However, their collisions with the seabed crush and recycle vast benthic communities, costing an estimated 4 × 104 t C yr−1. We calculate that giant iceberg formation (ice shelf disintegration) has a net potential of approximately 106 t C yr−1 sequestration benefits as well as more widely known negative impacts.Sea ice, including icebergs, has a complex relationship with the carbon held within animals (blue carbon) in the polar regions. Sea-ice losses around West Antarctica's continental shelf generate longer phytoplankton blooms but also make it a hotspot for coastal iceberg disturbance. This matters because in polar regions ice scour limits blue carbon storage ecosystem services, which work as a powerful negative feedback on climate change . This resets benthic biota succession  and also fertilizes the ocean with nutrients, generating phytoplankton blooms, which cascade carbon capture into seabed storage and burial by benthos. Small icebergs scour coastal shallows, whereas giant icebergs ground deeper, offshore. Significant benthic communities establish where ice shelves have disintegrated , and rapidly grow to accumulate blue carbon storage. When 5000 kmThis article is part of the theme issue ‘The marine system of the West Antarctic Peninsula: status and strategy for progress in a region of rapid change’. Over the last four decades, the seasonal sea-ice maximum area and duration have decreased considerably over the Arctic and West Antarctica . By cont2) are not uncommon, currently numbering 47 larger than 30 km2 in March 2017, including six that exceeded 1000 km2 in area . Duprat et al. [2, the passage and fate of which have also captured widespread concern, from collisions with rich life on the seabed or anthropogenic installations  than current ice shelf losses, which are producing giant tabular icebergs . The rapt et al.  were ablt et al. . The chaoil rigs ) to indiThe positive feedback on climate change and potential for serious societal impacts of increased ice shelf losses are considerable but these are by no means the only important effects. The continental shelves part-covered by ice shelves are biologically unproductive but their collapse may open up vast new bays to generate new primary productivity , secondaHere we consider recent progress in measuring and understanding more general relationships between sea ice, icebergs and blue carbon on the Southern Ocean's seabed. Our aim is to better understand what we know and do not know about sea ice and iceberg impacts on blue carbon around the Southern Ocean. From published papers we analysed recent (2014–2017) benthic blue carbon data from (i) the relationship between fast-ice duration and the frequency of ice scouring in shallow waters at Rothera and Carlini stations on the West Antarctic Peninsula (WAP), (ii) iceberg impacted and non-impacted sites on South Georgia's north shelf and (iii) how correlation changes with depth between sea ice and benthic blue carbon around Marguerite Bay. Fourthly, we also analysed US National Ice Center circumpolar data of giant iceberg status , grounding hotspots and tracks to assess the trade-offs of giant iceberg formation with respect to effects on benthic ecosystem services of blue carbon . Finally, we used these findings and literature to construct a polar projection map (based on 3° × 3° grid cells) of data illustrating benthic blue carbon change where we completely lack data.2.in situ manually operated or remote image capture. Around Antarctica these have proved to be complex, in terms of extent, geography and drivers of change. Overall, there have been marginal increases in Southern Ocean sea ice to date, but there have also been considerable losses around West Antarctic continental shelves and even larger gains, mainly over deeper water, in the Weddell and Ross Seas [a) to ground-truth coastal remote-sensed data, and even fewer are monitoring benthos responses . Circumpolar satellite-derived data have revealed that the WAP particularly is a hotspot for sea-ice losses. Iceberg scour can be very frequent in the shallows if the area above it is free of fast ice (because there is less obstruction to wind-driven travel) [figure 1b) shows even higher rates than at Rothera. As at Rothera, the site (at Carlini) closest to a glacier terminus showed some very high scouring rates, suggesting that local iceberg production as well as fast ice are important factors for seabed disturbance. We suggest that climate forcing has made the WAP a hotspot for coastal iceberg scouring. Results from the longest running (2003 onwards) iceberg scour-monitoring experiment near Rothera research station [4 tonnes of carbon per year [a), and sites 11 and 13 (which showed no evidence of having been scoured in the remote-sensing era). We found no significant difference between scoured and unscoured samples with respect to blue carbon stock mass , suggesting recovery is at least possible within approximately 7 years even at a 200 m depth on the South Georgia shelf. Part of the problem and complexity of analyses of iceberg blue carbon impact, at least coastally, is the confounding effects of sea ice and phytoplankton blooms. In shallow coastal waters, sea-ice losses drive both blue carbon decreases by increased scour probability, and increases through longer phytoplankton blooms sustaining benthic growth. To examine the relationship between sea ice and benthic biomass, we reanalysed data presented in Barnes [b). We found that, in the top approximately 10 m, there was a positive correlation between blue carbon and sea-ice duration . We interpret this as longer sea-ice durations reduce ice scour chances, which outweighs slower growth from the shorter phytoplankton blooms. There was no relationship between benthic blue carbon and sea-ice duration from 10–40 m depth , which is probably explained by the reduction in ice scour chances due to longer sea-ice durations approximately balancing slower growth from the shorter phytoplankton blooms. At typical continental shelf depths (200–500 m), there were negative correlations between blue carbon and sea-ice duration . Ice scour chances are so rare in deep water (because few icebergs are big enough to have keels hundreds of metres deep) that longer sea-ice durations have little impact on this, so the main effect is slower growth from shorter phytoplankton blooms. Between 40 and 200 m is a zone of very little benthic sampling, perhaps often too shallow or geographically close to shallows for safe ship operation but too deep for self-contained underwater breathing apparatus (SCUBA)—a data-poor zone that we need to address (b).The Southern Ocean sea surface features striking contrasts in albedo between open water and sea ice, and as such remote sensing is an ideal tool to detect trends in sea-ice spatial and temporal extent . Satellioss Seas . The conoss Seas ,21, and oss Seas ,22. Reseoss Seas ,24. Neaross Seas ,9,25. Suoss Seas . To dateper year  collecte zone A, b. We int zone B, b, which  zone C, b. Ice sc address b.Figure3.(a)Away from the shallows, much larger icebergs (by area) are easier to track from genesis to eventual exit from the Southern Ocean, yet much less is known of the subsurface impacts of these giants on blue carbon, although recent studies have highlighted their importance in their pelagic fertilization roles. Here, we discuss and quantitatively estimate the importance of blue carbon impacts of Southern Ocean icebergs.a). Giant icebergs leave a trail of trace nutrient enrichment, enhancing phytoplankton blooms in their path [Giant icebergs mainly originate from ice shelf disintegration and can remain fasted in sea ice for years in deep coastal water, roam on or off the continental shelf or run aground a. Giant eir path . This pheir path , and inceir path . Beyond eir path ,9. Signieir path , contrib4 km2 of ice shelf have collapsed, to open up continental shelf waters to primary production and secondary benthic carbon capture and storage [6 tonnes of carbon per year  within the skeletal matrix of the benthos) [et al. [2 iceberg from Larsen C could generate 7 × 104 tonnes of Cimmobilized per year . However, four years after the Larsen A and B ice shelf collapses, Fillinger et al. [et al. [2 iceberg creation may actually generate more than 105 tonnes of Cimmobilized per year . Such productivity could be increased because of (a) additional climate-forced sea-ice losses in some regions (e.g. Antarctic Peninsula) [5 tonnes of Cimmobilized per year estimate is more likely to be approximately 106 tonnes of Cimmobilized per year  as a conservative negative feedback consequence of giant iceberg formation along the Antarctic Peninsula area of the polar regions alone.Giant icebergs have another very important and different impact on enhancing blue carbon storage—their calving involves break-off from an ice shelf. As in current time all ice shelves occur over continental shelf, their break-up to form giant icebergs creates new areas of open water above potentially highly productive seabed. In the last few decades more than 10 storage . New ben [et al.  estimater et al.  reported [et al.  estimatek et al.  estimate [et al.  growth rninsula) . The magninsula) . Thus wek et al.  estimate [et al.  growth r [et al.  iceberg (b)b). When ice shelves collapse to form giant icebergs, the buttressing of land-based ice sheets is reduced [2 absorption, which could reduce benthic carbon immobilization and sequestration through ocean acidification [Small icebergs can halve incremental growth of benthic communities and thus the potential for carbon storage, as happened between 2007 and 2009 in Ryder Bay near Rothera research station (WAP) due to ice scour [ reduced ,29. Thus reduced ,9 or a r reduced . More opfication . This oc2) and B22A (1056 km2)) were grounded in 2016/2017 in the shelf seabed. Such grounding can be a prolonged occurrence; for example, A23A has been grounded for more than three decades, since 1986. The chances and location of any particular iceberg scouring depend not only on size (most 2016/2017 giants were 100–500 km2 in area) but also on shape, calving location and current and wind direction and strength, as well as stochastic factors such as which eddies entrain it. Larger icebergs can be tracked such that their motion and fate can be followed; for example, in early 2017 there were four icebergs (8.5%) larger than 30 km2 in area grounded, and thus scouring blue carbon. The extent of this could be ground-truthed using ship-based multibeam, swath bathymetry data collection. Benthic imaging and sampling could then quantify blue carbon losses and, as such scours can be many kilometres long, they are likely to crush and recycle hundreds–thousands of tonnes of blue carbon standing stock. For example, benthos in the vicinity of the grounded scouring giant icebergs A23A and B22A hold 8–16 tonnes of standing stock of carbon per square kilometre [a). The Lazarev, Cosmonauts, Davis and East Amundsen Seas were hotspots of deep shelf grounding, whereas there was little deep-water scouring in the Bellingshausen and Ross Seas. The paths of giant icebergs  also showed distinct patterns, firstly following the anticlockwise coastal current, then leaving the shelf from the Eastern Weddell and Ross Seas and being dispersed by the Antarctic Circumpolar Current (b).The major negative impact of giant icebergs on benthic blue carbon is if they collide with the seabed and scour. Giant icebergs would only ground and scour in deeper water, which may on average only occur once every few hundred years at any one location . Howeverd B22A 106 km2) weilometre . Imageryd B22A 106 km2) weilometre ). We fou Current b.Figure3 tonnes of benthic carbon per year could be recycled rather than stored and buried. To date, such losses have not been quantified in absolute frequencies at depths relevant to giant icebergs. Even scouring has some positive blue carbon aspects in terms of potential to bury benthos in sediment berms, resuspend food material (enhancing growth in nearby benthos) and open up space for faster-growing benthic pioneer species [Even if only one-fifth of giant icebergs of the size of A23A/B22A were grounded, more than 5 × 10 species .