Liu, Zhiyuan, Longjun Zhang, Wei-Jun Cai, Liang Wang, Ming Xue

Limnol. Oceanogr., 59(2), 2014, 413–426
2014, by the Association for the Sciences of Limnology and Oceanography, Inc.
doi:10.4319/lo.2014.59.2.0413
E
Removal of dissolved inorganic carbon in the Yellow River Estuary
Zhiyuan Liu,1 Longjun Zhang,1,* Wei-Jun Cai,2 Liang Wang,1 Ming Xue,1 and Xiangshang Zhang 1,3
1 Key
Laboratory of Marine Environment and Ecology, Ministry of Education, College of Environmental Science and Engineering,
Ocean University of China, Qingdao, P.R. China
2 School of Marine Science and Policy, University of Delaware, Newark, Delaware
3 Qingdao Maritime Safety Administration, Qingdao, P.R. China
Abstract
The Yellow River of China runs mainly through an arid and semiarid midlatitude region that has experienced
substantial anthropogenic and climatic change. This area includes the carbonate-rich Loess Plateau and carries
water of exceptionally high carbonate content. To investigate the processes by which dissolved inorganic carbon
(DIC) is biogeochemically modified as the river approaches the sea, a multipronged field investigation was
conducted in the Yellow River estuary, 2005–2009. The project included four research cruises (spring and fall), a
year of monthly sampling at a lower-river hydrological station (Lijin), and in situ bottle incubations. Our study
revealed that 4–11% of the Yellow River DIC was removed from the water column in the estuarine mixing zone
and thus was not transported to the sea. DIC removal was greater in the spring and occurred at a higher salinity
range than in the fall. As a unique feature of the Yellow River estuary, calcium carbonate (CaCO3) precipitation
was nearly as important as net biological production in the DIC removal. Longer freshwater–seawater mixing
distances (and times) and higher DIC concentrations in the freshwater end member also promoted net biological
production and CaCO3 precipitation, thus encouraging DIC removal.
Estuaries are regions of active land–ocean interaction.
Globally, a total of ,0.34 Gt (1015 g) of dissolved inorganic
carbon (DIC) is annually exported to the ocean by rivers
(Ludwig et al. 1998; Mackenzie et al. 2004; Lerman et al.
2007). DIC usually exhibits nonconservative behavior
during estuarine mixing (Cai and Wang 1998; Abril et al.
2003; Cai 2003) due to a number of processes: strong
internal biogeochemical activity (Wollast 2003); material
exchange with surrounding environments (Hans et al.
2011), in particular coastal wetlands (Cai 2011); and
complex sediment dynamics during mixing. As a result,
riverine carbon fluxes can be over- or underestimated if
these physical and biogeochemical processes are not
considered.
Previous research has shown that many factors regulate
the nonconservative behavior of DIC in river-dominated
estuaries. In most cases, biological production (resulting in
removal of DIC from the water column) or respiration
(resulting in addition of DIC) has been identified as the
dominant factor. For instance, in several estuaries along
the eastern coast of the United States, DIC has been
observed to increase nonconservatively with salinity due to
bacterial respiration during mixing (Cai and Wang 1998;
Cai et al. 1999; Raymond et al. 2000). Occasionally,
however, CaCO3 dissolution (addition of DIC) has been
reported to dominate (Abril et al. 2003; Ortega et al. 2005,
2008). Within the Mississippi River plume, Cai (2003)
observed that DIC losses occurred where salinity was , 30
and noted the cause might be a combination of biological
production and CaCO3 precipitation. Guo et al. (2012) and
Huang et al. (2012) subsequently concluded that biological
production was the main cause for the plume DIC removal
and the effect of CaCO3 precipitation was minor.
Identifying the biogeochemical mechanisms responsible
for DIC production and removal can be further complicated by the existence of time-varying or multiple riverine
end members (Officer 1979). These nonbiogeochemcial
influences can produce an apparently nonconservative
DIC–salinity relationship, and detailed mixing schedules
are often required to elucidate the contributing factors.
This type of complexity has been seen in the Mississippi
River and Atchafalaya River plumes (Guo et al. 2012;
Huang et al. 2012), the Pearl River estuary of China (Cai et
al. 2004), and the Scheldt River plume (Hellings et al.
2001). Ammonium oxidation, denitrification, and air–water
CO2 exchange may also contribute to nonconservative
behavior of DIC in estuaries (Cai et al. 2004; Dai et al.
2008; Guo et al. 2008). Therefore, understanding the
biogeochemical processes controlling DIC distributions
and quantifying net DIC fluxes and variations in estuaries
constitutes a rather challenging step in the study of carbon
transport from rivers to oceans (Vanderborght et al. 2002).
The Yellow River of China constitutes an especially
interesting and relevant case study with respect to estuarine
DIC dynamics. The river is located in a midlatitude zone
with an arid and semiarid climate, and its middle reach
flows through the carbonate-rich Loess Plateau. Its
drainage basin is characterized by high evaporation (Yang
et al. 2004; Chen et al. 2006), a high land-utilization rate,
and severe water and soil losses (Wang et al. 2006, 2007).
These characteristics and, most important, intense weathering of the carbonate-containing loess, result in Yellow
River HCO{
3 concentrations that are among the highest of
all the world’s large rivers (Cai et al. 2008).
In addition, the Yellow River has been dramatically
affected by human activities, including over 2000 years
of agriculture and, recently, several decades of intensive
irrigation and damming and a decade of direct human
* Corresponding author: [email protected]
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Liu et al.
Fig. 1.
Yellow River estuary: (A) study area location and (B) survey station locations.
regulation of water and sediment discharge (Wang et al.
2007). The water and sediment regulation (WSR) plan,
which is the only one regularly executed in the world, was
designed to flush out sediment accumulation in reservoirs
and riverbed in the lower reach of the river once a year. The
WSR activities constitute an extreme human disruption of
river-matter transport (Zhang et al. 2013). In particular,
estuarine mixing conditions differ greatly between the preand post-WSR periods. Understanding how WSR events
change estuarine DIC dynamics represents an important
step to understand and predict how human intervention of
natural systems affects carbon cycling and fluxes.
DIC dynamics in the Yellow River estuary have not been
well characterized. On the basis of data from a single
survey, Cauwet and Mackenzie (1993) speculated that DIC
removal occurred in the low-salinity (S , 5) area of the
estuary, probably caused by CaCO3 precipitation. However, no further explanation or justification was provided, nor
has subsequent fieldwork been reported to confirm this
preliminary observation and attribution. The effect of
biological production on DIC removal has been documented in the Mississippi River plume (Cai 2003; Guo et al.
2012; Huang et al. 2012), but the Yellow River is clearly
different from the Mississippi in several important aspects.
For instance, for the period 1950–2000, Mississippi River
discharge increased by 30%, whereas Yellow River discharge decreased by . 50% (Milliman et al. 2008). DIC
removal within the Mississippi River plume occurs where
salinity is , 30 (Cai 2003), whereas in the Yellow River
estuary, removal occurs in the lower-salinity zone (Cauwet
and Mackenzie 1993). Thorough study of DIC removal
mechanisms in the distinctive Yellow River estuary will
serve to broaden our understanding of inorganic carbon
distributions, production and removal processes, and fluxes
at the highly variable land–ocean interface.
To identify the major processes controlling DIC removal
in the Yellow River estuary, we collected lower-river and
estuarine concentration data for water-column DIC, total
alkalinity (TAlk), and other relevant parameters. DIC
consumption due to net biological production and CaCO3
precipitation was also measured by bottle incubation. These
data were analyzed and interpreted according to first principles of carbonate chemistry and estuarine mixing.
Methods
Site description—The Yellow River discharges into the
Bohai Sea, which is located north of China’s Shandong
Peninsula (Fig. 1). The lower reach of the river has only a
single channel with no distributary branches. Due to rapid
sediment deposition, the river mouth (delta) area has been
advancing seaward at a recent rate of 1.1–2.6 km yr21 (Hu
et al. 1998; Qiao et al. 2010). The estuary is a typical riverdominated estuary with weak tides; the tidally affected zone
extends approximately 10–20 km upriver (Huang and Lu
1995).
The Yellow River ranks low among the world’s large
rivers in terms of water discharge (with a 1919–1995
average of 59 3 109 m3 yr21), but it ranks first or second in
terms of suspended sediment load (with a 1950s–1970s
average of 1.6 3 1015 g yr21; Chen et al. 2005, 2006). In
recent decades, annual water and sediment fluxes of the
Yellow River have declined dramatically due to climate
change, reservoir construction, and irrigation-related withdrawals (Wang et al. 2006, 2007; Dai et al. 2009).
Since 2002, discharges have been deliberately manipulated to prevent sediment accumulation in the downstream
channel and reservoirs. In that year, the Yellow River
Conservancy Commission began conducting WSR trials in
the river’s middle reach during the early wet season (late
June and early July). In 2005, the program became
operational. The annual WSR scheme consists of two
periods: the water release period, during which a large
amount of water is released from the Xiaolangdi Reservoir
(,700 km upstream from Lijin) for the purpose of flushing
sediment from the downstream waterway; and the sediment
release period, during which water is discharged from two
upstream reservoirs (Sanmenxia and Wanjiazhai) for the
purpose of flushing sediment from the Xiaolangdi Reservoir. (For more details, see http://www.yellowriver.gov.cn/
and Zhang et al. 2013.)
DIC removal in the Yellow River Estuary
These activities constitute an extreme human disruption
of river-matter transport (Zhang et al. 2013). As a result,
estuarine mixing conditions differ greatly between the preWSR (spring) and post-WSR (fall) periods. The summertime reservoir releases deliver to the estuary large quantities
of sediment, most of which is deposited to form a cluster of
barrier bars (Wang et al. 2005). This major depositional
event temporarily confines estuarine mixing to short
distances and times. In subsequent months, the barrier
bars are gradually flushed away by less turbid river flow or
eroded by tidal action. By spring of the next year, the bars
have nearly disappeared and estuarine mixing conditions
have returned to normal.
Research cruises—Four cruises were carried out in the
Yellow River estuary over the first 5 yr of the WSR
operational program: September 2005, April 2006, May
2009, and September 2009. All investigations started from
the Xintan floating bridge (Fig. 1). Because the river
channel morphology changed yearly or even seasonally,
the sampling locations were not the same for every cruise.
The 2009 site locations were different from those in 2005
and 2006, but we do not expect this differnece to cause
cruise-to-cruise differences in chemical and biological
processes. Station locations during each cruise were chosen
according to salinity variations. The river’s most seaward
hydrological station (Lijin), where salinity was , 0, was
always sampled as the freshwater end member.
Average discharges during our cruises were as follows:
454 m3 s21 in September 2005, 372 m3 s21 in April 2006,
270 m3 s21 in May 2009, and 380 m3 s21 in September 2009.
On all four cruises, discrete surface-water samples were
collected and analyzed for salinity (S) and concentrations of
total suspended solids (TSS), chlorophyll a (Chl a), DIC,
TAlk, particulate inorganic carbon (PIC), and nutrients
(nitrate, ammonia, and phosphate). During the 2009 cruises,
surface-water salinity, temperature (T), saturation of dissolved
oxygen (DO%), and partial pressure of carbon dioxide ðPCO2 Þ
were also measured continuously. Two additional cruises were
conducted in April and September 2004, but those data are not
extensively discussed here because Chl a and freshwater endmember data are not available from those cruises.
Lijin monthly monitoring—Between July 2010 and
August 2011, monthly investigations of DIC were conducted at the Lijin hydrological station. Discharge data were
obtained from China Hydrology (http://www.hydroinfo.
gov.cn/). River discharges were lower in winter and spring
than in summer and fall.
Sampling and analytical methods—During the cruises of
September 2005 and April 2006, surface S was measured at
each discrete sampling station with a portable salinometer
(3SYA2-2) that had a precision of 0.005. In May and
September 2009, underway surface T and S were measured
continuously using an SBE 45 MicroTSG (Sea-Bird
Electronics). DO% was measured with a YSI 5000 oxygen
analyzer (YSI) that had been calibrated using the Winkler
titration method. CO2 partial pressure was measured using
a nondispersive infrared spectrometer (LI-7000, LI-COR)
415
coupled to a shower-head equilibrator. All four parameters
were automatically recorded each second and then averaged over 1 min, yielding nominal precisions of 0.002uC
for temperature, 0.005 for salinity, 0.1% for DO%, and
, 0.1%, for PCO2 (Zhang et al. 2010, 2012).
Discrete water samples were collected into 5 liter Niskin
bottles for later analysis of TSS, Chl a, DIC, TAlk, PIC,
and nutrients. All samples were filtered in situ and were
analyzed within 1 week. TSS samples were collected onto
preweighed cellulose acetate membrane filters (pore size
0.45 mm) and preserved at 220uC. The membranes were
later dried at 45–50uC and then weighed on an electronic
balance (AL104, Mettler Toledo) with a precision of
0.0001 g. Chl a samples were collected onto glass fiber
membrane filters (0.7 mm pore size; Whatman GF/F) at a
pressure of , 0.05 MPa. Saturated magnesium carbonate
solution was added to the membranes, which were
preserved at 220uC. Membrane samples were later
extracted with 90% acetone, and Chl a was determined
using a fluorescence spectrophotometer (F-4500, Hitachi
High-Technologies). The standard curve was constructed
using Sigma C-5753 chlorophyll (Sigma-Aldrich). DIC and
TAlk samples were filtered through cellulose acetate
membranes (0.45 mm), poisoned with saturated mercury
bichloride (final concentration , 0.02% by volume;
Dickson and Goyet [1994]), and preserved at 4uC. DIC
was measured using a total organic carbon analyzer (TOCVCPN, Shimadzu). TAlk was determined by Gran titration,
using an alkalinity titrator (AS-ALK2, Apollo SciTech).
Measurements of DIC and TAlk were both calibrated
against Certified Reference Material (provided by A. G.
Dickson from Scripps Institution of Oceanography) at a
precision and accuracy level of 0.1%. PIC samples were
collected onto glass fiber membrane filters (0.7 mm, Whatman GF/F) and preserved at 220uC. PIC was determined
using a total organic carbon analyzer coupled to a
combustion device for solid samples (TOC-VCPN and
SSM-5000A, Shimadzu). The method of phosphoric acid
extraction and combustion (200uC) has a measurement
precision of 1%. This method necessitated the use of the
GF/F filter (Aucour et al. 1999), which has a pore size that
is greater than the size that operationally defines ‘‘particulates’’ (0.45 mm). As a result, this method likely underestimates PIC concentrations but to an extent that we believe
is insignificant. Samples collected for nutrient analysis
(NO3-N, NH4-N, and PO4-P) were filtered through pretreated cellulose acetate membranes (0.45 mm; immersed in
0.1% HCl solution over 24 h, then flushed with Milli-Q
water). The filtrate was poisoned with chloroform and
preserved at 220uC. The nutrient samples were later
quickly thawed, then analyzed using an AutoAnalyzer 3
(Bran + Luebbe). Samples collected for determination of
calcium ion (Ca2+) concentrations were filtered through
cellulose acetate membrane filters (0.45 mm) and then
analyzed by ethylene glycol bis (2-aminoethyl) tetraacetic
acid titration. The precision of this method is 0.1%
(Tsunogai et al. 1968).
In situ incubations—During the two 2009 cruises, in situ
incubations were carried out using the dark–light bottle
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Liu et al.
Fig. 2. Distribution of continuously measured surface temperature, salinity, DO%, and
PCO2 in the Yellow River estuary: (A) May 2009 and (B) September 2009. The open circles
indicate where discrete water samples were collected.
method. Spring incubation salinities were 2.4, 5.1, and 10.0;
fall salinities were 0.5, 2.5, 5.4, 10.3, and 19.6. For each
salinity level, water was drawn into six light bottles and five
dark bottles. The bottles were then placed into 0.5 m deep
chests for incubation; temperature was maintained with
continuously flowing estuarine surface water. For the light
bottles, DO%, DIC, and TAlk were measured every 2 h
from 06:00 h to 18:00 h (light intensity was . 10,000 lux).
For the dark bottles, DO% was measured every 8 h. The
total incubation time was 40 h. During the day, in situ light
intensity was measured every 0.5 h with an illumination
photometer (TES-1339 Light Meter Pro, TES Electrical
Electronic).
Estimation of DIC and TAlk removal—To calculate in
situ removal of DIC and TAlk from the water column, we
first used observed end-member concentrations to obtain
theoretical dilution lines for constituents within the estuary
(Officer 1979; Regnier et al. 1998). Values on the dilution
lines (DICT and TAlkT) are the concentrations that would
have theoretically been observed in the case of conservative
mixing. Removal of DIC and TAlk from the water column
(DDIC and DTAlk) were then obtained as the difference
between the theoretical (T) and measured (M) constituent
concentrations. The PIC analyses were not used to calculate DIC removal rates.
Results
Underway temperature, salinity, DO%, and PCO2 —In
spring 2009, a downstream one-way survey of surface
conditions was conducted from the Xintan floating bridge
to offshore waters; in the fall, data were collected on the
return leg as well (Fig. 1). Distributions of surface T, S,
DO%, and PCO2 are shown in Fig. 2. Because all four
parameters were nearly constant within the first 7 km
downstream of the floating bridge, data for this stretch are
not shown. Downstream of the 7 km point, salinity
increased seaward within the estuarine mixing zone, with
dramatic fluctuations observed in the neighborhood of the
12 km mark (at 12–13 km in the spring and 11–12 km in the
fall). This pattern is typical of a river-dominated estuary
with weak tides and a clear interface between freshwater
and salt water. Within the mixing zone, seawater along the
channel bottom moves upstream against the river current
in a wedge shape (Huang and Lu 1995) and exhibits
reciprocating flow.
Variations of salinity, DO%, and PCO2 were similar in
the spring and fall (Fig. 2). In the freshwater end member,
DO% was relatively low, whereas PCO2 was high (though
still generally , 75 Pa; 1 Pa 5 10.1 matm). At the sharp
mixing interface (, 12 km), salinity, DO%, and PCO2 all
fluctuated markedly over short distances. In general, DO%
increased downstream (i.e., in the direction of increasing
salinity), whereas PCO2 decreased. On approach to the highsalinity end member (S . 25), DO% remained oversaturated and PCO2 decreased to , 40 Pa.
Some differences between the two cruises were also
observed. In spring 2009, DO% increased downstream
from 93% to 115%, becoming oversaturated where S , 18.
In the fall, DO% ranged from 85% to 105%, and saturation
was not observed until S . 25; at S , 18, DO% was only
95%. Spring temperatures decreased markedly downstream
(from 21.5uC to 17.4uC), showing a negative correlation
with salinity (i.e., the seawater was colder than the river
water). In the fall, however, surface temperatures were
nearly constant over the entire transect, except for a 1uC
increase at , 12.0–12.5 km, where the water was very
shallow. In the spring, mixing-zone fluctuations were more
drastic and the mixing zone extended over a much greater
distance than in the fall.
DIC and TAlk—DIC and TAlk exhibited strong spatial
and seasonal variation (Fig. 3), as well as some features
unique to the Yellow River: (1) The Yellow River has
extremely high DIC and TAlk values (,3178 mmol L21 and
DIC removal in the Yellow River Estuary
417
Fig. 3. DIC, TAlk, and salinity in the Yellow River estuary. The solid lines are theoretical
dilution lines for DIC and TAlk under conditions of conservative behavior; the dotted lines
represent extrapolation of the high-salinity line segments back to S 5 0. (Lines are not shown for
2004 because freshwater end-member data are not available from these cruises.) (A, C, E) show
spring data and (B, D, F) show fall data. Note the difference in vertical scale between the spring
and fall graphs.
3242 mmol L21, respectively), likely the highest among large
rivers worldwide (Cai et al. 2008). (2) DIC and TAlk
concentrations were consistently higher in spring than in
fall, opposite the pattern of river freshwater discharge
(which is low in spring and maximal in fall). Spring DIC
ranges were 2567–3718 and 2411–3384 mmol L21 (in 2006
and 2009, respectively); fall ranges were 2482–2682 and
2155–2927 mmol L21 (in 2005 and 2009). (3) DIC and TAlk
decreased seaward with increasing salinity but with a
nonlinear relationship that clearly indicates removal of
DIC and TAlk in the low-salinity zone. This phenomenon
was observed during all our surveys.
The solid lines in Fig. 3 show the theoretical dilution
lines for DIC and TAlk, calculated assuming conservative
mixing of the end-member waters (Officer 1979; Regnier et
al. 1998). Constituent removals and additions (DDIC and
DTAlk) at any survey station can be obtained as the
difference between the theoretical (T) values on the line and
the measured (M) concentrations—i.e., on Fig. 3, the
vertical distances between the dilution lines and the DICM
and TAlkM data points. DDIC and DTAlk values thus
calculated for the Yellow River estuary indicate that
removals were consistently greater than additions, with
net changes being larger in the spring than in the fall. The
salinity ranges where DIC removal occurred were higher
(i.e., more marine in character) in the spring than in the fall:
S ,18 and 13 during the spring cruises (2006 and 2009,
respectively) but , 10 during the fall cruises. DIC removal
and TAlk removal were not coincident in space: the salinity
ranges where DIC removal occurred were slightly higher
(more marine) than where TAlk removal occurred (Fig. 3).
At the seaward end of the mixing zones, DIC and TAlk
varied linearly with salinity, indicating no removal or
addition of these constituents in the later (more downstream) stages of mixing (Fig. 3). Extrapolation of the DIC
conservative-mixing line segments from the high-salinity
areas back to S 5 0 (shown as dotted lines in Fig. 3) yields
the so-called effective concentrations of DIC exported from
freshwater end members (DICe) when nonconservative
removals and additions are considered. The difference
between DICe and actual DIC measured in the freshwater
end member is equal to the DIC that was removed during
mixing and therefore not transported to the sea (Boyle et al.
1974; Cai and Wang 1998). On the basis of this approach,
we estimate that approximately 5.3% (6 0.09%) and 11%
(6 0.35%) of the Yellow River DIC was removed from the
water column (not transported to the sea) for the fall 2005
and spring 2006 cruises, respectively. Approximately 5.7%
(6 0.04%) and 3.8% (6 0.03%) of river DIC was removed
for the spring and fall 2009 cruises. Thus, in the Yellow
River estuary, DIC in situ percentage losses are greater in
spring than in fall.
TSS and PIC—Concentrations of TSS were consistently
higher in the freshwater end members than in seawater
(Fig. 4A), but clear differences in the TSS–S relationships
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Liu et al.
Fig. 4. Correlations between (A) TSS and salinity and (B) PIC% and salinity in the Yellow
River estuary.
were observed between the earlier and later cruises. During
the first two cruises (fall 2005 and spring 2006), TSS
concentrations in the freshwater end members were very
high (1407 mg L21 and 1970 mg L21, respectively);
concentrations during the final two cruises were significantly lower (900 mg L21 and 708 mg L21 in spring and fall
2009). On all four cruises, TSS declined sharply where S ,
0.5, then decreased almost conservatively downstream as
salinity increased.
PIC% (i.e., PIC : TSS 3 100%) also decreased with
increasing salinity (Fig. 4B) but with a different downstream pattern than TSS. Also unlike TSS, PIC% exhibited
strong seasonal differences. During the spring cruises,
PIC% remained high until appreciable levels of salinity
were reached (S . 23 for spring 2006 and S . 15 for spring
2009). This pattern of sustained high PIC% suggests that
PIC production was occurring in the low- and mid-salinity
regions of the estuary. During the fall cruises, in contrast,
PIC% deceased rapidly where S , 5. (As noted in the
Methods section, the PIC% numbers may be slight
underestimates. The patterns and trends are, however,
robust because the same method was used on all cruises.
We do not use absolute PIC% values to quantify DDIC or
other derived quantitites.) We have no direct evidence that
spring PIC production in the low-to-moderate salinity area
was caused by CaCO3 precipitation, but the overlapping
salinity ranges of PIC production and DIC and TAlk
Fig. 5.
removal strongly suggest in situ CaCO3 precipitation as a
common cause.
Nutrients—Nitrogen, phosphorus, and other nutrients
are important in influencing phytoplankton growth.
During the 2009 cruises, nitrate (NO3-N) concentrations
were as high as 265–284 mmol L21 in the freshwater end
members, but then decreased with increasing salinity
(Fig. 5A). Concentrations of ammonia (NH4-N) and phosphate (PO4-P) were very low (Fig. 5B, C). The fall distributions indicate that some in situ generation of these
constituents may have been occurring in conjunction with
mixing (Fig. 5B, C). Fall NH4-N and PO4-P concentrations
were higher than spring concentrations, consistent with the
observation of lower DO% in the fall (Fig. 2). The highest
NH4-N concentration was 5.52 mmol L21 (much lower than
NO3-N); NH4-N could not have been the main nitrogen
source for phytoplankton growth. PO4-P concentrations
ranged from 0.03 to 0.40 mmol L21, sufficient for
phytoplankton photosynthesis.
Chl a—During the fall 2005 and spring 2006 cruises, the
observed Chl a concentration ranges were 0.35–5.96 and
0.88–10.5 mg L21, respectively (Fig. 6A). Concentrations in
the mixing zones were higher than in either end member,
with the highest values observed at S 5 5 (fall 2005)
and S 5 8 (spring 2006). For the 2009 cruises, Chl a
Nutrient concentrations as a function of salinity in the Yellow River estuary: (A) nitrate, (B) ammonia, and (C) phosphate.
DIC removal in the Yellow River Estuary
419
Fig. 6. Chlorophyll a concentrations as a function of salinity in the Yellow River estuary:
(A) cruises of September 2005 and April 2006, and (B) cruises of May and September 2009.
concentrations were much higher in freshwater than in
seawater. Spring values generally declined with increasing
salinity; fall values remained relatively high until the point
where S .12 (Fig. 6B).
DO%, DIC, and TAlk changes during in situ incubations—
From variations in measured DO%, DIC, and TAlk in
the 2009 light and dark bottle incubations, rates can be
calculated for these processes: oxygen production in the light
bottles (FDO(light); net primary production), oxygen uptake in
the dark bottles (FDO(dark); respiration), and DIC and TAlk
consumption in the light bottles (FDICM and FTAlkM).
(Day lengths for the calculated rates are 12 h for light-bottle
quantities and 24 h for dark-bottle quantities.) Several
interesting features were observed (Table 1): (1) net primary
production rates, FDO(light), were consistently much higher than
respiration rates, FDO(dark); (2) rates of DIC consumption, FDICM, were quite high in low-salinity waters and
decreased with increasing salinity; (3) TAlk removal,
FTAlkM, was observed only in low-salinity (S , 5) waters;
and (4) for a given salinity, the DIC consumption rate
attributable to net biological production was greater in the
spring than in the fall.
Seasonal DIC variations—DIC concentrations were
measured at the Lijin hydrological station from July 2010
to August 2011 (Fig. 7). These concentration data agree very
well with DIC concentrations measured in freshwater end
members during our four cruises (open circles on Fig. 7). In
general, DIC was negatively correlated with discharge except
during the midsummer WSR periods, when large quantities
of water were being released from upstream reservoirs. DIC
concentrations were low during the wet, high-discharge
season of late summer and early fall, then increased steadily.
Highest DIC was observed during the dry, low-discharge
period of winter and early spring. Concentrations thereafter
decreased into the summer. Negative correlations of DIC
with discharge have been observed in many rivers (Cai et al.
2008).
WSR releases occurred on 19 June–02 July (water release
period) and 03 July–10 July (sediment release period) in
2010 and 19 June–03 July (water release) and 04 July–12
July (sediment release) in 2011. In association with these
releases, DIC concentrations at Lijin were observed high on
06 July 2010 and again the following year on 25 June and
08 July (Fig. 7). During the 2011 WSR period, for instance,
a large quantity of water (, 2600 m3 s21) was released from
the Xiaolangdi Reservoir on 19 June. This water reached
the Lijin station on 25 June, as seen in the roughly fivefold
increase in discharge. Sediment that had been previously
deposited in the Yellow River watercourse was flushed out
and resuspended, leading to a sharp increase in TSS on 25
Table 1. Results of the 2009 in situ incubation experiments. FDO(light) and FDO(dark) give measured net primary production
and respiration rates. FDICM and FTAlkM give measured consumption rates of DIC and TAlk in the light bottles. Dashes represent no
experimental data. (Day lengths: 12 h for light bottles and 24 h for dark bottles.)
Salinity
Spring
Fall
2.4
5.1
10.0
0.5
2.5
5.4
10.3
19.6
FDO(light)
(mmol O2 L21 d21)
188
151
119
123
91
68
48
13
(610.0)
(63.1)
(612.5)
(60.8)
(60.4)
(60.5)
(61.0)
(60.9)
FDO(dark)
(mmol O2 L21 d21)
223
223
211
217
215
214
210
24
(60.2)
(60.2)
(60.2)
(60.4)
(60.2)
(60.4)
(60.2)
(60.2)
FDICM
(mmol C L21 d21)
2202
2153
2108
2212
2118
272
249
212
(61.7)
(60.7)
(60.9)
(62.5)
(66.6)
(66.2)
(62.9)
(60.9)
FTAlkM
(mmol C L21 d21)
—
—
—
2207 (63.7)
284 (60.4)
223 (61.4)
27 (61.0)
22 (60.7)
420
Liu et al.
Fig. 7. Discharge rates and DIC concentrations measured at the Lijin station from July 2010
to August 2011. The open circles (with dates) show DIC concentrations measured in the
freshwater end members of the four earlier cruises.
June (up to 6685 mg L21, about 10 times as much as preWSR). DIC increased 25%, to achieve a concentration of
3458 mmol L21. Discharge was still high on 08 July, but
sediment and DIC concentrations had by then begun to
decline. After the WSR ended on 12 July, discharge rates
and DIC concentrations returned to normal. This pattern
of high DIC in association with high discharge is contrary
to the pattern of natural seasonal variability.
On 01 August 2010, very low DIC was observed at Lijin,
probably due to a dilution effect associated with heavy
rainfall. During 24 July–03 August, seven heavy downpours
were recorded at the middle or lower reaches of the Yellow
River. At the Huayuankou gauging station (in the middle
Yellow River basin) and at the Lijin station, associated
discharges were 68% and 108% greater than during the flood
period of 2009 (http://www.yellowriver.gov.cn/).
The annual average concentration of DIC at Lijin was
3158 mmol L21. This value is higher than the 2591 mmol L21
estimate given in the synthesis of Cai et al. (2008), which
was based on earlier observations.
estuaries (e.g., those of the Changjiang, Pearl, Mississippi,
and Hooghly rivers; Mukhopadhyay et al. 2002; Zhai et al.
2005, 2007). WSR events, which can change DIC and TAlk
transport patterns within a short period, all fell outside the
times of our four cruises. Thus, we conclude by the process
of elimination that internal removal processes such as net
biological production and CaCO3 precipitation must be
responsible for the observed nonconservative behavior of
DIC in the Yellow River estuary.
Due to the chemical changes that occur when river water
encounters seawater—i.e., changing pH and ionic strength
and especially rapidly increasing Ca2+—calcium carbonate
precipitation is likely to occur within the estuary. Gan et al.
(1983), on the basis of calculations of calcite saturation
index (SICalcite), noted that calcite was oversaturated in the
Yellow River mainstream. We calculated SICalcite from our
2005–2009 estuarine cruise data:
Discussion
where [ ] represents concentration and KCalcite is the
solubility product of calcite (a polymorph of the mineral
calcium carbonate). For water in an equilibrium state with
respect to calcite, log10(SICalcite) 5 0. A value , 0 indicates
undersaturation, and a value . 0 indicates oversaturation
(Mucci 1983; Neal 2002).
On all four cruises, CaCO3 was oversaturated in lowsalinity areas (Fig. 8). Highest values of SICalcite were
observed during the spring cruise of 2006, when calcite
oversaturation occurred over a broad range of salinity,
from 0 to 18. DIC removal was also greatest during this
cruise (Fig. 3). During the two fall cruises (2005 and 2009),
SICalcite and DIC removal were much lower, with the zones
of calcite oversaturation (Fig. 8) and DIC removal (Fig. 3)
restricted to much lower (less marine) salinity ranges. These
Causes of estuarine DIC removal—Many processes can
contribute to nonconservative DIC behavior in estuaries
(Abril et al. 2003; Cai et al. 2004)—e.g., net biological
production and respiration, CaCO3 dissolution and precipitation, ammonium oxidation, denitrification, tributary or
local anthropogenic inputs (e.g., sewage), and CO2 air–sea
exchange. For the Yellow River estuary, some of these
processes can be eliminated from consideration. This estuary
has no tributaries and no direct anthropogenic inputs. Its
waters are characterized by high DO% and relatively low
CO2 degassing rates. During the 2009 cruises, DO% ranged
between 85% and 115%, whereas PCO2 ranged between 35 Pa
and 75 Pa, far less than observed in some other large
log10 ðSICalcite Þ~
{log10 ðKCalcite Þ
log10 Ca2z | CO2{
3
ð1Þ
DIC removal in the Yellow River Estuary
{
z
106CO2 z16NO{
3 zH2 PO4 z122H2 Oz17H
~ðCH2 OÞ106 ðNH3 Þ16 ðH3 PO4 Þz138O2
Fig. 8. Relationship between the log of the calcite saturation
index (SICalcite) and salinity. The dotted line, where log10(SICalcite)
5 0, represents saturation with respect to calcite. Values . 0
indicate oversaturation; values , 0 indicate undersaturation.
results suggest that CaCO3 precipitation is probably
responsible for DIC removal in the low-salinity areas of
the estuary. However, the salinity zones of linked processes
were not always coincident. The salinity range within which
DIC removal occurred was a little broader than the salinity
range of calcite oversaturation. The salinity zones for TAlk
removal and calcite oversaturation were almost the same.
Other factors besides chemical precipitation of CaCO3
(e.g., biological production) clearly contributed to the
removal of DIC from estuary waters.
The ratio DDIC : DTAlk can be used to determine which
processes—CaCO3 precipitation or net biological production or both—may be influencing the removal of DIC from
the water column (Cai et al. 2004; Ortega et al. 2005, 2008).
CaCO3 precipitation alone removes DIC and TAlk in a
ratio (DDIC : DTAlk) of 1 : 2. When CaCO3 precipitation
and net biological production co-occur, DDIC : DTAlk .
1 : 2 because net biological production results in a sharp
decrease in DIC but very little change in TAlk (Redfield
et al. 1963):
421
ð2Þ
The greater the DDIC : DTAlk ratio, the greater the
contribution from net biological production.
In the Yellow River estuary (Fig. 9), DDIC : DTAlk
values were 0.946 (fall 2005), 1.29 (spring 2006), 1.70
(spring 2009), and 1.88 (fall 2009). The fact that DDIC :
DTAlk was consistently greater than the benchmark value
of 0.5 suggests that not only CaCO3 precipitation but also
net biological production contributed to DIC removal
during all four cruises. The relative effect of net biological
production vs. CaCO3 precipitation was strongest during
the 2009 cruises.
Biological production consumes not only DIC but also
NO3-N, with a DIC : NO3-N ratio of 106 : 16 (Eq. 2). In the
Yellow River estuary, evidence of NO3-N removal was seen
in the low-salinity DIC-removal area (Fig. 5A), but the
NO3-N effect was subtle because river-water concentrations
were so high.
If DIC and TAlk removal in the low-salinity area (i.e.,
where S , 18) is attributed to net biological production
and CaCO3 precipitation, the following equations can be
derived:
DDIC~DDICOC zDDICCaCO3
ð3Þ
DTAlk~DTAlkOC zDTAlkCaCO3
ð4Þ
CaCO3 precipitation removes DIC and TAlk in a 1 : 2 ratio,
DDICCaCO3 =DTAlkCaCO3 ~1=2
ð5Þ
whereas according to Eq. 2, the DIC and TAlk change ratio
due to biological production is:
DDICOC =DTAlkOC ~{106=17
ð6Þ
Thus, DIC and TAlk variations caused by biological production (DDICOC and DTAlkOC) and CaCO3 precipitation
Fig. 9. Relationship between DDIC and DTAlk: (A) fall 2005 and spring 2006; (B) spring
and fall 2009.
422
Liu et al.
Fig. 10. Contributions of net biological production and CaCO3 precipitation to changes in
DIC concentration: (A) April 2006, (B) September 2005, (C) May 2009, and (D) September 2009.
Negative DDIC indicates DIC removal.
(DDICCaCO3 and DTAlkCaCO3) can be calculated by solving
Eqs. 3–6. It is important to note that DDICOC and
DDICCaCO3 are accumulative properties that include removal
signals inherited from processes occurring upstream of the
station location (i.e., between the riverine end member and
the station in question).
Calculated DDICCaCO3 and DTAlkCaCO3 values (Fig. 10)
demonstrate that DIC removal was strongest during the
spring cruises, with DDIC decreasing in the following
order: April 2006, May 2009, September 2009, and
September 2005. The salinity ranges over which DIC
removal occurred also decreased in the same order. These
rankings are consistent with our qualitative assessment
above (Fig. 3). In the first two cruises (Fig. 10A,B), the
contribution of net biological production to DIC removal
was nearly equal to or only slightly greater than that of
CaCO3 precipitation. During the latter two cruises,
however, the contribution of net biological production to
DIC removal was almost twice that of CaCO3 precipitation
(Fig. 10C,D).
Insights from in situ incubations—Our in situ bottle
incubations provide direct evidence for a mechanistic
interpretation of the DIC removal signals (Table 1).
Measured consumption rates of DIC (FDICM) and TAlk
(FTAlkM) decreased with increasing salinity, indicating
more intense DIC and TAlk removal in the low-salinity
area of the estuary. The fact that FDO(light) was always
larger than FDO(dark) suggests that biological production
(which consumes DIC and produces DO) was the primary
biological process. TAlk removal, like DIC removal, was
greatest in the low-salinity (S , 5) incubations, indicating
that CaCO3 precipitation also plays an important role in
DIC removal.
The rate of DIC consumption due to net biological
production (FDICOC) can be calculated from measurements
of FDO(light) and the 106 : 138 ratio of Eq. 2:
FDICOC ~{ð106=138ÞFDOðlightÞ
ð7Þ
The rate of DIC consumption due to CaCO3 precipitation (FDICCaCO3 ) can be similarly calculated from measurements of FTAlkM, the 1 : 2 ratio of DDIC : DTAlk caused by
CaCO3 precipitation (Eq. 5), and the 17 : 138 ratio of
DTAlk : DDO caused by photosynthesis (Eq. 2):
FDICCaCO3 ~FTAlkCaCO3 =2
~ðFTAlkM {FDICOC Þ=2
ð8Þ
~FTAlkM =2{ð17=276ÞFDOðlightÞ
For our Yellow River estuary incubations, these calculations indicate that the total DIC removal rate (FDICM) and
both of its component rates (FDICOC and FDICCaCO3 ) were
highest in the low-salinity bottles (where S , 5). All three
rates decreased with increasing salinity (Fig. 11). This
result further indicates that both net biological production
and CaCO3 precipitation were responsible for the removal
of DIC from the waters of the Yellow River estuary.
Figure 11 also shows that as salinity increased, the relative
importance of CaCO3 precipitation decreased and the
relative importance of biological removal increased. In the
freshest waters, the two component contributions were
nearly equal, but by the location (condition) where S , 20,
biological removal was clearly dominant.
Our incubation conditions were more similar to surfacewater conditions as water was pumped from the surface
into the 0.5 m deep incubation chamber. Nevertheless, solar
DIC removal in the Yellow River Estuary
Fig. 11. DIC consumption during the in situ incubations of
fall 2009. FDICCaCO3 (white) is the rate of DIC consumption attributed
to CaCO3 precipitation, and FDICOC (black) is the rate attributed to net
biological production. FDICM (gray) is the DIC consumption rate
(total) measured in the light bottles.
insolation at the chamber was greater than at lower water
column deeper than 0.5 m. Thus, depth-integrated O2
production rates based on incubation-derived light-bottle
rates may be higher than actual water-column O2 production rates. If so, then our incubation-derived assessment of
the relative importance of CaCO3 precipitation in DIC
removal would be a conservative estimate. Overall, the
FDICCaCO3 : FDICOC ratio (or DDICCaCO3 : DDICOC ratio) is
somewhat lower from the incubation data (Fig. 11) than
from the field data (Fig. 10).
Primary productivity (as reflected in FDO(light)) was much
lower in fall 2009 than in the preceding spring (Table 1),
consistent with the extent of water-column DIC removal
measured during the two cruises (Fig. 3). Light intensity
differed little between the cruises (13,000–92,950 lux vs. 10,100–
93,280 lux), and water temperatures were actually higher in the
fall than in the spring (Fig. 2). The reasons for the lower
primary productivity in the fall merit further investigation.
Seasonal variations of DIC—Because net biological
production and CaCO3 precipitation are important mechanisms of DIC removal in the Yellow River estuary, DIC
distributions were strongly correlated with Chl a concentrations and, as noted above, SICalcite.
Observed Chl a concentrations give only the standing
crop and may have a complicated relationship with in situ
biological production rates. Nevertheless, Chl a distributions (Fig. 6) were consistent with distributions of DIC
removal due to biological production (Fig. 10). For
instance, Chl a concentrations were highest in spring
2006, especially at the mixing interface. DIC removal due
to net biological production (DDICOC) was also highest
during this cruise. Chl a concentrations were lowest in fall
2005, as was the contribution of net biological production
to DIC removal. Looking at spatial (salinity) distributions
in 2009, DDICOC can be seen to reach a maximum where
Chl a concentrations were highest (i.e., where S , 5).
423
As mentioned above, the salinity range where DIC
removal occurred was generally a little broader than the
range where calcite oversaturation was observed (Figs. 3,
8). On three of the four cruises, there were locations in the
estuary where waters were undersaturated with respect to
calcite and TAlk behavior was nearly conservative (where S
. 6 in September 2005, S . 18 in April 2006, and S . 5 in
September 2009). This combination indicates that chemical
precipitation of calcite was unlikely in these areas.
However, DIC removal due to biological production (as
indicated by high Chl a) continued in the undersaturated,
mid-salinity waters. On the fourth cruise (May 2009), where
CaCO3 was undersaturated, Chl a was relatively low.
Our study suggests that biological and chemical removals of DIC are also regulated by DIC concentrations in the
freshwater end members and by freshwater–seawater
mixing distances (and times). Higher DIC concentrations
in freshwater were accompanied by more severe DIC
removal. Data from our four core cruises plus two earlier
ones (Fig. 3) and the monthly Lijin observations (Fig. 7)
show that freshwater DIC concentrations were higher in
spring than in fall. The proportion of DIC removal was
also greater in the spring, as were the corresponding salinity
ranges. This pattern is probably due to the fact that higher
freshwater DIC concentrations resulted in higher SICalcite
values and more spatially extensive zones (wider salinity
ranges) of calcite oversaturation (Fig. 8).
Mixing distances and times are influenced mainly by
estuary geomorphology (shape) and magnitude of discharge.
Longer mixing distances and times provide conditions
favorable for net biological production and CaCO3 precipitation, thus leading to greater DIC removal and a wider
range of salinity encompassed by the DIC removal zone. For
a given season, greater discharges alone can lead to longer
mixing distances. For comparing across different seasons,
though, the effect of estuary shape cannot be ignored. For
example, during our fall cruises, barrier bars in the estuary
limited mixing to a very short distance and time. In spring,
the bars were absent, and mixing distances and times were
much longer (e.g., May 2009, Fig. 2).
In summary, the Yellow River estuary data indicate that
higher Chl a concentrations and higher SICalcite values
generally correlate to greater DIC removal through the
processes of net biological production and CaCO3 precipitation. To the extent that freshwater end-member DIC concentrations and freshwater–seawater mixing distances (and
times) influence production and precipitation, these factors
may also influence DIC removal from estuarine waters.
Causes of DIC increase during WSR events—An important question is why DIC increased during the WSR flood
periods but decreased during natural flood periods (Fig. 7).
This phenomenon is most likely caused at least in part by
the different water origins for the different types of floods.
During the WSR period, Lijin floodwaters came mainly
from upstream reservoirs, which held waters of relatively
high DIC content. Natural floodwaters, however, originated primarily as rainwater of very low DIC. As observed in
other rivers (Cai et al. 2008), dilution effects associated with
heavy precipitation can be significant.
424
Liu et al.
High-DIC source water alone is not sufficient to explain
the DIC increase observed at Lijin during the WSR period.
DIC in the originating Xiaolangdi Reservoir is generally
lower than DIC at Lijin. In July 2007, for example,
Xiaolangdi DIC was 3370 mmol L21, whereas Lijin DIC
was 3540 mmol L21. In July 2009, Xiaolangdi DIC was
3040 mmol L21 and Lijin DIC was 3180 mmol L21 (our
unpubl. data). Other factors must be invoked to explain the
increase in DIC between the reservoir and the river mouth.
During the initial water-release phase of the annual WSR
program, a large amount of clear water is discharged from
the Xiaolangdi Reservoir, and downstream TSS concentrations increase sharply due to resuspension of riverbed
sediments. Suspended sediments in the Yellow River
originate mainly from the Loess Plateau (Chen et al. 2005),
which has a carbonate content as high as 15% (Cai et al.
2008). Near-bottom and sediment-interstitial waters in the
watercourse must therefore have very high DIC concentrations due to prolonged contact with carbonate-rich particles.
Vigorous flushing associated with the WSR release also
likely accelerates weathering and encourages the transformation of PIC to DIC. Thus, we attribute the high DIC
observed at Lijin in association with the Xiaolangdi clearwater release to not only high-DIC source water but also to
flushing out of high-DIC bottom and riverbed water and to
enhanced PIC dissolution. This scenario is similar to
situations in which severe soil loss and erosion lead to
intensified chemical weathering (Van Oost et al. 2007).
During the subsequent sediment-release phase, Xiaolangdi Reservoir sediments and near-bottom and sedimentinterstitial waters are flushed from the reservoir and
transported downriver. These sediments and waters are
high in DIC content, and intense chemical weathering
during the long-distance transport serves to keep riverwater DIC concentrations relatively high. At Lijin,
sediment discharge started to decline on 08 July 2011, even
though water discharge was still high. DIC concentration
declined at that same time, further linking high DIC
concentrations to sediment-related processes. Following
the shutoff of the WSR release on approximately 12 July,
Lijin discharge rates and concentrations of TSS and DIC
declined to normal levels.
We considered the possibility that organic matter
degradation may contribute to the DIC increase observed
at Lijin during WSR events but concluded that this is
unlikely. The ratio of particulate organic carbon (POC) to
Chl a during recent releases was high—as high as , 4000 in
the Yellow River mainstream and , 960 in the reservoirs—
which indicates low biological production and POC
degradation (Zhang et al. 2013). Considerations of composition also argue against an organic contribution to high
DIC. The POC content of suspended sediments during WSR
periods was nearly the same as that of loess (0.6% ; Zhang
et al. 2013), and POC is typically composed mainly of
refractory natural humus of low degradability (Chen et al.
2003, 2004). Indirect contributions from the pool of
dissolved organic carbon (DOC) also seem unlikely. DOC
concentrations in the Yellow River reservoirs (228–
337 mmol L21 ; Zhang et al. 2013) were much lower than
the observed DIC increase at Lijin (,600 mmol L21; Fig. 7).
Possible WSR effects on estuarine DIC removal—Reservoir construction is often called for along rivers in arid and
semiarid regions due to scarce water resources and the
demands of agricultural activity. In the Yellow River,
turbidity is extremely high and a great deal of transported
particulate material is retained within its reservoirs (Zhang
et al. 2013). In recent decades, sediment discharge from the
river has declined dramatically, from average 1.6 3
1015 g yr21 in the 1950s–1970s (Chen et al. 2005, 2006) to
as low as average 0.49 3 1015 g yr21 during the 1980s–
1990s (http://www.yellowriver.gov.cn/).
During WSR release events, Lijin water discharges and
DIC concentrations increase. These factors undoubtedly
result in greater DIC removal within the estuary and
removal over a wider range of salinity. DIC removal due to
chemical precipitation of CaCO3 in particular increases,
but removal due to biological production may be limited
during these periods by the extremely high TSS concentrations. The sharply increased discharge shifts the mixing
zone and the DIC removal zone seaward. On the other
hand, most of the sediment flushed from the reservoirs and
the watercourse is deposited in the river mouth or on
reclaimed tidal flats, thereby reducing the opportunity for
the newly precipitated CaCO3 to be transported to the sea.
In conclusion, we speculate that Yellow River WSR
activities likely enhance DIC removal in the estuary. More
research is needed to understand the effects of this acute
human intervention within the context of simultaneous
long- and medium-term human interventions (e.g., land use
changes, irrigation, damming) on land–river–ocean carbon
cycling.
Implications—DIC removal in other large estuaries has
been attributed primarily to biological production or
mixing. In the Mississippi River plume, for example, DIC
removal has been controlled mainly by biological production (Cai 2003; Guo et al. 2012; Huang et al. 2012). In the
Scheldt Estuary, removal processes have been dominated
primarily by the mixing of waters contributed from
tributaries (Hellings et al. 2001). To our knowledge, the
Yellow River estuary is the only example where CaCO3
precipitation has been demonstrated to play a very
important role in estuarine DIC removal.
Milliman et al. (2008) concluded that 34 representative
world rivers experienced 30% declines in discharge during
the last half of the 20th century due to damming,
irrigation, and interbasin water transfers. Of those, 18
(including the Indus and Yellow rivers) experienced
discharge declines of . 50%. Most of the 34 rivers drain
arid and semiarid regions in Africa, Asia, and Australia,
and all of them are characterized by high DIC concentrations (Cai et al. 2008). If the patterns and processes
observed in the Yellow River are generally applicable,
these decreased discharges (due to irrigation and more
evaporation than precipitation) may lead to yet higher
DIC concentrations and thus enhanced CaCO3 precipitation in the estuarine zone. This alteration may in turn lead
to a general trend of increasing importance of CaCO3
precipitation in DIC removal at the river–ocean interface
in arid and semiarid regions. Understanding DIC and
DIC removal in the Yellow River Estuary
TAlk removal mechanisms in the Yellow River estuary
may aid our understanding of DIC distributions at the
land–ocean interface in general and of DIC fluxes to the
ocean in an important category of rivers.
Acknowledgments
We thank Dongmei Liu, Qizhen Liu, Peng Yin, Min Wang,
Baosen Wang, Chunchao Xiao, and Yan Cui for sampling and
measurement work. We also thank Tonya Clayton for her help
in language editing. Extensive and thoughtful comments and
guidance from the reviewers and editors are highly appreciated.
This work was supported by the National Science Foundation
of China (grant 41173107) and the Foundation for Innovative
Research Groups of the National Science Foundation of China:
Marine Organic Biogeochemistry (grant 41221004).
References
ABRIL, G., H. ETCHEBER, B. DELILLE, M. FRANKIGNOULLE, AND
A. V. BORGES. 2003. Carbonate dissolution in the turbid and
eutrophic Loire estuary. Mar. Ecol. Prog. Ser. 259: 129–138,
doi:10.3354/meps259129
AUCOUR, A. M., S. M. F. SHEPPARD, O. GUYOMAR, AND J.
WATTELET. 1999. Use of 13C to trace origin and cycling of
inorganic carbon in the Rhône river system. Chem. Geol. 159:
87–105, doi:10.1016/S0009-2541(99)00035-2
CAI, W. J. 2003. Riverine inorganic carbon flux and rate of
biological uptake in the Mississippi River plume. Geophys.
Res. Lett. 30: 1–4, doi:10.1029/2002GL016312
———. 2011. Estuarine and coastal ocean carbon paradox: CO2
sinks or sites of terrestrial carbon incineration? Annu. Rev. Mar.
Sci. 3: 123–145, doi:10.1146/annurev-marine-120709-142723
———, L. R. POMEROY, M. A. MORAN, AND Y. C. WANG. 1999.
Oxygen and carbon dioxide mass balance for the estuarine–
intertidal marsh complex of five rivers in the southeastern
U.S. Limnol. Oceanogr. 44: 639–649.
———, AND Y. C. WANG. 1998. The chemistry, fluxes, and
sources of carbon dioxide in the estuarine waters of Satilla
and Altamaha Rivers, Georgia. Limnol. Oceanogr. 43:
657–668, doi:10.4319/lo.1998.43.4.0657
———, AND OTHERS. 2004. The biogeochemistry of inorganic
carbon and nutrients in the Pearl River estuary and the
adjacent northern South China Sea. Cont. Shelf Res. 24:
1301–1319, doi:10.1016/j.csr.2004.04.005
———, AND ———. 2008. A comparative overview of weathering
intensity and HCO 3 flux in the world’s major rivers with
emphasis on the Changjiang, Huanghe, Zhujiang (Pearl) and
Mississippi Rivers. Cont. Shelf Res. 28: 1538–1549,
doi:10.1016/j.csr.2007.10.014
CAUWET, G., AND F. T. MACKENZIE. 1993. Carbon inputs and
distribution in estuaries of turbid rivers: The Yang Tze and
Yellow Rivers (China). Mar. Chem. 43: 235–246, doi:10.1016/
0304-4203(93)90229-H
CHEN, J. S., D. W. HE, AND Y. ZHANG. 2003. Is COD a suitable
parameter to evaluate the water pollution in the Yellow
River? Environ. Chem. 22: 611–614.
———, F. Y. WANG, AND D. W. HE. 2006. Geochemistry of water
quality of the Yellow River basin. [In Chinese.] Earth
Sci. Front. 13: 58–73, doi:10.3321/j.issn:1005-2321.2006.
01.009
———, ———, M. MEYBECK, D. W. HE, X. H. XIA, AND L. T.
ZHANG. 2005. Spatial and temporal analysis of water
chemistry records (1958–2000) in the Huanghe (Yellow River)
basin. Global Biogeochem. Cycles 19: GB3016, doi:10.1029/
2004GB002325
425
———, Y. ZHANG, T. YU, AND D. W. HE. 2004. A study on
dissolution and bio-degradation of organic matter in sediments from the Yellow River. Acta Sci. Circumst. 24: 1–5.
DAI, M. H., L. F. WANG, X. H. GUO, W. D. ZHAI, Q. LI, B. Y. HE,
AND S. J. KAO. 2008. Nitrification and inorganic nitrogen
distribution in a large perturbed river/estuarine system: The
Pearl River Estuary, China. Biogeosciences 5: 1227–1244,
doi:10.5194/bg-5-1227-2008
DAI, S. B., S. L. YANG, AND M. LI. 2009. The sharp decrease in
suspended sediment supply from China’s rivers to the sea:
Anthropogenic and natural causes. Hydrol. Sci. 54: 135–146,
doi:10.1623/hysj.54.1.135
DICKSON, A. G., AND C. GOYET [EDS.]. 1994. U.S. Department of
Energy (DOE). Handbook of methods for the analysis of the
various parameters of the carbon dioxide system in sea water,
v. 2. ORNL/CDIAC-74. Available from http://cdiac.esd.ornl.
gov/oceans/DOE_94.pdf
GAN, W. B., H. M. CHEN, AND Y. F. HAN. 1983. Carbon transport
by the Yangtze (at Nanjing) and Huanghe (at Jinan) Rivers,
People’s Republic of China, p. 459–470. In E. T. Degens [ed.],
Transport of carbon and minerals in major world rivers (Part
2). SCOPE/UNEP.
GUO, X. H., W. J. CAI, W. D. ZHAI, M. H. DAI, Y. C. WANG, AND
B. S. CHEN. 2008. Seasonal variations in the inorganic carbon
system in the Pearl River (Zhujiang) estuary. Cont. Shelf Res.
28: 1424–1434, doi:10.1016/j.csr.2007.07.011
———, AND OTHERS. 2012. Carbon dynamics and community
production in the Mississippi River plume. Limnol. Oceanogr.
57: 1–17, doi:10.4319/lo.2012.57.1.0001
HANS, H. D., G. L. GOULVEN, M. K. CHERYL, P. S. CAROLINE,
M. MICHEL, AND M. HANS. 2011. Worldwide typology of
nearshore coastal systems: Defining the estuarine filter of
river inputs to the oceans. Estuaries Coasts 34: 441–458,
doi:10.1007/s12237-011-9381-y
HELLINGS, L., F. DEHAIRS, S. V. DAMME, AND W. BAEYENS. 2001.
Dissolved inorganic carbon in a highly polluted estuary (the
Scheldt). Limnol. Oceanogr. 46: 1406–1414, doi:10.4319/
lo.2001.46.6.1406
HU, C. H., Z. W. JI, AND T. WANG. 1998. Dynamic characteristics
of sea currents and sediment dispersion in the Yellow River
Estuary. Int. J. Sed. Res. 13: 20–30.
HUANG, S., AND Q. M. LU. 1995. Estuary hydraulics. Water
Resources and Electric Power Press. [In Chinese.]
HUANG, W. J., W. J. CAI, R. T. POWELL, S. E. LOHRENZ, Y. WANG,
L. Q. JIANG, AND C. S. HOPKINSON. 2012. The stoichiometry of
inorganic carbon and nutrient removal in the Mississippi
River plume and adjacent continental shelf. Biogeosciences 9:
2781–2792, doi:10.5194/bg-9-2781-2012
LERMAN, A., L. WU, AND F. T. MACKENZIE. 2007. CO2 and H2SO4
consumption in weathering and material transport to the
ocean, and their role in the global carbon balance.
Mar. Chem. 106: 326–350, doi:10.1016/j.marchem.2006.04.
004
LUDWIG, W., P. AMIOTTE-SUCHET, G. MUNHOVEN, AND J. L.
PROBST. 1998. Atmospheric CO2 consumption by continental
erosion: Present-day controls and implications for the last
glacial maximum. Glob. Planet. Change 16–17: 107–120,
doi:10.1016/S0921-8181(98)00016-2
MACKENZIE, F. T., A. LERMAN, AND A. J. ANDERSSON. 2004. