Temporal changes in surface partial pressure of carbon dioxide and carbonate saturation state in the eastern equatorial Indian Ocean during the 1962 – 2012 period

Information on changes in the oceanic carbon dioxide (CO2) concentration and air–sea CO2 flux as well as on ocean acidification in the Indian Ocean is very limited. In this study, temporal changes of the inorganic carbon system in the eastern equatorial Indian Ocean (EIO, 5 N–5 S, 90–95 E) are examined using partial pressure of carbon dioxide (pCO2) data collected in May 2012, historical pCO2 data since 1962, and total alkalinity (TA) data calculated from salinity. Results show that sea surface pCO2 in the equatorial belt (2 ◦ N–2 S, 90–95 E) increased from ∼ 307 μatm in April 1963 to ∼ 373 μatm in May 1999, ∼ 381 μatm in April 2007, and∼ 385 μatm in May 2012. The mean rate of pCO2 increase in this area (∼ 1.56 μatm yr ) was close to that in the atmosphere (∼ 1.46 μatm yr). Despite the steady pCO2 increase in this region, no significant change in air–sea CO2 fluxes was detected during this period. Ocean acidification as indicated by pH and saturation states for carbonate minerals has indeed taken place in this region. Surface water pH (total hydrogen scale) and saturation state for aragonite (arag), calculated from pCO2 and TA, decreased significantly at rates of−0.0016± 0.0001 and −0.0095± 0.0005 yr, respectively. The respective contributions of temperature, salinity, TA, and dissolved inorganic carbon (DIC) to the increase in surface pCO2 and the decreases in pH and arag are quantified. We find that the increase in DIC dominated these changes, while contributions from temperature, salinity, and TA were insignificant. The increase in DIC was most likely associated with the increasing atmospheric CO2 concentration, and the transport of accumulated anthropogenic CO2 from a CO2 sink region via basin-scale ocean circulations. These two processes may combine to drive oceanic DIC to follow atmospheric CO2 increase.


Introduction
Over the past decade, the global ocean took up atmospheric carbon dioxide (CO 2 ) at a rate of about 2.5 Pg C yr −1 (1 Pg = 10 15 g), roughly a quarter of all the anthropogenic CO 2 released from fossil fuel burning, cement production, and land-use change (Le Quéré et al., 2014).The oceans' uptake of atmospheric CO 2 plays an important role in slowing down the increase of atmospheric CO 2 (Sabine et al., 2004;Takahashi et al., 2009) and hence the global climate change.Therefore, it is important to accurately document changes of the oceanic CO 2 sink in order to accurately project future atmospheric CO 2 levels and global climate change (Takahashi and Sutherland, 2013).
The oceanic CO 2 sink is mainly controlled by the gradient of the partial pressure of CO 2 (pCO 2 ) between the atmosphere and the ocean ( pCO 2 = pCO 2 water − pCO 2 air ).
Published by Copernicus Publications on behalf of the European Geosciences Union.

