Carbonate system biogeochemistry in a subterranean estuary ... · Carbonate system biogeochemistry in a subterranean estuary – Waquoit Bay, USA Qian Liua,b, Matthew A. Charetteb,⇑,

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Geochimica et Cosmochimica Acta 203 (2017) 422–439

Carbonate system biogeochemistry in a subterraneanestuary – Waquoit Bay, USA

Qian Liu a,b, Matthew A. Charette b,⇑, Crystaline F. Breier b, Paul B. Henderson b,Daniel C. McCorkle c, William Martin b, Minhan Dai a

aState Key Lab of Marine Environmental Science, Xiamen University, Xiamen, Fujian, ChinabDepartment of Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution, Woods Hole, MA, USA

cDepartment of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, MA, USA

Received 24 March 2016; accepted in revised form 23 January 2017; available online 3 February 2017

Abstract

Quantifying carbon fluxes associated with submarine groundwater discharge (SGD) remains challenging due to the com-plex biogeochemistry of the carbonate system in the subterranean estuary (STE). Here we conducted time series measurementsof total alkalinity (TAlk) and dissolved inorganic carbon (DIC) in a well-studied coastal aquifer (Waquoit Bay, Mas-sachusetts, USA). Groundwater samples were collected monthly from May 2009 to June 2010 across the freshwater-saltwater mixing zone of the Waquoit Bay (WB) STE. The concentrations of both TAlk and DIC in zero-salinity groundwaterwere variable, but were lower than those in the bay water (S � 28). DIC underwent slightly non-conservative mixing betweenlow and intermediate salinities while there was an apparent additional DIC source at high salinity (>20) in all seasons. TAlkconcentrations exhibited even stronger variations, with evidence of both production and consumption in high salinity zones,and consistent TAlk consumption at intermediate salinity in summer and fall (June–December, 2009). The increases in DICand TAlk at high salinity were attributed to aerobic respiration and denitrification in WB sediments during bay waterrecharge of the STE. We infer that the loss of TAlk at intermediate salinity reflects H+ production as reduced compounds(e.g. Fe2+) are oxidized within the STE. In terms of impacts on surface water inorganic carbon budgets, the SGD-derivedDIC flux was mainly controlled by seasonal changes in SGD while a combination of TAlk concentration variability andSGD drove the TAlk flux. SGD-derived DIC, aqueous CO2, and H+ fluxes to the bay were �40–50% higher in summervs. in winter, a result of enhanced marine groundwater flux and significant TAlk removal (proton addition) during periodsof high seawater intrusion. Furthermore, the SGD-derived DIC flux was consistently greater than TAlk flux regardless of sea-son, indicating that SGD serves to reduce the CO2 buffering capacity of surface water. Our results highlight the importance ofseasonality and subsurface biogeochemical processes on the subterranean estuary carbonate system and the resulting impacton SGD-derived TAlk, DIC, aqueous CO2, and H+ fluxes to the coastal ocean.� 2017 Elsevier Ltd. All rights reserved.

Keywords: Groundwater; Subterranean estuary; Dissolved inorganic carbon; Alkalinity; Ocean acidification

http://dx.doi.org/10.1016/j.gca.2017.01.041

0016-7037/� 2017 Elsevier Ltd. All rights reserved.

Abbreviations: SGD, submarine groundwater discharge; STE,subterranean estuary; DIC, dissolved inorganic carbon; TAlk, totalalkalinity; WB, Waquoit Bay⇑ Corresponding author. Fax: +1 508 457 2193.E-mail address: [email protected] (M.A. Charette).

1. INTRODUCTION

Submarine groundwater discharge (SGD) has beenincreasingly recognized as a significant carbon source tothe coastal ocean (Lee and Kim, 2015a; Liu et al., 2012,2014; Moore, 2010; Moore et al., 2011; Santos et al.,2011; Reckhardt et al., 2015; Wang et al., 2015; Yang


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Q. Liu et al. /Geochimica et Cosmochimica Acta 203 (2017) 422–439 423

et al., 2015) thereby playing a role in the marine carbonatesystem. In some coral reef lagoons, wetlands and mangrovesystems, the SGD flux has been shown to dominate estuar-ine CO2 dynamics (Santos et al., 2012; Maher et al., 2013;Cyronak et al., 2014). SGD has also been reported as amain source of acidity contributing to coastal acidificationin both coral reef and wetland ecosystems (de Weys et al.,2011; Santos et al., 2011; Wang et al., 2014). In the coastalmargin of a volcanic island, however, seawater pH actuallyincreased due to enhanced biological productivity sup-ported by SGD-derived nutrients (Lee and Kim, 2015a).Together, these studies illustrate the complexity and signif-icance of SGD to coastal ocean carbon dynamics.

Similar to river-estuary systems, terrestrial groundwaterundergoes significant modifications in its chemical compo-sition before discharging into the coastal ocean due to mix-ing with seawater and both biotic and abiotic reactions withaquifer sediments. This reaction zone has been termed the‘subterranean estuary’ (STE; Moore, 1999). In contrastwith surface estuaries, however, STEs generally have longerwater residence times, enhanced sediment–water interac-tions, wider redox gradients, and higher microbial activities(Moore, 1999). The CO2 system in the STE has been shownto be significantly modified by diagenetic reactions such ascarbonate dissolution and organic matter remineralization(Cai et al., 2003; Dorsett et al., 2011; Liu et al., 2012; Leeand Kim, 2015b).

Given the STE’s dynamic hydrology and biogeochem-istry, groundwater CO2 system measurements must be col-lected on fine spatial and temporal scales to better evaluatethe relevant biogeochemical reactions and their impacts onthe coastal ocean inorganic carbon cycle. Waquoit Bay,Massachusetts (WB) has a well defined salinity transitionzone occurring over a narrow region (ca. 20 m wide, 1–2 m thick, Charette et al., 2005; Gonneea and Charette,2014) that hosts a suite of biogeochemical processes(Charette and Sholkovitz, 2006; Spiteri et al., 2008). Sea-sonal movement of the salinity transition zone in thisSTE responds to oscillations in the aquifer hydraulic gradi-ent and sea level (Michael et al., 2005; Gonneea et al.,2013a,b). Through year-long monthly carbonate chemistryobservations in the WB STE, we sought to quantify CO2

system dynamics in the context of seasonal hydrogeologicchanges and to determine factors that drive CO2 biogeo-chemistry within three main aquifer zones: terrestrial, inter-mediate salinity, and high salinity groundwater. Finally weestimated the seasonal variability in SGD-derived totalalkalinity (TAlk) and dissolved inorganic carbon (DIC)fluxes to the bay.

2. MATERIALS AND METHODS

2.1. Study site

Waquoit Bay is a shallow semi-enclosed estuary (aver-age water depth 1 m) in Cape Cod (MA, USA, Fig. 1); itexperiences semi-diurnal tides with an average tidal rangeof 1.1 m (Charette et al., 2001). The freshwater contributedto the bay includes rainfall (11%), groundwater discharge(34%), and rivers (55%) (Cambareri and Eichner, 1998).

Importantly, the rivers are largely groundwater fed:approximately 90% of the river flow is supplied by ground-water recharge. Therefore, groundwater is the major fresh-water source to the bay.