(c)2 would typically be expected to generate 2 × 106 tonnes of Cimmobilized per year in total (estimate from §3a (×2 for two icebergs)) and ‘cost’ up to 16 tonnes for each km2 scoured (estimate from §3b) [2 which becomes recycled) = up to 4 × 104 tonnes of Cimmobilized. This cost is likely to be highly overstated because (a) the largest giants can be grounded in the same location for many years  and (b) it assumes icebergs become grounded at random locations on the continental shelf. It is more likely that there will be hotspots of scour on rises or sills, such as the continental shelf break, which will be hit more often and have less time to recover (and thus less stored carbon is recycled). We calculate that the net carbon gain of giant icebergs is at least 2 × 106 − 4 × 104 = 1.9 × 106 tonnes of Cimmobilized per year. This is equivalent of the carbon exhaust output of 400 000 cars per year, or nearly four million cars in terms of total zoobenthic carbon  ,32 at a obilized ). This iThus, there is a strong argument that giant iceberg formation has a net positive climate benefit by growing important negative feedbacks ,18. Of c(d)−2 yr−1 in the shallows [−2 yr−1 at 100–300 m and approximately 1.5 g C m−2 yr−1 below 300 m depth [−2 yr−1 is used, scaled up to 3 t C km−2 yr−1 and the same amount of carbonate (which varies enormously and of which only 12% is carbon), this gives a total of approximately 3.4 t C km−2 yr−1. If the shallows below 100 m are of similar value, then 3.4 × 2 = 6.8 t C km−2 yr−1. Multiplied by the area of Antarctica's continental shelf (4.4 × 106 km2), it is approximately 30 Mt C yr−1. This figure is considerably less than 1% of annual anthropogenic emission; however, production has doubled (at least in some biota) since the 1990s around many Antarctic shelves due to sea-ice changes [−1. This has to be further increased to account for ice shelf losses opening up large new areas of benthic production, estimated to generate as much as 10 Mt C yr−1 [2 iceberg calved. Each of these increases to the initial estimate adds considerable uncertainty, but may be approximately 80 Mt C yr−1  or held in pelagic or forest biomass, which has much more time and space to be broken down on death and recycled in the microbial loop. Increases in sea temperature could yet increase zoobenthic carbon storage further [Zoobenthic production around Antarctic continental shelves varies considerably with geography and depth, but can be approximately 13 g C mshallows , through m depth ,35. Prod m depth ,9. Most  m depth ,35. If at C yr−1 . Yet suct C yr−1 . The curf losses  × 2 for f losses  than thof losses ). Doubli further , but thi further .4.Thus we conclude that sea ice interacts strongly with benthic blue carbon ecosystem services, partly through iceberg activity.et al. [It is clear that sea ice and icebergs have complex influences on benthos and their blue carbon ecosystem services—and thus on the power of feedbacks on climate change. Pooling literature data sources ,26,27,32et al.  have shoet al. . The rec6 tonnes of Cimmobilized per year (around the Antarctic Peninsula), which matches similarly sized gains in the same region driven by sea-ice losses [Our calculations suggest that giant icebergs alone contribute a net gain of 10e losses . The cone losses ), which 6 tonnes for increases from ice shelf losses  in addition to 107 tonnes for increases driven by sea-ice losses [et al. [6 km2 of Kerguelen Plateau) may also be increasing blue carbon storage—and when quantified could prove the single largest negative feedback (on climate change). All three  of these sinks are expected to increase in response to climate change, in contrast to shrinking forests  and become progressively less efficient at ocean sponging [On the scale of terrestrial forests and oceanic water absorption, Antarctic benthos carbon cycling is small, at just 10e losses . Recentl [et al.  demonstrsponging . Sub-Arcsponging  but of s


We interpret these lags as indicating that Great Basin climate is sensitive to and indirectly forced by changes in the cryosphere, as evidenced by fast and strong linkages to global ice volume and Arctic paleoclimate indicators. Mid-latitude atmospheric circulation was likely impacted by a northward shifted storm track and higher pressure over the region arising from decreased sea ice and snow cover. Because anthropogenic warming is expected to reduce northern hemisphere snow and ice cover, continued increase in atmospheric greenhouse gases is likely to result in warming and drying over coming centuries that will amplify a warming trend that began ~2400 years ago.Although Great Basin paleoclimate history has been examined for more than a century, the orbital-scale paleoclimate forcings remain poorly understood. Here we show – by a detailed phasing analysis of a well-dated stalagmite δ The orbital paradigm suggests that Northern Hemisphere Summer Insolation on June 21 at 65°N latitude (hereafter NHSI) controls global climate via variations in ice sheet extent, which in turn exert responses in temperature and atmospheric and oceanic circulation. However, in many continental regions the role of orbital variations in paleoclimate remains poorly known due to the difficulty of obtaining long absolutely dated paleoclimate records. For example, controls on Great Basin paleoclimate have been debated since the classic 19th century monographs on Pluvial Lakes Lahontan and Bonneville7 . We use insolation sub-stage (ISS) nomenclature, which labels peaks and troughs in NHSI analogously to extrema in Marine Isotope Stage (MIS) numbering wherein, for example, the last glacial maximum is ISS 2.0 and the penultimate interglacial is ISS 5.5  component of cave infiltration12, and so we interpret the speleothem δ18Oivc variations in the past as a proxy for wintertime temperature and moisture source variations, and our vadose-zone cave locations are unaffected by changes in groundwater levels.Following ref.18Oivc values for the Leviathan chronology show pronounced orbital-scale variability that exhibit clear correlative peaks to NHSI  to NHSI, and the lags exceed the 2σ age uncertainties determined from time series age model uncertainties  between forcing and paleoclimate in the Great Basin  to 135,100 yr B2k (δ18Oivc = −13.35‰), lagging NHSI by 4400 years  at 126,400 yr was delayed 1000 years behind insolation. We also show that the onset of the last interglacial lagged insolation: Termination II (TII), as defined by the mid-point δ18Oivc value of −12.03‰, dates to 131,100 yr B2k  of 9800 years using the TII δ18Oivc threshold , providing a last interglacial duration (sensu lato) of 10,800 years. These durations are close to those predicted for Milankovitch forcing, being approximately ½ the 19,000 and 23,000 year precession cycles, respectively. The timing of glacial and stadial periods also fits well with precessional-scale orbital forcing. We lack continuous data for the Last Glacial Maximum and are thus unable to precisely estimate phase relationships, but a likely timing based on the filtered δ18Oivc of 19,200 yr B2K suggests a 4300 year lag behind NHSI , filtered Leviathan δ18Oivc peaked at 7600 yr B2K, 3500 years after NHSI at 11,100 yr B2K. Unfiltered δ18Oivc peaks earlier (8380 yr B2K), giving a minimum lag of ~2700 years. In an unusual departure compared to past behavior where δ18Oivc tracked insolation closely, late Holocene δ18Oivc has been increasing since ~2400 yr B2K while insolation has decreased.The filtered Great Basin δs Figure , and could Table . We also18Oivc lag behind NHSI suggests that the Great Basin climate response is instead indirect and mediated through one or more fast-response  components of the climate system. We investigate several pathways by which insolation forcing is linked to Great Basin climate: global ice volume13, northern hemisphere temperature16, regional sea surface temperature (SST) in the California Current system off WNA17, Arctic sea-ice variations18, radiative forcing from greenhouse gases20, and high latitude temperatures from Greenland21 and Antarctic ice core data22.If insolation were directly forcing Great Basin paleoclimate via local temperature, we would expect to see nearly synchronous speleothem climate variations in response to changing NHSI. The millennial-scale δ18O stack13, a proxy for global ice volume, salinity, and deep water temperature  also followed NHSI, demonstrating that precessional-scale insolation forced deglaciation in WNA over TII. The peak cold conditions during the ISS 5.4 δ18Oivc minimum align closely with increased global ice volume (a benthic δ18O maximum). We also observe a strong similarity between the Leviathan chronology and an estimate of northern hemisphere temperature (40–80°N) derived from an ice sheet model and the LR04 stack16. In contrast, we observe several major discrepancies with California Current alkenone SST (core ODP-1012). We conclude that the Leviathan chronology’s δ18Oivc variations are strongly linked to global climate evolution on glacial/interglacial timescales, but are disconnected from regional ocean temperatures off WNA, at least on the orbital scale.The Leviathan chronology shares strong similarities with the global benthic foraminifera δ18O ice core record (GICC05 extended chronology)21. Peak cold conditions in ISS 5.4 were synchronous in both locations, and both records show a small warming ‘shoulder’ at ~114,600 yr B2k , and decreasing Holocene δ18Oivc. How could climate be so apparently cold in the Great Basin in ISS 5.4 when global ice volume was small and northern hemisphere temperatures were warm? Rising greenhouse gas concentrations and reduced North American ice sheet extent , which characterizes ostracod density in Arctic ocean sediments18, which in turn were forced by sea-ice mediated productivity changes. The APIo shows high productivity during high NHSI, and vice versa for low NHSI. A major productivity decrease during ISS 5.4 suggests a glacial-like sea-ice expansion despite intermediate global temperatures and ice volume, and is supported by cold climate conditions in the Siberian Arctic26 at this time when permafrost ice wedge growth was vigorous. A decrease in high latitude temperature27 and expanded snow fields28 coincident with the ISS 5.4 insolation minimum may have driven sea ice expansion and amplified Arctic cooling, and both cryosphere indicators exhibit a fast response to orbital forcing. The low ISS 5.4 δ18Oivc in Leviathan could be interpreted as forced by an increase in Arctic sea ice extent and high-latitude snow fields driven by the period’s lowest temperatures and summer insolation. Such a contention is supported by the strong influence between NHSI and Arctic summer sea ice extent (and weak influence from greenhouse gas forcing) in modeling studies29. Stadial conditions during ISS 4.0 and 2.0 were associated with a cessation of growth in stalagmite LC-1 from possible periglacial conditions overlying the high-altitude (2400 m) Leviathan Cave8, and both hiatuses also coincided with pronounced Arctic productivity minima and high sea-ice extent  and pollen evidence for decreasing Holocene summer temperatures in the Western Canadian Arctic31  decrease in δc31 Fig. . An incr41. Modern observations suggest that the impact of increasing snow cover may have resulted in decreased atmospheric pressure and increases in storm track intensity and cyclogenesis in the eastern North Pacific Ocean42. Such forcings would result in stronger storms that deliver greater winter precipitation to the Great Basin. In contrast, a reduction in modelled sea ice extent resulted in pronounced winter precipitation deficits, by up to 50%, in the Great Basin and Western U.S.41. The reduced sea ice extent increased the 500-mb geopotential height, diverting storms away from the southwest to higher latitudes, resulting in a northward shift of the mid latitude storm track43. During