L. Xue et al.: Temporal changes in surface pCO 2 and carbonate saturation state
Considering that the spatiotemporal variability of atmospheric pCO 2 is much smaller than that of the surface water pCO 2 , one could assume that the magnitude of pCO 2 and hence the net air-sea CO 2 flux are governed primarily by oceanic pCO 2 (Takahashi and Sutherland, 2013).While atmospheric CO 2 levels increased from ∼ 280 ppm (parts per million by volume) in the preindustrial era to ∼ 393 ppm in 2012 almost homogeneously across the globe (Tans and Keeling, 2013), oceanic pCO 2 showed different rates of change depending on local oceanographic processes (e.g., lateral mixing, upwelling, and biological activity) (Takahashi et al., 2006).For instance, Le Quéré et al. (2009) found that the pCO 2 increase rate in the North Atlantic was larger than that in the atmosphere during the 1981-2007 period, while the pCO 2 increase rate in the North Pacific was smaller than that in the atmosphere.Even a decrease in sea surface pCO 2 was observed in the vicinity of the Bering and Okhotsk seas between 1970 and 2004 (Takahashi et al., 2006).Consequently, the oceanic CO 2 sink also exhibited different trends.For example, a decrease in the sink for atmospheric CO 2 was observed in the North Atlantic subpolar gyre (50-70 • N, 80-10 • W) from 1982 to 1998 (Lefèvre et al., 2004), whereas a large increase of the CO 2 sink was found in the western tropical North Atlantic (19-20 • N, 65-68 • W) from 2002 to 2009 (Park and Wanninkhof, 2012).It is therefore crucial to determine changes in sea surface pCO 2 in local regions to better understand changes and variability in the global oceanic carbon sink.
While effectively alleviating the global climate change, the oceans' uptake of atmospheric CO 2 is taking a toll on the world's oceans (Doney et al., 2009).It causes decreasing pH, carbonate ion concentrations, and carbonate mineral saturation states, a process commonly termed "ocean acidification" (OA) (Caldeira and Wickett 2003;Feely et al. 2004;Orr et al. 2005).The pH decline can induce speciation shifts of major and minor elements in seawater, affecting their bioavailability to phytoplankton (Doney et al., 2009).The decrease of calcium carbonate (CaCO 3 ) saturation state could affect the ability of marine calcifying organisms to form their CaCO 3 shells and skeletons.
The Indian Ocean, strongly influenced by seasonal monsoonal forcing (Schott and McCreary, 2001), is a unique basin with highly variable oceanic circulation and multiscale air-sea interaction processes (Schott et al., 2009).It plays an important role in the global biogeochemical cycling of carbon and nutrients (e.g., Wiggert et al., 2009).Much effort has been devoted to this region to understand distributions of inorganic carbon parameters and their controlling processes as well as the region's role in the global carbon cycle.Activities have been primarily propelled by several national or international programs, such as the Joint Global Ocean Flux Study (JGOFS) and the World Ocean Circulation Experiment (WOCE) (e.g., Naqvi, 2002;Sabine et al., 2002;Bates et al. 2006).
Nevertheless, the oceanic pCO 2 change and its influence on the oceanic CO 2 sink in the Indian Ocean are far from well documented when compared to those in the Pacific and Atlantic oceans (Takahashi et al., 2009;Lenton et al., 2012;Fay and McKinley, 2013;Sarma et al., 2013).Furthermore, many of the existing studies in the region are based on model results (e.g., Sarma et al., 2013;Valsala and Maksyutov, 2013), which may not well reflect the real situation.For instance, a recent study, based on ocean biogeochemical models, suggested a small enhancement in the CO 2 sink during the 1990-2009 period in the southern Indian Ocean (Sarma et al., 2013), even though Metzl (2009) observed that oceanic pCO 2 increased at a rate of 2.11 ± 0.07 µatm yr −1 during the 1991-2007 period in the southwestern Indian Ocean and the corresponding Antarctic sector (implying a reduction in the oceanic CO 2 sink).It is worthwhile to conduct more studies using in situ observations in the Indian Ocean to better understand the change of surface pCO 2 and the oceanic CO 2 sink in the Indian Ocean.
Furthermore, surface water pH in the Indian Ocean is relatively low, when compared to other oceans (Takahashi and Sutherland, 2013).This, combined with the fact that corals are widely distributed in the region (Allen and Adrim, 2003), makes the Indian Ocean one of the most vulnerable regions in terms of OA.However, relatively little information on OA is available in the Indian Ocean.It is urgent to examine the changes of pH and CaCO 3 saturation state in the Indian Ocean and their potential influence on marine organisms.
In this paper, we report an oceanic pCO 2 increase in the eastern equatorial Indian Ocean (EIO, Fig. 1) for the first time, detected using pCO 2 data collected during May 2012, together with historical data since 1962 as integrated by Takahashi et al. (2013) (i.e., LDEO_Database_V2012).We examine temporal changes in air-sea CO 2 flux and OA indicators (pH and aragonite saturation state, arag ), and explore the factors responsible for the pCO 2 increase and OA.

Study site
The study region was located to the west of Sumatra, covering the area from 5 • N to 5 • S and 90 to 95 • E (Fig. 1).The local climate is characterized by the seasonal monsoon, which exhibits a weak westerly annual mean wind and has a strong semiannual cycle.Consequently, the equatorial currents are quite unique and different from those in the other equatorial oceans (Schott et al., 2009).The ocean currents are mainly in the zonal direction and the most distinguished are the strong surface eastward flows, known as Wyrtki jets (Wyrtki, 1973), which occur during two intermonsoon periods in spring (April-May) and fall (October-November).There is no climatological equatorial upwelling due to lack of steady equatorial easterlies.Another important current is the Equatorial Undercurrent (Knauss and Taft, 1964), which exists at the thermocline depth and occurs mainly during the later winter to spring.The strong equatorial zonal currents link the EIO with the western equatorial Indian Ocean, through which it joins the basin-scale circulation.

