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Spectrophotometric measurements of the carbonate
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ion concentration: CaCO3 saturation states in the
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Mediterranean Sea and Atlantic Ocean
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Noelia M. Fajar§, Maribel I. García-Ibáñez§, Henar SanLeón-Bartolomé†, Marta Álvarez† and
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Fiz F. Pérez*§
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§
†
Instituto Investigaciones Marinas (IIM-CSIC), Eduardo Cabello 6, 36208 Vigo, Spain
Instituto Español de Oceanografía, Centro de A Coruña, Apdo. 130, 15080 A Coruña, Spain
Corresponding author
*email: fiz.perez@iim.csic.es
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Keywords: carbonate system, acidification, pH, aragonite, saturation state, chemical
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oceanography
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Submitted to Environmental Science and Technology
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ABSTRACT
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INTRODUCTION
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About one third of the total anthropogenic CO2 (Cant) emitted since the onset of
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industrialization has been absorbed by the global ocean (Khatiwala et al., 2009) leading to rapid
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changes in the carbonate chemistry of the upper layers of the ocean, revealed by a decrease of
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0.11 in pH and 13% in carbonate ion concentration ([CO32–]). Natural dynamic processes such
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as the Meridional Overturning Circulation (MOC) in the Atlantic Ocean and the double closed
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active overturning cells observed in the Mediterranean Sea (Talley, 2013; Schroeder et al., 2013)
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convey heat, salt and oxygen to the deep-ocean (Hall and Bryden, 1982; Talley et al., 2011).
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They also provoke strong accumulations of Cant in the intermediate and deep waters of the
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Atlantic Ocean (e.g., Wallace, 2001; Álvarez et al., 2003; Pérez et al., 2013) and the
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Mediterranean Sea (e.g., Schneider et al., 2010); favouring a rapid decrease in pH of the
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intermediate waters enhanced by the low buffering capacity in the Atlantic Ocean (Vázquez-
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Rodríguez et al., 2012; Resplandy et al., 2013) and despite the high buffering capacity in the
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Mediterranean Sea (Álvarez et al., 2014). If the current rate of CO2 emissions is maintained, pH
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reductions exceeding 0.2–0.3 units are expected by the year 2100 in about 23% of the North
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Atlantic deep-sea canyons and 8% of the seamounts (Gehlen et al., 2014). These areas are of
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special interest because in them inhabit cold-water corals (CWC) with calcareous skeleton
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(aragonite) as Lophelia pertusa or Madrepora oculata (Movilla et al., 2013; Lunden et al., 2014).
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Lophelia pertusa reefs and deep-water coral carbonate mounds are important hotspots of
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biodiversity in the Atlantic Ocean and the Mediterranean Sea located between 700 and 1200 m
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depth (Form and Riebesell, 2012) many of them proposed as sites of marine protected areas. The
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average global reduction in [CO32-] of 56% projected by the year 2100 (Gattuso et al., 2014) may
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be critical for the existence of CWC reefs, since 95% of these reefs are located above the
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aragonite saturation horizon (Guinotte et al., 2006). For instance, in the intermediate and deep
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waters of the Iceland Basin (1200 ± 300 m depth), the observed decrease in pH of 0.0008–0.0013
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units per year during the last three decades (Vázquez-Rodríguez et al., 2012; García-Ibáñez,
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2015) permits to estimate that these waters would be undersaturated when the CO2 concentration
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in the atmosphere would be higher than 530 ± 30 ppm (2035–2060). These changes are likely
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to impact on the structure and functioning of marine ecosystems including reduced growth and
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net erosion of coral reefs (Gattuso et al., 2014).
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CaCO3 saturation states in seawater are chiefly determined by [CO32–] because the calcium
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concentration ([Ca2+]) is conservative and depends only on salinity (0.01028*S/35 mol·kg–1)
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(Sarmiento and Gruber, 2006). Prior to 2008, the [CO32–] could only be estimated using the
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thermodynamic equations of the carbonate system in seawater (Dickson et al., 2007) and two of
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the main measurable variables of the system in seawater (pH, total alkalinity -AT-, total inorganic
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carbon -CT- and fugacity CO2). In 2008, Byrne and Yao (2008) published the first shipboard
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routine technique to directly measure [CO32–] in seawater by using spectrophotometric
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measurements of Pb(II) complexation with CO32–. Recently, Easley et al. (2013) applied this
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technique for the first time on field measurements in the North Pacific and Arctic Ocean, also
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improving the method by refining the molar absorbance ratios previously obtained by Byrne and
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Yao (2008) by comparing directly measured [CO32–] in seawater with those indirectly estimated
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by conventional means (i.e., calculated from the pairings pH–CT or pH–AT).
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The new parameterization by Easley et al. (2013) produced [CO32–] measurements consistent
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with the thermodynamics of the relatively low salinity seawater from the Pacific and Arctic
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coastal waters. In this study, we increase the range of salinity, pH and AT used by Easley et al.