winters following sea ice lows from the previous summer, the atmosphere may ‘remember’ the summer sea ice extent because of a lowered mid- to high-latitude temperature gradient, which weakens the Aleutian Low44 and results in fewer and less intense storms reaching the Great Basin. A strengthening of the subtropical highs arising from decreased sea ice and snow cover extents43 also results in fewer storms impacting the southwestern United States. This atmospheric configuration typically produces high pressure over the Great Basin and inhibits delivery of Pacific moisture to the Great Basin45 today, and we suggest that similar climate dynamics applied in the past. Similarly, increased snow cover extent results in cooling via the snow-albedo feedback46, and advection of cold high-latitude moisture to the Great Basin during periods of greater sea-ice and snow-field extent is a plausible explanation for the ISS 5.4 and Holocene cooling in the absence of significant ice sheet extent.Because Great Basin climates lagged insolation even during periods of low global ice volume , it is difficult to attribute forcing solely to ice sheet effects on northern hemisphere paleoclimate. Changes in both high latitude sea ice and snow cover are likely to influence the mid latitudes via atmospheric teleconnections27 driven by changes in the meridional temperature gradient. Our data strongly support the hypothesis that, on orbital time scales, reduction in sea-ice extent is associated with higher temperature and less precipitation over the Great Basin. Arctic sea ice extent likely exerts control on mid-latitude atmospheric circulation via a weakening of the jet stream associated with a decreased latitudinal temperature gradient.On long orbital timescales, similar relationships may also be operative: changes in atmospheric circulation as a result of insolation changes are largely mediated through sea ice and positive feedbacks in the Arctic18Oivc at 38,500 yr B2k leads June 21 insolation by 900 years), it lags peaks in April 21 and May 21 insolation which arrive at 39,000 to 40,000 yr B2k, respectively. July 21 and August 21 insolation would reduce the lags at other time intervals. Clearly, reducing the complex Earth climate system to a single month of insolation is an oversimplification, a fact noted by previous workers10. In addition to variable seasonal sensitivity, the variable phase lag may also arise from multiple feedback processes in the climate system, each of which may have different time-variable response times to insolation forcing that depends on changing boundary conditions, e.g., ice-sheet size, the carbon cycle, Arctic Ocean temperatures, and the state of the biosphere. To test for this possibility, we find a weak but apparent negative correlation  between Leviathan δ18Oivc lags and the correlative NHSI value  in the δ18Oivc lag times. We would expect to see shorter lag times during pluvial climates if effective moisture variations in the Great Basin were driving faster infiltration rates into the cave systems, or if the δ18Oivc values were influenced by local pluvial lakes. That we do not see such effects is likely due to the rapid response time of drips into the shallow vadose zone caves and a lack of any significant local control on precipitation δ18O values. Lag times in phreatic zone caves, such as Devils Hole, however, may have been influenced by such local hydroclimatic variations.The large and variable phase lag (−900 to 6600 yr) appears to be a robust feature of our data, as the lags are considerably larger than the age model uncertainties Table . A possie Figure : the lag47. First, maximum expansion of mountain glaciers during the LGM was reached around 20,000 yr BP48, some four thousand years after minimum NHSI at 24,000 yr BP. Second, the formation of the iconic Great Basin pluvial lakes typically reached highest levels around 17,000 to 18,000 yr B2k associated with Heinrich stadial 150, and which also lagged insolation. Both observations suggest that lake and glacier hydroclimatic indicators lagged insolation to similar degrees as Leviathan δ18Oivc record, but the lake and glacier records were also associated with abrupt millennial-scale forcing associated with Heinrich stadials and the Bolling/Allerod interstadial49.Our observations also have implications for broader Great Basin paleoclimate and the interpretation of the phreatic zone Devils Hole calcite record18Oivc record (see Methods) in Devils Hole is significantly smaller , and the Holocene δ18Oivc peaks at ca. 5000 yr B2K, compared to 7600 yr B2K in the filtered Leviathan δ18Oivc and 11,100 yr B2K for insolation. Further support for dispersion of the δ18O signal in the aquifer is that the last interglacial optimum (MIS 5e) duration is half again as long  in Devils Hole as in surface climate. The timing of lags in Devils Hole may also be related to changes in hydroclimate via effective recharge rates in the phreatic aquifer, a possibility that merits further investigation. We conclude that Devils Hole δ18O represents a delayed and dampened response to surface climate variations; ages and durations of isotopic anomalies are thus minima and maxima, respectively.Finally, our vadose-zone cave climate record constrains atmospheric circulation changes that are later recorded in phreatic zone calcite at Devils Hole. The amplitude of the ice-volume corrected δr Figure  than Lev51 in the 21st century as it likely did during the mid-Holocene52. Such a sea ice reduction will perturb wintertime atmospheric circulation in the Great Basin, leading to dynamic warming and drying of winters similar to that observed in the instrumental record45. Additionally, human-caused warming will amplify the warming trend that began 2400 years ago, with mean annual temperature increases at our study area of 3–4 °C by 2080 under a moderate (RCP 4.5) warming scenario53. Model results indicate Great Basin winter precipitation decreases of up to 50% as a result of a northward migration of the winter jet stream45. Considering that the peak warming interval in the Great Basin of the early to mid-Holocene ‘altithermal’54 was one of intensely hot and arid conditions and the region was only able to sustain sparse human populations, the future of Nevada climate looks to be one of increasing warmth and drought. Without aggressive action to decrease the rate of greenhouse gas concentrations in the atmosphere, the future of the Great Basin is likely to be one of heating and drying.Our conclusion of high sensitivity of Great Basin paleoclimate to the state of the Arctic cryosphere has important implications for future climate projections. The modern climate is characterized by rapidly rising greenhouse gas concentrations, and the Arctic is expected to rapidly transition to sea-ice-free summers8. The time series of the vadose zone stalagmite δ18O values were adjusted for the change in δ18O of the ocean (denoted δ18Oivc) using the reconstruction of55, and for differences in the isotopic composition of precipitation falling at Lehman and Pinnacle Caves  relative to Leviathan Cave following8. This adjustment results in a small change in δ18O values that does not affect the shape or timing of the composite isotope curve. The Leviathan chronology δ18Oivc record presented here includes a correction for the age of nine stable isotope samples (and the derivative interpolated time series) beneath a hiatus at 290.5 mm. This changes the age of the last isotope sample (291 mm) before the hiatus to 73,647 from 69,622 yr B2k. The age correction does not change any interpretations of this or previous work. We estimated age uncertainties for individual stalagmite age models using the COPRA software56, and interpolated age uncertainties at the 95‰ level for each climate event in Table Ages on Nevada stalagmites were determined on a ThermoElectron Neptune multi-collector inductively coupled plasma mass spectrometer at the University of New Mexico Radiogenic Isotope Laboratory, and stable isotope values were determined on a Kiel IV automated carbonate preparation device at the Las Vegas Isotope Science laboratory at UNLV directly coupled to a ThermoElectron Delta V stable isotope ratio mass spectrometer in dual inlet mode18O values by changes in δ18O of the ocean that were scaled by a ratio of 0.4551, which is the ratio of the standard deviations of the Leviathan and DH2-D records, respectively. This ice volume correction more accurately reflects how a given precipitation isotopic signal is filtered and dampened via transit through the aquifer feedings Devils Hole better than the undampened correction47. Because of our different ice volume correction, the timings and amplitudes of DH2-D δ18Oivc anomalies are slightly different than previously reported.Because the DH2-D record is significantly dampened relative to the Leviathan signal due to aquifer mixing and dispersion  with zero phase lag using the filtfilt function in Matlab. The low-pass filters were then used to constrain timings of isotopic events based on a peak and trough detection algorithm matched by correlation to June 21st 65°N insolation for various insolation substages  of the Berger 1978 solution11. The reason for using filters instead of raw data is to minimize the possible effects of noise on the choice of start and end points to provide more objective and reproducible results, though the latter approach may also be valid. Relative timings of insolation sub-stages were determined by differencing correlative ages. Phasing diagrams show the timing differences between the forcing (insolation) and the climate response in the Leviathan and DH2-D filtered records.We then fit low-pass precessional-scaleSupplementary Information


Microbiota and fauna were generally similar to other sites in the OT, although with some different characteristics. In terms of microbiota, the H2-rich vent fluids in Neuschwanstein chimney resulted in the dominance of hydrogenotrophic chemolithoautotrophs such as Thioreductor and Desulfobacterium. For fauna, the dominance of the deep-sea mussel Bathymodiolus aduloides is surprising given other nearby vent sites are usually dominated by B. platifrons and/or B. japonicus, and a sponge field in the periphery dominated by Poecilosclerida is unusual for OT vents. Our insights from the Yokosuka site implies that although the distribution of animal species may be linked to depth, the constraint is perhaps not water pressure and resulting chemical properties of the vent fluid but instead physical properties of the surrounding seawater. The potential significance of these preliminary results and prospect for future research on this unique site are discussed.Since the initial discovery of hydrothermal vents in 1977, these ‘extreme’ chemosynthetic systems have been a focus of interdisciplinary research. The Okinawa Trough (OT), located in the semi-enclosed East China Sea between the Eurasian continent and the Ryukyu arc, hosts more than 20 known vent sites but all within a relatively narrow depth range (600–1880 m). Depth is a significant factor in determining fluid temperature and chemistry, as well as biological composition. However, due to the narrow depth range of known sites, the actual influence of depth here has been poorly resolved. Here, the Yokosuka site (2190 m), the first OT vent exceeding 2000 m depth is reported. A highly active hydrothermal vent site centred around four active vent chimneys reaching 364°C in temperature, it is the hottest in the OT. Notable Cl depletion (130 mM) and both high H Af. Afb). Sts depth g. This ant site) . The occnt site) ). We the‘Shisa’ , but no sampling was carried out. Carrying on the same direction past Shisa, we at last rediscovered Neuschwanstein chimney and commenced hydrothermal fluid sampling (maximum temperature 356.9°C). Although we further travelled eastward up to 400 m away from the four discovered chimneys, the seafloor was consistently rocky with occasional sponges after leaving Neuschwanstein and no further signs of hydrothermal activity were found.For the final dive at the Yokosuka site (no. 719), the ROV landed on sediment-covered seafloor southwest of Neuschwanstein. After travelling for about 100 m northeast, we encountered sediments with brown discoloured mat followed shortly by white microbial mat and broken sulfide pieces. Continuing along the same direction, we soon found a large mound-like chimney that only had concentrated venting at a bulge-like flange structure at the side. We gave it a name, 3.2.Concentrations of each chemical species dissolved in the fluid samples are shown in the magnesium diagram  and all Mg diagrams of each chemical species did not display a single mixing line from Mg-rich ambient seawater to Mg-depleted endmember fluid composition , suggest2O in the host rock. This seems to be consistent with basaltic andesite and its low K2O content (0.9%) exhibited by the Yaeyama Knoll body, which is probably identical to the subseafloor host rock of the Yokosuka site.The endmember K/Cl value, which is a representation of the K content in the primitive upwelling fluid without influence from the phase separation, is significantly lower in the Yokosuka site fluids (0.055) than fluids of other OT sites (greater than 0.09) ,46. As t2S, than the other two chimneys , potentially provide large yields of bioavailable energy for hydrogenotrophic and methanotrophic metabolisms in the mixing zone [1/C2 ratio) ranges 75–140, lower than typical range in the OT fields .Figure4 in sediment-associated hydrothermal sites, including the Yokosuka site, is thought to result from the thermal decomposition of sedimentary organic matter and microbial methanogenesis in the anoxic sediment [δ13C–CH4 value of −25.7‰ of the Yokosuka fluid is the uppermost value in the δ13C–CH4 range observed in the sediment-associated sites (−58 to −25‰) and suggests a predominance of thermogenic CH4 rather than biogenic one. The C1/C2 ratio supports the thermogenic origin of methane. As abundant H2 exceeding 1 mM is only seen with ultramafic-rock- and fresh basalt-hosted hydrothermal activities [2 and CH4 concentrations to as low as 0.1 mM in the subseafloor fluid reservoir  and H2 (−356‰) are in accordance with the hydrogen isotope equilibrium with H2O having δ2H–H2O of +0‰ at the endmember fluid temperature of 364°C [δ2H–H2 value in the Neuschwanstein chimney fluid, containing significant sulfate (no. 719-W1), may be attributable to microbial H2 oxidation after fluid cooling in the sampler bottle. Previously, hydrogenotrophic sulfate reducers, in particular, were reported to decrease the δ2H–H2 value approximately below −600‰ [The abundance of CHsediment ,74. The tivities ,68,72, teservoir . As therng fluid , thermogof 364°C . The lowow −600‰ ,77. Inde2 concentration of Yokosuka site fluids (approx. 300 mM) was as high as vent fluids from other SOT hydrothermal sites . Inputs of thermogenic CO2 with δ13C–CO2 of approximately 25‰ after branching of the upwelling fluid into each chimney, in addition to the primitive CO2 having δ13C–CO2 of −6 to 0‰ [δ13C–CO2 variation in the Yokosuka site. The isotope effect on phase separation is another possible source of this variation, but it seems inconsistent with small variations in isotope ratios exhibited by other volatiles. Microbial production/consumption is not considered to be an attributable source due to the lack of carbon species comparable with CO2, from a stoichiometric viewpoint.The COal sites . Althouges e.g. ,78), the, the2 co−6 to 0‰ , may be 3.3.7 ± 2.9 × 106 cells g−1, 1.2 × 106 ± 2.0 × 105 cells g−1 and 1.4 × 107 ± 3.3 × 106 cells g−1, respectively. These cell abundances were similar with those previously reported from other deep-sea vents  in chimneys from Neuschwanstein, Hohenschwangau and Heidelberg vents were 2.3 × 10ells g−1 –83), desThioreductor were frequently detected in the Neuschwanstein chimney  were abundant in the Heidelberg chimney. All of these Epsilonproteobacteria are mesophilic chemoautotrophs [Thioreductor species lack the ability to use sulfur compounds as the energy source and instead uses H2 as the primary energy source [Thioreductor had a relatively limited distribution compared with those of Sulfurovum and Sulfurimonas [Thioreductor were dominantly detected only in the in situ cultivation device deployed into the gas-rich fluid flow [Desulfobacterium species of the class Deltaproteobacteria, were abundantly detected only in Neuschwanstein chimney. These suggested that more H2-rich vent fluids in Neuschwanstein resulted in the dominance of hydrogenotrophic chemolithoautotrophs within the associated chimney structure.Microbial community structures were assessed via 16S rRNA gene amplicon sequencing. The PCR amplicon was not obtained from Hohenschwangau chimney, which had the lowest cell density. A total of 1169 different OTUs were identified from Neuschwanstein and Heidelberg chimneys on the basis of classification with greater than or equal to 97% of identity. Although the proportions of archaeal reads were low , Methanococcales and Thermoplasmatales members were the most frequently detected archaea in Neuschwanstein and Heidelberg, respectively. Members of the order Methanococcales are hydrogenotrophic methanogens widely distributed in the OT hydrothermal fields ,84,85. T chimney . In conttotrophs ; howevery source . In prevurimonas ,81,84,853.4.Alvinocaris sp. sensu [In the hydrothermally active areas of the Yokosuka vent site, 21 species of megabenthos were identified , includip. sensu ). The vea), Shinkaicaris leurokolos shrimps, Branchinotogluma sp. scale worms and Paralvinella aff. hessleri polychaetes lived in closest proximity to high-temperature effluents (b), with all three being present on Neuschwanstein but P. aff. hessleri being absent in Hohenschwangau. These three species found to dominate the area close to vent effluent are known to be the most heat-tolerant of vent animals found in OT, and are usually found close to a vent orifice [Alvinocaris shrimps  was obsffluents ,36 of thAlvinocaris being equally abundant but with S. crosnieri being present at much higher abundance (c). Paralvinella was not present, and Branchinotogluma was present only in low abundance. Three more peripheral molluscs including two snails in genus Provanna and one limpet in genus Bathyacmaea were present too, albeit in low abundances . The top of the chimney appeared to be devoid of other megabenthos, although alvinocaridid shrimps became abundant in the lower parts. As Shisa chimney was not observed in detail and no samples were taken, it is currently not known why only Branchinotogluma sp. was abundant there.At Heidelberg chimney where venting was less robust, zonation was not as clear with the two species of bundance c. Paralvle worms d. The toAlvinocaris shrimps and Munidopsis squat lobsters . Shinkaia crosnieri is often present too but in low abundance, the same can be said for the shrimp Lebbeus shinkaiae. Upon closer examination, a high abundance of Provanna clathrata snails was revealed, occasionally together with low abundances of P. subglabra and Bathyacmaea sp. limpets. Very rarely, one or two Bathymodiolus aduloides mussels were seen. Further away from the chimney base, an unidentified demosponge begins to cover sulfide deposits.The bases of all four chimneys were similar in faunal composition. Most notably, images of the bases are dominated by a high abundance of f) and large, weakly hydrothermally influenced areas dominated by a branching poecilosclerid sponge (g) that often engulfed the tubeworms. With the exception of Lamellibrachia being more abundant in tubeworm bush regions where poecilosclerid sponge was not at all present, megafaunal communities observed at these two peripheral regions were generally similar in composition and relative abundance. The mussel B. aduloides was common, as well as gastropods Provanna spp., Lepetodrilus nux and Bathyacmaea sp. No other Bathymodiolus species were found to inhabit the Yokosuka site, although more thorough exploration in the future may recover other species. The dominance of B. aduloides is nevertheless of interest because at all other OT vents it is never present in high abundance [B. platifrons and B. japonicus. Perhaps, the fact that B. aduloides host sulfur-oxidizing endosymbiont , the doaponicus ) contribS. crosnieri completely, whereas the latter allows some intrusion of these taxa. Within the periphery cluster, the tubeworm-dominated diffuse flow and sponge-dominated periphery were almost identical in species composition but the chimney base habitat was slightly more dissimilar (60% similarity).The results of similarity analysis of megabenthos composition among the different habitat types identified  revealed4.2 and CH4, probably resulting from inputs of thermogenic volatiles and the preferential venting of vapour-rich phase after subseafloor fluid boiling. Hydrogenotrophic sulfate-reducing bacteria were found to be abundant only on the Neuschwanstein chimney, which exhausts much more H2 than the other chimneys analysed. Although the potential of available energy yield from a unit of hydrothermal fluid for chemolithotrophic microbes is high, the total microbial cell density and the overall microbial composition of chimney habitats were not distinct from those known from other OT vent sites. The same is true for megafaunal composition, with species inhabiting the Yokosuka site being in common with those from other OT vents [At a depth of 2190 m the Yokosuka hydrothermal vent site discovered in this study is the deepest OT vent known so far, extending the bathymetric range of hydrothermal activities within the OT by approximately 300 m. The highest fluid temperature recorded in the OT is also extended, as anticipated from the depth, up to 364°C. Neither the depth nor the temperature resulted in large changes in vent fluid compositions at this site compared with other OT sites. However, these fluids were distinct from other OT sites due to the notably high concentrations of HOT vents  despite The preliminary results presented here suggest that there is probably no biogeographic barrier between the 2190 m deep Yokosuka site and other OT vents between 800 and 1650 m deep, and that the Yokosuka site is also part of the well-mixed gene pool of deeper OT vents . HoweverAccording to the bathymetric data available  the deep


