The vertical distribution of water vapour in the atmosphere is highly variable. It is shaped by source, transport, and sink processes, such as evaporation and condensation, horizontal and vertical advection, and mixing as determined by the atmospheric stability. The water vapour concentration itself does not provide information on the role different processes may play in its variability. The measurement of passive trace gases, such as carbon monoxide and methane, can provide insight into atmospheric mixing processes if significant sources are present at the surface but are not specific to water vapour. Considering the stable isotopic composition of atmospheric waters can support the interpretation of the phase change history of water vapour in an air mass. The isotopic composition in water vapour is here quantified by ratios between the concentration of the heavy stable isotopologues H218O and HDO and the most common isotopologue, H216O, reported in δ-notation as δD and δ18O in the water vapour . It should be noted that the isotope ratios determined by laser spectroscopy are molecular isotope ratios. In practical cases it can be shown that the difference between the molecular ratios and the more commonly used atomic ratios derived from isotope-ratio mass spectrometry is much smaller than the measurement precision .

The temperature at which phase changes occur is the main driver of isotope fractionation that alters δD and δ18O due to the heavier isotopes preferring the more tightly bound phase. As more and more water vapour condenses and rains out from a cooling air mass, the remaining vapour thus becomes increasingly depleted of heavy isotopes, as described by the Rayleigh model e.g.. The isotopic composition of water vapour is also closely related to atmospheric transport as well as source (evaporation) and sink (condensation) processes and can be seen as an integrating tracer of the water's Lagrangian transport history . In the past, some of the most prominent uses of the stable isotope composition of water have been in a palaeoclimate context to infer past temperatures and moisture sources from natural archives e.g., for ground water studies e.g., and in studies investigating the stratospheric water budget e.g.. The process-based insight provided by the isotope composition, however, also extends to synoptic and sub-diurnal timescales, and to the lower troposphere, where most atmospheric water vapour resides. In light of the key role of atmospheric moisture for the uncertainties in weather forecasts and climate predictions, understanding processes affecting isotopes on these timescales promises to serve as an additional constraint for future model improvements e.g..

Historically, the isotopic composition of atmospheric water vapour has mainly been observed at the surface using cold traps or inferred indirectly from precipitation samples e.g.. Due to the difficult sampling procedures, the vertical distribution of stable isotopes in atmospheric water vapour has in the past been sampled only on very few occasions. were the first to report δD from airborne samples made inside Hurricane Faith in 1966. Their measurements of water vapour and ice crystals collected in several vertical layers allowed developing a hypothesis on the horizontal separation of updraft and downdraft shafts in tropical storms. Shortly thereafter, published a key data set of the isotopic composition in atmospheric water vapour from 20 flights made during 1967/68 over southern Germany. These profiles showed substantial vertical and temporal (daily to seasonal) variation. Atmospheric layers with remarkably different isotopic composition were interpreted as resulting from the combined influence of turbulent mixing within the boundary layer, as well as advection and subsidence above an inversion-capped boundary layer. used a Rayleigh-type condensation model to interpret an average profile of Taylor's (1972) data. Their approach was disputed by , noting that a model fit to averaged profiles will be meaningless unless the variability of each individual weather situation is taken into account. This controversy emphasized both the importance of the meteorological conditions during individual flights as well as the potential explanatory power of reliable climatological mean isotope profiles . devised a sophisticated multi-level adiabatic trajectory model to interpret the data by taking into account the pseudo-adiabatic ascent and dry adiabatic descent of the measured air masses. His study provided important insight into the control of cloud processing and vertical motion on the observed isotope profiles.

A second major stable isotope data set was obtained by (revised by ) from repeated sampling flights during 1966–1967 and 1971–1973 over different parts of the US. The averaged profiles, obtained during different seasons and reaching up to 13 km altitude, revealed a general decrease in the heavy stable isotopes with height throughout the troposphere. Both season and location were significant for the variability of the mean profiles. Data points in and above the upper-troposphere–lower-stratosphere (UT/LS) region were less depleted than in the troposphere below, which was confirmed from balloon-based isotope sampling up to 20 km altitude . Much of the further research was then focused on the UT/LS region, such as the collection of ice crystals in midlatitude storm tops , and airborne transects in the European Arctic at 400 hPa to investigate the moisture budget of the stratosphere . The isotope distribution in the atmospheric boundary layer and lower troposphere has received less attention from airborne measurement campaigns. obtained vertical profiles from three flights in the boundary layer and the lower troposphere 20 years after the Taylor and Ehhalt flights. Albeit limited to a few data points, their results allowed deriving a first estimate of the stable isotope flux from the surface to the atmosphere. A similar analysis was done by for a number of flights in the boundary layer and lower troposphere over Mongolia.