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(2013) by compiling [CO32–] measurements from three cruises in the Atlantic Ocean and one in
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the Mediterranean Sea. The ranges of salinity, pH and AT in the presented data base practically
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cover those found in the global ocean. Spectrophotometric [CO32–] directly measured at sea is
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compared with that indirectly estimated by conventional means (calculated from the pH–AT
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pair). The spatial variability of the in situ aragonite saturation state derived from measured
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carbonate and AT is also commented.
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METHODS
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Discrete seawater samples for [CO32–], pH and AT were collected and analysed on board (see
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Supporting Information, hereafter SI) during three oceanographic cruises in the Atlantic Ocean
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and one in the Mediterranean Sea (Fig. 1). The CO32– samples were analysed
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spectrophotometrically at 25ºC following the method established by Byrne and Yao (2008) and
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further reformulated by Easley et al. (2013). It consists on the addition of a PbCl2 solution to the
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thermostated seawater sample so that the Pb(II) reacts with the dissolved CO32– obtaining the
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complex PbCO3. The [CO32–] is calculated in terms of UV absorbance ratio (R) using equation 1
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(Byrne and Yao, 2008),
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(1)
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where
, in which λ1 is the UV absorbance wavelength at the isobestic point of
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PbCO3 (234 nm), λ2 is the mean value of the wavelength with high absorbance variation
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(250 nm) and λ3 is a non-absorbing wavelength to correct the absorbance due to sample
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manipulation (350 nm). The Pb(II) UV absorbance spectra is directly dependant on salinity (S)
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and therefore the formation constant 1 and the coefficients e1, e2 and e3 were determined as
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second-order polynomial functions of S (Byrne and Yao, 2008; Easley et al., 2013). We used the
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most recent parameterizations of Easley et al. (2013) to compute e1, e2 and β1. The resulting
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[CO32–] (reported at 25ºC) have an uncertainty of 2.4 µmol·kg–1 at [CO32–] of 210 µmol·kg–1
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(precision 1.1%).
Figure 1.- Hydrographic stations of the four cruises where carbonate ion concentration ([CO32–]), pH and
total alkalinity (AT) measurements were taken. The table shows the ranges of the main variables: Salinity,
pHT25 (pH on the total scale at 25ºC), AT in µmol·kg–1 and [CO32–] reported at 25ºC in µmol·kg–1. SG =
Strait of Gibraltar, SS = Strait of Sicily, DWBC = Deep Western Boundary Current.
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The pH samples were measured using the spectrophotometric method described in Clayton
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and Byrne (1993) (see SI). The pH values are reported at 25ºC and on the total scale (pHT25). The
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reproducibility of pHT25 measurements was lower than 0.001, with an accuracy of 0.0055 (Carter
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et al., 2013).
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The AT samples were measured by potentiometric titration and determined by a double
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endpoint method (Pérez and Fraga, 1987; Mintrop et al., 2000; Pérez et al., 2000). Measurements
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of AT of Certified Reference Materials (CRM) of CO2 were also performed following the same
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procedure. The uncertainty of this potentiometric method is less than 2 µmol·kg–1 (Dickson et
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al., 2007).
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RESULTS AND DISCUSSION
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CO32– variability. Since the parametrization of equation 1 is S dependent, it is important to
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keep in mind that the range of S of the samples here studied differs from that of the data used by
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Easley et al. (2013) (Fig. 2). The S range in our sample collection (34.5–39.4; Figs. 1, 2) is wider
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and towards higher values compared with the salinity range in the data set compiled by Easley et
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al. (2013) from the Pacific and Arctic Oceans (26.6–34.9; Fig. 2). Besides, the bulk of samples
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with high [CO32–] used by Easley et al. (2013) are in the low-salinity and cold waters of the
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Arctic Ocean, whereas our samples with high [CO32–] are extremely saline and warm, mainly
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found in the Mediterranean Sea and in the Surface Atlantic Water (SAW). However, the samples
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with low [CO32–] present similar temperature/salinity ranges in both sets of samples: around
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2–8ºC and 34–34.7 of S in the Pacific Ocean associated with the intermediate and deep waters,
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and around 4–10ºC and 34.3–35 of S in the Atlantic Ocean associated with the Antarctic
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Intermediate Water (AAIW) and the South Atlantic Central Water (SACW). Therefore, the
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samples here analysed are good benchmarks to validate the parametrizations of Easley et al.
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(2013).