Data sources
Data in 2012 were from the Monsoon Onset Monitoring and its Social and Ecosystem Impact (MOMSEI) project cruise conducted during the period 1-9 May 2012.During this cruise, sea surface pCO 2 was continuously measured every 15 min with a HydroC ™ CONTROS sensor.The CO 2 mole fraction in the headspace behind a membrane equilibrator was measured using a two-wavelength nondispersive infrared detector (NDIR).The equilibrator consists of a flat silicone composite membrane, and additional sensors for pressure, temperature, and relative humidity measurements.Regular zeroings are automatically performed to correct instrument drift with time by scrubbing CO 2 from the internal gas stream (Saderne et al., 2013).More details on pCO 2 measurements can be found in Saderne et al. (2013).Fietzek et al. (2013) presented the detailed in situ calibration of the data, and asserted that the average difference between sensor reading and reference pCO 2 was −0.6 ± 3.0 µatm with a root-mean-square error (RMSE) of 3.7 µatm.Before the cruise in May 2012, a comparison study between this sen-sor and the Apollo CO 2 instrument (e.g., Jiang et al., 2008) indicated an accuracy of better than 5 µatm (see Fig. s1 in the Supplement).During the MOMSEI cruise, sea surface temperature (SST) and salinity (SSS) data were also collected every 15 min using a SBE 21 Seacat thermosalinograph.Surface water pCO 2 and its associated parameters (temperature and salinity) before 2012 were extracted from the Lamont-Doherty Earth Observatory (LDEO) Database (version V2012; Takahashi et al., 2013) (Tables 1 and 2).The pCO 2 results of the LDEO database are based on measurements made using air-seawater equilibration methods (Takahashi et al., 2013).All data points have been individually quality controlled before they were integrated into this database.The uncertainty of the pCO 2 data is estimated to be about ± 2.5 µatm on average, given differences in equilibrator designs, calibration methods, and some interpolated parameters (Takahashi et al., 2013).
In addition, several ancillary parameters including wind speed, atmospheric CO 2 concentration, mixed layer depth (MLD), SSS and chlorophyll a (Chl a) (Table 2) were used.Because long-term atmospheric CO 2 data since 1962 are not available in the Indian Ocean, we used atmospheric CO 2 data measured as the mole fraction in dry air at Mauna Loa, Hawaii (Table 2), as an alternative.Some limited observations collected during the 2004-2010 period at the BKT (Bukit Kototabang, Indonesia; Table 2) atmospheric CO 2 station, close to the study area (Fig. 1), showed that the atmospheric CO 2 level at Mauna Loa was at least 4.3 ppm larger than that at BKT, Indonesia (see Fig. s2).Therefore, we  a This mixed layer depth is in fact a temperature-mixed layer depth, or isothermal layer depth.Mixed layer depths are computed as the depth with a 0.2 • C absolute temperature difference from 10 m temperature (Keerthi et al., 2013).b See specific cruise information in Table 1.
corrected the atmospheric CO 2 concentration from Mauna Loa by subtracting 4.3 ppm.The atmospheric CO 2 mole fraction was then converted to pCO 2 by correcting to 100 % humidity at the mean SST and SSS during the investigation period, following Jiang et al. (2008).

Spatial averaging and seasonal correction
We first grouped all the data points (temperature, salinity, and pCO 2 ) into their individual 0.1 • latitudinal bands, then calculated the average for each band, and took the average of all mean values from each band as cruise mean.Considering that very few of the cruises covered the entire study area well (Fig. 1), we just used the data in the equatorial belt (2 • N-2 • S, 90-95 • E) to determine the trends of SST, SSS, and pCO 2 .We also corrected the effects of the seasonal cycle by using the climatological data from Takahashi et al. (2009) (Fig. s3) before the trend analysis.