All of the above flights suffered from the low temporal and spatial resolution available from a discrete sampling procedure. Samples were collected in cold traps, yielding only few samples per flight with a low temporal and spatial resolution. After the flight, stable isotope concentration was then determined in a laboratory from the collected water vapour with isotope-ratio mass spectrometry. Airborne in situ isotope measurements have only become possible with the advent of laser spectrometry. Since the first airborne deployment of a laser spectrometer by in the tropical UT/LS, a number of studies have investigated the isotopic composition in stratospheric water vapour and the tropical tropopause layer . In the context of validating remote-sensing products of the isotopic composition of water vapour, Dyroff's instrument has recently been deployed in the troposphere during seven flights in 2013 in the vicinity of the Canary Islands. The δD profiles acquired during these flights reveal very sharp vertical gradients related to air mass origin all the way through the troposphere . Similar features were apparent in δD measured during several airborne profiles up to ∼ 4500 m a.s.l. with a commercial laser spectrometer over interior Alaska . However, the isotopic composition in the lower troposphere and boundary layer themselves has so far not been interpreted with respect to the unprecedented information it can provide about the atmospheric transport history of water vapour.

The deuterium excess (d-excess =δD-8⋅δ18O), a second-order isotope parameter, is a widely used measure of the degree of non-equilibrium conditions during stable isotope fractionation. The parameter quantifies the deviation of the isotopic composition from the empirical global meteoric water line (GMWL) which relates δ18O and δD in global precipitation to one another with a ratio of 1:8 . During phase changes under non-equilibrium conditions, the different advection speeds of the HDO and H218O molecules lead to deviations from this empirical relationship. One example of this is evaporation into an unsaturated air mass, where strong winds prevent the system from reaching equilibrium. In such a case, there would be relatively more of the light and fast HDO molecules in the vapour phase than one would expect from the GMWL, hence the name deuterium excess. On a global average, the precipitation has a d-excess of about 10 ‰, indicating that evaporation and precipitation generally take place under non-equilibrium conditions. Under strong evaporation conditions characterized by atmospheres with low relative humidity, high d-excess of up to 40 ‰ has been observed in the vapour, in line with theoretical considerations , for example in the Mediterranean or close to the sea-ice edge . Note, however, that in highly depleted conditions with respect to HDO and H218O, as encountered at higher tropospheric levels, the d-excess can take high values because of the non-linearity of the scale it is defined on , even in the absence of kinetic fractionation (the d-excess becomes 7000 ‰ in the limiting case of -1000 ‰ for both HDO and H218O). Therefore, careful interpretation of this parameter is required.

Under the first-order approximation that these properties remain unaffected by the condensation history, the d-excess can serve as a tracer of moisture origin and source conditions . Moisture over the Mediterranean Sea is known to have relatively high values of d-excess. During a cruise in January 1995, acquired samples with a d-excess ranging from 10 to 34 ‰, which they related to intense evaporation into dry continental air along the coast. Based on measurements of d-excess at Reh¯ovot, Israel, modelled the range of d-excess over Mediterranean moisture sources as 15–40 ‰ for individual cases. concluded from his early airborne data that the d-excess would not provide any additional value. reported both δD and δ18O but did not interpret the d-excess in their data. From airborne sampling of hydrometeors in Hurricane Olivia, found a pronounced day-to-day change in the d-excess, which was hypothesized to be due to evaporation condition changes. Current in situ instrumentation is in principle capable of providing d-excess measurements, yet this has not been exploited for interpreting the d-excess of lower-tropospheric water vapour.

In this study, we report on airborne measurements of δD, δ18O, and the d-excess obtained from a custom-modified commercial cavity-ring down spectrometer (CRDS) with an enhanced laser setup. Measurements were performed during a 30-day period in the boundary layer and the lower troposphere around Corsica, France, in the western Mediterranean, yielding 21 successful flights with about 59 flight hours. The aircraft was based at Solenzara, Corsica, during the first HyMeX Hydrological cycle in the Mediterranean experiment, special observations period (SOP1) in September–October 2012, providing many additional observations from different platforms in the region .