The upper unconfined aquifer (10 m) is mainly com-posed of coarse and fine sand as well as gravel (Charetteet al., 2005). Precipitation (114 cm yr�1) has relatively lowseasonal variation (Walter and Whealan, 2005), however,the groundwater table is lower in summer compared withwinter and spring due to evapotranspiration (Michaelet al., 2005). Inland, the maximum annual water table vari-ability can reach as high as 1 m. Multiple measurementtechniques indicate that the most intensive direct SGDsource lies along the northern shoreline of the bay(Charette et al., 2001; Mulligan and Charette, 2006). Here-after, the salinity transition zone will be referred to as the‘‘mixing zone”. The mixing zone oscillates with changes inhydraulic gradient between the aquifer and sea level, whichis mainly driven by the latter in this STE (Gonneea et al.,2013a). A higher hydraulic gradient will lead to seawardmovement of the mixing zone while a lower hydraulic gra-dient will result in landward movement. Our sampling per-iod (May 2009 to June 2010) generally had higher sea levelcompared with previous years based on monthly mean sealevel (MSL) anomaly data from the nearby Woods Holetidal gauge (Fig. 2). The average MSL in summer is�4 cm higher in 2009 vs. 2005 to 2007. This pattern waslikely due to the combined effects of El Nino and a negativemode of the North Atlantic Oscillation (Gonneea et al.,2013a).

2.2. Field sampling

We conducted a time series sampling program forgroundwater and bay water at the head of the bay(Fig. 1). Monthly groundwater profiles were collected fromMay 2009 to June 2010 at PZ3, which is a series of nestedfixed-depth well points that span the mixing zone of subter-ranean estuary. There were two sample collection dates inDecember 2009 separated by �4 weeks (12a/2009 and12b/2009 in Figs. 3–5, 8). In total there are ten well pointsranging in depth from 0.5 to 5.5 m below the beach surface,with intervals of between 0.15 and 1 m. Samples were col-lected during the same tidal cycle and phase (four hourspast high tide, �3 days before the monthly spring tide) toreduce potential variability associated with tidal fluctua-tions. We simultaneously collected surface waters at fourstations equally spaced along the head of the bay,�100 m offshore, with a water depth of � 0.8 m. DuringNovember 2009 and June 2010, we collected two additionalpiezometer profiles at PZ10 (13.5 m inland of PZ3) and PZ5(3 m seaward of PZ3). Piezometer profile PZ10 includednine depths from 0.9 to 5.5 m while PZ5 included sixteendepths ranging from 0.3 to 6.1 m.

The groundwater samples were obtained with a stainlesssteel drive point piezometer system as described in Charetteand Allen (2006). Groundwater was extracted using peri-staltic pumps at 10–50 mL min�1. Groundwater samplingcommenced after purging the wells of �1 L or 6–13 wellvolumes. Ancillary water chemistry data including salinity,

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Fig. 1. Map of Waquoit Bay off Cap Cod in Massachusetts with a schematic figure showing the sampling stations in the surface bay water andsubterranean estuary. The STZ is the salinity transition zone where mixing between terrestrial groundwater and marine groundwater occurs.Monthly sampling was conducted from May 2009 to June 2010 both in the surface water and groundwater at PZ3, which is a series of nested,fixed-depth wells installed at 10 depths ranging from 0.5–5.5 m centered on the STZ of subterranean estuary. In addition, we sampled twomore piezometer wells at PZ10 (inland, 13.5 m away from PZ3 at depths ranging from 0.9 to 5.5 m) and PZ5 (seaward, 3 m away from PZ3, atdepths ranging from 0.3 to 6.1 m) in November 2009 and June 2010.

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Fig. 2. Monthly mean sea level (MSL) anomaly from Woods Hole tidal gauge (NOAA, #8447930). The shaded region indicates our samplingperiod from May 2009 to June 2010.

424 Q. Liu et al. /Geochimica et Cosmochimica Acta 203 (2017) 422–439

temperature, dissolved oxygen (DO), pH, and oxidation–reduction potential (ORP, mV) were measured with a YSI600XLM sonde in a flow-through cell. ORP is a qualitativemeasure of the tendency of solutes to acquire electron andthereby be reduced in aquatic environments. Separate

125 mL samples were collected for salinity analysis. Dis-crete samples for TAlk and DIC were collected in 150 mLborosilicate glass vials as follows: after overflowing 1.5times, water was removed to create a � 3 mL headspace,then poisoned with 50 lL of a saturated HgCl2 solution.


Laboratory tape was wrapped around the cap to maintainan air tight seal. Trace metal samples were filtered into acidcleaned LDPE bottles using a 0.2 lm capsule filter (PallAcropak). Samples were acidified to pH 1–2 with 20 lLof Optima ultrapure trace metal grade nitric acid (8 M).Nutrient samples were similarly filtered into acid washed20 mL polyethylene sample bottles, and stored on ice untiltheir return to the laboratory. A second nutrient vial wasfilled and acidified to �pH 2 with 8 M sulfurous acid toeliminate scavenging of phosphate by precipitation of dis-solved iron prior to analysis.

2.3. Sample analysis

Salinity was analyzed by a Guideline AutoSal instru-ment. DIC and TAlk were measured within two weeks ofcollection. DIC was determined by acidification of 0.5 mLof a water sample and the subsequent quantification ofCO2 with a non-dispersive IR detector (Li-Cor 6252). Theanalytical precision is ±3 lmol kg�1. TAlk was determinedusing Gran titration (Metrohm 808 Titrando with 1 mLburette) with a precision of ±4 lmol kg�1. Both DIC andTAlk were calibrated by certified reference materials fromA.G. Dickson of the Scripps Institution of Oceanography.

From the measured T and S, and DIC and TAlk concen-trations, we used the CO2SYS.XLS v.14 software (Lewisand Wallace, 1998) to derive the concentrations of theDIC components (aqueous CO2, bicarbonate, and carbon-ate), pCO2, and pH (total scale). Dissociation constants forcarbonic acid were those of Millero et al. (2006). The CO2

solubility coefficient was taken from Weiss (1974) and thesulfate dissociation constant was derived from Dickson(1990); both are the default values in the CO2SYS.XLSv.14 program. The calculated calcite or aragonite satura-tion indexes (O) were based on O = [Ca2+][CO3

2�]/Ksp,whereby O < 1 indicates undersaturation for CaCO3,O = 1 represents the equilibrium CaCO3 phase, and O > 1is CaCO3 supersaturation (Ksp is the CaCO3 solubility pro-duct; Mucci 1983). For surface water, Ca2+ was derivedfrom an empirical equation involving measured salinity(Ca2+ (mmol kg�1) = 10.28 � S/35; Millero 2005). Forgroundwater samples, we estimated Ca2+ concentrationsbased on Waquoit Bay STE Ca measurements from thesummer of 2004 (Ca (mmol kg�1) = 0.28 � S + 0.11;Appendix Fig. S1).

