While the mechanisms described above are generally constrained to glacial climates due to the requirement of large ice sheets for the rapid input of melt water and change in large orographic barriers; there have been recent indications that interglacials may be susceptible to rapid climate change as well [Nicholl et al., 2012]. Freshwater input into the North Atlantic could occur via melting of the Greenland Ice Sheet (GrIS), and such a melting event would be tied to a rise in sea level. Estimates of sea level change indicate that the GrIS was indeed reduced during the Last Interglacial (LIG) [Dutton and Lambeck, 2012, Dutton et al., 2015b]. One difﬁculty here lies in constraining the timing and magnitude of GrIS disintegration during the LIG, as Dutton and Lambeck  indicate uncertainties in the sea level reconstructions. These are based largely upon examining fossil coral reef terraces, which lie above the modern sea level, and can be dated using U-series geochronometry. However, the sites which were examined to reconstruct LIG sea level changes were not all tectonically stable, and geologic uplift cannot be ruled out for all locations. Such a geophysical effect would conceivably lead to changes in the actual measured elevation relative to the paleo-elevation. Additionally the paleo-water depth at which the corals lived is associated with signiﬁcant vertical uncertainties [Dutton et al., 2015b] of up to 2 m. Therefore, the magnitude of freshwater contribution that might originate from GrIS melting cannot be precisely constrained. Timing of GrIS is also an issue, as there are some challenges interpreting the potential alteration of U-Th isotopic signatures [Stirling and Andersen, 2009].
To date, most ice core studies on the Greenland ice sheet have been carried out point-wise (e.g., Dye 3, GRIP, GISP2, NGRIP), which begs the question of how representative one single long ice core record is for deriving a comprehensive record of past climate. A study of ice cores from south- ern Greenland revealed that winter season stablewater iso- topes are largely influenced by the North Atlantic Oscilla- tion (NAO) and are strongly related to southwestern Green- land air temperatures. On the other hand, summer season sta- ble waterisotope ratios show higher correlations with North Atlantic sea surface temperature conditions (Vinther et al., 2010). In particular, northern Greenland has been little inves- tigated so far. The summit in Greenland’s center is the high- est site and separates Greenland into a northern and southern part. Northern Greenland differs significantly from the south in terms of lower air temperatures and lower snow accumula- tion rates (Fischer et al., 1998c). Thus, the results from south- ern Greenland are not directly transferable to the northern part.
In future research, two main issues will have to be addressed in order to establish an appropriate description of the dependence of the non-equilibrium fractionation factor k on wind velocity (or, alternatively, to show in a more definite way that this dependence is negligible). First, more measurements of isotope ratios in atmospheric water vapor should be made available that al- low to test newly developed parameterizations of k. As isotope ratios in the evaporation flux cannot be measured directly, complex models, which include processes like the advection of water vapor (e.g. GCMs or the Lagrangian approach used here) have to be applied to compare theoretical predictions from a Craig-Gordon model with measurements. Second, more recent parameterizations of water evaporation from the ocean (see e.g. Fairall et al., 2003) might pro- vide the theoretical basis for the description of k. These parameterizations have the advantage that they are grounded on measurement data in a much stronger way than the mostly theoretical Brutsaert model applied by MJ79. However, they usually do not contain an explicit formulation of molecular diffusion, but subsume the properties of the diffusive surface layer in a parameter called moisture roughness length, which is then parameterized with an empirical equation. Ba- sically, the moisture roughness length can also be expressed in terms of a diffusion coefficient or Schmidt number (and thus calculated for the different water isotopes) (cf. Liu et al., 1979; Brutsaert, 1982). But, owing to its empirical formulation, it is not straightforward to employ these more recent parameterizations of evaporation for the deduction of the isotope fractionation factor. In our opinion, this issue will have to be addressed with the help of a comprehensive ex- perimental (e.g. wind tunnel) study analyzing the dependence of the moisture roughness length on the Schmidt number, extending the work of Merlivat (1978a).
(2600–3500 m water depth) are among the lowest fluxes recorded worldwide. These low flux values are a result of strongly stratified and nutrient-depleted upper waters in the gyre. Such oligotrophic conditions lead to low primary production rates in a relatively homogeneous and isolated ocean region. Consequently, we observe an almost constant rain of POC fluxes in space and time, although minor variations in the net primary production (NPP) and in the sea surface temperature (SST) are seen in satellite surveys and model estimations. Factors contributing to the lack of seasonality in the POC fluxes are intense organic matter degradation, variations in the ocean mixed layer depth (OMLD), and impacts of physical mixing (surface wind stress, cyclonic eddies). Preliminary estimates indicate that the average POC export efficiency (𝜀 = 0.03 ± 0.01) is extremely low in the IOSG. Assuming that the IOSG, as well as comparable ocean regions, will expand under climate warming conditions, it is of major importance to investigate POC export fluxes to the deep ocean in order to predict changes in the global carbon cycle during the next decades.
The experimental Wüstebach headwater catchment (39 ha; WU14, Figure 1) belongs to the Lower Rhine/Eifel Observatory of the TERENO network (Bogena et al., 2012; 2018) located in the Eifel National Park (50 30 0 16 00 N, 06 20 0 00 00 E). A smaller tributary catchment (11 ha) located northeast of the Wüstebach catchment serves as an uncut ref- erence site for the clear-cut experiment ( ‘reference catchment’, WU17, Figure 1). The Wüstebach catchment covers altitudes ranging from 595 m a.s.l. in the northern part to 628 m a.s.l. in the south. The subsoil consists of Devonian slate that is covered by a periglacial solifluction layer with a maximum depth of 2 m. Cambisols and Planosols mainly occur on the slopes of the hills, while Gleysols and Histosols have developed under the influence of the groundwater in the riparian zone in the valley (Figure 1). The catchment is mainly covered by Norway spruce (Picea abies L.) and characterized by a humid, temperate climate with warm summers and mild winters with an annual average tempera- ture of approximately 7 C. The average annual precipitation and runoff are approximately 1200 and 700 mm, respectively.
Atmospheric precipitation is the primary source of water for ter- restrial and aquatic ecosystems (Brázdil, 1992; Galloway & Cowling, 1978). Hence, the detailed understanding of circulation processes and background mechanisms controlling precipitation of a given region and its temporal and spatial variations is an important issue for assess- ment and forecast of weather and water regimes worldwide (Bailey, Klein, & Welker, 2019; Brázdil, 1992; Tang et al., 2017; Trenberth, 2011). In particular, hydrogen and oxygen isotopes in precipitation are natural tracers, which help to understand the atmospheric moisture cycle (Araguás-Araguás, Froehlich, & Rozanski, 2000). Variations in stableisotope composition ( δD and δ 18 O) of precipitation are caused by isotope fractionation processes occurring at phase transitions in the hydrological cycle (Dansgaard, 1964). They correlate with climatic parameters such as air temperature, humidity, and precipitation amount (Dansgaard, 1964; Rozanski, Araguás-Araguás, & Gonfiantini, 1993) and are defined by air mass trajectories (Kurita, 2011; Merlivat & Jouzel, 1979). Consequently, stable isotopes have a great potential to provide unique information on atmospheric circulation patterns and climate changes. In recent years, the stableisotope composition of precipitation has become one of the most reliable tools for meteo- rological, climatological, and hydrological studies (Bowen, 2010; Tang et al., 2017; Wei, Lee, Liu, Seeboonruang, & Koike, 2018; Wu, Zhang, Xiaoyan, Li, & Huang, 2015; Yang et al., 2019) and modelling (Bowen, 2008; Butzin et al., 2014; Gryazin et al., 2014; Werner, Langebroek, Carlsen, Herold, & Lohmann, 2011; Yao et al., 2013). In addition, data on isotope composition of modern precipitation are widely used for decoding information about past climate conditions stored in natural archives (Rozanski, Johnsen, Schotterer, & Thompson, 1997) such as lake sediments (Kostrova, Meyer, Chapligin, Tarasov, & Bezrukova, 2014; van Hardenbroek et al., 2018), ground ice (Meyer et al., 2015; Meyer, Dereviagin, Siegert, Hubberten, & Rachold, 2002), firn/ice cores (Casado, Orsi, & Landais, 2017; Fernandoy, Meyer, & Tonelli, 2012; Pang, Hou, Kaspari, & Mayewski, 2014), tree rings (e.g., Leonelli et al., 2017), and cave stalagmites (Liang et al., 2015; Partin et al., 2012).
Rasch , 1994 ; Williamson and Olson , 1994 ). This formulation uses a shape-preserving interpolation method ( Williamson and Rasch , 1989 ), which avoids the generation of spurious minima or maxima through supersaturation by the transport of water vapor ( Williamson and Olson , 1994 ). The scheme has been found to be suﬃciently accu- rate for conserving isotopic ratios during advection to low-temperature environments ( Noone and Simmonds , 2002 ; Noone and Sturm , 2010 ), however, it does not guar- antee mass conservation ( Staniforth and Côté , 1991 ; Rasch and Williamson , 1990 ; Williamson and Olson , 1994 ; Williamson and Rasch , 1994 ). A “mass ﬁxer” that re- peatedly restores global mass is used in CAM3.0 to account for this imbalance ( Collins et al. , 2004 ). Studies with the MUGCM ( Noone and Simmonds , 2002 ) ﬁnd that the application of such a mass ﬁxer leads to ﬁctitious changes in the isotope distribution, as the mass restoration is not local and the mass is not balanced where the spuri- ous sinks/sources have removed/added the mass, which aﬀects especially the polar regions. The avoidance of mass-ﬁxing causes an annual global energy imbalance at the top of the model (TOM) and the surface (cf. Table 3.5) in comparison with the high-resolution model simulations using CAM3.0 ( Hack et al. , 2006 ). The sur- face temperature and precipitation patterns in the simulations conducted were nearly identical to the fully-coupled Community Climate System Model (CCSM3.0). Since IsoCAM gives a better tracer-tracer correlation without the application of a posteriori mass ﬁxer, our simulations were carried out without mass ﬁxing.
To investigate the global development of climatic changes in the past the oceans are an essential source of information. Due to the ideally steady sedimentation at the sea floor, past environmental conditions are recorded via proxies in marine sediments. As one component of a various number of different biogeochemical and paleontological proxies, shelled microfossils, e.g. foraminifera, are a widely used tool, because they are diverse, often highly abundant and have a wide distribution in marine environments. Foraminifera are single-celled heterotrophic protists that often build up shells, so called tests. These tests can be made of organic material, agglutinated components or calcium carbonate (calcite, aragonite). Planktonic foraminifera, which have always calcitic tests, can provide information about temperature or salinity of the sea surface. The more diverse benthic foraminifera reflect bottom water conditions from coastal areas down to the abyssal sea floor (Sen Gupta, 1999). They can occupy different microhabitats, which can be attached on surfaces or directly at the sediment-water interface, so called epifaunal microhabitats, or at different depths up to several cm within the sediment, referred as infaunal microhabitats (Corliss, 1985). Besides the utilization of benthic faunal assemblages for the reconstruction of past oceanic and climatic conditions, the geochemical composition of their carbonate tests offer additional information. Especially the stable oxygen and carbon isotope ratios are common and well used proxies to reconstruct past climate conditions (reviews in Rohling & Cooke, 1999; Pearson, 2012). The conventional agreement to work with the different ratios of 16 O to 18 O as well as 12 C to 13 C the δ-notation (δ 18 O & δ 13 C) is used, referring
Hydrogen is the most abundant element in the Universe. But the utilization of the H isotopic composition (?H-2 value) of soil to elucidate biogeochemical processes or to serve as a palaeo climate proxy is still in its infancy. In our research, we will focus on the ?H-2 value of nonexchangeable H in the clay fraction of soils. The ?H-2 value of structural H in clay minerals – mainly from C-poor subsoils - has been studied since the 1970s. The ?H-2 value of clay minerals mainly depends on (a) the average ?H-2 value of ambient water at the site and time of formation, and on (b) the size of the equilibrium isotopic fractionation factor between water and clay mineral at the temperature of formation. In our research, we will focus on the ?H-2 value of nonexchangeable H in the clay fraction of soils. Only nonexchangeable H in in structural water of minerals preserves its inherited ?H-2 value and does not exchange with water at temperatures usually occurring in soil environments at the Earth’s surface. Nonexchangeable H is bound in crystal water, which integrates the ?H-2 value of soil water over several millennia. This is in turn determined by palaeoclimatic variations of the precipitation’s ?H-2 signal with distinguishable shifts e.g., from Pleistocene to Holocene. For a global data set, Ruppenthal (2014) reported a close correlation of bulk soil ?H-2 values with those of the mean local precipitation and confirmed this for organic matter, while the clay fraction of soils was up to now not studied. We will adapt a steam equilibration method with water vapor of known H isotopic composition – formerly applied by Ruppenthal (2014) on SOM and bulk soil – to clay fractions and compare our results to the hitherto used heating treatments (200-250°C) under vacuum. We expect that the ?H-2 signal of the clay fraction of Bt horizons will serve to differentiate soils developed in different climatic epochs (e.g., Holocene, last interstadial, last interglacial) by analyzing dated palaeo soil samples. To test the hypothesis that there is a similar global regression line of the ?H-2 values in structural water of clay as up to now reported for bulk soils and soil organic matter, we will analyze the clay fraction in a global set of soil samples.
Multi-collector inductively coupled plasma mass spec- trometry (MC ICP-MS) has been applied routinely for S iso- tope ratio analysis [ 5 , 8 , 9 ]. While high sensitivity (<0.1 μmol S required for analysis [ 5 ]) and low measurement uncertainty (<0.03 % [ 5 ]) can be achieved with MC ICP-MS, non-spectral interferences caused by matrix elements (mainly K, Na, and Ca) have been shown to be major limitations [ 5 , 9 ]. Sample purification procedures have been applied successful- ly for overcoming matrix interferences in measurements of the sulfate-S isotopic composition in soil extracts and soil porewaters [ 5 , 9 ]. Although post-sampling separation proce- dures are effective, they represent a time-consuming step with the potential to cause method-related isotope fractionation. A targeted sampling procedure for soil sulfate, that separates potential interferents already during the sampling step, would be an ideal alternative to conventional separation procedures. In a recent study, we developed a novel technique for passive sampling of labile soil sulfate [ 10 ], based on the diffusive
consists of two parts: an equilibrium fractionation occurring between the liquid water and a thin, saturated vapor layer above the water mass, plus a kinetic fractionation process occurring during the diffusion of the water molecules from the saturated vapor layer into the undersaturated free atmo- sphere (Gat, 1996). For the equilibrium fractionation we per- form sensitivity studies to distinguish between three differ- ent approaches. First, we assume that no fractionation during evapotranspiration occurs at all, similar to the approach used in the ECHAM5-wiso model (Werner et al., 2011). Second, we assume that fractionation only occurs during evaporation from bare soil but not during transpiration. Last, we consider that fractionation processes take part during both evapora- tion and transpiration of water from land surface. For the im- pact of the kinetic fractionation factor, we additionally ana- lyze two different formulations given by Merlivat and Jouzel (1979) as well as Mathieu and Bariac (1996).
As already mentioned, the currently Eastern Mediterranean Sea (EMS) is a highly oligotrophic oceanic environment (Antoine et al., 1995; Béthoux, 1989). The primary production presents, approximately, half of the values observed in the mid- ocean gyres such as the Sargasso Sea or the Northeast Pacific (Krom et al., 2003). This low productivity is caused by the anti-estuarine circulation in the EMS which was extensively described in Chapter 1. The Levantine Intermediate Water (LIW) which is fed by the flows of sinking water masses that originate from the Atlantic Ocean, is characterized by a high temperature and salinity pattern at depths from 200- 500 m. Below the surface circulation, at a depth of > 500m, the Eastern Mediterranean Deep Water (EMDW) 3 (Malannote-Rizzolli and Bergamasco, 1989; Wüst, 1961) gains sufficient density after winter cooling (Lascaratos et al, 1999).
Observations from glacier icecores have shown that cyclic changes in atmospheric CO2 levels occurred over the last 420,000 years with glacial periods displaying about 80 ppmv lower values compared to interglacials (~ 280 ppmv) (Fischer et al., 1999; Petit et al., 1999). Isotope paleothermometry on the Vostok ice core revealed significant covariation between air temperature and pCO2 of the past glacial cycles (Cuffey and Vimeux, 2001), suggesting that CO2 may be an important forcing factor for climate. In contrast, Fischer et al. (1999) observed that the pCO2 increase lags the warming of the last three deglaciations by 600 ± 400 years, rather arguing for an important feedback mechanism than a real climate forcing function. However, the cyclicity between glacial and interglacial pCO2 cannot be simply explained by higher oceanic CO2 solubility due to lower temperatures because the concomitant sealevel decrease and salinity increase (e.g. Fairbanks, 1989) largely compensate the pCO2 decrease due to cooling. Although many approaches have been made to determine the major processes that control the state of the glacial ocean (e.g. Archer and Maier-Reimer, 1994; Boyle, 1988b; Broecker, 1997; Broecker and Clark, 2001b; Martin, 1990), contradictions between theories and observations could not yet be excluded so that the interactions between glacial-interglacial shifts in atmospheric CO2 and oceanic carbon sequestration remain elusive (e.g. Anderson and Archer, 2002; Elderfield, 2002; Maher and Dennis, 2001).
Stable isotopes in soil water and plant stem water (usually assumed to be xylem water) have been invaluable tools in elucidating ecohydrological interactions over the past decade (Penna et al., 2018). Earlier work by Ehleringer and Dawson (1992) and Ehleringer and Dawson (1992) explained the isotope content of xylem water in trees in terms of potential plant water sources. Building on that, Brooks et al. (2010) showed that the isotope characteristics of xylem water did not always correspond to bulk soil water sources as plant xylem water was fractionated and offset relative to the global meteoric water line (GMWL) compared to mobile soil water, groundwater and stream flow signatures. This led to the “Two Water Worlds” hypothesis which speculated that plant water was drawn from a “pool” of water that was “ecohydrologically separated” from the sources of groundwater recharge and stream flow (McDonnell, 2014). Research at some sites has found similar patterns of ecohydrologic separation (e.g., Goldsmith et al., 2012; Sullivan et al., 2016) and suggested it may be a ubiquitous characteristic of plant-water systems (Evaristo et al., 2015). Others have found that differences between plant water and mobile water may be limited only to drier periods (e.g., Hervé-Fernández et al., 2016; McCutcheon et al., 2017; Zhao et al., 2016), or may be less evident in some soil-vegetation systems (Geris et al., 2015). Direct hypothesis test- ing of potential processes that may explain the difference between the isotopic composition of xylem water and that of potential water sources has been advanced by detailed experiments in controlled environments, often involving the use of Bayesian mixing models which assume all potential plant water sources have been sampled (e.g., Stock et al., 2018). However, as field data become increasingly available from critical zone studies, more exploratory, inferential approaches can be insightful in terms of quantifying the degree to which xylem water isotopes can or cannot be attributed to measured soil water sources (Amin et al., 2020).
An alternative way of transporting fluids from lower to higher crustal levels might be convective transport by ascending melts. The solubility of H 2 O in granitic melts is strongly pressure dependent. At P-T- conditions as they prevailed at the level of migmatisation (5 kbar, 700-800°C) a granitic magma would be able to dissolve up to 10 wt. % H 2 O (e.g. Holtz et al. 1995, McMillan & Holloway 1987). Upon ascend magmas might exsolve water, due to the pressure dependent decrease in water solubility, even before cooling. In the same way magmas may have transported water from lower crustal levels to the observed level of beginning migmatisation, contributing to the inferred excess fluid conditions. On the basis of stableisotope data this scenario seems rather unlikely as interaction with magma would modify the isotopic composition of the fluid. The extent of the isotopic variations would depend on the composition and temperature of the magma. From the Navachab gold deposit (see this work) as well as from the migmatites (G. Stevens pers. comm.) there is no evidence for the input of fluids with isotopic signatures external to those of the metapelitic rocks.
Dissimilatory sulphate reduction (DSR) leads to an overprint of the oxygen isotope composition of sulphate by the oxygen isotope composition of water. This overprint is assumed to occur via cell-internally formed sulphuroxy intermediates in the sulphate reduction pathway. Unlike sulphate, the sulphuroxy intermediates can readily exchange oxygen isotopes with water. Subsequent to the oxygen isotope exchange, these intermediates, e.g. sulphite, are re-oxidised by reversible enzymatic reactions to sulphate, thereby incorporating the oxygen used for the re-oxidation of the sulphur intermediates. Consequently, the rate and expression of DSR-mediated oxygen isotope exchange between sulphate and water depend not only on the oxygen isotope exchange between sulphuroxy intermediates and water, but also on cell-internal forward and backward reactions. The latter are the very same processes that control the extent of sulphur isotope fractionation expressed by DSR. Recently, the measurement of multiple sulphur isotope fractionation has successfully been applied to obtain information on the reversibility of individual enzymatically catalysed steps in DSR. Similarly, the oxygen isotope signature of sulphate has the potential to reveal complementary information on the reversibility of DSR. The aim of this work is to assess this potential.
Jayne et al. [ 9 ] discussed the possibility that isotopic anomalies encountered for waters close to the Juan de Morales basin (CP 3), might be caused by hydrothermal water-rock interactions. In fact, the Juan de Morales basin exhibits some hot springs (water temperature ~50 ◦ C) at altitudes around 2800 m a.s.l. (among others in Mamina) (Figures 2 B, 5 C and 6 ). However, Uribe et al. [ 8 ] sampled cold springs (<15 ◦ C) in the Juan de Morales basin at elevations of ~4000 m a.s.l. (Figures 2 B and 5 C). Despite the absence of a hydrothermal imprint on these cold springs, the respective waters show likewise relatively high δ 18 O values that are in perfect accord with values measured at nearby hot springs. Hence, it is not likely that hydrothermal isotope exchange processes affect the isotopic composition of groundwater in CP 3 despite a possible circulation along deep faults into fractured segments. Also, thermal waters at Pica which circulate along disruption zones within the Andean slope to depths of ~900 m b.g.l. [ 10 ], show no anomalous deviation in their isotopic composition from surrounding waters in CP 4 (Figure 2 A,B and Figure 6 ). Overall there is, therefore, no evidence for the assumption that hydrothermal water-rock interactions affect the isotopic composition of groundwater west of the Andean Precordillera. Hence, there must be other effects involved that cause catchment-dependent isotopic value ranges.
As far as the relationship between the number of OTUs and salinity values is concerned, it has been shown that it is negative, i.e. diversity decreases with the increase of the salinity. Thus, the sediment microbial communities in the Amvrakikos Gulf salinity gradient do not appear to follow Remane’s concept but rather the linear model proposed by Attrill (2002) with the species minimum at the point of maximum salinity range. This is in contrast with the constant relationship observed by Herlemann et al. (2011) for the Baltic Sea bacterioplankton along the salinity gradient. The divergence from the species-minimum concept could be attributed to the different life strategies of micro- and macroorganisms, as has been suggested by Telesh et al. (2013). In addition, microorganisms can be transported with water movement, apart from experiencing adaptation only at the molecular and cellular level (Telesh et al., 2015). Thus, they experience salinity stress differently from benthic animals with reduced mobility (Skarlato and Telesh, 2017), which causes their deviation from the recognized models for macrobenthic organisms.
Ren and co-workers developed an in vitro SPT activity test system utilizing yeast microsomes treated with palmitoyl-CoA and deuterated L -serine as SPT substrates. By means of HPLC- TQ MS, they were able to follow the incorporation of the labeled amino acid into 3KS and d18:0 Sph, which they used to determine SPT kinetics in yeast ( Ren et al., 2018 ). More recently, Harrison et al. used differentially labeled L -serine isotopologs and either palmitoyl-CoA or pentadecanoyl-CoA as SPT substrates to study the impact of stable-isotope label number and position in L -serine on the kinetics of recombinant human and bacterial SPT. For this, they applied HPLC equipped with a fluorescence detector to quantify C17 LCBs, and direct infusion high-resolution MS for detection of 3KS products ( Harrison et al., 2019 ). Both studies used state-of-the-art techniques but limited their scopes to the first two steps of sphingolipid de novo synthesis. Another recently published study presented a more complex approach that allows comprehensive monitoring of sphingolipid metabolism, not restricted to the ER. The authors tracked the flux of d17:0 Sph through the sphingolipid metabolism of MCF-7 human breast adenocarcinoma cells in an one-step in situ assay using HPLC-TQ MS. Data on complex sphingolipids carrying a C17 sphingoid base backbone, such as dhCer, Cer, sphingomyelin, and hexosylceramide, were presented ( Snider et al., 2018 ). However, the choice of d17:0 Sph as a starting material inevitably excludes the first two steps of de novo synthesis, catalyzed by SPT and KDSR. Strategies for monitoring sphingolipid metabolism in mammalian cells have recently been reviewed ( Snider et al., 2019 ). To the best of our knowledge, the here presented method is the first reported cell-free assay capable of assessing de novo sphingolipid synthesis in its entirety. This results in two major innovations applicable to this field of research. First, the complete sphingolipid de novo synthesis can now be monitored in tissues or cells of choice, ex vivo. Second, our assay is the first that allows investigation of the complete ER-based metabolism of atypical or chemically tailored sphingoid bases, molecules that are of growing interest to the field.