After describing the measurement campaign, data acquisition system, and data processing (Sect. ), we present and interpret a quasi-climatological autumn vertical isotope profile of water vapour over the western Mediterranean from all available isotope flight data (Sects.  and ). Then, we examine the temporal evolution of the isotope profiles on several days along a repeated flight track, exemplifying the tight coupling of the isotopic composition to synoptic processes (Sect. ). For a single flight, detailed back-trajectory analysis shows that the isotopic composition, including d-excess, provides additional insight into moisture origin and evaporation processes (Sects.  and ). Finally, we provide a summary and conclude with the main results of the study (Sect. ).

Figure  displays the averaged isotopic composition in the western Mediterranean from all corrected and calibrated data from the 21 flights. The measurements reveal an unprecedented level of detail of the isotope distribution in the western Mediterranean lower troposphere. The 200 m binned vertical profiles of the main isotopes δD and δ18O have distinct, coherent shapes, showing an overall decrease from close to the surface (δD = -96.7 ‰ and δ18O = -14.1 ‰ between 0 and 200 m a.s.l.) to the free troposphere (δD = -225.1 ‰ and δ18O = -30.2 ‰ between 3400 and 3600 m a.s.l.; Fig. a and b, thick red lines). The vertical gradients were -36.7 ‰km-1 for δD and -4.2 ‰km-1 for δ18O. At all elevations, the distribution is skewed towards the minimum values, as is apparent from the horizontal stripes of data points acquired during longer horizontal transects. For this reason, the bin minimum and maximum values are discussed, rather than the standard deviation. The bin maximum values decrease less strongly than the bin means with vertical gradients of -16.7 ‰km-1 for δD and -2.1 ‰km-1 for δ18O (Fig. a and b, red dashed lines). The bin minimum values show a distinct decrease between 1700 and 2300 m a.s.l., which is to some extent reflected in the mean and also leads to stronger vertical gradients of the minimum values (-51.9 ‰km-1 for δD, -6.0 ‰km-1 for δ18O). While there are indications for further decreasing mean values above 3700 m a.s.l., data are only available from one flight and are thus not sufficient to extend the bin averaging to these elevations.

The shape of the composite vertical profile of the d-excess is different from that of the primary stable isotope parameters (Fig. c). Mean values show a maximum close to the surface (16.0 ‰ between 0 and 200 m a.s.l.), decrease to an overall minimum bin mean of 4.9 ‰ between 1200 and 1400 m a.s.l., and then increase again to bin mean values of 8–9 ‰ above 2800 m a.s.l. The bin minimum values are negative throughout the profile (overall minimum -15.7 ‰ at 1300 m a.s.l.). The d-excess maximum values show a marked linear decrease from 60.2 ‰ close to the surface (between 0 and 200 m a.s.l.) to about 21.2 ‰ at 1500 m a.s.l. and again a marked increase above 1700 m a.s.l. to relatively uniform maximum values around 60 ‰ above 2000 m a.s.l. Many of the high d-excess values aloft have been sampled during relatively dry conditions (q< 2 gkg-1, indicated by the blue dots). The overall minimum in d-excess at 1300 m a.s.l. is an interesting finding, which is further investigated in Sect. .

It is known from several other studies that upper-tropospheric water vapour can have high d-excess values . They are partly due to the definition of the d-excess as a function of δD and δ18O with a constant slope and to the non-linearity of the delta scale. This provides further evidence that the high d-excess encountered above the marine boundary layer is not an indication of insufficient data quality but a real feature of the isotopic composition in the atmosphere. That said, the precision of the d-excess measurements from the CRDS instrument (in terms of signal-to-noise ratio) decreased substantially at very low humidity during the flight (see Appendix ). This is also where the humidity dependency response is most influential (see Appendix ).