Nutrient analyses (nitrate, phosphate, ammonium) wereperformed at the Woods Hole Oceanographic InstitutionNutrient Analytical Facility using standard methods on aLachat QuickChem 8000 Flow Injection Analyzer. Mnand Fe were analyzed on a Finnigan Element 2 high-resolution inductively coupled plasma mass spectrometerat the Woods Hole Oceanographic Institution ICP-MSFacility. Briefly, samples were diluted 20-fold with 5%Optima nitric acid spiked with Indium (In) as an internalstandard to correct for instrument drift and matrix interfer-ences of the solution. Standards were prepared in the samemanner as samples (In-spiked) and six-point standardcurves were used to calculate sample metal concentrations(Charette and Sholkovitz, 2006).

3. RESULTS

3.1. Salinity distribution in the subterranean estuary

A decreasing hydraulic gradient will cause the landwardmovement of the mixing zone (Gonneea et al., 2013a), lead-ing to generally higher salinities within our fixed depthpiezometers. Conversely, lower salinities are observed withan increasing hydraulic gradient. Although the hydraulicgradient from May 2009 to June 2010 was relatively invari-ant compared with previous observations, we still observeda shallowing of the mixing zone (salinity 5–10) in summerand fall (June to December 2009, Fig. 3) relative to winterand spring (January to May 2010, Fig. 3). Salinity anoma-lies in the freshwater portion of the aquifer were observedin June–July 2009 and May–June 2010 (Fig. 3) and werelikely associated with the mixing between tide and waveinduced overtopping of bay water and terrestrial ground-water (Rogers, 2010; Gonneea et al., 2013b).

3.2. Redox processes in the subterranean estuary

Dissolved oxygen saturation ranged from 3% to 90%(median: 9%), indicative of persistent low-oxygen condi-tions in the STE (Fig. 4a), while surface bay water hadnearly saturated and supersaturated DO ranging 70–154%(Appendix Table S1). In the STE, DO saturation was gen-erally lowest below the mixing zone. The high DO (33–90%)layer at 2.5 m was especially pronounced from June toDecember, 2009.

Dissolved Mn from the monthly time series (Fig. 4b)was elevated below 4 m, ranging from 9.1–31.9 lmol kg�1,but depleted at shallow depths, ranging from 0–3.6 lmol kg�1. By comparison, the bay water had low dis-solved Mn (average = 0.4 lmol kg�1; Appendix Table S1).Dissolved Fe ranged from 0–542 lmol kg�1, with the high-est values observed in the shallow layer (<2.5 m) in associ-ation with the freshwater plume (Fig. 4c). Like Mn, surfacewater Fe was generally low (average = 0.5 lmol kg�1;Appendix Table S1). During the STE cross-section sam-pling in November 2009 and June 2010, the highest Mnand Fe concentrations were associated with high salinitygroundwater at the seaward PZ5 well (AppendixTable S1). This is the most reducing region within theSTE (Charette and Sholkovitz, 2006), and these elevateddissolved Mn and Fe values are the result of Mn and Fe(oxy)hydroxide reduction associated with organic matteroxidation, with the organic matter being supplied by seawa-ter intrusion (Charette et al., 2005; Spiteri et al., 2008).These microbial reaction pathways are favored upon deple-tion of oxygen and nitrate as the primary electronacceptors.

Ammonium (Fig. 4e) and phosphate (Fig. 4f) follow asimilar pattern to Mn, though there were elevated concen-trations of both at some shallow depths (less than 2.5 m)and in some seasons, e.g. NH4

+ was as high as 43 lmol L�1

at 1.5 m in September 2009. In contrast with ammonia,nitrate plus nitrite (N + N) was almost nil or below detec-tion in the deeper groundwater (>3.5 m, Fig. 4d). Elevated

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N + N (163–245 lmol L�1), centered at a depth of 2.5 m,corresponded with the high DO layer.

3.3. Temporal variability of the carbonate system1 in the

subterranean estuary

TAlk and DIC followed a consistent pattern in the STE(Fig. 5a and b): terrestrial groundwater (salinity < 1) hadlower TAlk (40–873 lmol kg�1, median: 97 lmol kg�1)and DIC (512–1039 lmol kg�1, median: 695 lmol kg�1) ascompared with marine groundwater (1 � salinity � 25.6),which had a TAlk range of 84–1985 lmol kg1 (median:1311 lmol kg�1) and DIC range of 598–2046 lmol kg�1

(median: 1609 lmol kg�1).Groundwater carbonate ion concentrations were low

across all salinities, less than 1% of the DIC. In terrestrialgroundwater, DIC was dominated by aqueous CO2,accounting for 16–93% (median: 85%) of the DIC whilein the marine groundwater DIC was dominated by bicar-bonate (11–95%, median: 83%). Both TAlk and DIC in ter-restrial groundwater were lower than in bay water (TAlk:1364–2063 lmol kg�1; DIC: 1320–1936 lmol kg�1, Appen-dix Table S1). The TAlk/DIC ratio in all groundwater sam-ples was below 1, which is in contrast with the bay watersamples, which had a TAlk/DIC of greater than 1. Seasonalvariations of TAlk and DIC at fixed depths were consistentwith the observed salinity variations, i.e. TAlk and DICconcentrations in summer and fall were greater than thosein winter and spring at fixed depths, suggesting that TAlkand DIC within the STE was influenced by physical mixingbetween terrestrial and marine groundwater.

The entire STE had between one and two orders of mag-nitude higher pCO2 (�1044–22,670 latm, median:6243 latm, Fig. 5c) relative to atmospheric CO2

(�400 ppm) and bay water pCO2 (138–1113 latm, Appen-dix Table S1). Marine groundwater had lower pCO2 thanterrestrial groundwater. The calculated pH values at

1 Four typical parameters represent carbonate system, pH,partial pressure CO2 (pCO2), total alkalinity (TAlk, [HCO3

�]+2[CO3

2�]+[B(OH)4�]+[OH�]+[HPO4

2�]+2[PO43�]+[SiO(OH)3

�]+[NH3 ]+[HS�]+ � � ��[H+ ]F�[HSO4

�]�[HF]�[H3PO4 ]�� Dickson, 1981), and dissolved inorganic carbon (DIC, [CO2

*]+[HCO3

�]+[CO32�]), [CO2

*] represents the concentration of allunionized carbon dioxide including [CO2] and [H2CO3], i.e.aqueous CO2.

* is the sum of [CO2] and [H2CO3].

in situ temperature were low in terrestrial groundwater,ranging from 5.23–7.12; pH increased across the mixingzone reaching maximum values of 7.39 at high salinity.The pH was lowest at 2.5 m in association with low TAlk(Fig. 5d). Bay water pH was greater than 7.47 with a max-imum of 8.36 (Appendix Table S1). The STE calcite satura-tion index increased from 0 at the surface to 0.61 at 5.5 m(Fig. 5e and Appendix Table S1), while aragonite satura-tion ranged from 0 to 0.38 (Appendix Table S1), suggestingthe system was undersaturated with respect to calcium car-bonate across the entire STE during all seasons.