Here we provide multiple lines of evidence suggesting that the ∼20 p.p.m.v. overshoot in CO2 at the end of Termination 2 (T2) ∼129 ka was associated with an abrupt (≤400 year) deepening of Atlantic Meridional Overturning Circulation (AMOC). In contrast to Termination 1 (T1), which was interrupted by the Bølling-Allerød (B-A), AMOC recovery did not occur until the very end of T2, and was characterized by pronounced formation of deep waters in the NW Atlantic. Considering the variable influences of ocean circulation change on atmospheric CO2, we suggest that the net change in CO2 across the last 2 terminations was approximately equal if the transient effects of deglacial oscillations in ocean circulation are taken into account.Large amplitude variations in atmospheric CO 2 rise. Here, the authors argue that late ocean circulation recovery during the penultimate deglaciation (T2) led to a larger rise in CO2 compared with T1.Differences in the sequence and timing of ocean circulation changes across glacial terminations could affect the magnitude of deglacial atmospheric CO For example, the change in atmospheric CO2 across the most recent deglaciation  was ∼80 p.p.m.v. compared with a rise of ∼100 p.p.m.v. across T2  effects on CO2  together with benthic foraminiferal δ13C and δ18O from the same core , planktic foraminiferal faunal abundance, ice rafted debris (IRD) counts, and additional carbonate preservation indices, all measured on the same samples from ODP Site 1063. In addition we present new sortable silt (SS) measurements (Methods section) from ODP Site 983 , to assess the strength of deep water overflows (at that site) emanating from the Nordic Seas2 change across the last two terminations can be explained, at least in part, by differences in the sequence and timing of deglacial events and their relative influence on atmospheric CO2.Recently published records of seawater Nd isotopes (measured on bulk sediment leachates) and sedimentary Pa/Th from ODP Site 1063 retrieved from the NW Atlantic18O, fauna and IRD counts (Neogloboquadrina pachyderma plus Neogloboquadrina incompta) reveals several intervals of colder surface conditions, which are also reflected by more positive planktic δ18O values and the presence of IRD. The records suggest that the interval assigned to HS11 experienced the coldest conditions within the period of interest, with the warmest conditions  being attained directly following HS11, during the earliest part of MIS 5e.Our new records of planktic foraminiferal δD counts  provide 4 at ∼128.7 ka (128.9–128.5 ka) on the ice-core timescale, AICC2012  of ∼−13, reflecting the exported mixture of deep waters formed in the eastern and western North Atlantic (NADW) and those emanating from the Southern OceanɛNd>−11) at the end of MIS 6 and throughout HS11  than their southern counterparts2013C and poor carbonate preservation (low % coarse fraction and high fragmentation) we observe during HS11 ) in the abyssal North Atlantic during HS11 relative to today. This is analogous to the millennial-scale cold events of the last glacial cycle13C shows considerably more scatter than might be expected for an epifaunal species such as Cibicidoides wuellerstorfi, an observation that has also been made for other δ13C records from this region13C record from ODP Site 1063.Modern bottom water at ODP Site 1063 has a Nd isotopic composition  and that from intermediate depth ODP Site 983 13C from ODP Site 1063  attained during early MIS 5e demand a different deep water mass configuration relative to that of the present-day North Atlantic, and a source of bottom water with no modern analogue. At present, NADW represents a predominant mixture of deep waters formed in the NE Atlantic (overflows from the Nordic Seas) and the NW Atlantic ɛNd surface waters found in the NW Atlantic  are observed north and south of the modern convection areas in the Labrador Sea and it is feasible that convection during early MIS 5e was shifted to a more southerly location compared to today, tapping into surface waters with more negative Nd isotopic composition. It is also possible that retreat of the North American ice sheet increased the weathering and supply of old continental material (with very negative ɛNd) to a broader area of the surface NW Atlantic. Irrespective of the exact location of convection though our results suggest that AMOC recovery following HS11 was accomplished (at least in part) by a drastic deepening of deep waters formed in the NW Atlantic. Similarly negative ɛNd values have been documented at the location of ODP site 1063 during the early Holocene21et al.ɛNd values might reflect the ‘re-labelling' of deep waters in the Labrador Sea by interaction with particularly un-radiogenic sediments, followed by their southward advection to abyssal depths. A key question remains as to whether these deep waters originated in the NW or NE Atlantic.Surface waters with sufficiently negative Nd isotopic compositions (ɛNd lasts for thousands of years.Results from an Earth System Model experiment (run in coarse resolution)Our results require that deep waters formed in the NW Atlantic were dense enough (at least relative to other deep waters present at that time) to influence the deepest parts of the Atlantic basin throughout the earliest part of MIS 5e. This would be possible only if other sources of deep water (which today represent the densest waters within the Atlantic) were diminished or possessed lower densities than those forming in the NW Atlantic. Indeed, a partial solution to this conundrum is hinted at by a recent proxy reconstruction from the Southern OceanFurthermore, evidence from the deep NE Atlantic suggests that deep water formation in the Nordic Seas and its overflow into the Atlantic as L-NADW may not have resumed its typical interglacial mode until ∼124 ka ref. . Our newIn 4018O provides further evidence for the timing and rapidity of ocean circulation change at the onset of MIS 5e in the NW Atlantic  and occurs after the main phase of deglacial sea-level rise18O. More likely it reflects a change in water mass geometry and the relative dominance of water masses with very different temperature/salinity characteristics within the abyssal North Atlantic. Equally rapid changes in deep ocean circulation in the same area across MIS 5e/d were reported previouslyThanks to its high temporal resolution our new record of benthic foraminiferal δAtlantic  and 3. T18O can reflect changes both in bottom water temperature and the oxygen isotopic composition of seawater (δ18Osw or δw), which is related to salinity18O implies a warming of 3.6 °C given no change in δw. Modern deep waters formed in the Southern Ocean are fresher and have lower δw than more northerly intermediate waters  southern deep waters may have been significantly more saline (with higher δw by 0.12–0.42‰) than northern water masses18O. Furthermore, since the deglacial rise in sea level across T2 was only just complete by 129 ka . If the ka refs ; Fig. 3b18O, as well as the very negative ɛNd values ∼129 ka without invoking subsidence of a ‘low density' water mass, is the formation of dense brines possibly through wind action over coastal polynyas in Baffin Bay  across the last two glacial terminations are shown in 2) and interglacial (high CO2) conditions has led to a plethora of hypotheses as to the mechanisms controlling atmospheric CO2 on orbital timescales . This may be contrasted with two distinctly more abrupt increases (∼10–15 p.p.m.v. in 100–200 years) that occurred on recovery to a stronger mode of AMOC following HS1 and the YD2 during times of particularly weak or shallow AMOC (HS events), followed by an abrupt increase on recovery is not unique to glacial terminations and is observed repeatedly throughout the last glacial period. For example, the transition from HS4 into Dansgaard-Oeschger (D-O) event 8 was marked by an abrupt increase in atmospheric CO2 of ∼10 p.p.m.v.  are thought to be linked to resumption of the AMOC52 observed while the AMOC is in a strong mode is thought to reflect the reversal of processes driving its increase during intervals of weakened circulation2 associated with a strengthening of AMOC is only a transient feature, reflecting the different timescales of the mechanisms involved .The much more rapid increases in atmospheric CO2 and the AMOC described above suggests that when ocean circulation is in quasi equilibrium  constant. Of course, additional drivers such as carbonate compensation  and HS11 (T2). Following the continuous rise in CO2 throughout HS11 a transient maximum was attained on recovery and (inferred) overshoot of the AMOC during early MIS 5e, after which CO2 stabilized at an interglacial level as the AMOC resumed its interglacial mode by ∼124 ka. In contrast, the transient maximum in atmospheric CO2 associated with the B-A occurred within its overall deglacial rise across T1 and its effect is therefore obscured  within the net change in CO2. Moreover, the abrupt rise in CO2 associated with AMOC recovery following the YD was apparently smaller than that following HS11 (perhaps reflecting the shorter duration of the YD) and in combination with the very long duration of HS11 the overall change in CO2 across T2 was larger (by ∼20 p.p.m.v.) than that across T1. On the other hand, allowing for the transient maxima in CO2 associated with AMOC recovery, the net change in CO2 between glacial and interglacial conditions was actually quite similar across both terminations  may affect the magnitude of the transient maximum in CO2. Our arguments here revolve around the timing of recovery only. We also acknowledge that the majority of our discussion is based on findings from a single core site in the NW Atlantic. Future validation of our results will require equivalent reconstructions from a variety of sites across a much broader region.Building on these arguments we now compare the last two terminations . Atmosphinations . We ther2 had reached its interglacial level.A final question concerns why there was no YD-like event during T2. It is thought that recovery of the AMOC during deglaciation may occur with the cessation of freshwater release across the North AtlanticG. inflata picked from the 300 to 355 μm fraction. Due to very low abundances of benthic foraminifera, 3 species  were used to obtain a more complete record. Measurements were performed at Cardiff University stable isotope facility using a Thermofinnigan MAT-252 mass spectrometer  for benthic samples and a Delta Advantage V  for planktics. Offsets in δ18O between benthic species were accounted for by correcting to C. wuellerstorfi by subtracting the average offset between species as measured in samples where multiple species were present . All results are reported within ODP Site 1063 core was resampled every 4 cm along the shipboard splice across the interval of interest. Sediment core samples were washed and sieved to 63 μm before drying and weighing. Planktonic foraminiferal species and fragment counts were performed on splits of the >150 μm fraction containing ∼300 individual tests. Per cent fragmentation is calculated following Le and ShackletonFossil fish teeth and debris were handpicked from the >63 μm sediment fraction of 91 samples at ODP Site 1063 between 34.0 and 39.9 metres composite depth (mcd). To obtain enough material for Nd isotope analyses, up to five samples were combined as indicated in 146Nd/144Nd ratio of 0.7219 to correct for instrumental mass bias following the exponential law. 144Sm interferences can be adequately corrected if the 144Sm contribution is <0.1% of the 144Nd signal, which was the case for all samples. Measured 143Nd/144Nd ratios of the JNdi standard yielded ratios of 0.512133±0.000013  and 0.512056±0.000015  during two separate sessions. Measurements on the Thermal Ionisation Mass Spectrometer (TIMS) were carried out as Nd oxides (NdO+) following the method outlined by Crocket et al.i143Nd/144Nd ratios of 0.512101±0.000007 . Accuracy was achieved by correcting all sample results from both machines to the published JNdi143Nd/144Nd ratio of 0.512115 ±0.000007 surface map of seawater Nd isotopic compositions in the North Atlantic  were assmall map . Water dmall map .18O (measured on G. inflata) and the Greenland ice core temperature record. Abrupt shifts in planktic δ18O (including for G. inflata) in the Northwest Atlantic are thought to have been synchronous to the shifts in Greenland ice core δ18O across Termination 1 and throughout MIS 3  with rapid shifts in Greenland temperature during the last 120 kyr  as the warming implied by our planktic δ18O and faunal records at the end of the HS11, which we tie to the abrupt rise in CH4 at ∼128.7 ka .Age uncertainties in our approach derive from the precision of alignment between the various records and the absolute uncertainty of the ice core age model. Because we are here interested in the relative timing of marine events with respect to the ice core record our error analysis does not consider the additional uncertainty of the ice core chronology. In the case of the implied resumption of deep overturning circulation following HS11, we note that the −0.9‰ shift in benthic δ128.7 ka . The abrSamples from ODP Site 983 were prepared for SS analysis following established protocols24All data generated or analysed during this study are included in this published article and its .How to cite this article: Deaney, E. L. et al. Timing and nature of AMOC recovery across Termination 2 and magnitude of deglacial CO2 change. Nat. Commun.8, 14595 doi: 10.1038/ncomms14595 (2017).Publisher's note: Springer Nature remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.Excel file containing all results from this study and published data from ODP sites 983 and 1063 as used in Figures 2 and 3.