We now explore two hypotheses on how the observed mean isotope profiles could have been caused. The composite profiles (Fig. ) likely reflect the substantially different isotopic composition of water vapour in the (mostly marine) boundary layer and the depleted free troposphere above. An earlier interpretation of the vertical gradient of the stable isotopes is based on the increasing depletion due to a Rayleigh fractionation process . Thereby, fractionation during a moist adiabatic ascent leads to progressive depletion higher up in the atmosphere, as can be described in terms of a simple model. The mean values in the free troposphere (blue values above 1700 m a.s.l.) decrease in line with what can be expected from Rayleigh fractionation (see below). Alternatively, adopting an air mass mixing perspective, the maximum values of the primary isotope parameters could reflect the end member of water vapour from the sea surface, close to equilibrium conditions, which due to typically well-mixed conditions is almost constant up to 1200 m a.s.l. The slowly decreasing maximum values above are then a result of the (increasingly less likely) detrainment of such end-member boundary-layer air to the free troposphere and vice versa. The minimum values of δD and δ18O above 1700 m a.s.l. would then reflect air masses of a different, higher-elevation origin, for instance due to descending air masses. Then, the low values below 1700 m a.s.l. might reflect entrainment and mixing of such depleted water vapour into the boundary layer. As descending air masses are typically dry, the impact on the isotope composition will be lower during downward mixing than upward mixing. This is in agreement with the larger vertical gradient of the minimum values in the boundary layer compared to the maximum values. The plausibility of these two hypotheses is further investigated in the following sections.

The contrast between weakly depleted water vapour in the boundary layer and moderately to strongly depleted water vapour in the free troposphere is further explored now by considering the atmospheric environment and airborne measurements during specific flights. As will be exemplified below, diurnal variability and larger-scale processes jointly shape the vertical stratification reflected in the isotope profiles.

A sequence of three flights was executed following pattern 3 (Fig. ) between 20 and 21 September 2012 (flights 09–11). The pattern consisted of several east–west transects across the island, a low-level leg over the sea to the west, and vertical profiles over the island and to the east and west of it (Fig. a, white line; Fig. , green line). The vertical profile northwest of Corsica was also included in pattern 1, which was flown after the sequence, on 23 September 2012 (flight 12). This sequence of flights allows for investigating the development of the vertical isotope profiles for four flights and on 4 days. The period from 20 to 23 September 2012 was marked by pronounced differences in the vertical stratification of the lower and middle troposphere. On 20 September, ex-Hurricane Nadine was located at the upstream side of an upper-level ridge near 30∘ W and 35∘ N, in close proximity to an equatorward extending potential vorticity streamer (Fig. , thick black contour). Interaction between these two dynamical features determined the synoptic evolution over the western Mediterranean during the following days.

Initially, the eastern flank of a large upper-level ridge was located over Corsica, leading to the advection of air masses from the northwest with wind speeds of ∼ 16 ms-1 on 500 hPa (Fig. a, white arrows; Fig. , black contour). Central Europe was under the influence of a high-pressure system, and the Mediterranean experienced intermediate temperatures, with 282–286 K on 850 hPa (Fig. , colour shading). The region of interest was almost cloud-free (Fig. a, shading). Winds were from northerly directions with speeds of about 7 ms-1 on 700 hPa and associated with cold-air advection throughout the lower troposphere (Fig. a, red arrows). During the first flight of this sequence (flight 09) at 20 September 2012, a vertical profile was flown at the northwest end of flight pattern 3, west of Corsica over the Mediterranean Sea (Fig. a, cyan line).

Figure a–d displays the temperature, specific humidity, and the stable isotope parameters at the vertical profile during flight 09. The temperature profile depicts two pronounced inversions between 900–1000 and 2300–2500 m a.s.l., which are well reproduced during ascent and descent (Fig. a, black and grey lines). The boundary layer below 900 m a.s.l. and the layer above 2500 m a.s.l. are both well-mixed, with uniform lapse rates of 7.3 and 6.0 Kkm-1, respectively. The intermediate layer has a lapse rate of -7.6 Kkm-1 and an approximately isothermal lower end that shows enhanced temperature variability. The profile of specific humidity shows changes corresponding to the observed thermal structure (Fig. b). The boundary layer has a specific humidity of ∼ 9–11 gkg-1, which drops to substantially lower values (3–5 gkg-1) at and above the capping inversion. As for temperature, humidity variations are enhanced between 1000 and 1700 m a.s.l., possibly related to the partial cloud cover observed in the satellite image (relative humidity reached 90 % between 600 and 900 m a.s.l.). Specific humidity drops to uniformly very low values (1–2 gkg-1) at the upper inversion. Some difference between the upward and downward profile is apparent for humidity (Fig. b, black and grey lines).