4. DISCUSSION

4.1. Physical and biogeochemical control on the carbonate

chemistry in the subterranean estuary

Terrestrial groundwater (salinity < 1) displayed a widerange of TAlk and DIC concentrations(Figs. 6a, 7a and b); elevated values were generally associ-ated with a terrestrial groundwater plume that also carriedelevated concentrations of ammonia, phosphate, and ironin PZ10 (Fig. 6c and d). Isotopic data such as d56Fe(Rouxel et al., 2008) and d15N-NH4

+ (Kroeger andCharette, 2008) suggest that this plume originates fromaquifer recharge through a freshwater pond to the northof the study location. In addition, organic matter degrada-tion produces DIC, ammonia, and phosphate (Table 2).The observed elevated TAlk and a negative correlationbetween TAlk and ORP suggest that TAlk and DIC areproduced by organic carbon decomposition processes inthe terrestrial groundwater (Fig. 6a and b, and AppendixFig. S2). The positive correlation of both TAlk and DICwith ammonia is also consistent with this idea(Fig. 6b and d, and Appendix Fig. S2). It is unlikely thatcarbonate minerals are exerting a significant control onSTE TAlk and DIC concentrations as aquifer sedimentsin this region are primarily composed of coarse grainedsands with quartz (Rouxel et al., 2008). In addition, terres-trial groundwater was observed in the shallow layer(�2.5 m) of the monthly time-series piezometer (PZ3) sam-ples. At 2.5 m, we observed elevated DO and nitrate butdepleted ammonia, particularly from June to September2009 (Fig. 4), which may be an indication of nitrification.

In order to better understand and quantify the pro-cesses that are operating on the inorganic carbon cycle

Fig. 4. Time-series contours of pore water chemistry for (a) dissolved oxygen (DO), (b) total dissolved Mn, (c) total dissolved Fe, (d) nitrateand nitrite (NO3

� + NO2�), (e) ammonium (NH4

+), and (f) phosphate (PO43�) at PZ3 from May 2009 to June 2010. The dashed line represents

salinity contour line of 1. The x-axis indicates time and the y-axis indicates depth below the surface.


in the STE during mixing of terrestrial and marinegroundwater, we must first define the low and high salinityendmembers to be used in our analysis. On the freshwaterside, groundwater flow is generally laminar such thatmuch of the mixing takes place along the STE salinity gra-

dient. As such, the plume containing high DIC and TAlkthat coincides with a persistent terrestrial groundwaternutrient and trace metal plume (2–4 m at site PZ10,Fig. 6; see also Kroeger and Charette (2008) and Spiteriet al. (2008)) does not interact or mix with deep STE

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Fig. 5. Time-series contours of pore water chemistry for (a) TAlk, (b) DIC, (c) pCO2 (in situ temperature), (d) pH (total scale, in situ

temperature), and (e) calcite saturation index at PZ3 from May 2009 to June 2010. The dashed line represents salinity contour line of 1. The x-axis indicates time and the y-axis indicates depth below the surface.


groundwater. This plume is characterized by relatively lowORP (Fig. 6a), hence, we define our terrestrial groundwa-ter endmember as samples with an ORP > 150 mV, whichincludes those samples that lie just above the mixing zonebut below the terrestrial nutrient plume (phosphate and

ammonium). On the high salinity side, we use averagebay water concentrations since this is the water that isbeing recharged into the STE during seawater circulationthrough the aquifer. In this way, any increases ordecreases in DIC or TAlk across the mixing zone are

(a) (b) (c) (d)

Fig. 6. Depth profiles showing (a) salinity, oxidation reduction potential (ORP), pH, (b) TAlk, DIC, (c) dissolved Mn, Fe, (d) nitrate,ammonium, and phosphate for the inland well (PZ10) sampled in November 2009. The general overlap in the distribution of elevated TAlk,DIC, Fe, Mn, and ammonia at 2–4 m with low ORP and high pH is indicated by the gray shading.


assumed to be exclusively a result of biogeochemical pro-cesses that occurred in the subsurface. A summary of theendmembers for terrestrial groundwater and bay water areshown in Table 1.

We used the endmember definitions above to constructtwo-endmember mixing diagrams for both the time seriesPZ3 profiles (Appendix Fig. S3) and the high resolutionprofiles collected in November 2009 and June 2010(Fig. 7). We also introduce DTAlk and DDIC, where Dis the difference between the calculated conservativetwo-endmember mixing and the measured value(Fig. 8); positive values are non-conservative additionwhile negative values are removal. DIC in the STE dis-played addition at high salinity (>20) for all seasons,while it exhibited approximately conservative mixingthrough intermediate salinities (1–20). TAlk appeared tobe added at high salinity in some seasons, but this TAlkexcess is smaller and less consistent than the DIC excess.At intermediate salinity, TAlk displayed a large seasonalvariation from nearly conservative mixing (December 28,2009, January, February, April, and June 2010) to non-conservative removal (June to August 2009, November,December 1, 2009, and May 2010) and slightly non-

conservative addition in March 2010 (Fig. 8 and Appen-dix Fig. S3). These DIC and TAlk patterns are evidencefor the existence of temporally-variable physical and bio-geochemical controls on their concentration during trans-port through the STE. To examine in more detail thepotential driving forces on DIC and TAlk behavior inthe STE, we will separately discuss the transformationsoccurring at mid salinity (1–20) and high salinity (>20)with a focus on the more detailed November 2009 andJune 2010 datasets.

4.1.1. Processes driving DIC and TAlk concentrations at mid

salinity

DIC displayed only slight curvature (consumption) inthe mid salinity zone (1–20) from November 2009 and June2010 (Fig. 7b and d). This indicates that DIC was predom-inately controlled by physical mixing between terrestrialgroundwater and bay water in these two seasons. However,while TAlk was approximately conservative in the samezone during June 2010, it was significantly depleted inNovember 2009 (Fig. 7a and c). The TAlk removal waslocated primarily between 0.3 and 0.9 m in PZ5 (Fig. 9b).Previous studies have observed oxidation of reduced Mn

(a) (b)

(c) (d)

Fig. 7. Plots showing TAlk versus salinity (a, c), and DIC versus salinity (b, d) in the Waquoit Bay subterranean estuary during November2009 and June 2010 in piezometer wells PZ10, PZ3, and PZ5. TAlk and DIC for terrestrial groundwater in Nov are also shown as inset graphswith an expanded salinity scale (0–0.5). The solid lines represent the conservative mixing lines for TAlk and DIC in deep STE, which was usedto quantify the variations of the carbonate system in high salinity zone. Note the scale difference on the y-axis for (a) and (b) vs. (c) and (d).

Table 1Salinity, TAlk, DIC, and pH in the endmembers of terrestrial groundwater and bay water collected in November 2009 and June 2010.

Sampling date Endmember Salinity TAlk DIC pH

(lmol kg�1)

Nov-09 *Terrestrial groundwater 0.1 67 747 5.4Baywater 28.8 1888 1754 8.0

Jun-10 #Terrestrial groundwater 0.4 75 666 5.4Baywater 27.6 1849 1747 7.7

* Average from stations at 4.9 and 5.5 m in piezometer well PZ10.# Average from stations at 4.3, 4.9, and 5.5 m in piezometer well PZ10, 2.4 m in PZ3. The locations of these stations were marked in Figs. 9

and 12. The pH is calculated from the measured TAlk and DIC with in situ temperature.


and Fe in terrestrial groundwater upon transiting the STE,and Fe/Mn oxide coated subsurface sands have beenobserved in the mixing zone down to the depth of at least6 m around PZ3 and PZ5, the so-called ‘‘iron curtain”(Charette and Sholkovitz, 2002; Charette et al., 2005;Gonneea et al., 2008). There is also a Fe and Mn redoxfront along the high salinity boundary of STE, whichaccounts for the large vertical extent of the ‘‘iron curtain”(Charette and Sholkovitz, 2002; Charette et al., 2005;Charette and Sholkovitz, 2006).