Calculation of air-sea CO 2 fluxes
We calculated the air-sea CO 2 flux based on Eqs. ( 1) and ( 2): where F is the air-sea CO 2 flux (mmol m −2 d −1 ), in which a positive value represents CO 2 releasing from the ocean to the atmosphere; k is the gas transfer velocity (cm h −1 ), calculated using Eq. ( 2) based on the parameterization of gas transfer velocity with wind speed proposed by Wanninkhof (1992) and recently revised by Ho et al. (2011); s is the solubility coefficient of CO 2 (mol L −1 atm −1 ) (Weiss, 1974); and pCO 2 water and pCO 2 air are the pCO 2 in the surface ocean and in the atmosphere (µatm), respectively.In Eq. ( 2), U 10 (m s −1 ) is the wind speed at a height of 10 m above the sea surface.We used monthly mean wind speeds from the National Centers for EnvironmentalPrediction (NCEP) with a spatial resolution of 2.5 • × 2.5 • (http://www.esrl.noaa.gov/psd/data/gridded/data.ncep.html)for calculating of CO 2 fluxes.Sc is the Schmidt number, which was calculated based on the formula of Wanninkhof (1992).This relationship produces an RMSE of 1.7 µmol kg −1 , which was less than that produced by the Lee et al. ( 2006) formula (±10.6 µmol kg −1 ).

Estimate
The DIC (dissolved inorganic carbon), pH, and CaCO 3 saturation state were calculated from pCO 2 , TA, temperature, and salinity using the CO2sys program (Lewis and Wallace, 1998) and adopting the CO 2 system coefficients of Mehrbach et al. (1973) as refitted by Dickson and Millero (1987).Because the saturation state of calcite is usually about 50 % greater than that of aragonite at 25 • C (Mucci, 1983), hereafter we only discuss the results for arag .

Respective contribution of temperature, salinity, TA, and DIC to the changes of pCO 2 , pH and arag
We used the method of Wakita et al. (2013) to quantify the contributions of temperature, salinity, TA, and DIC to changes of pCO 2 , pH, and arag .For example, the contributions of these properties to pCO 2 change can be expressed as the sum of the individual contributions as follows: + (αpCO 2 /αT A) TA + (αpCO 2 /αDIC) DIC, where T and S are SST and SSS, respectively; and T , S, DIC, and TA denote the changes in SST, SSS, DIC, and TA, respectively.
During the calculation, we evaluated the rate of pCO 2 change by allowing one parameter to vary while using mean values for the other parameters.For example, we estimated the contribution of DIC change to pCO 2 ((αpCO 2 / αDIC) DIC) by calculating pCO 2 using the DIC trend from 1962 to 2012 and mean values for the other parameters, and then we computed the impact of DIC change on pCO 2 .The contributions of these properties to pH and arag changes were calculated similarly.

Temporal change of surface water pCO 2
The in situ sea surface pCO 2 data in the EIO starting from 1962 are shown in Figs. 3 and 4. Spatially, pCO 2 distributions were relatively homogeneous in each cruise, with a standard deviation of less than 9 µatm (Fig. 3).A gradual increase in sea surface pCO 2 with time is the most evident feature (Figs. 3, 4).The mean value of sea surface pCO 2 in the equatorial belt (2 • N-2 • S, 90-95 • E) increased from ∼ 307 µatm in April 1963 to ∼ 373 µatm in May 1999, ∼ 381 µatm in April 2007, and ∼ 385 µatm in May 2012 (Fig. 4, Table 1).After seasonal correction, we find that sea surface pCO 2 increased at a mean rate of 1.56 ± 0.08 µatm yr −1 from 1963 to 2012 (Fig. 4a).
For comparison purposes, we also estimated the surface pCO 2 trend along the Equator from 89.5 to 94.5 • E, and found that since the International Indian Ocean Expedition (IIOE, 1960(IIOE, -1965)), surface pCO 2 increased from 307 ± 4 µatm in April 1963 to 392 ± 6 µatm in May 2012 (Fig. 4b), with a mean rate of ∼ 1.64 µatm yr −1 (after seasonal correction).This rate of surface pCO 2 increase is not significantly different from that obtained using all data measured during the 1962-2012 period (Fig. 4a).The pCO 2 increase rate in the EIO is similar to that in the subtropical North Atlantic (Bates, 2007), higher than that in the western tropical North Atlantic (Park and Wanninkhof, 2012), and lower than that in the equatorial Atlantic and Pacific (Oudot et al., 1995;Takahashi et al., 2006) (Table 3).