The vertical profile of δD (Fig. c) in turn corresponds to the humidity profile. In the boundary layer, values of weak depletion (-90 to -100 ‰ δD) reflect proximity to the evaporation source. At the capping inversion (900 m a.s.l.) δD drops to -150 ‰, with increased variability. Above 1700 m a.s.l. δD increases again to -120 ‰. This isotopic inversion together with the specific humidity information confirms that the layer between 1000 and 2300 m actually consists of two layers of different origin or air mass history. Without the isotope measurements this distinction could hardly be made. Repeatability is lower for the upward and downward profiles for δD, in particular above 2300 m a.s.l. The downward profile (black line) is expected to be more reliable, as memory effects are smaller when moving from dry to moister conditions during descent (Appendix ). Note that the coherent variability across the three available specific-humidity measurements point to at least a partial contribution from spatial variability (not shown). A further influence may come from the pressure dependency of the CRDS measurements during upward and downward motion of the aircraft (Appendix ).

The shape of the d-excess profile (Fig. d) does not reflect the thermodynamic structure observed in temperature and humidity. In the lowermost 300 m a.s.l. values of up to 40 ‰ are measured for the d-excess. This likely reflects non-equilibrium fractionation during intense evaporation in the marine surface layer at the aircraft location or upstream. As a proxy for the d-excess during evaporation, relative humidity with respect to SST (RHSST) has been calculated from qHCLY at flight altitude and the skin temperature measured from the infrared radiometer (Table ). The observed d-excess of ∼ 30 ‰ corresponds to an RHSST of ∼ 40 % (Pfahl and Sodemann, 2014), slightly lower than the measurements along a low-level transect immediately following the profile (RHSST of ∼ 48 % at 150 m a.s.l.; Table , second row). It is noteworthy that the SST is more than 5 K warmer than the air temperature at 150 m a.s.l., which suggests that the at the elevation of the aircraft measurements, the atmosphere may not yet be fully coupled to the surface conditions at this time of day. Nonetheless, the d-excess indicates strong non-equilibrium fractionation conditions at the underlying sea surface during intense evaporation. The d-excess decreases to ∼ 15 ‰ in the layer between 300 and 1500 m a.s.l. and shows values near 10 ‰ between 1500 and 2300 m a.s.l., before increasing to more than 60 ‰ at the highest part of the profile. The d-excess is fairly well reproduced for the upward and downward profile, with most pronounced differences in the very dry part of the profile above 2300 m a.s.l.

The same flight pattern was repeated 23 h later during flight 10 on 21 September 2012 (Fig. b). The large-scale circulation remained very similar to the previous flight (Fig. ). Around Corsica, a weak southeasterly flow advected drier air towards the island at lower-troposphere levels and reduced cloudiness (Fig. b). Upper-level winds had strengthened and turned to westerly directions. The vertical profile of temperature had changed considerably, now showing an inversion between 1700 and 2200 m a.s.l. (Fig. e). The upper layer had warmed by about 3 K and dried uniformly to a specific humidity of 1–2 gkg-1 (Fig. f). The vertical contrasts in the isotopic composition had increased markedly. Below the inversion δD was slightly more depleted compared to 23 h before (-110 to -120 ‰; Fig. g). This is rather different from equilibrium with the given SST (δD = -80 ‰ for an SST of ∼ 20 ∘C), even though the sea surface had cooled by ∼ 6 K (Table ). Across the inversion, the δD decreased to -380 ‰ within only 300 m vertically during the downward profile (Fig. g, solid line). At the highest elevations (3700–3300 m a.s.l.) the δD profile showed very good reproducibility of the upward and downward profiles with about -250 ‰ depletion. The d-excess profile had a similar shape as during the previous flight, albeit at lower elevation, and agreed very well for both profiles (Fig. h). A minimum d-excess of about 10 ‰ was measured just below the temperature inversion at ∼ 1700 m.

The flight pattern was repeated a third time another 3.5 h later during flight 11 (11:00–15:00 UTC, 21 September 2012) with largely unchanged large-scale circulation and winds around Corsica (not shown). The temperature inversion had descended to 1400 m a.s.l. (Fig. i). Contrasts in specific humidity were very pronounced (Fig. k), with a moister boundary layer than during flight 10 (∼ 9 gkg-1; compare Fig. f), a very dry intermediate layer (1 gkg-1), and a moderately moist upper layer (3–4 gkg-1). The drastic humidity difference between the two lowermost layers was also apparent in the δD profile, with a jump from -100 to -200 ‰ within about 200 m vertical distance (Fig. k). Reproducibility was very good for this δD profile, except for the dry intermediate layer, which indicates possible memory effects in the inlet system when entering from humid into dry air layers. The d-excess showed quite a different profile for this flight with values of up to 30 ‰ in the dry layer (Fig. l). Possibly, low d-excess water vapour from the surrounding may have reduced the d-excess by entrainment. The absence of a clear maximum near the bottom indicates that high RH conditions (70–80 %) and weak evaporation prevailed in this situation (d-excess of ∼ 14 ‰).