Transactivation response element RNA-binding protein (TRBP or TARBP2) initially identified to play an important role in human immunodeficiency virus (HIV) replication also has emerged as a regulator of microRNA biogenesis. In addition, TRBP functions in signaling pathways by negatively regulating the interferon-induced double-stranded RNA (dsRNA)-activated protein kinase (PKR) during viral infections and cell stress. During cellular stress, PKR is activated and phosphorylates the α subunit of the eukaryotic translation factor eIF2, leading to the cessation of general protein synthesis. TRBP inhibits PKR activity by direct interaction as well as by binding to PKR’s two known activators, dsRNA and PACT, thus preventing their interaction with PKR. In this study, we demonstrate for the first time that TRBP is phosphorylated in response to oxidative stress and upon phosphorylation, inhibits PKR more efficiently promoting cell survival. These results establish that PKR regulation through stress-induced TRBP phosphorylation is an important mechanism ensuring cellular recovery and preventing apoptosis due to sustained PKR activation. Although IFNs induce expression of PKR at a transcriptional level, PKR’s kinase activity stays latent until it binds to one of its activators leading to its autophosphorylation and catalytic activation4. The best-characterized cellular substrate of PKR is the translation initiation factor, eIF2α, the phosphorylation of which on serine 51 (S51) results in an inhibition of protein synthesis6. An immediate response of cells exposed to various forms of stress is a general inhibition of protein synthesis, which is mainly caused by the increased S51 phosphorylation of eIF2α7. The eIF2α phosphorylation thus serves an important function to block the general protein synthesis and allow cells to either recover from stress or undergo apoptosis when damage is beyond repair8. PKR plays an important role in regulating apoptosis after exposure to several diverse stress signals that include viral pathogens, oxidative stress, endoplasmic reticulum (ER) stress, and growth factor or serum deprivation10.The double-stranded RNA (dsRNA)-activated protein kinase (PKR) is an interferon (IFN)-induced serine/threonine protein kinase expressed ubiquitously in mammalian cells11, activates PKR by a direct interaction. The dsRNA binds to PKR via the two dsRNA-binding motifs (dsRBMs) present at the N terminus15, changing the conformation of PKR to expose its ATP-binding site17 and consequent autophosphorylation18. The two dsRBMs also mediate dsRNA-independent protein-protein interactions with other proteins that carry similar domains20. Among these are proteins inhibitory for PKR activity such as TAR RNA-binding protein (TRBP)21, and also a PKR activating protein (PACT)23. PKR activation in response to stress signals is tightly regulated by PACT and TRBP, both acting to regulate its catalytic activity by a direct interaction with PKR as well as with each other25. As the dsRBMs in PKR, PACT, and TRBP mediate protein-protein interactions26, these three proteins form both heterodimers as well as homodimers and the stress-dependent phosphorylation of PACT changes the relative strengths of PKR-PACT, PACT-TRBP, and PACT-PACT interactions to bring about a timely and transient PKR activation with precise control27. This regulates the general kinetics as well as level of eIF2α phosphorylation thereby influencing the cellular response to stress either as recovery and survival or elimination by apoptosis28.During viral infections, the double-stranded (ds) RNA, which is a replication intermediate for several viruses30. TRBP inhibits PKR by interacting with dsRNA and sequestering it away from PKR as well as by forming PKR-TRBP heterodimers31. In the absence of viral infections and stress signals, TRBP forms heterodimers with both PKR and PACT, preventing their association and PACT-mediated PKR activation32. Importantly, the stress-induced serine 287 phosphorylation of PACT decreases its interaction with PKR inhibitory protein TRBP thereby further aiding in rapid PKR activation following exposure to stress signals25. In contrast, not much is known about how similar post-translational modifications may affect TRBP’s interaction with PKR and consequently, its ability to inhibit PKR during cellular stress. Previous reports indicate that TRBP is phosphorylated by the two MAPKs; ERK 1/2 and JNK, with specific effects on RISC component stability and PKR activation by endogenous Alu transcripts during mitosis respectively34. In this study, we used various biochemical assays to determine if TRBP undergoes stress-induced phosphorylation, and if this affects TRBP’s ability to inhibit PKR during oxidative stress. Our findings implicate MAPKs (ERK1/2 and JNK) in oxidative stress-induced TRBP phosphorylation, and show that TRBP phosphorylation significantly enhances TRBP’s ability to interact with and inhibit PKR during oxidative stress to regulate apoptosis.TRBP has three dsRBMs; the first two are true dsRBMs and interact with dsRNA, while the third carboxy-terminal dsRBM mediates TRBP’s interactions with other proteins such as Dicer, and MerlinTo evaluate TRBP’s effect on the cellular response to oxidative stress, we established a stable HeLa-Tet off cell line that would conditionally overexpress Flag-TRBP only when doxycycline was absent from the growth medium. A HeLa-Tet off cell line with stably transfected empty vector pTRE2pur was established as a control. We initially characterized 20 individual puromycin resistant clones and selected one clone that showed the least expression of Flag-TRBP in the presence of doxycycline and showed a good induction of Flag-TRBP expression in the absence of doxycycline. As seen in Fig. 35. As seen in Fig. In order to compare the relative apoptosis in control and TRBP overexpressing cells we used DNA fragmentation analysis. DNA fragmentation is a late marker of apoptotic cells as the DNA is cleaved by caspase-activated DNases (CADs) into nucleosomal fragments of 180 bp36. We measured and quantified PARP1 cleavage in both sets of cells after treatment with arsenite  in response to arsenite. The 116 kDa protein PARP1 is cleaved into an 89 kDa fragment by Caspase-3 in response to apoptosis-inducing stimuliite Fig. . As seenite Fig. , 24 hr, ite Fig. , 24 hr. In order to determine if TRBP undergoes post-translational modifications in response to stress signals with any functional implications on TRBP’s ability to inhibit PKR. Western blot analysis of extracts from TRBP-HeLa cells over 24 hours of exposure to sodium arsenite revealed the presence of an additional Flag-TRBP band with reduced electrophoretic mobility as indicated by an asterisk Fig.  that incMitogen-activated protein kinase (MAPK) signaling pathways are activated in response to diverse stimuli37, and elicit either pro-apoptotic or pro-survival cellular responses. Previous studies have demonstrated that MAPKs such as the Extracellular-signal regulated Kinase (ERK 1/2) and c-Jun N-terminal kinase (JNK) play important roles in mediating the cellular response to oxidative stress38. To test if ERK 1/2 phosphorylates TRBP in response to oxidative stress, we pretreated the Flag TRBP overexpressing cells with the MEK inhibitor PD0325901, and exposed the cells to sodium arsenite. In the samples not pretreated with PD0325901, we observed the p-TRBP band at 8 and 12 hrs after treatment  was also generated by substituting aspartic acid for serine at the same four sites  which contains alanine for serine substitution at 4 sites  previously identified as MAPK/ERK 1/2 substrate sitestes Fig. .Figure 3To examine the effect of TRBP phosphorylation on oxidative stress-induced apoptosis, we transfected HeLa cells with each TRBP phospho-mutant as well as wt TRBP and observed changes in the induction of apoptosis in response to oxidative stress. As seen in Fig. 40. It has been previously observed by us and others that a PKR-EGFP fusion construct encodes a constitutively active PKR, which induces apoptosis when transfected in mammalian cells41. Thus, cells were transfected with a constitutively active PKR expression plasmid (wt-PKR-EGFP+EV) or in combination with the TRBP AAAA phospho-defective or DDDD phospho-mimic mutant and assayed for changes in apoptosis induced by active PKR. We used nuclear condensation as the hallmark sign of apoptosis as indicated by the intense DAPI nuclear fluorescence42. There is a significant amount of apoptosis at 87.5%  on apoptosis induced by constitutively active PKR-EGFP was measured. Similar to our result in Fig. We further assayed apoptosis by using the mitochondrial membrane depolarization as an early marker for apoptotic cellsS. cerevisiae yeast strain. The expression of active PKR in S. cerevisiae suppresses yeast growth, and this growth inhibition can be reversed by co-expression of PKR inhibitors such as the dominant negative PKR mutant, K296R44. We introduced a galactose-inducible wt PKR yeast expression plasmid (wt PKR/pYES2) in combination with K296R, wt TRBP, AAAA TRBP, or DDDD TRBP expression plasmids (pYES3CT) into INVSc1 yeast cells. As seen in Fig. 45. Interestingly, when we co-expressed the phospho-deficient TRBP mutant (AAAA TRBP), it was unable to reverse the growth phenotype  suggesting that TRBP phosphorylation is crucial for TRBP’s ability to inhibit PKR. On the other hand, co-expression of the phospho-mimic TRBP mutant (DDDD TRBP) reversed the PKR-mediated growth inhibition more efficiently as compared to wt TRBP. It is interesting that the inhibition of PKR dependent growth suppression by wt TRBP is less efficient than DDDD TRBP but more efficient than AAAA TRBP. This would indicate that the wt TRBP is being phosphorylated in yeast and depending on what percentage of wt TRBP is phosphorylated, it is able to inhibit PKR more efficiently than AAAA TRBP but less efficiently that DDDD TRBP. These results indicate that phosphorylated TRBP inhibits PKR’s kinase activity in a more efficient manner.To determine if protection from apoptosis was a direct result of enhanced PKR inhibition by the phosphorylated TRBP isoform, we performed a yeast growth inhibition assay using the INVSc1 To understand how stress-induced phosphorylation of TRBP affects its interaction with PKR, we used the yeast-two hybrid system to test the strength of interaction between PKR and TRBP phospho-mimic and phospho-defective point mutants. The TRBP DDDD and TRBP AAAA point mutants were expressed in yeast as GAL4 DNA binding domain fusion proteins (pGBKT7) with PKR was expressed as a GAL4 activation domain fusion protein. In this system, a stronger interaction between TRBP and PKR is indicated by increased yeast growth in media lacking tryptophan, leucine, and histidine in the presence of the imidazole glycerol-phosphate dehydratase competitive inhibitor, 3-Amino-1,2,4-triazole (3-AT). As seen in Fig. We further confirmed the effect of phospho-deficient or phospho-mimic mutations on TRBP’s interaction with PKR using a co-immunoprecipitation assay from mammalian cell extracts. As seen in Fig. Since we showed enhanced interaction with PKR resulting in more efficient PKR inhibition with the phospho-mimic DDDD TRBP mutant, we tested if we observe significantly increased TRBP-PKR interaction in response to arsenite at 8 and 12 hours post-treatment when we observe the strongest TRBP phosphorylation Fig. . We test27, we investigated how phosphorylation affects TRBP-TRBP interactions. We expressed both TRBP phospho-mutants in yeast as GAL4 DNA-binding domain fusion proteins (pGBKT7) and GAL4 activation domain fusion proteins (pGADT7) and assayed for the strength of TRBP AAAA and TRBP DDDD homomeric interactions by the amount of yeast growth on nutrient deficient media in the presence of 3-AT. As seen in Fig. Since stress-induced phosphorylation