Temporal changes of air-sea CO 2 flux, pH, and arag
The increase in surface pCO 2 could potentially affect air-sea CO 2 flux by changing the gradient of air-sea pCO 2 (Eq.1).  a "sink" means that the ocean absorbs atmospheric CO 2 , while "source" indicates that the ocean releases CO 2 to the atmosphere.b The growth rate here refers to the mean difference between 1982 and 1992.
The mean rate of pCO 2 increase in the EIO from 1962 to 2012 (∼ 1.56 µatm yr −1 ) was only slightly greater than that in the atmosphere (∼ 1.46 µatm yr −1 ; Fig. 4a), suggesting a weak trend of increase for the air-sea CO 2 gradient (Fig. 5a).Equation ( 1) showed that the air-sea CO 2 flux could also be affected by the gas transfer velocity, which is typically expressed as a power function of wind speed (e.g., Wanninkhof, 1992).Figure 5 showed no obvious changes in wind speed.All factors considered, air-sea CO 2 fluxes showed a weak but insignificant increasing trend during the 1962-2012 period (Fig. 5).
OA has taken place in this region during the past 50 yr, as indicated by pH and arag .Surface water pH (in the total hydrogen scale) decreased significantly at a rate of  −0.0016 ± 0.0001 yr −1 from 1962 to 2012 (Fig. 6), similar to the rate observed at the time-series stations of Bermuda (BATS) and Hawaii (HOT) (Takahashi and Sutherland, 2013).An average decline of 0.08 pH units over the past 50 yr  in the EIO is surprising, considering that the average surface ocean pH has just declined by about 0.1 units since the 1700s due to the absorption of anthropogenic CO 2 (Raven et al., 2005).A rapid reduction in arag with a rate of −0.0095 ± 0.0005 yr −1 in this region was also observed (Fig. 6).This reduction rate is faster than that in the subsurface waters in the subarctic western North Pa-cific Ocean (−0.004 to −0.005 yr −1 ) (Wakita et al., 2013), and similar to that in the southern California Current system (0.009 ± 0.006 yr −1 ) (Leinweber and Gruber, 2013) and in the North Atlantic Ocean (−0.0100 ± 0.0012 yr −1 ) (Bates et al., 2012).the increase in surface pCO 2 and decreases in pH and arag .

Contributions
The results show that the DIC increase played the most portant role in elevating surface water pCO 2 and decreasing pH and arag .In contrast, the contributions from temperature, salinity, and TA were insignificant (Fig. 8).These results (Fig. 8) are in good agreement with the trends of temperature, salinity, TA, and DIC (Fig. 7).From 1962 to 2012, the changes of temperature, salinity, and TA were not significant, while DIC increased significantly during the 1962-2012 period (Fig. 7).In addition, the good consistency between the sum of the decomposed individual contributions (Tot) and the observed trend (Obs) verified the robustness of the method of Wakita et al. (2013) in our study (Fig. 8).

Air-sea CO 2 exchange
It is obvious that the DIC increase with time was not due to local uptake of CO 2 via air-sea exchange, given that the EIO was almost always a CO 2 source to the atmosphere (Fig. 5, Bates et al., 2006;Takahashi et al., 2009).However, the rapidly rising atmospheric CO 2 since 1962 created the potential to reduce or even reverse the CO 2 release in this region from the ocean to the atmosphere (Fig. 4), and can directly induce DIC increase in the mixed layer.Figure 9 gives a schematic of DIC increase in the mixed layer induced by   From time t 0 to t when atmospheric CO 2 increases, the driving force of air-sea exchange (pCO 2 water -pCO 2 air ) and correspondingly the CO 2 outgassing from the ocean to the atmosphere would decrease.This may induce DIC increase in the mixed layer due to reduction in the magnitude of CO 2 source (Schneider et al., 2012).F denotes air-sea CO 2 fluxes, k the gas transfer velocity, s the solubility coefficient of CO 2 , and pCO 2 water and pCO 2 air are the pCO 2 in the surface ocean and in the atmosphere.Details on air-sea CO 2 fluxes can be found in Sect.2.3.H and D are mixed layer depth and seawater density, respectively.
atmospheric CO 2 increase in a CO 2 source region with respect to the atmosphere.When atmospheric CO 2 concentration increases, CO 2 outgassing in this region from the ocean to the atmosphere would be reduced or even reversed, and would increase DIC concentration in the mixed layer.Thus, oceanic DIC increase in a CO 2 source region to the atmosphere could be caused via reduction in the magnitude of CO 2 source (Schneider et al., 2012)  See Table 2 for more details on data sources.All data are from the same month (April).
was directly absorbed by the ocean from the atmosphere due to higher atmospheric CO 2 concentrations.