Since Fe2+ and Mn2+ oxidation reactions produce H+ butno DIC, these processes would serve to reduce TAlk and pHyet have no effect on DIC. Spiteri et al. (2006) reported thatthe pH gradient between PZ3 and PZ5, driven by seawaterintrusion with a relatively high pH, caused a 7-fold increasein the rate of Fe2+ oxidation relative to fresh groundwater.We observed shallowing of the salinity mixing zone in sum-mer and fall (June to December 2009, Fig. 3) relative to win-ter and spring (January to May 2010, Fig. 3), suggestingenhanced seawater intrusion occurred in summer and fall

4/2010 5/2010 6/2010

Dep

th(m

)

5/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

-800-700-600-500-400-300-200-1000100200300400500600700800

4/2010 5/2010 6/2010

Dep

th(m

)

5/2009 6/2009 7/2009 8/2009 9/2009 11/2009 12a/2009 12b/2009 1/2010 2/2010 3/2010-5.5

-4.5

-3.5

-2.5

-1.5

-0.5

-800-700-600-500-400-300-200-1000100200300400500600700800

Fig. 8. Time-series contours of pore water chemistry for (a) DTAlk and (b) DDIC (D means the difference between the calculated conservativemixing and the measured value) at PZ3 from May 2009 to June 2010. Dashed lines are the salinity contour lines of 1, 5 and 20. The x-axisindicates time and the y-axis indicates depth below the surface.

Fig. 9. Distributions of (a) salinity, (b) DTAlk, (c) DDIC, fraction of (d) aerobic oxidation, (e) denitrification, and (f) Mn & Fe reduction inthe piezometer wells sampled in November 2009. In panel (a), triangles show the sites of terrestrial groundwater endmember, rectangles denoteareas of high salinity (>20) but with negative DTAlk/DDIC ratios, which are beyond those predicted for various OM degradation reactions.The box in panel (b) represents area with maximum negative DTAlk.


concurrent with our observation of TAlk removal duringJune through early December 2009 (Appendix Fig. S3 andFig. 8). The pH-salinity relationship for November 2009,which showed a reduction of up to 1 pH unit across the mix-

ing zone, also supports our observation (Fig. 10). Con-versely, pH was largely conservative through the midsalinity mixing zone in June 2010, consistent with theobserved conservative distributions of both TAlk and DIC.

5.0

5.5

6.0

6.5

7.0

7.5

8.0

0 10 20 30

pH

Salinity

observation-Nov.09

observation-Jun.10

conservativemixing-Nov.09

conservativemixing-Jun.10

Fig. 10. The pH versus salinity within subterranean estuarysamples collected in November 2009 and June 2010 from piezome-ter wells PZ10, PZ3, and PZ5. The conservative mixing lines for pHwere calculated as follows: conservative mixing concentrations forTAlk and DIC were estimated as a function of salinity using theterrestrial groundwater and bay water endmembers from Table 1.These data were entered in CO2SYS, which was used to obtain thecorresponding mixing-derived pH. The observed pH in November2009 from intermediate and high salinity are much lower than thepredicted values based on conservative mixing. In June 2010, theobserved pH is nearly consistent with the theoretical mixing line atmid salinity (1–20), but is slightly smaller than the calculatedconservative pH value at high salinity (>20).


Of the entire time-series dataset (PZ3 and PZ5; Figs. 8and 9), the maximum observed TAlk depletion was780 lmol kg�1 at 0.5 m in PZ3 (salinity 12.0, Fig. 8a,May-2010) and 723 lmol kg�1 at 0.9 m in PZ5 (salinity20.4, Fig. 9b, November-2009), which would require a cor-responding proton addition of 780 and 723 lmol kg�1, withthe actual H+ increase determined by the carbonate chem-istry and hence buffering capacity. Based on the reactionformulas (Burdige 2006):

Fe2þ þ 0:25O2 þ 2:5H2O ! FeðOHÞ3 þ 2Hþ ð1ÞMn2þ þ 0:5O2 þ H2O ! MnO2 þ 2Hþ ð2Þwhere Fe2+:H+ and Mn2+:H+ equal 1:2, the total amountof reduced Fe and Mn should be �390 and 361 lmol kg�1

assuming that these redox processes are the sole source ofacidity driving the TAlk decrease. Manganese concentra-tions reached a maximum of only �4 mmol kg�1 in the ter-restrial groundwater, while peak dissolved ironconcentrations were well in excess of 400 lmol kg�1. Ele-vated ferrous iron in terrestrial groundwater was alsoobserved in prior studies (Charette and Sholkovitz, 2006),therefore, a sufficient level of reduced Fe exists to supportthe proton production needed to balance the TAlk con-sumption in this shallow zone. Further, this level of Fe2+

oxidation would have consumed 97 and 90 lmol kg�1 ofDO, well within the amount available within the TAlkdepletion zone (up to 225 lmol L�1).

While Fe-associated redox reactions can largely explainthe observed changes in TAlk, other contributing reactionscannot be entirely ruled out. For example, a slight additionof dissolved Ca in the mixing zone hints that calcium car-bonate dissolution may play a small role (AppendixFig. S1). Also, nitrification could serve to reduce TAlk with

no effect on DIC. For the large TAlk depletion at PZ5(0.9 m), low nitrate (0.6 lmol L�1) and elevated ammonia(19.5 lmol L�1) preclude a significant contribution fromnitrification. However, this process may have contributedslightly to the TAlk removal in PZ3 at 2.5 m (Fig. 4),though this is well above the zone of significant TAlkremoval (3.5 m, Fig. 8). The ‘‘iron curtain” is within thisdepth range (Gonneea et al., 2008), suggesting that Fe2+

oxidation is likely the major controlling factor even at3.5 m. We note, however, that these estimates of STE TAlkremoval depend on the choice in TAlk endmembers. Ifslightly shallower terrestrial groundwater with higher TAlkwere used, then these calculated values would representlower limit estimates.

4.1.2. Processes driving DIC and TAlk concentrations at high

salinity

In June 2010, non-conservative addition of DIC andTAlk was observed in the high salinity groundwater thatwas characterized by elevated ammonia, phosphate, anddissolved Mn (Fig. 11). Because changes in DIC relativeto TAlk would be differentially modified by various organiccarbon decomposition reactions (Table 2), changes in theirratio can be combined with changes in the products oforganic matter decomposition (e.g. NH4, Mn, and Fe) toquantify the relative importance of the geochemical pro-cesses occurring within this zone of the STE (Froelichet al., 1979; Bender and Heggie, 1984; Cai et al., 2003).During our study, we did not measure SO4

2�, H2S or CH4

though previous studies indicated that SO42� is conservative

and H2S has never been detected (Gonneea and Charette,2014). We therefore only considered aerobic respiration,denitrification, manganese reduction, and iron reductionas potential driving mechanisms for the observed TAlkand DIC distributions in the high salinity zone (Table 2).Further, Ca2+ data measured in 2004 for these samepiezometer locations (Appendix Fig. S1) indicate conserva-tive mixing through the high salinity region. Hence, we alsoassume that no calcium carbonate precipitation/dissolutionoccurred in this zone. The details of this approach are pre-sented in Appendix.