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Figure 2.- Potential temperature/Salinity diagram showing the [CO32–] variability in our data collection
from the Atlantic Ocean (CAIBOX, MOC2 and OVIDE) and the Mediterranean Sea (Med. Sea;
HOTMIX) and the North Pacific and Artic Ocean data discussed in Easley et al. (2013), all reported at
25ºC. SMW (Surface Mediterranean Water), SAW (Surface Atlantic Water), NACW (North Atlantic
Central Water), SACW (South Atlantic Central Water), LIW (Levantine Intermediate Water), AAIW
(Antarctic Intermediate Water), MW (Mediterranean Water), LSW (Labrador Sea Water), MDW
(Mediterranean Deep Waters), NADW (North Atlantic Deep Water).
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Based on the typical uncertainties in pHT25 (0.005) and AT (0.1%), we can state that
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[CO32-]calc is much more sensitive to pHT25 than to AT. The effect of the uncertainty of pHT25 in
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[CO32-]calc is ten times higher than that derived from AT. At high pHT25 (8.1), an error of 0.005
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in pHT25 results in an error of 2 µmol·kg–1 in [CO32-]calc, while at low pHT25 (7.65) it only yields
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to an error of 1 µmol·kg–1 in [CO32-]calc. For the [CO32–]meas, its main source of the variability is
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also associated with pHT25 (97%) with a minor influence of AT (6%).
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CO32– calculated versus measured. We calculated [CO32–] from the thermodynamic
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equations of the carbonate system combined with the dissociation acid constants of Mehrbach et
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al. (1973) refitted by Dickson and Millero (1987), using pHT25 and AT as input measurements.
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The correlation between the measured ([CO32–]meas) and estimated ([CO32–]calc) carbonate
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concentrations (both reported at 25ºC) is very high (r2 = 0.992) with a mean and standard
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deviation of the differences (µd, d) of 0.1 ± 4.5 µmol·kg–1 (Fig. 3A,B). By cruises, all the
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Atlantic ones show slopes very close to Y = X with intercepts indistinguishable from zero
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(p-level < 0.001) and high r2 (≥ 0.992). However, in the Mediterranean Sea (HOTMIX cruise)
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the correlation between the [CO32–]calc and the [CO32–]meas presents a slope < 1 (0.963 ± 0.008;
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p-level <0.01) with a positive intercept.
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The distribution of the [CO32–] differences (Fig. 3B-F) gives insights about the goodness of the
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model fit proposed by Easley et al. (2013) within a wider range of S and pH. The histograms of
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the differences (Fig. 3B) show a quite well centred difference for the MOC2 and OVIDE cruises,
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with very low d (3.0 and 3.3 µmol·kg–1, respectively) and with ~61% of the samples fell within
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2 µmol·kg–1 of [CO32–]calc, which is comparable with the results by Easley et al. (2013) (73%).
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The randomness of the spatial vertical distribution of the differences in the MOC2 and OVIDE
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cruises (Fig. 3D,E) corroborates the good agreement between [CO32–]calc and [CO32–]meas. The
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other Atlantic cruise, CAIBOX, shows a small bias of µd (-2.9 µmol·kg–1) with a quite low d
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(2.8 µmol·kg–1) (Fig. 3B). The spatial vertical distribution of the [CO32-] differences along this
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cruise (Fig. 3C) is random and almost 0 for the northern part of the CAIBOX cruise, but the
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[CO32–]meas is slightly higher than the [CO32–]calc in the southern part of the section, south of 33ºN
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(Fig. 3C). These biases may be related to the fact that those where the first shipboard
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spectrophotometric [CO32–] measurements in the Atlantic Ocean within our research group.
Figure 3.- A) Calculated ([CO32–]calc) versus measured ([CO32–]meas) carbonate ion concentration at 25ºC.
B) Histogram of differences between [CO32-]calc and [CO32-]meas at 25ºC. C-F) Vertical distribution with
longitude or latitude of these differences along the cruise tracks. The Y axis (depth in meters) is expanded
in the upper 2000 meters. Units are µmol·kg–1.