Another 36 h later, during the final flight in this sequence (flight 12 at 09:00 UTC, 23 September 2012), the potential vorticity streamer in the large-scale circulation had evolved into a cut-off with rapid cyclogenesis near the Bay of Biscay, leading to the advection of air with high equivalent potential temperature from the Sahara to the western Mediterranean (not shown). Southwesterly flow dominated throughout the troposphere, associated with widespread broken cloud cover (Fig. c). The single downward vertical profile available on 23 September 2012 shows that the vertical structure of the boundary layer had changed significantly in all variables (red profiles in Fig. i–l). Temperature had increased by about 2–3 K and the layer above 1500 m a.s.l. had moistened significantly, leading to the observed cloudiness. The δD profile had shifted to uniformly less depleted conditions of around -140 ‰ above 1400 m a.s.l, while the d-excess gradually decreased from values around 20–30 ‰ below 500 m a.s.l. towards about 0 ‰ at the top of the profile.

The temporal variability of the vertical structure of the isotopic composition observed above has implications for the relative importance of vertical mixing and large-scale advection for understanding the vertical profiles. It appears that the two processes go hand in hand: the inversions delineate layers that are internally well-mixed but evidently have weak mixing across the inversion. Inversions thus also separate layers with a very different origin and transport history. The combination of the two can lead to strongly variable and well-structured vertical profiles as observed during the HyMeX flights, which show surprisingly strong day-to-day variability also in a situation where the large-scale conditions are seemingly rather similar. One implication of this is that a single vertical profile in a high-pressure situation might not be representative of this weather situation. The presence of the complex topography around Corsica with various influences in upstream and downstream regions further complicates the situation. The strongest isotope gradient and temperature inversion were encountered during flight 10; the following sections discuss in more detail how stable isotopes allow interpreting the transport history of an air mass.

We now present the high-resolution horizontal and vertical variability obtained during flight 10 on 21 September 2012 (the second flight discussed in Sect. ). During this flight clouds were almost absent, and a strong vertical wind shear prevailed, with southeasterly winds in the boundary layer and westerly winds in the free troposphere (Fig. b). A deep inversion of ∼ 3 K between 1700 and 2200 m a.s.l. characterized the atmospheric stratification during this flight, both at locations west (Fig. e) and east (not shown) of Corsica, separating a moist marine boundary layer from a very dry free troposphere (Fig. ).

Figure  shows the flight altitude and meteorological and isotope parameters during the first 2.5 h of flight 10 (flown along pattern 3; compare Fig. ). The flight period has been subdivided into segments A–H in Fig. . During the vertical profile in segment A, the aircraft encountered a dramatically different meteorological and stable isotope composition between flight altitudes of 1500 and 2500 m a.s.l. At 07:00 UTC, during ascent from about 1700 to 2000 m a.s.l., specific humidity decreased abruptly from 8 to about 1 gkg-1 (Fig. c, red and black lines). Humidity was in fact below the detection limit for the dew point mirror hygrometer. Associated with the drop in specific humidity were an increase in potential temperature from 299 to 305 K (Fig. b, black line), an increase in air temperature from 10 to 12 ∘C (panel b, red line), and a drop in relative humidity (Fig. d) from 80 to ∼ 15 %. Winds gradually changed from southerly to northwesterly directions and back during segment A (Fig. e).

The stable isotope composition shows similar strong changes across the inversion at the boundary-layer top. δD (δ18O) shows an initial drop from -100 to -225 ‰ (-13 to -35 ‰) and then a further depletion before the re-entry into the boundary layer at 07:11 UTC with a minimum of -275 ‰ (-41 ‰ for δ18O) at the end of segment A (Fig. f and g). The d-excess mirrors the jump of the primary isotope parameters, increasing from 12 to 65 ‰ across the inversion at 07:02 UTC (the d-excess is drawn in red for q< 2 gkg-1). It then decreases to about 40 ‰ for 5 min when the aircraft is above 3000 m a.s.l., before increasing again to about 55 ‰ (the period when the primary isotope parameters are lowest), before finally returning to about 15 ‰. The d-excess thus hints at a different air mass above 3000 m a.s.l. at this location, which is supported by close inspection of the wind speed, relative humidity, and specific humidity. This could have been overlooked easily without considering the d-excess.