In June 2010, both excess TAlk and DIC were observedin the high salinity region (Figs. 12b, c, and 13); theseexcesses at 2.4–6.1 m of PZ5 and deep in PZ3 (4.0–5.5 m,Fig. 12) are within the range of reactions from aerobic res-piration to iron reduction, but are closest to aerobic respi-ration and denitrification (Fig. 13). Our calculationssuggest that aerobic respiration was responsible for 24–100% (average 55 ± 19%; Fig. 12d) of the organic carbonremineralization, while denitrification was in the range of0–79% (average 42 ± 18%; Fig. 12e). Both Mn and Fereduction were much smaller (Mn reduction; 0–11%, aver-age 3 ± 3%; Fe reduction; 0–1%, average 0.1 ± 0%;Fig. 12f). If we assume that bay water in the STE rechargezone had an initial O2 concentration of 240–448 lmol L�1

(the range in our surface water measurements) and theSTE O2 in the high salinity zone was in the range of 52–136 lmol L�1, then the oxygen utilization should be 188–312 lmol L�1, which is within the range of our modeledDO consumption (40–457 lmol L�1). Therefore, aerobic

(a) (b)

(c) (d)

Fig. 11. Relationship between (a) ammonia, (b) phosphate, (c) dissolved Mn, (d) dissolved Fe and salinity for PZ10, PZ3, and PZ5 wellssampled in November 2009 and June 2010.

Table 2Organic matter degradation reactions (Froelichetal., 1979; Bender and Heggie, 1984; Cai et al., 2003).

Process Geochemical reactions

Aerobic respiration OM + 138O2 ? 106CO2 + 16HNO3 + H3PO4 + 122H2O (1)Denitrification OM + 0.8 � 106NO3

� + 0.8 � 106H+ ? 106CO2 + 0.4 � 106N2 + 16NH3 + H3PO4 + 1.4 � 106H2O (2)Manganese reduction OM + 2 � 106MnO2 + 4 � 106H+ ? 106CO2 + 2 � 106Mn2++16NH3 + H3PO4 + 318H2O (3)Iron reduction OM + 4 � 106FeOOH + 8 � 106H+ ? 106CO2 + 4 � 106Fe2++16NH3 + H3PO4 + 583H2O (4)

The stoichiometric ratios of changes in TAlk and DIC per mole organic matter degradation(1) dTA/dDIC = (�16–12)/(106) = �0.160(2) dTA/dDIC = (0.8 � 106 + 16–12)/(106) = 0.932(3) dTA/dDIC = (4 � 106 + 16–12)/(106) = 4.132(4) dTA/dDIC = (8 � 106 + 16–12)/(106) = 8.132

OM = (CH2O)106(NH3)16(H3PO4).d represents the change in TAlk and DIC associated with individual organic matter degradation as it proceeds from left to right.


respiration and denitrification are the dominant reactionsbehind organic matter oxidation at high salinity in June.In addition, aside from denitrification, elevated NH4

+ mightresult from ammonification within the organic carbon richsediments of the outer bay (Gonneea and Charette, 2014)where saline groundwater recharge originates (Michaelet al., 2005). However, previous work suggests that denitri-fication does occur within this high salinity zone (Kornerand Zumft, 1989).

The DTAlk and DDIC distributions for November 2009at high salinity are more complex. While TAlk and DICadditions were observed at the base of PZ3 (4.3–5.5 m)and PZ5 (4.9–6.1 m), above these depths there was DICaddition in the presence of TAlk depletion(Fig. 9b and c). Restricting our model calculations to thesegreater depths, we obtained an aerobic respiration fractionof 27–93% (average 67 ± 26%), a denitrification fraction of1 to 70% (average 31 ± 26%), Mn reduction from 0 to 4%

Fig. 12. Distributions of (a) salinity, (b) DTAlk, (c) DDIC, fraction of (d) aerobic oxidation, (e) denitrification, and (f) Mn & Fe reduction inthe piezometer wells sampled in June 2010. The triangles mark the locations used in generating the terrestrial groundwater endmember.

Fig. 13. The DTAlk versus DDIC within subterranean estuary in mid salinity (1–20) and high salinity (>20) zone samples collected inNovember 2009 and June 2010 from piezometer wells PZ3 and PZ5. The arrows indicate how DTAlk and DDIC will change in response todifferent biogeochemical processes including aerobic respiration, denitrification, manganese reduction, iron reduction, Fe2+ or Mn2+

oxidation, and CaCO3 production/dissolution. Numbers inside parentheses represent stoichiometric ratios of changes in TAlk and DIC permole organic matter degradation. The detailed biogeochemical reactions are presented in Table 2.



(average 1 ± 2%), and Fe reduction of 0 to 5% (average 1± 2%). Similar to June 2010, the modeled DO consumption(181–441 lmol L�1) is within the range of the observed DOloss (48–430 lmol L�1). In general, the high salinity zone inPZ5 is predominately controlled by aerobic oxidation bothin November 2009 and June 2010 (Figs. 9 and 12).

Since our approach is based on net changes in waterchemistry, it should be noted that anaerobic decompositionprocesses and subsequent re-oxidation of the reduced com-pounds could be interpreted in our model as being due toaerobic decomposition. Potential evidence for this is thenegative DTAlk/DDIC ratios (�3.1 to �29.3) observedfor November 2009 at PZ3 from 3.4 to 3.6 m and PZ5 from1.5 to 4.2 m (Fig. 9a). These ratios are beyond those pre-dicted for various OM decomposition processes (�0.2 to8, for aerobic respiration to iron reduction, Fig. 13). Thesecould include oxidation of Fe2+ and Mn2+ (Fig. 13), whichare believed to be responsible for the large TAlk sink in themid salinity zone. Other possibilities include Mn2+

mediated-oxidation of ammonia as suggested by Kroegerand Charette (2008) for this same STE.

4.2. Seasonal variability of SGD-derived TAlk and DIC

fluxes to surface water

To assess seasonal variability of SGD-associated TAlkand DIC fluxes into the bay, we combined a previouslypublished hydrological flow model for the Waquoit BaySTE (Gonneea et al., 2013a,b) with our time-series TAlkand DIC concentrations. The model simulated SGD fluxwas derived from a two-dimensional cross sectionalgroundwater flow and salt transport model (SEAWAT)(Mulligan et al., 2011). The model was bounded by anupland margin (CCC1 in Fig. 1) and extended 125 m sea-ward of CCC1. Since the groundwater head at CCC1 wasnot measured during our sampling period, we used theaverage SGD water flux for summer and winter (2005–2007) from Gonneea and Charette (2014). As withGonneea et al. (2013a,b), we divided SGD into five salin-ity groups: 0–5, 5–10, 10–15, 15–20, and >20.Groundwater-derived TAlk and DIC fluxes were then cal-culated by multiplying the average groundwater TAlk andDIC concentrations by the water flux for the same salinitygrouping (Table 3). The same approach was utilized toestimate fluxes of aqueous CO2 and H+ discharged intothe bay from groundwater (Table 3). Since the simulatedgroundwater flux was from 2005 to 2007, while our car-bonate system parameters were sampled during 2009–2010, we limit our discussion to assessing the dominantcontrolling factors for seasonal variability in SGD-derived carbonate system fluxes from a typical unconfinedcoastal aquifer.