is essential for efficient PACT-PACT interactionsWe also examined TRBP AAAA and TRBP DDDD homomeric interactions in mammalian cells using a co-immunoprecipitation assay. As seen in Fig. 25. Initially at early time points after stress, PACT activates PKR to aid inhibition of protein synthesis via phosphorylation of eIF2α46. Our results presented here, demonstrate that TRBP is phosphorylated by ERK 1/2 and JNK in response to oxidative stress at late time points and the phosphorylated TRBP inhibits PKR’s kinase activity more efficiently to protect cells from apoptosis. In addition, the enhanced PKR inhibition and protection from apoptosis by phospho-TRBP is brought about by an increased interaction between phospho-TRBP and PKR as well as decreased phospho-TRBP homomeric interactions. The timely downregulation of PKR activity and eIF2α phosphorylation is achieved in part by induction of GADD34, a regulatory subunit of protein phosphatase 1 (PP1)48. Our data indicates that TRBP also plays an important role in the downregulation of PKR activity. Thus, PKR activity during cell stress is dictated not only by stress-induced changes in interactions between PACT and PKR, but also by interactions between PKR and TRBP. TRBP interacts with both PACT and PKR and although TRBP phosphorylation enhances its affinity for PKR while reducing the TRBP-TRBP interactions, we observed no effect of TRBP phosphorylation on TRBP-PACT interactions (data not shown). On the contrary, stress-induced PACT phosphorylation reduces the PACT-TRBP interactions while increasing PACT-PACT interactions and PACT-PKR interactions, thereby leading to PKR activation46. Strikingly, although TRBP and PACT are very homologous, the stress-induced phosphorylation affects the protein-protein interaction properties of PACT and TRBP quite differently.Activation of PKR during cellular stress is regulated by PACT and TRBP, PACT acting positively to activate PKR and TRBP acting negatively to suppress excessive PKR activityBased on our data, we present a schematic model for TRBP-mediated downregulation of PKR activity at late time points after cellular stress. As depicted in Fig. 49. This motif is widely distributed in eukaryotic proteins, as well as in proteins from bacteria and viruses and the dsRBM-containing proteins are involved in a variety of cellular processes ranging from RNA editing, nucleocytoplasmic transport, RNA localization, protein phosphorylation in translational control, and contain a variable number of dsRBM domains50. In addition, dsRBMs can also recognize non-RNA targets (proteins and DNA), and act in combination with other dsRBMs and non-dsRBM motifs to play a regulatory role in catalytic processes51. Our work presented here add one more layer of complexity to the versatility of the dsRBM as it demonstrates that phosphorylation sites residing outside the dsRBM influence the strength of protein-protein interactions mediated via the dsRBMs. This also underscores the importance of optimal juxtaposition of the multiple dsRBM motifs relative to each other in regulating protein interactions as the phosphorylation of specific serines outside the motif can possibly bring about significant changes in overall protein conformations. In other members of the diverse family of dsRBM-containing proteins, the role of phosphorylation and post-translational modifications in regulating interactions with RNA or proteins remains to be investigated in future.The dsRBM motifs present in PKR, TRBP, and PACT possess the characteristic alpha-beta-beta-beta-alpha fold that has two well-characterized functions to bind structured RNA molecules and to mediate protein-protein interactionset al. who reported that phospho-TRBP efficiently inhibits PKR during M-G1 transition to regulate cell cycle52. In this study, the phosphorylation of TRBP during M-G1 transition was shown to be mediated by JNK. Of the four sites we studied, S142 and S152 were also identified by Kim et al. to be phosphorylated by JNK during M phase. ERK mediated phosphorylation of TRBP on S142, S152, S283, and S286 was reported in response to mitogenic signaling and was accompanied by a coordinated increase in the levels of growth-promoting miRNAs and a reduction in the levels of tumor suppressor let-7 miRNA34. TRBP phosphorylation was has been reported to occur in response to metabolic stress and inhibition of TRBP phosphorylation during metabolic stress reduced inflammation and improved systemic insulin resistance and glucose metabolism54. However, the exact functions of TRBP and PKR in high-fat diet-induced obesity and associated metabolic and inflammatory complications remains unclear and controversial55. TRBP plays an important pro-viral function in HIV infected cells by regulating PKR activity and promoting HIV replication57. Any effect of TRBP phosphorylation in HIV-infected cells also remains unexplored at present. Our work on the impact of TRBP phosphorylation on the stress signaling pathway and cellular survival thus presents an additional paradigm for exploring the existence and importance of such TRBP-mediated regulatory mechanisms in virus infected cells as well as miRNA expression and function in response to cellular stress.Our results on the effect of TRBP phosphorylation on PKR activity are in agreement with Kim HeLaM and HeLa Tet off cells were cultured in Dulbecco’s Modified Eagle’s Medium (DMEM) containing 10% fetal bovine serum and penicillin/streptomycin. Transfections were performed with Effectene Transfection Reagent (Qiagen) according to the manufacturer’s protocol.Tetracycline inducible cell lines stably expressing Flag TRBP were generated by transfection of HeLa Tet off cells with 500 ng of Flag TRBP/pTRE2pur expression plasmid. Selection of puromycin-resistant colonies was carried out 24 hours after transfection by the addition of 700 ng/ml puromycin. Another cell line was also established using the pTRE2 puro plasmid as a control. Doxycycline inducibility was quantified in Flag TRBP/pTRE2pur cell clones after removal of doxycycline by western blot analysis.Sodium arsenite, phosphatase inhibitor cocktail (Phosphatase Inhibitor Cocktail 2 – P5726), and the JNK inhibitor  were purchased from Sigma Aldrich. MEK1/MEK2 inhibitor (PD0325901) was purchased from Calbiochem (444968).Antibodies used are as follows: mouse monoclonal anti-FLAG M2 HRP (Sigma-Aldrich A8592), mouse monoclonal anti-c-Myc HRP 9E10 (Santa Cruz Biotechnology SC-40), mouse monoclonal anti-polyhistidine clone His-1 HRP (Sigma-Aldrich A7058), mouse monoclonal anti-PKR (R&D systems MAB1980), rabbit polyclonal anti-phospho PKR Thr 451 , rabbit monoclonal anti-Phospho-p44/42 MAPK (Erk1/2) (Thr202/Tyr204) (D13.14.4E) XP® , rabbit monoclonal anti-p44/42 MAPK (Erk1/2) (137F5) , rabbit polyclonal anti-PARP , mouse monoclonal anti-β-Actin HRP (Sigma-Aldrich A3854), mouse monoclonal anti-GAPDH HRP (Sigma-Aldrich G9295), goat anti-mouse IgG HRP (Sigma-Aldrich A3682), and goat anti-rabbit IgG HRP (BioRad 170-6515).+, wt TRBP/pGBKT7, myc wt TRBP/pcDNA 3.1−, Flag K296R PKR/pcDNA 3.1−, K296R PKR/pGAD424 and PKR /pYES2 expression plasmids were prepared as described previously59. Full length TRBP1 ORF with an N-terminal Flag tag from Flag TRBP/BSIIKS+ was inserted into the NotI and EcoRV restriction sites of the tetracycline-responsive vector, pTRE2pur (Clontech) to generate Flag TRBP/pTRE2pur. The phospho-defective (TRBP AAAA) and phospho-mimic (TRBP DDDD) point mutants were generated at S121, S131, S262, and S265 by substituting each serine with alanine or aspartic acid using the following primers:The Flag TRBP/BSIIKSTRBP sense:5′-GCTCTAGACATATGGAAATGCTGGCCGCCAACC-3′S121D antisense:5′-GTTCCATGGCGGGGTCCCTGGTTAGGACTACAGATGGAACTGGGG-3′S121A antisense:5′-GTTCCATGGCGGGGGCCCTGGTTAGGACTACAGATGGAACTGGGG-3′S131D sense:5′-CGCCATGGAACTGCAGCCCCCTGTCGACCCTCAGC-3′S262D S265D antisense:5′CGGAGCTCACTGAGGACACGGCAGCAGGCAGGGCCCAGGGCACCCAGGTCGCCCAGGTCGCAACTGC-3′S131A sense:5′-CGCCATGGAACTGCAGCCCCCTGTCGCCCCTCAGC-3′S262A S265A antisense:5′CGGAGCTCACTGAGGACACGGCAGCAGGCAGGGCCCAGGGCACCCAGGGCGCCCAGGGCGCAACTGC-3′The PCR products were sub-cloned into the pGEMT-Easy vector (Promega) and each sequence was verified. Full length AAAA TRBP and DDDD TRBP point mutants were generated in the pGBKT7 yeast expression vector (Clontech) by three-piece ligation of NdeI-NcoI restriction fragment from S121A or S121D/pGEMT-Easy, NcoI-SacI restriction fragment from S131D S262D S265D or S131A S262A S265A/pGEMT Easy and NdeI-SacI cut TRBP/pGBKT7. Each point mutant was subsequently introduced into the pGADT7 yeast expression vector (Clontech) by insertion of the NdeI-BamHI restriction fragment from AAAA TRBP or DDDD TRBP/pGBKT7 into the NdeI-BamHI restriction sites in pGADT7.− (Invitrogen) were generated by first introducing the NdeI-BamHI restriction piece from AAAA TRBP/pGBKT7 or DDDD TRBP/pGBKT7 into NdeI-BamHI cut Flag/TRBP BSIIKS+, and then inserting the XbaI-BamHI restriction fragment from Flag/AAAA TRBP BSIIKS+ or Flag/DDDD TRBP BSIIKS+ into the XbaI-BamHI sites in pcDNA 3.1−. Myc-tagged full length AAAA and DDDD TRBP point mutants were generated by introducing HincII-BamHI restriction fragments from AAAA TRBP/pGBKT7 and DDDD TRBP/pGBKT7 into the EcoRV-BamHI restriction sites in pcDNA 3.1−.Flag-tagged full length AAAA and DDDD TRBP point mutants in pcDNA 3.16 Flag TRBP/pTRE2pur and pTRE2pur HeLa tet off cells described in “Reagents, Cell Lines and Antibodies” were treated with 10 μM sodium arsenite for the indicated time points. Cells were collected and washed with ice cold 1 × PBS, and lysed in 100 μl of lysis buffer  for 5 minutes on ice. Lysates were centrifuged at 13200 rpm for 5 minutes, and were incubated with 100 μg Proteinase K at 37 °C for 2 hours. 5 μl of 6 M NaCl and 110 μl of isopropanol were subsequently added to the lysates which were then incubated at −20 °C overnight. The precipitated DNA was then collected by centrifugation at 14,000 rpm for 5 minutes. After the isopropanol was removed from each sample, the DNA was dissolved in 20 μl TE Buffer . The DNA was incubated with 20 μg/ml RNase A at 37 °C for 1 hour before analysis on a 1.5% agarose gel.5 × 102, 40% Glycerol, and phosphatase inhibitor cocktail 2 (Sigma) at 1:100 dilution) for 5 minutes on ice. Lysates were centrifuged at 13,200 rpm for 2 minutes. Protein concentration in the supernatant was quantified using Bradford reagent. Western blot was performed with the indicated antibodies and western blot images were analyzed using the Typhoon FLA 7000 and ImageQuant LAS 4000 .Cells were treated with sodium arsenite alone or in combination with 10 μM MEK1/MEK2 inhibitor PD0325901  or 10 μM JNK inhibitor SP600125 and harvested at indicated time points. Cells were washed twice with ice cold 1 × PBS. Harvested cells were lysed in western lysis buffer  in 100 μl co-immunoprecipitation buffer at 4 °C for 1 hour. The beads were washed in 500 μl of co-immunoprecipitation buffer three times and bound Flag-TRBP was analyzed by western blot analysis using anti-Flag antibody. Blot was then stripped and re-probed with anti-His antibody to ascertain equal His-PKR pull down. 25 μg aliquots of whole cell-lysate were analyzed by western blot analysis with anti-Flag and anti-GAPDH antibodies to ensure that equal amounts of cell lysate were used for immunoprecipitation.Flag TRBP/pTRE2pur HeLa Tet off cells grown to 50% confluency in 100-mm dishes were treated with 25 μM sodium arsenite for the indicated time points. Cell extracts were prepared in 100 μl co-immunoprecipitation buffer , 1% Triton-X 100, 2% Glycerol, and phosphatase inhibitor cocktail 2 (Sigma) at 1:100 dilution). 25 μg of cell extract was bound to 500 μg of recombinant, hexahistidine-tagged PKR (His-PKR) protein immobilized on Ni− (ii) 100 ng myc AAAA TRBP/pcDNA 3.1− and 250 ng Flag K296R PKR/pcDNA 3.1− (iii) 100 ng myc wt TRBP/pcDNA 3.1− and 250 ng Flag K296R PKR/ pcDNA 3.1− (iv) 100 ng myc DDDD TRBP/pcDNA 3.1− and 250 ng Flag K296R PKR/pcDNA 3.1− using the Effectene reagent (Qiagen). 