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Quite different behaviour is found in the first [CO32–] measurements in the Mediterranean Sea
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during the HOTMIX cruise. They present slightly high [CO32–]calc compared to [CO32–]meas
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(4.0 ± 5.0 µmol·kg–1) (Fig. 3A,B). The spatial vertical distribution of the [CO32–] differences
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(Fig. 3F) is random and around 0 for the Atlantic part of the section (similar to the distributions
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in the MOC2 and OVIDE cruises), whereas inside the Mediterranean Sea it presents positive
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values, even reaching more than 10 µmol·kg–1 high differences. This is related to the high
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salinity values within the HOTMIX cruise (Figs. 1, 2). In fact, in the Eastern and Western
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Mediterranean Basins, where the differences reach the highest values, the differences are well
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correlated with S ([CO32–] = [(1.7 ± 0.2) * (S – 35.66)]; p-level < 0.001 with a standard error of
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4.5 µmol·kg–1). Moreover, the high pHT25 and AT values of the Mediterranean Sea (Figs. 1, S1)
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would propitiate noisier [CO32–]calc in the HOTMIX cruise. To double check the causes of the
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positive bias of the [CO32–]calc, [CO32–] were also calculated using coulometric CT measurements
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(N = 154) and AT ([CO32–]calc_CT). The difference between the [CO32–]calc_CT and the [CO32–]meas
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shows the same spatial pattern and correlation versus S ([CO32–]calc_CT = [(2.1 ± 0.3) * (S –
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34.5)]; p-level < 0.001) than the differences resulting from the pair pHT25-AT. This suggests that
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both calculated [CO32–] are consistent regardless which pair of carbonic variables is chosen, and
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that the recent parametrizations of Easley et al. (2013) produce low [CO32–]meas at S > 36 and
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[CO32–] > 150 µmol·kg–1. In fact, Easley et al. (2013, their Fig. 4d) also reported low [CO32–]meas
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values (5–10 µmol·kg–1 lower than the [CO32–]calc) at high [CO32–] values, which might indicate
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that the spectrophotometric model fit by Easley et al. (2013) would underestimate [CO32-], and
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consequently saturation states, at high saturation levels of CaCO3 in the ocean. The same
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conclusions are attained if using the set of CO2 constants by Millero et al. (2006) or the new
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reported total borate concentration (Lee et al., 2010). The previous parametrization by Byrne and
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Yao (2008) produce slight but significantly different results from those obtained using the
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parametrization by Easley et al. (2013), depending strongly on pHT25. The new parametrization
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corrects the deviation of the measured values from the calculated ones for pHT25 < 7.9 and
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pHT25 > 8.02, leading to a better agreement between the [CO32–]calc and the [CO32–]meas. Recently,
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Patsavas et al. (2015) improved the methods of Byrne and Yao (2008) and Easley et al. (2013)
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by using (Pb(ClO4)2) as titrant instead of PbCl2, which is more soluble, resulting in better signal-
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to-noise ratios. Their procedure improved the agreement between the [CO32–]calc and the [CO32–
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]meas for [CO32–] > 180 μmol·kg−1. However, we cannot test the Patsavas et al. (2015)’s
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parametrization due to the differences in the titrant solution.
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Aragonite saturation. Typically, the in situ degree of aragonite saturation (ΩA) is given by
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(2)
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where KA is the CaCO3 aragonite solubility product and ‘is’ stands for the [CO32–] at in situ
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conditions of temperature and pressure. The [Ca2+] behaves conservatively (Sarmiento and
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Gruber, 2006) and can be obtained from [Ca2+] = 0.01028*S/35 mol·kg–1.
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An often convenient measure of the CaCO3 aragonite saturation state is simply the difference
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between the observed in situ [CO32–] ([CO32–]is) and the saturation [CO32–] ([CO32–]sat; ΩA = 1),
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i.e., the excess carbonate ion concentration ([CO32–]xs):
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(3)
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Positive (negative) [CO32–]xs means the water is supersaturated (undersaturated) with respect to
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CaCO3. The absolute value of the [CO32–]xs is a measure of the tendency for the mineral CaCO3
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to precipitate/dissolve. The [CO32–]xs has the advantage over ΩA that it is directly comparable to
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[CO32–]is. The [CO32–]sat only depends on temperature, S and pressure, and tends to be in steady
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state except for the upper waters affected by long term warming. Before development of the
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spectrophotometric technique by Dr. Byrne’s group, [CO32–]is was typically computed from pairs
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of two conventional measured carbonate system variables (pH-CT or pH-AT). Taking advantage
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of the [CO32–]meas, we also determined the [CO32–]is using [CO32–]meas and AT. Both pairings
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(pH-AT or CO32–-AT) give very close results (r2 = 0.995) with similar µd and d at 25ºC and 1
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atm (0.4 ± 4.5 µmol·kg–1). Easley et al. (2013) described a very similar behaviour in the North
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Pacific in the carbonate equilibrium equations needed to compute the [CO32–]is.
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In figure 4 (panels A,C,E,G) we present the first results of directly measured [CO32–] in the
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Atlantic Ocean and the Mediterranean Sea. Our [CO32–] measurements were performed in
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oceanic waters within a wide range of temperature, S, pH and AT, while the previous studies of
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Easley et al. (2013) and Patsavas (2015) were performed in relatively coastal areas. Our
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[CO32–]xs (computed from [CO32–]meas and AT) shows a large range of variability (Fig.
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4B,D,F,H): from -55 µmol·kg–1 in the deep waters of the Atlantic Ocean to 196 µmol·kg–1 in the
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surface waters of the Western Tropical Atlantic, being greater than 150 µmol·kg–1 in the surface
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waters of the Eastern Mediterranean.
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Figure 4.- Left panels (A, C, E, G): measured carbonate ion concentration reported at 25ºC ([CO32–]meas).