Gonneea et al. (2013a) found that the fresh groundwaterflux (salinity from 0 to 5) was relatively constant betweensummer and winter (2.2 ± 0.6 vs. 3.2 ± 0.4 m3 m�1 day�1;Table 3). They also showed that landward movement ofthe mixing zone in summer was due to the relatively lowhydraulic gradient compared with winter (Fig. 3). As such,the saline groundwater flux was ca. 5 times greater in sum-mer than winter (3.9 ± 1.3 vs. 0.8 ± 0.2 m3 m�1 day�1). This

contrast is largely responsible for the �50% higher esti-mated SGD-derived net DIC and aqueous CO2 fluxes tothe bay in summer (3.8 ± 2.5 vs. 2.7 ± 0.4 mol m�1 day�1

for DIC flux, 2.5 ± 0.6 vs. 1.7 ± 0.2 mol m�1 day�1 foraqueous CO2 flux, Table 3). The estimated SGD-associated H+ flux followed the same trend (9.4 ± 2.5 vs.6.7 ± 1.1 mmol m�1 day�1, Table 3). Hence, SGD couldbe an important contributor to coastal ocean acidificationin addition to other factors that are driving ocean pH lower.

Despite the 5� higher marine SGD flux in summer, thenet bay water SGD-TAlk flux was comparable between thetwo seasons (0.9 ± 2.3 vs. 0.9 ± 0.3 mol m�1 day�1,Table 3), a result of low summertime STE TAlk concentra-tions. Here the multi-salinity zone model we employed wascrucial: because of the strong mixing zone gradient in TAlk,an approach that divided SGD water fluxes into just twocomponents (fresh and marine groundwater), would haveresulted in a significant overestimate in the SGD-TAlk flux(nearly 5� for the summer period).

Regardless of season, the SGD-derived DIC flux isalways greater than TAlk flux, indicating that SGD wouldserve to increase the surface water DIC/TAlk ratio. In gen-eral, our bay water DIC concentrations were lower thanTAlk (Appendix Table S1). Since the CO2 buffering capac-ity of seawater is weakened as the DIC/TAlk ratioapproaches 1 (Egleston et al., 2010), our results for this sys-tem indicate that SGD serves to reduce the CO2 bufferingcapacity of surface water, which is consistent with otherrecent studies suggesting that this may be a common effectof SGD on the global carbon cycle (Liu et al., 2014; Sadat-Noori et al., 2016; Wang et al., 2015).

4.3. Potential SGD effects on the surface water CO2 system

on annual time scales

To examine the net annual effect of SGD on the coastalocean carbon cycle, we compared the yearly inputs and out-puts of DIC and TAlk to the Waquoit Bay STE. The inputterms include terrestrial groundwater and the bay waterthat enters the STE as marine groundwater recharge, whilethe output is the groundwater discharged into the bay.Using the model average water flux for 2005, the DIC out-put was greater than the combined inputs which translatedinto a net STE DIC production rate of 420 mol m�1 yr�1,which is �19 % of the DIC flux to the bay via groundwaterdischarge (Table 4). This is consistent with the DIC excessthat was observed in groundwater during almost all sea-sons. In contrast, the TAlk input was roughly equivalentto the output (5% difference, Table 4), suggesting that TAlkremoval in the mid salinity zone was offset by TAlk produc-tion at high salinity on an annual basis. We would reach thesame conclusions if we used the modeled mean water fluxesfor 2006.

In Waquoit Bay, mixing zone dynamics appear to be lar-gely controlled by seasonal variations in mean sea level.Gonneea et al. (2013a) noted that the period of our studywas characterized by a positive sea level anomaly drivenby regional climate oscillations. Their model results indicatethat such an anomaly would result in enhanced STE mixingand marine groundwater recharge/discharge in summer.

Table 3A comparison of carbonate chemistry and SGD-related TAlk and DIC fluxes in summer and winter in the Waquoit Bay subterranean estuary.

Season Salinity zone SGR Net flux

0–5 5–10 10–15 15–20 >20

Summer (July (July 1 and July 22) and August, 2009)

Water Flux* (m3 m�1 day�1) 2.2 ± 0.6 0.7 ± 0.7 0.3 ± 0.3 0.5 ± 0.4 2.4 ± 1.0 �2.4 ± 0.8TAlk (mmol m�3) 170 ± 141 230 ± 55 411 ± 72 732 ± 272 1648 ± 62 �1748 ± 135DIC (mmol m�3) 724 ± 63 1019 ± 68 1187 ± 56 1255 ± 202 1826 ± 26 �1670 ± 55Aqueous CO2 (mmol m�3) 551 ± 77 552 475 524 ± 143 205 ± 34 �17 ± 3H+ (mmol m�3) 3.47 ± 0.52 1.03 0.59 0.77 ± 0.36 0.15 ± 0.03 �0.01 ± 0.004TAlk flux (mol m�1 day�1) 0.4 ± 0.3 0.2 ± 0.1 0.1 ± 0.1 0.4 ± 0.3 4.0 ± 1.6 �4.2 ± 1.5 0.9 ± 2.3DIC flux (mol m�1 day�1) 1.6 ± 0.4 0.8 ± 0.7 0.3 ± 0.4 0.7 ± 0.5 4.4 ± 1.8 �4.0 ± 1.4 3.8 ± 2.5Aqueous CO2 flux (mol m�1 day�1) 1.2 ± 0.4 0.4 ± 0.4 0.1 ± 0.1 0.3 ± 0.2 0.5 ± 0.2 �0.04 ± 0.02 2.5 ± 0.6H+ flux (mmol m�1 day�1) 7.8 ± 2.3 0.8 ± 0.7 0.2 ± 0.2 0.4 ± 0.4 0.4 ± 0.2 �0.03 ± 0.01 9.4 ± 2.5

Winter (December (December 1 and December 28), 2009; January and February, 2010)

Water flux* (m3 m�1 day�1) 3.2 ± 0.4 0.4 ± 0.1 0.2 ± 0.1 0.1 ± 0.1 0.1 ± 0.0 �0.4 ± 0.06TAlk (mmol m�3) 284 ± 43 512 ± 141 701 ± 197 1116 ± 183 1719 ± 74 �1843 ± 256DIC (mmol m�3) 750 ± 23 875 ± 77 1062 ± 167 1424 ± 100 1863 ± 43 �1753 ± 214Aqueous CO2 (mmol m�3) 465 ± 38 344 ± 156 335 ± 141 313 ± 92 167 ± 63 �22 ± 6H+ (mmol m�3) 2.08 ± 0.25 0.40 ± 0.19 0.31 ± 0.10 0.26 ± 0.14 0.10 ± 0.05 �0.01 ± 0.003TAlk flux (mol m�1 day�1) 0.9 ± 0.2 0.2 ± 0.1 0.2 ± 0.1 0.2 ± 0.1 0.1 ± 0.02 �0.7 ± 0.2 0.9 ± 0.3DIC flux (mol m�1 day�1) 2.4 ± 0.3 0.4 ± 0.1 0.2 ± 0.1 0.2 ± 0.1 0.1 ± 0.02 �0.7 ± 0.1 2.7 ± 0.4Aqueous CO2 flux (mol m�1 day�1) 1.5 ± 0.2 0.1 ± 0.1 0.08 ± 0.04 0.04 ± 0.02 0.01 ± 0.005 (�9 ± 3)�10�3 1.7 ± 0.2H+ flux (mmol m�1 day�1) 6.6 ± 1.1 0.2 ± 0.1 0.07 ± 0.03 0.04 ± 0.02 0.01 ± 0.004 (�4 ± 1)�10�3 6.7 ± 1.1