Throughout the flight, the aircraft returned four times into this dry air layer above 1700 m a.s.l. (Fig. , segments C, D, F, H). Very similar behaviour was observed for the meteorological and stable isotope parameters at the same altitudes. The profile at segment D is shown in Fig. e–h, and underlines the sharpness of the transition at the inversion. From the downward profiles the vertical gradient of the isotopic composition across the inversion is estimated as -25.4 ‰100m-1 for δD and 24.0 ‰100m-1 for the d-excess.

After completion of the vertical profile at Segment A, the aircraft performed a low-level leg at the minimum safe altitude (ca. 155 m above the surface) over open water (07:12–07:25 UTC, segment B). Specific humidity and relative humidity remained narrowly at about 8 gkg-1 and 57 %, respectively, during this transect (Table ). RHSST was about 57 %. The expected non-equilibrium fractionation conditions during evaporation do not fully correspond to the high measured d-excess of ∼ 37 ‰ (depending on the non-equilibrium fractionation factor, one would expect here RHSST to be about 18–25 % according to the Craig–Gordon model) and point towards evaporation sources of at least part of the sampled water vapour elsewhere. With average wind speeds of 4 ms-1 from southeasterly directions, only weak evaporation would be expected.

A second low-level transect performed west of Corsica over the open ocean (08:05–08:25 UTC, segment E) is marked by lower d-excess of 19 ‰ compared to segment B, which was performed east of Corsica (Table  and Fig. ). Specific and relative humidity are slightly higher (10 gkg-1 and 65 %, respectively), while RHSST was on average 72 %, matching the SST (19.5 ∘C), which was lower than air temperature (20.6 ∘C). Again, the RHSST does not explain the d-excess, and thus the main sources of the moisture measured by the aircraft may be located elsewhere (in a northwesterly direction, as indicated by the mean wind direction of 320∘). Throughout the flight, the lowest values of the d-excess occur consistently at high RH conditions, at elevations where the BL (boundary-layer) top is expected, such as during segments A, C, D, and F. This points again to the role of cloud and precipitation evaporation processes in shaping the d-excess minimum observed in the overall profile.

Air mass origin and the transport history for a part of the air sampled during flight 10 are now investigated with a backward trajectory analysis using the LAGRANTO (Lagrangian Analysis Tool) model based on ECMWF operational analysis data at a 0.5∘×0.5∘ horizontal grid spacing. Trajectories were calculated 5 days backward in time every 10 s along the flight track for the time interval 07:54–08:06 UTC, i.e. during the transition from the free troposphere into the boundary layer (second part of segment D; Fig. a). The trajectory analysis confirms a very different origin of the air masses in the two vertical layers. In the lower troposphere (700–800 hPa), the aircraft encountered air masses that 4–5 days back in time had been located over southern Greenland (Fig. a, red trajectory segments) in the upper troposphere at pressure altitudes of ∼ 400 hPa (Fig. b, black lines). These air masses descended gradually by ∼ 100 hPaday-1 before being encountered by the aircraft. Large-scale descent of upper-tropospheric air due to its southward displacement along sloping isentropes with adiabatic warming is thus the most plausible cause for the low relative humidity observed during flight 10. This is also supported by a trajectory analysis for flight segments A, C, and F, which show consistent source regions and descent from similar altitudes (not shown).

The trajectories descending from high elevation were very dry (0.3±0.2 gkg-1) 4 days before encounter and hardly moistened according to the ECMWF analysis data at -12 h (0.5±0.4 gkg-1). These values agree well with the humidity observed by the aircraft (0.4–0.7 gkg-1 during the downward profile in segment D; Fig. c). The potential temperature of the air masses from the upper troposphere decreased due to radiative cooling by about 5 K during the 5 days of descent and was 303–309 K during encounter (Fig. c), in very good agreement with the observations (Fig. b). In contrast, the air mass encountered during the lower parts of segment D was much more humid (6 gkg-1; Fig. a), lower than the range of the observations (7–10 gkg-1). Trajectory analysis shows an origin of these moist air masses over Italy, France, Corsica, and nearby areas of the western Mediterranean. The air masses gradually warmed (Fig. c, grey trajectories) and moistened (not shown) while being advected in the lower troposphere from the northwest during the 5 days before encounter (Fig. a, blue trajectory segments). This underlines that the isotope composition is related to the transport history of air masses in this situation.