Right panels (B, D, F, H): excess carbonate ion concentration over aragonite saturation at in situ
conditions ([CO32–]xs). All panels are vertical distributions with longitude or latitude according to the
cruise. The Y axis (depth in meters) is expanded in the upper 2000 meters. Units are µmol·kg–1. Red
(yellow) dashed lines indicate salinity maximum (minimum) values.
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The Mediterranean Sea presents an eastward trend of decreasing AT and increasing pH, the
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eastern basin presents higher AT (2560–2644 μmol·kg−1) and pHT25 (7.935–8.032) and less
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variable CT (2247–2331 μmol·kg−1) than the western basin (with AT of 2388–2608 μmol·kg−1,
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pHT25 of 7.861–7.988 and CT of 2110–2336 μmol·kg−1) (Álvarez et al., 2014; Fig. S1). This leads
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to an eastward [CO32–] increase (Fig. 4A,B). The saline and warm Surface Mediterranean Water
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(SMW; Fig. 2), in the Eastern Basin of the Mediterranean Sea, presents very high [CO32–]meas
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(201–252 µmol·kg–1; Fig. 4A) and [CO32–]xs (> 150 µmol·kg–1; Fig. 4B) being associated with
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high pHT25 (8.00–8.05; Fig. S1) caused by the close equilibrium with the atmosphere of these
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waters. Below SMW, the Levantine Intermediate Water (LIW; 13.55 ± 0.06ºC and
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38.755 ± 0.015 of S; Fig. 2) presents values of [CO32–]meas (213 ± 3 µmol·kg–1) and [CO32-]xs
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(125 µmol·kg–1) similar to those of SMW. This is due to the high buffering capacity (low
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Revelle factor; Egleston et al., 2010; Álvarez et al., 2014) in these high AT (Fig. S1) and warm
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Mediterranean waters, which contributes to keep high CaCO3 supersaturation for long timescales
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although the convergence circulation in this basin transports important amounts of Cant to the
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deep waters (Schneider et al., 2010). The transition from the Eastern to the Western
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Mediterranean Basin at the Strait of Sicilia leads to a clear decrease in the [CO32–]meas and the
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[CO32–]xs, mainly due to the decrease in S, pHT25 and AT (Figs. 2, S1). In fact, the Western
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Mediterranean Deep Water (WMDW; 12.92 ± 0.05ºC and 38.489 ± 0.013 of S; Fig. 2), presents
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averaged values of [CO32–]meas (192 ± 4 µmol·kg–1) and [CO32–]xs (> 50 µmol·kg–1) lower than
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those observed in the deep waters of the Eastern Basin. The transit from the Mediterranean Sea
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to the Atlantic Ocean through the Strait of Gibraltar is marked by a decrease in the [CO32–]meas
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and in the [CO32–]xs in the deep waters, which reflects their low pHT25. The surface layer in the
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Atlantic part of the HOTMIX cruise keeps high [CO32–]meas and [CO32–]xs associated with high
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pHT25 result of the close equilibrium with the atmosphere of these waters. The effect of the
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pressure in the aragonite saturation provokes a strong vertical gradient of the [CO32–]xs below the
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upper layer (500 m depth), with undersaturated waters below 2500 m depth. The vertical
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gradient in the [CO32–]meas is interrupted by a local minimum of 103 ± 2 µmol·kg–1 around 800 m
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depth close to the Canary Islands, associated with the vestiges of AAIW (weak S minimum of
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35.25 in Fig. 2) coming from the South.
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The CAIBOX cruise, which covers the Iberian Basin (30–45ºN), also presents a strong vertical
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gradient of [CO32–]xs (Fig. 4C), with supersaturated surface waters that present a northward
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decrease of the [CO32–]xs due to the northward decrease in temperature, and undersaturated deep
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waters (North Atlantic Deep Water; NADW; 2.64 ± 0.55ºC and 34.95 ± 0.05 of S; Fig. 2). The
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vertical gradient is also interrupted by the weak minimum of [CO32–]meas (120 ± 3 µmol·kg–1;
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south of 32ºN and at ~1000 m depth) due to the influence of AAIW; and also by a weak
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maximum of [CO32–]meas (135 ± 3 µmol·kg–1; between 37ºN and 41ºN and at ~1000 m depth)
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associated with the eastward moving Mediterranean Water (MW; 9.8 ± 0.7ºC and 35.86 ± 0.08;
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Fig. 2) in the Atlantic Ocean (Carracedo et al., 2011, and here in). In fact, this slight maximum is
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concomitant with a core of S maximum (> 35.8; red dashed line in Fig. 4D). The moderate
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values of pHT25 (7.771 ± 0.006; Fig. S1) of this water mass provoke a downward displacement of
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the 50 µmol·kg–1 isoline of [CO32–]xs. Interestingly the levels of aragonite saturation of the deep
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waters shoal slightly northwards because of the deep penetration of Cant coming from the North
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(Pérez et al., 2010; Fajar et al., 2011).
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The [CO32–]meas of the surface waters continuously decreases from the Iberian Peninsula to
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Greenland (OVIDE cruise; Figs. 1, 4E), which responds to the temperature decrease. The most
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striking feature in the OVIDE cruise is the core of minimum [CO32–]meas (114 ± 3 µmol·kg–1)
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located between 600 and 3000 m depth denoting the presence of the recently strong-ventilated
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Labrador Sea Water (LSW) coming from the Labrador Sea. The [CO32–]xs of the core of LSW in
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the Iceland Basin (S < 34.93; dashed yellow line in Fig. 4F) range from 0 to 20 µmol·kg–1. This
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rapid decrease of the [CO32–]xs is due to the effect of the high Cant levels here found and the low
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buffering capacity of these waters (Resplandy et al., 2013). At the pressure range of the LSW
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core (1000–1800 m depth) important ecosystems are sustained by the presence of CWC that
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probably would be in danger in the future decades (Gehlen et al., 2014) due to the rapid decrease
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in pH and the low [CO32–]xs.
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The MOC2 cruise (7.5ºN) presents a very narrow surface layer with the highest [CO32–]meas
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(246 ± 4 µmol·kg–1) and very high [CO32–]xs (> 175 µmol·kg–1) in the West Tropical Atlantic
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(Fig. 4G,H), which corresponds to SAW (Fig. 2). Those values decrease very sharply down to
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500 m depth. Below this layer of high [CO32–]meas and high [CO32–]xs, a strong minimum of
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[CO32–]meas is present along the entire section between ~250–1250 m depth associated with
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minimum values of [CO32–]xs. This layer of minimum [CO32–]meas is derived from the sustained
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accumulation of CT derived from the mineralization of biogenic matter during the long transit
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time of AAIW from its formation region (the sub-Antarctic convergence zone, 60–55ºS)
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(Stramma and England, 1999; Stramma and Schott, 1999). The AAIW layer, centred at 800 m
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depth (S < 34.68; yellow dashed line in Fig. 4H), presents an average [CO32–]meas of
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80 ± 4 µmol·kg–1, and it is characterized by a S minimum (34.60 ± 0.06; Fig. 2; Mémery et al.,
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2000) and low pHT25 (7.539 ± 0.021; Fig. S2). Although AAIW is more important in the western
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part of the MOC2 cruise (lower S values), the minimum of [CO32–]meas (70 ± 1 µmol·kg-1) occurs
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close to the African Coast (400–800 m depth), where the layer of minimum [CO32–]xs is thicker.
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This minimum of [CO32–]meas could be related to the strong depletion in oxygen and the increase
268
in CT in the Guinea Dome (East of 22ºW). This is in agreement with the strong carbon fixation in
269
the photic layer characteristic of the eastern upwelling systems, which enhances the presence of
270
strong-mineralized subsurface waters with very low pHT25 and [CO32–] in the tropics, like those
271
found in the MOC2 cruise in the layer of oxygen minimum (centred at 400 m depth) coinciding
272
with SACW (Fig. 2). These levels of corrosive (aragonite undersaturated) old waters have also
273
been observed over the shelf of the California Coast (Feely et al., 2008; Easley et al., 2013). This
16
274
processes would be enhanced in future global changes scenarios reaching the surface layers close
275
to the African upwelling systems (Gruber et al., 2012).
276
Another striking feature of the MOC2 cruise is the maximum of [CO32–]xs below the minimum
277
related with AAIW, which is associated with the upper component of NADW detected by its
278
relative S maximum (S > 34.95; red dashed line in Fig. 4H) related to some contributions of MW
279
(Mémery et al., 2000) that also increase slightly the [CO32–]xs. The NADW is transported
280
southwards by the Deep Western Boundary Current (DWBC) (Dickson and Brown, 1994; Talley
281
et al., 2011), which ventilates the deep ocean from the North Atlantic towards the Antarctica, all
282
along the western basin (Fig. 1). A clear signal of it is detected in the MOC2 cruise between
283
2000 and 4000 m depth, West of 40ºW, by a relative deep maximum of [CO32–]meas (Fig. 4D).
284
The NADW core at this position (2.74 ± 0.63ºC and 34.941 ± 0.017) presents relatively high
285
[CO32–]meas (120 ± 2 µmol·kg–1), which is slightly higher than in northern North Atlantic
286
(OVIDE) because of the lower contribution of Cant.
287
CONCLUSIONS
288
In this work we performed a critical assessment of the rapid and robust spectrophotometric
289
method for direct CO32- determination proposed by Byrne and Yao (2008) and refined by Easley
290
et al. (2013) using field data of the Pacific and Arctic Oceans. By collecting shipboard
291
measurements of pH, AT and CO32- over a wide range of salinity, temperature and CO2 chemistry
292
conditions in the Atlantic Ocean and Mediterranean Sea, we conclude that the parameterizations
293
of e1, e2 and β1 given by Easley et al. (2013) provide better agreement between the observed and
294
calculated [CO32–] (from pHT25 and AT) than the earlier ones from Byrne and Yao (2008). The
295
Atlantic Ocean cruises where [CO32-] ranged the values in Easley et al. (2013) (< 250 µmol·kg–1)
17
296
provide very good agreement between measured and calculated [CO32–] with negligible biases.
297
About a 61% of the samples in the Atlantic cruises fell within 2 µmol·kg–1 of the calculated
298
[CO32–], in agreement with the results in Easley et al. (2013). However, in the warm, salty, high
299
AT and pH Mediterranean waters, the parameterization by Easley et al. (2013) seems to
300
underestimate the carbonate ion concentration and consequently the saturation state, thus
301
suggesting a refinement of the proposed equations or the application of the new method of
302
Patsavas et al. (2015).
303
Our CO32– measurements allow describing a large variety of environments in terms of
304
saturation states of aragonite. The Mediterranean Sea has very high excess of [CO32–] despite its
305
Cant inventory is quite high. This is related to its very high buffering capacity, which allows the
306
Mediterranean Sea waters to remain over the saturation level of aragonite for long periods of
307
time. Therefore the Mediterranean Sea is a favourable place for the development of CWC. In the
308
opposite side, the relatively thick layer of undersaturated waters between 500–1000 m depth in
309
the Tropical Atlantic is expected to progress even to more negative [CO32–]xs. As Easley et al.
310
(2013) described in the California Upwelling System, this undersaturated subsurface waters are
311
prone to be advected to the upper twilight and shelf water layers dramatically affecting
312
ecosystems (Gruber et al., 2012). However, the already undersaturated Atlantic waters below
313
3000 m depth, which have very low concentrations of Cant, are not expected to experience large
314
changes in the future decades.
315
The northern North Atlantic (> 45ºN) presents slightly positive [CO32–]xs in a large fraction of
316
its waters. However, the expected increase of the Cant content in the intermediate waters and the
317
subsequent acidification will result in a reduction their aragonite saturation levels, and possibly
318
reaching undersaturation levels in the next decades. Most of the CWC communities already live
18
319
in aragonite saturated waters (Guinotte et al., 2006), being very abundant along the North
320
Atlantic below than 1200 m depth (Maier et al., 2009; Form and Riebesell, 2012; Sánchez et al.,
321
2014). The long-term monitoring of in situ [CO32–] using automated spectrophotometric
322
techniques based in the Pb(II) method may help to monitor those expected future changes, and
323
should be promoted considering the good results here reported.
324
325
326
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337
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Supporting Information (SI)
480
METHODS
481
CO32- sampling and measurements: Unfiltered seawater samples were directly taken from the
482
Niskin bottles into cylindrical quartz Perkin Elmer cells of 28 mL of volume and 100 mm of path
483
length. These cells were carefully stored in a thermostatic bath at 25ºC at least one hour before
484
the
485
spectrophotometers (Table S1). First the absorbance measurements of the thermostated sample
486
were directly performed, and then 225 µL of stock solution of PbCl2 (1.1 mM) were added to the
487
sample and the absorbance measurements were performed again. To test the precision of the
488
spectrophotometer, six CO32- samples of open ocean surface water were analysed, resulting in a
489
standard deviations of 1.9 µmol·kg-1 (precision 0.9%) for the Perkin Elmer spectrophotometer
490
when [CO32-] = 210 µmol·kg-1, of 2.9 µmol·kg-1 for the Shimadzu UV 2401 spectrophotometer
491
when [CO32-] = 230 µmol·kg-1 (precision 1.3%) and of 2.2 µmol·kg-1 for the Shimadzu UV-2600
492
spectrophotometer when [CO32-] = 194 µmol·kg-1 (precision 1.13%).
493
analysis.
Then
absorbance
measurements
were
performed
with
double
beam
Table S1
CAIBOX/OVIDE
CO32Spectrophotometer
Perkin Elmer λ 800
Precision
(µmol·kg-1)
1.9
CAIBOX/MOC2
pH
Spectrophotometer
Shimadzu UV 2401
MOC2
Shimadzu UV 2401
2.9
OVIDE
Perkin Elmer λ 800
HOTMIX
Shimadzu UV-2600
2.2
HOTMIX
Shimadzu UV-2600
Cruise
Cruise
494
pH sampling and measurements: Unfiltered seawater samples were directly taken from the
495
Niskin bottles into special optical glass spectrophotometric Hellma cells of 100 mm of path
496
length. These cells were carefully stored in a thermostatic bath at 25ºC around one hour before
23
497
the analysis. Then, the absorbance measurements were performed with double beam
498
spectrophotometers (Table S1). The methodology consists of measuring the absorbance values at
499
two wavelength (λ = 434 and 578 nm) before and after the addition of 75 µL of m-cresol purple
500
(mCP; ~ 0.2 mM) to the seawater sample. The parameterization of the molar absorptivity ratios
501
needed to determine pH from absorbance ratio (ΔR) was done using Kodad mCP by Clayton and
502
Byrne (1993) and the mCP used in this study was provided by Sigma-Aldrich. The small effect
503
over ΔR due to the dye was evaluated in each cruise (Clayton and Byrne, 1993; Dickson et al.,
504
2007). Recently, Yao et al. (2007) showed that the impurities in the indicator dye of different
505
manufacturers cause uncertainty in the measured pH values. Thus, the equation provided by Yao
506
et al. (2007) was also applied to harmonize the spectrophotometric pH measurements (pHmeas)
507
with the Clayton and Byrne (1993) parameterizations. The equation SI.1 shows the pH values
508
used in this work.
509
pH= pHΔR – [0.0010 + 0.0008·(pHmeas –7.2) + 0.0042·(pHmeas–7.2)2]
(S1)
510
A more recent parameterizations (Liu et al., 2011) using purified mCP allowed us to estimate the
511
effect of impurities of Sigma-Aldrich mCP used in the shipboard pH measurements here
512
reported. The pH measured using Sigma-Aldrich mCP are in average 0.004 ± 002 (N = 8) higher
513
than those using purified mCP and the parameterizations reported by Lui et al. (2011) in the pH
514
range between 7.5 to 8.15, which are within our estimate of the uncertainty of pH. This bias
515
leads to a small effect (+ 1.2 µmol·kg-1) in the computed [CO32-] from pH and AT.
516
AT sampling and measurements: Unfiltered seawater samples were directly taken from the
517
Niskin bottles to 600 mL borosilicate glass bottles (Chanson and Millero, 2007). Sampling
518
bottles were washed twice with sample before filling the bottle from the bottom using a silicone
519
pipe, overflowing half the equivalent volume of the bottle, and immediately stoppered. The
24
520
samples were stored for at least 24 hours before the analyses. Measurements of AT were done by
521
a one endpoint method using AT was measured using an automatic potentiometric titrator. A
522
gravimetrically calibrated Knudsen pipette was used to transfer the seawater samples from the
523
borosilicate glass bottles to an open Erlenmeyer flask in which the potentiometric titration was
524
carried out with HCl (0.1 M) to a final pH to 4.40 (Pérez and Fraga, 1987). In order to estimate
525
the accuracy of the AT method, Certified Reference Materials (CRMs; distributed by A.G.
526
Dickson from the Scripps Institution of Oceanography; batches 84, 99, 108, 118 in CAIBOX,
527
MOC2, HOTMIX and OVIDE, respectively) analysis were also performed.In addition, an extra
528
calibration (substandard) was made by using a closed container of 50 L of filtered and nutrient
529
exhausted open ocean surface water in each cruise. This substandard was analysed to assess
530
possible daily drifts of the electrode and to double check the CRM single analysis measurements.
531
Two replicates of each AT sample, CRM and substandard were measured obtaining a standard
532
deviation not greater than 2 µmol·kg-1 between replicates.
533
Perturbation CO32- measures: During the measurement of the [CO32-] small changes in the
534
[CO32-] of the sample may occur. Byrne and Yao (2008) and Easley et al. (2013) found no
535
discernible perturbation. However, Fajar (2013) detected a small acidification of the samples
536
(0.01 pH units) that could imply a change in [CO32-] between 1.8 and 4 (0.4 %) depending of
537
the sample pH. The parametrizations of Easley et al. (2013) include the effect of this perturbation
538
because they fit the parameters of equation (1) using the absorbance measurements and the
539
[CO32-]calc from pH and AT. Any further changes in the set of equations and constants used to
540
compute the [CO32-]calc from pH and AT should require the reformulation of the parameters of the
541
equation (1). In fact, the universal ratio of boron to salinity has recently increased in 4%, which
542
implies changes in the [CO32-]calc less than 1µmol·kg-1. Besides, the use of the new equilibrium
25
543
constants of the carbonic acid of Millero et al. (2006) would lead to low changes (0.4 %) in the
544
[CO32-]calc.
545
SUPPLEMENTARY GRAPHS
Figure S1.- Vertical distributions (Y axis is depth in meters, X axis is longitude or latitude depending on
the cruise) for measured pH on the total scale at 25ºC (A, C, E, G) and measured total alkalinity in
µmol·kg-1 (B, D, F, H) during the HOTMIX, CAIBOX, OVIDE and MOC2 cruises.
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De
ep
in
m
et
er
s
Figure S2.- Excess of the [CO32-] over the aragonite saturation at in situ conditions in µmol·kg-1 along
the MOC2 cruise.
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