Ocean circulation
Transport via basin-scale ocean circulation also contributed to DIC increase.This can be verified by the high contents of anthropogenic CO 2 in the mixed layer of the EIO (e.g., Sabine et al., 1999;Sabine et al., 2004), which has no direct uptake of atmospheric CO 2 (Fig. 5, Bates et al., 2006;Takahashi et al., 2009).The observed increase in anthropogenic CO 2 in the EIO is likely due to accumulation of anthropogenic CO 2 in CO 2 sink regions and subsequent transport to the equatorial belt via basin-scale ocean circulation.For instance, carbon in the region between 15 and 50 • S in the Indian Ocean could be finally transported to the EIO.On the one hand, this region (between 15 and 50 • S) is a major subduction zone (Schott et al., 2009), serves as a significant sink of atmospheric CO 2 (Fig. 1a, Takahashi et al., 2009;Valsala et al., 2012), and hosts the largest inventories of anthropogenic CO 2 across the Indian Ocean (Sabine et al., 2004).Furthermore, it is reported that the oceanic increase in carbon storage roughly kept pace with atmospheric CO 2 increase (Sabine et al., 1999).
On the other hand, the upper ocean horizontal circulation and the meridional overturning cells are believed to account for carbon transport towards the EIO (Fig. 1; Schott et al., 2002;Schott et al., 2009;Valsala et al., 2012).In the upper layer, water masses move westward in the South Equatorial Current (SEC) and partly merge into the East African Coast Current (EACC); they can move further to the EIO along with the eastward flows, including the South Equatorial Countercurrent (SECC), the Wyrtki jets (Wyrtki, 1973), and the Equatorial Undercurrent (EUC) (Knauss and Taft, 1964).At the thermocline depth, shallow overturning cells, including the cross-equatorial cell (CEC) and the Southern Hemisphere subtropical cell (STC), bring water masses from the subduction zone to the off-equatorial upwelling zone, where they upwell to the surface.Thus, the southern subtropical Indian Ocean may act as a window for carbon uptake (Valsala et al., 2012), and the basin-scale circulation provides the route to transport the absorbed anthropogenic CO 2 ultimately to the EIO.A similar mechanism was also proposed in the equatorial Pacific by Feely et al. (1999), who pointed out that entrained subtropical water was injected into upwelled water at the Equator.
Nevertheless, the pathways of ocean circulation are very complicated, and there must be other ways to increase DIC in the equatorial Indian Ocean.For instance, the Red Sea-Persian Gulf Intermediate Waters formed in the northwestern Indian Ocean carry anthropogenic CO 2 signals and spread equatorward (Sabine et al., 2004;Alvarez et al., 2009, and references therein), which also contributes to the increase of equatorial waters' DIC.Overall, ocean circulation may play an important role in transporting carbon accumulated in the CO 2 sink region to the equatorial belt on a basin scale.To a large extent, this could account for the paradox that the increase in anthropogenic CO 2 occurred in the CO 2 source region, where CO 2 was emitted to the atmosphere from the ocean.

Vertical mixing and biological activity
Vertical mixing with deep waters rich in CO 2 can elevate surface DIC content.For instance, the enhanced vertical mixing, usually accompanied by salinity rise and MLD deepening, will bring more CO 2 -rich waters to the surface layer, leading to higher DIC levels (e.g., Takahashi et al., 2006;Dumousseaud et al., 2010).However, there was no significant trend for SSS and MLD during the study period (Figs. 7,10), indicating an insignificant influence of vertical mixing on DIC changes.
Biological activity could also affect DIC (e.g., Zhang et al., 2010).Model studies indicate that there were no significant changes in net primary production, particle export and export efficiency from 1960 to 2006 in this region (Laufkötter et al., 2013).Satellite data also do not show a significant trend for Chl a (a proxy for biological activity) from 1998 to 2012 (Fig. 10).Therefore, biological activity was not the main factor leading to DIC rise, either.