Error bar is standard deviation, representing monthly variation in one season within one year or three years; all fluxes are normalized to permeter of shoreline.Negative values denote water and material transport from bay water to the subterranean estuary.SGR means submarine groundwater recharge, here we used the average surface samples as the endmember of SGR.* Water flux was modeled from 2005 to 2007 (Gonneea and Charette, 2014).

Table 4Annual mass balance of water, total alkalinity (TAlk) and dissolved inorganic carbon (DIC) in the Waquoit Bay subterranean estuary.

Water Flux# TAlk Flux DIC Fluxm3 m�1 year�1 mol m�1 year�1 mol m�1 year�1

STE inputs

Terrestrial groundwater* 1200 276 509Submarine groundwater recharge (from Waquoit Bay) 730 1305 1216

STE output

Submarine groundwater discharge (into Waquoit Bay) 1930 1468 2177Net STE Production – 68 (5%) 419 (19%)

# Water flux was modeled from January 2005 to December 2005, and normalized to per meter of shoreline.* Terrestrial groundwater TAlk is 0.19 mol m�3 and DIC is 0.72 mol m�3.


Such an enhancement would serve to increase the DIC fluxvia enhanced groundwater flux and organic matteroxidation. Further, since nutrients are regenerated in thisprocess, they might serve to reduce the net effect of theSGD-delivered DIC flux through biological uptake insurface waters. However, these two processes may not per-fectly offset each other as some nitrogen is removed in thehigh salinity zone of the STE (Kroeger and Charette,2008), and phosphate may be retained in the subsurfacethrough sorption to Fe oxides (Charette and Sholkovitz,2002, 2006; Kroeger and Charette, 2008). Therefore, sea-sonal oscillations in MSL may still result in a net DIC fluxto surface water by SGD.

We estimated that TAlk removal was offset by TAlkexcess in the STE on an annual basis in a ‘‘normal” year

(2005–2007). However, our data suggest that net STE TAlkremoval could occur in association with positive sea levelanomaly as was observed from 2009 to 2010. In this scenar-io, the SGD-derived DIC flux would be even greater thanthe SGD-TAlk flux, resulting in (1) enhanced surface wateraqueous CO2 inputs thereby enhancing air-sea exchangeand reducing the seawater buffering capacity and (2)enhanced proton fluxes that would exacerbate coastalocean acidification. Our results highlight the need for in-depth studies on estuarine acidification in systems likeWaquoit Bay, where organisms including many commer-cially valuable shellfish may already be experiencing car-bonate chemistry conditions at least as severe as thosepredicted for the open ocean by the end of the 21th century(Doney et al., 2009).


5. CONCLUSIONS

This study demonstrates the complexity of the inorganiccarbon cycle in subterranean estuaries due to time-varyinghydrologic and biogeochemical processes. Our main con-clusions are as follows:

(1) The major component of DIC in the terrestrialgroundwater was aqueous CO2 (83% of DIC); theentire STE had one to two orders of magnitude high-er pCO2 levels relative to atmospheric CO2. The STEhad lower pH relative to surface water and it wasundersaturated with respect to calcite and aragoniteduring all seasons.

(2) DIC concentrations in brackish groundwater (salin-ity 0–20) were lower than in surface water. DICwas slightly consumed at low and mid salinities whilethere was an apparent DIC source at high salinity(>20) in all seasons. In November 2009 (fall) andJune 2010 (summer), TAlk and DIC additionoccurred in high salinity groundwater. Stoichiometriccalculations suggest that these additions were mainlydriven by aerobic oxidation and denitrification.

(3) Within brackish groundwater during summer andfall, TAlk generally decreased due to proton produc-tion from Fe2+ oxidation reactions, in part due tonitrification; TAlk largely displayed conservativemixing in other seasons. Hence, oscillations in thelocation of the mixing zone, which is also a redox/pH boundary, are driving Fe-based oxidation–reduc-tion cycles that in turn are driving the STE TAlkdistribution.

(4) The net SGD-derived DIC flux to the bay exceededthe TAlk flux by a factor of �3–5; the former was�50% higher in the summer vs. winter while the latterwas constant year-round. The lower TAlk flux wasdue to lower TAlk vs. DIC concentrations in theSTE. The magnitude of the groundwater flux con-trolled SGD-DIC flux, while, depending on the sea-son, groundwater flux and TAlk concentration wereco-drivers of the SGD-TAlk flux.

(5) On an annual cycle, the subterranean estuary was anet producer of DIC while the TAlk cycle was in bal-ance, i.e. consumption and production of TAlk wasoffset in the mid and high salinity zones, respectively.

Future studies aimed at quantifying SGD-carbonatechemistry fluxes should take into account potential non-conservative processes in the STE as well as seasonal vari-ability. In particular, our finding of significant TAlkremoval (H+ addition) during high seawater intrusion peri-ods implies that SGD’s role in the carbon cycle for receivingwater bodies may increase during seasonal periods of highsea level. In Waquoit Bay and the greater northeastern U.S., Gonneea et al. (2013a,b) postulated that sea level is amajor controlling factor for SGD; recent studies supportthis idea for other areas of the world’s coastline (Leeet al., 2013; Wood and Harrington, 2015). Enhancedseasonal or annual variability in sea level due to climate

oscillations like the El Nino-Southern Oscillation (ENSO)and North Atlantic Oscillation (NAO) could disrupt theyear to year balance in net inorganic C fluxes from theSTE, which may lead to further enhancements in the DICand aqueous CO2 discharge from groundwater and theassociated impacts on seawater’s CO2 buffering capacity.

ACKNOWLEDGEMENTS

The research was financially supported by the United StatesNational Science Foundation Chemical Oceanography program(OCE-0425061 and OCE-0751525 to M.C.) and a China Scholar-ship Council to Q.L. We thank Dr. Weijun Cai and Dr. ChristopheRabouille for their thoughtful comments on early versions of themanuscript. We are also grateful to Dr. Meagan Eagle Gonneeafor helpful suggestions related to the groundwater flux model andto Dr. Caroline Slomp for the calcium measurements. We wouldalso like to acknowledge three anonymous reviewers for their crit-ical and constructive comments, which resulted in substantialimprovements to the manuscript.

APPENDIX A. SUPPLEMENTARY DATA

Supplementary data associated with this article can befound, in the online version, at http://dx.doi.org/10.1016/j.gca.2017.01.041.

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