Vertical cross sections using ECMWF analyses centred at Corsica approximately at the time of flight 10 (06:00 UTC, 21 September 2012) provide insight into the spatial extent of different air masses and the sharpness of the transition between the free troposphere and the boundary layer in the ECMWF model (Fig. ). Dry air reached from the upper troposphere down to 750–850 hPa directly over Corsica (upper white areas in Fig. a and b), corresponding to the downward sloping isentropic flow from the northwest along the 300 and 305 K isentropes. Further south, a fairly moist (∼ 5–7 gkg-1) air mass with potential temperatures between 305 and 315 K was located above the dry air, and east of 6∘ E at higher elevations (above 850 hPa; Fig. a). This subtropical air mass associated with ex-Hurricane Nadine and the cut-off low displaced the dry upper-tropospheric air during the next 2 days, apparent also as increased cloudiness (Fig. c).

Red segments of the flight path during flight 10 mark where the aircraft encountered the very dry upper-tropospheric air, in very good correspondence with the operational analysis. The highest branches of the flight track correspond to the measurements during segments A and D. The less dry conditions, but in particular also the lower depletion of the δD and δ18O, indicate that the front of the advancing tropical air mass had been encountered during segment D (Fig. a), explaining the marked isotopic inversion between 2300 and 3500 m a.s.l. (Fig. g) as the result of the complex layering of air masses with very different water vapour transport histories.

In summary, we present here results from an extensive airborne survey of the isotopic composition of water vapour above the western Mediterranean during HyMeX SOP1. The data set represents one of the first extensive, well-documented airborne measurement campaigns of the isotopic composition lower troposphere since the 1970s with a high spatial and temporal resolution. During the 21 successful flights with 59 flight hours in total, several distinct weather situations were probed, yielding high day-to-day variability and often strong vertical and horizontal gradients in the stable isotope signals and confirming similar findings from another recent airborne study . In general, the spatial variability was much larger than anticipated from earlier low-resolution data, which causes challenges for remote-sensing observations of stable isotopes that have a low vertical resolution and for the comparison with model simulations that average over larger spatial regions than the aircraft measurements are representative of.

A sequence of flights performed in a situation of subsidence of upper- and middle-troposphere air over Corsica showed strong vertical gradients in the isotopic composition of water vapour, reaching up to -25.4 ‰100m-1 for δD across an inversion layer about 700 m thick. Mixing between the descending air mass and boundary-layer air influenced by evaporation appears as the dominant process responsible for the observed variability. Immediate, local stable isotope fractionation that could be described by a Rayleigh fractionation model, for example in dense, precipitating clouds, was not observed during the flights and must thus have taken place elsewhere in regions with saturated conditions beforehand. Obtaining this kind of information would not be possible from water vapour concentration measurements alone, which emphasizes the additional use of stable isotopes in studies of the atmospheric water cycle.

Our high-resolution measurements of d-excess clearly show that this second-order parameter adds information on the origin and processing of atmospheric water vapour. A careful design of the measurement installation and calibration routines is required to obtain interpretable data at a temporal resolution of 30–60 s and a spatial resolution of 2–4 km horizontally and 150–300 m vertically. High d-excess near the sea surface and above the boundary-layer inversion are contrasted with relatively low d-excess near the top of the boundary layer. These characteristics are consistent with the prevalence of non-equilibrium evaporation conditions near the sea surface and during evaporation of rain in unsaturated regions and cloud dissolution near the boundary-layer top. At high elevations, non-linearities in the delta scale may have contributed to high values of d-excess in highly depleted isotope conditions. Because of limitations of the instrumentation in low-humidity conditions, the remaining uncertainties call for further studies to confirm these findings. Nonetheless, the data reported here do demonstrate that the d-excess is a valuable, high-resolution process and source tracer for the meteorological situations encountered during the aircraft campaign.

The comprehensive data set of the stable isotope composition of atmospheric water vapour acquired together with many other observations during the HyMeX SOP1 can be further exploited for process studies using both the primary isotopes and the d-excess as well as for model evaluation. In the future, more extensive data sets for other regions on the atmospheric isotopic composition will be highly valuable in constraining the atmospheric water cycle in numerical weather and climate prediction models. Future campaigns would benefit in particular from complementary sampling, either ship-based or land-based, of the near-surface water vapour in the evaporation region, providing further insight into the mixing between boundary layer and free troposphere.