Beneath West Antarctica low viscosity mantle approaches isostatic equilibrium more quickly compared with global-average timescales. The consequence of this is that the present-day rate of adjustment will depend heavily on recent, and relatively localised, surface loading changes56, and these require quantification of mantle temperature, grain size, and water content, which have varying measurement uncertainties. Mantle temperatures, and hence viscosities, may be inferred from seismic velocity perturbations, but different approaches yield different results. When using a power-law rheology, the stress in the mantle prior to the change in surface load also plays a role in defining the viscosity, but this is often taken as zero114. Such stresses are expected to be largely a function of long-term mantle convection processes but could be more complex in the shallow mantle and asthenosphere due to ongoing isostatic relaxation or earthquake-related deformation.Across East Antarctica, spatial variations in Earth rheology are currently poorly constrained, as they are across all offshore regions, due to significant uncertainty in Earth structure resulting from the absence of seismic stations. Across all Antarctica, uncertainties also exist in relation to the rheological law that should be used to describe mantle deformation. Although traditional GIA models do not parameterise viscosity to be directly related to the physical properties of the mantle, physically-based approximations to mantle creep processes are being implemented in new models115, although there is evidence of a dynamic West Antarctic Ice Sheet119 and large changes in net accumulation during this period121. Progress can be made by studying the internal structure of the ice sheet to determine past flow patterns123, but at present, global and Antarctic-focused GIA models generally assume no ice-load change over the last 1–2 ka . Combined with low mantle viscosities in much of West Antarctica, this means that the predicted upper-mantle component of present-day deformation is likely erroneous. For vast sections of East Antarctica almost no data exist on past ice extent and retreat history125 . It is possible to extend data to earlier time-periods through the study of sediment from submerged offshore basins127. Opportunities exist for new absolute gravity128, InSAR129 and, especially, GNSS measurements at rock outcrops with large sections of East Antarctica sparsely or un-observed. Continuous measurements allow both increased precision but also, in regions of low mantle viscosity, probing of mantle properties by time-varying surface loading59. The largely unexploited horizontal deformation field promises new insights into local or regional-scale deformation patterns of both elastic and viscoelastic processes130, although a robust approach to remove tectonic plate motion and post-seismic deformation prevents continent-wide analyses at present132. A methodology is not yet available to precisely measure bedrock displacement under the present ice sheet or offshore, and this precludes model validation for these regions.Data with which to validate or constrain the GIA signal across Antarctica are also sparse Fig. . Less th43; such estimates are useful in separating competing conventional ‘forward model’ predictions of the GIA signal124. While possessing their own inter-solution variation due largely to differences in altimeter snow-densification corrections, inverse estimates tend to differ more from the spatial patterns of deformation predicted by current forward models . Given that there is just one real Earth, a standardised viscoelastic software framework that allows the consistent treatment of GIA and post-seismic viscoelastic deformation would have distinct advantages; it would pave the way for new insights into Earth’s rheology, provide a framework to advance Earth’s viscoelastic predictability, and help resolve ongoing debates regarding model robustness142.A practical problem for those wishing to employ viscoelastic models is the absence of open source software that includes state-of-the-art model physics. While open source software are available and widely used for purely elastic36. It has long been understood that the growth and decay of an ice sheet will alter the shape of the solid Earth and result in spatially variable sea-level change7. We are beginning to understand how those changes affect the dynamics of ice sheets71, but further work is needed to quantify feedbacks between ice sheet evolution and landscape evolution before we can fully explain the long-term evolution of the Antarctic Ice Sheet.Recent modelling advances have highlighted the importance of understanding the role of the solid Earth in moderating the response of the global ice sheets to past and future climate change71. This mechanism has the potential to slow or even halt grounding line retreat along a reverse bed slope, depending on the rate of solid Earth rebound. Such negative feedbacks have been modelled in association with ice loss across West Antarctica10, but crucially the rheology of the underlying mantle is poorly constrained, and so the strength of the feedback cannot currently be quantified.Coupled models that consider interactions between ice sheets and the solid Earth via dynamic modelling of both systems represent a new type of model that can provide insight into the extent, timing, and mechanisms of past ice sheet evolution and sea-level change. Such models also provide a new tool for understanding the link between past climate and global sea-level change. Coupled models have been used to demonstrate that the solid Earth response to ice loss, and the accompanying fall in sea level, will have a stabilising effect on grounding line dynamics56, with low viscosity mantle inferred to lie beneath West Antarctica. The short mantle relaxation time associated with such regions means that estimates of the contemporary GIA signal derived by forward modelling will be biased if they do not account for ice-sheet change over the last few millennia, with implications for estimates of current ice mass balance8. Improved quantification of late Holocene ice history, as well as tighter constraints on mantle rheology, are needed to reduce uncertainty on the contemporary GIA signal across Antarctica. Low viscosity mantle will also enhance the strength of the stabilising effect of GIA on grounding line dynamics, highlighting the importance of considering such feedbacks when modelling the future evolution of the Antarctic Ice Sheet. A number of studies have successfully incorporated 3-D Earth structure into GIA models beneath modern ice sheets144, and it is now apparent that coupled modelling and inclusion of 3-D Earth structure should both be considered when modelling solid Earth-cryosphere feedbacks15. Such models are being developed87 but further interdisciplinary work, combining modelling and observational approaches, is needed to calibrate such models and better understand controls on the evolution of the Antarctic Ice Sheet.Interdisciplinary approaches to determining mantle rheology reveal large differences between East and West AntarcticaPeer Review File