24 hours after transfection, cell extracts were prepared in co-IP buffer . Myc AAAA TRBP, wt TRBP, and DDDD TRBP were immunoprecipitated with anti-c-myc agarose beads (Santa Cruz Biotechnology) in co- IP buffer overnight on a rotating wheel at 4 °C. The agarose beads were washed 5 times in co-IP buffer and the bound proteins were analyzed by western blot analysis with the anti-c-myc (Santa Cruz Biotechnology) and anti-Flag (Sigma) antibodies.HeLa cells were transfected in 6-well culture dishes with (i) 250 ng Flag K296R PKR/pcDNA 3.1− and Flag TRBP DDDD/pcDNA 3.1−, (ii) myc TRBP DDDD/pcDNA 3.1− and pcDNA 3.1− (iii) myc TRBP AAAA/pcDNA 3.1− and Flag TRBP AAAA/pcDNA 3.1−, (iv) myc TRBP AAAA/pcDNA 3.1− and pcDNA 3.1− using the Effectene reagent (Qiagen). 24 hours after transfection, cell extracts were prepared in co-IP buffer . Flag TRBP AAAA and Flag TRBP DDDD were immunoprecipitated with anti-Flag mab-agarose (Sigma) in co- IP buffer. The agarose beads were washed 5 times in co-IP buffer. The bound proteins were then analyzed by western blot analysis with the anti-c-myc (Santa Cruz Biotechnology) and anti-Flag (Sigma) antibodies.HeLa cells were co-transfected in 6-well culture dishes with 250 ng each of (i) myc TRBP DDDD/pcDNA 3.1600 of 2 in YPD media . 500 μl of each culture was pelleted and resuspended in an appropriate amount of distilled water to yield an OD600 of 10. Serial dilutions were then made to yield OD600 values of 1, 0.1, and 0.01. 10 μl of each dilution was then spotted onto synthetic medium lacking uracil and tryptophan and containing either glucose or galactose as a carbon source (Clontech).Wild-type and TRBP phospho-mimic and phospho-defective point mutants were subcloned into the pYES3CT yeast expression plasmid (Invitrogen). Wild-type PKR was subcloned into the pYES2 yeast expression vector (Invitrogen) as previously described for galactose inducible PKR expression. The constructs were introduced into InvSc1 yeast cells (Invitrogen) using the Clontech Yeast Transformation Kit. Transformed yeast cells were grown to an OD600 of 2 in YPD media . 500 μl of each culture was pelleted and resuspended in an appropriate amount of distilled water to yield an OD600 of 10. Serial dilutions were then made to yield OD600 values of 1, 0.1, and 0.01. 10 μl of each dilution was then spotted onto triple dropout SD minimal media lacking histidine, tryptophan, and leucine in the presence of 10 mM 3-amino-1,2,4-triazole (3-AT). Plates were incubated at 30 °C for 3 days for the PKR-TRBP interaction assay and 5 days for the TRBP-TRBP interaction assay.To test TRBP-PKR interaction, full length K296R PKR was expressed as a GAL4 DNA-activation domain fusion protein from the pGAD424 vector and wt TRBP and the AAAA and DDDD TRBP point mutants were expressed as GAL4 DNA-binding domain fusion proteins from the pGBKT7 vector. Full length AAAA and DDDD TRBP point mutants were expressed as GAL4 DNA-activation domain fusion proteins from the pGADT7 vector and GAL4 DNA-binding domain fusion proteins from the pGBKT7 vector to test TRBP-TRBP interaction. The AAAA TRBP and DDDD TRBP pGBKT7/pGADT7 construct pairs and PKR/pGAD424 and AAAA (DDDD) TRBP/pGBKT7 construct pairs were co-transformed into AH109 yeast cells (Clontech) and the transformed yeast cells were plated on double dropout SD minimal medium lacking tryptophan and leucine. In order to check for the transformants’ ability to grow on triple dropout media, transformed yeast cells were grown to an OD− and 200 ng of pEGFPC1 (Clontech) using Effectene (Qiagen). Cells were also co-transfected with 200 ng BSIIKS+ (Agilent) and 200 ng pEGFPC1 as a control. The cells were observed for GFP fluorescence 24 hours after transfection using an inverted fluorescence microscope (EVOS® FL Imaging System). Cells were treated with 25 μM sodium arsenite, and cellular morphology was monitored at 1 hour intervals. 12 hours after treatment, the cells were rinsed with ice-cold phosphate buffered saline (PBS) and fixed in 2% paraformaldehyde for 10 minutes. Cells were washed twice in ice-cold PBS and permeabilized with 0.1% Triton-X for 10 minutes, after which the cells were washed twice in ice-cold PBS. Cells were stained with the DAPI nuclear stain  at 0.5 μg/ml in PBS for 10 minutes at room temperature in the dark. The cells were rinsed once with PBS and viewed under the fluorescent microscope. At least 300 GFP-positive cells were counted as apoptotic or live based on their morphology. Cells showing normal flat morphology were scored as live, while cells showing cell shrinkage, membrane blebbing, rounded morphology and nuclear condensation with intense fluorescence as apoptotic. The percentage of cells undergoing apoptosis (Percent apoptosis) was calculated using the formula:  × 100.HeLa cells were grown to 50% confluency in six-well plates and co-transfected with 200 ng of Flag wt TRBP, TRBP AAAA or TRBP DDDD/pcDNA 3.1−, Flag TRBP AAAA/pcDNA 3.1−, or Flag TRBP DDDD/pcDNA 3.1− using Effectene (Qiagen). 24 hours after transfection, the cells were rinsed with ice-cold phosphate buffered saline (PBS) and fixed in 2% paraformaldehyde for 10 minutes. Cells were washed twice in ice-cold PBS and permeabilized with 0.1% Triton-X for 10 minutes, after which the cells were washed twice in ice-cold PBS. The cover slips were mounted in Vectashield mounting medium containing DAPI (Vector Laboratories). Cells were then viewed under the fluorescence microscope (EVOS® FL Imaging System). At least 500 EGFP-positive cells were scored as live or apoptotic as described in ‘Apoptosis Assay’.HeLa cells were grown on coverslips and transfected with 500 ng of wt PKR pEGFPC1 and 20 ng pcDNA 3.1−, Flag TRBP AAAA/ pcDNA 3.1−, or Flag TRBP DDDD/pcDNA 3.1−. MitoPT® TMRM assay was performed using the manufacturer’s instructions (ImmunoChemistry Technologies MitoPT® TMRM Assay Kit). Green fluorescence (EGFP-PKR) and changes in red fluorescence  were observed under an inverted fluorescence microscope (EVOS® FL Imaging System). At least 500 PKR expressing cells (GFP positive cells) were scored as live or dead based on decreased or absent red fluorescence. The percentage of cells undergoing apoptosis (Percent apoptosis) was calculated using the formula:  × 100.HeLa cells were grown to 50% confluency in six-well plates and transfected with 500 ng of wt PKR pEGFPC1 and 20 ng pcDNA 3.1Statistical significance of western blot quantifications and percent apoptosis were determined by two-tailed Student’s T-test assuming equal variance or one-way ANOVA followed by post-Hoc Tukey test respectively. Figure legends indicate the statistical test used, and p values are denoted by brackets and special characters. Alpha level was p = 0.05.All data generated during this study are included in this published article (and its Supplementary Information files).Supplementary material


Related research article Badheka D, Yudin Y, Borbiro I, Hartle CM, Yazici A, Mirshahi T, Rohacs T. 2017. Inhibition of transient receptor potential melastatin 3 ion channels by G-protein βγ subunits. eLife6:e26147. doi: 10.7554/eLife.26147Related research article Quallo T, Alkhatib O, Gentry C, Andersson DA, Bevan S. 2017. G protein βγ subunits inhibit TRPM3 ion channels in sensory neurons. eLife6:e26138. doi: 10.7554/eLife.26138Related research article Dembla S, Behrendt M, Mohr F, Goecke C, Sondermann J, Schneider FM, Schmidt M, Stab J, Enzeroth R, Leitner MG, Nuñez-Badinez P, Schwenk J, Nürnberg B, Cohen A, Philipp SE, Greffrath W, Bünemann M, Oliver D, Zakharian E, Schmidt M, Oberwinkler J. 2017. Anti-nociceptive action of peripheral mu-opioid receptors by G-beta-gamma protein-mediated inhibition of TRPM3 channels. eLife6:e26280. doi: 10.7554/eLife.26280Many of the cells in our body communicate by releasing small molecules that bind to receptors on the surface of target cells. These molecules include hormones and, in the case of nerve cells, neurotransmitters. Signal transduction pathways then relay the information from the receptor to inside the cell and either activate or inhibit ‘effector’ proteins that cause the cells to respond appropriately. Heterotrimeric G proteins – protein complexes that consist of three different subunits named α, β and γ – provide one such pathway, and work with cell surface receptors called G protein-coupled receptors GPCRs; .α) binds to a molecule called GDP and is tightly associated with the β and γ subunits, forming a heterotrimer. When the complex interacts with an activated GPCR, the molecule of GDP is exchanged for GTP, and the G protein complex dissociates into two parts: Gα-GTP and a stable Gβγ dimer. Both Gα and Gβγ contain ‘anchors’ that keep them attached to the cell membrane, but allow them to diffuse laterally along the membrane to find their target effector proteins. Eventually, Gα breaks down the GTP to form GDP, and Gα-GDP associates with Gβγ to reform the heterotrimer.The structure and mechanism of heterotrimeric G proteins has been studied at atomic resolution . In the α-GTP activates or inhibits a target enzyme, depending on which class of α subunit is involved (αi) inhibits the enzyme that produces a chemical messenger called cyclic AMP (or cAMP). Heterotrimeric G proteins may also regulate ion channels within the membrane via a different pathway. For instance, the GIRK channels (which are responsible for slowing the heart rate) are activated by Gβγ directly binding to them  did not detect any interaction between Gαi and TRPM3. Moreover, they and Tibor Rohacs of New Jersey Medical School and co-workers – who include Doreen Badheka and Yevgen Yudin as joint first authors – show that wild-type inhibitory Gαi subunits do not decrease the activity of TRPM3, and neither can mutant subunits that cannot break down GTP and are therefore permanently active. Furthermore, Talisia Quallo and colleagues at King’s College London show that GPCR-mediated TRPM3 inhibition is unaffected by an inhibitor that selectively acts upon the inhibitory Gαi subunit. Dembla et al. and Badheka et al. also show that altering the concentration of the chemical messenger cAMP (which is decreased by the activity of inhibitory Gαi subunits) has no effect on either the activity or inhibition of TRPM3.The studies accumulate strong evidence showing that the Gβγ in signaling to TRPM3. Badheka et al. and Dembla et al. both show that TRPM3 activity is strongly inhibited by the overexpression of Gβγ, whereas the overexpression of engineered proteins that bind to Gβγ (and so prevent it from interacting with TRPM3) eliminates GPCR-mediated TRPM3 inhibition. Co-immunoprecipitation experiments demonstrate a direct interaction between Gβγ and TRPM3. Finally, Badheka et al. show that the flow of ions through TRPM3 channels is strongly and reversibly inhibited when the cell membrane is flushed with purified Gβγ, but not with Gα or boiled Gβγ. These elegant studies thus reveal that Gβγ inhibits TRPM3 by directly binding to the channel  lead to the localized depletion of PIP2  contribute to this process in any native cell.Interesting mechanistic questions remain. Not all G to open . Because of PIP2 , heterot of PIP2 . HoweverFinally, in vivo experiments by all three groups highlight the practical relevance of the uncovered pathway for pain signaling in the peripheral nervous system: molecules that bind to and activate two types of GPCRs – the GABA-B and μ opioid receptors – significantly reduce TRPM3-dependent pain. Of note, the strongest peripheral painkillers currently available activate the μ opioid receptor, but cause severe adverse effects in the brain such as addiction, tolerance or respiratory depression. The new findings suggest that peripheral pain might be better treated by drugs that inhibit TRPM3 directly.