Summary and conclusions
We compiled sea surface pCO 2 data in the eastern equatorial Indian Ocean over the past 50 yr to document the longterm trends in the inorganic carbon system.Results show that sea surface pCO 2 in the EIO increased significantly from 1962 to 2012.The mean rate of oceanic pCO 2 increase (∼ 1.56 µatm yr −1 ) was close to the rate of atmospheric CO 2 increase (∼ 1.46 µatm yr −1 ).Despite the steady pCO 2 increase in this region, no significant change in air-sea CO 2 fluxes during this period is detected.Ocean acidification as indicated by pH and arag did take place during this study period.Surface pH (total hydrogen scale) and arag decreased significantly at rates of −0.0016 ± 0.0001 and −0.0095 ± 0.0005 yr −1 , respectively.The increase in DIC is the main driver contributing to the increase in surface pCO 2 and decreases in pH and arag .The increase in DIC was most likely associated with the increasing atmospheric CO 2 concentration, and the transport of accumulated anthropogenic CO 2 from a CO 2 sink region via basin-scale ocean circulations.These two processes may combine to drive oceanic DIC to follow atmospheric CO 2 increase.
This study improved the understanding of the temporal changes of the inorganic carbon system in the Indian Ocean, demonstrating that big changes in carbon chemistry (e.g., OA) and other marine biogeochemical processes may have taken place in the Indian Ocean since the IIOE.However, the influences of climate events on oceanic pCO 2 and hence the oceanic carbon sink were not taken into account, due to spatial heterogeneity and scarcity of data.Long-term and continuous CO 2 observations with good resolutions are needed in the future in the Indian Ocean, where climate events such as the Indian Ocean Dipole (IOD) and El Niño-Southern Oscillation (ENSO) often occur (e.g., Currie et al., 2013;Valsala and Maksyutov, 2013).
The Supplement related to this article is available online at doi:10.5194/bg-11-6293-2014-supplement.
of TA, DIC, pH and CaCO 3 saturation state The conservative behavior of TA allows us to estimate TA using the salinity data collected.We used data in the upper 20 m collected in April 2007 in this region during the CLIVAR/CO 2 (Climate Variability and Predictability) section I9N cruise (http://cchdo.ucsd.edu) to build the relationship between TA and SSS (Fig. 2) as follows: TA = 75.31(±5.15)× SSS − 332.95(±175.35)(3)(r= 0.94, n = 28, p < 0.0001) .

Figure 9 .
Figure9.A schematic of DIC increase in the mixed layer induced by atmospheric CO 2 increase in a CO 2 source region with respect to the atmosphere.DIC changes in the mixed layer can be attributed to vertical entrainment ( DIC Ent ), vertical and horizontal advection and diffusion ( DIC Adv ), biological activities ( DIC bio ) and air-sea exchange ( DIC as ).From time t 0 to t when atmospheric CO 2 increases, the driving force of air-sea exchange (pCO 2 water -pCO 2 air ) and correspondingly the CO 2 outgassing from the ocean to the atmosphere would decrease.This may induce DIC increase in the mixed layer due to reduction in the magnitude of CO 2 source(Schneider et al., 2012).F denotes air-sea CO 2 fluxes, k the gas transfer velocity, s the solubility coefficient of CO 2 , and pCO 2 water and pCO 2 air are the pCO 2 in the surface ocean and in the atmosphere.Details on air-sea CO 2 fluxes can be found in Sect.2.3.H and D are mixed layer depth and seawater density, respectively.

Table 1 .
Summary of cruise information and mean values of surface temperature, salinity, and surface pCO 2 in the equatorial belt (2 • N-2 • S, 90-95 • E), which are reported as mean ± standard deviation.These mean values in this table are calculated as described in Sect.2.2.2, and are not deseasonalized.The observation period in this table refers to the period when the data used in this study were collected.bWeusedsalinitydata from the simple ocean data assimilation(SODA)since in situ data are not available during these cruises (http://coastwatch.pfeg.noaa.gov/erddap/griddap/hawaii_d90f_20ee_c4cb.graph).c We did not use data from this cruise to determine the trend due to the poor coverage latitudinally.

Table 2 .
Sources of data used in this study.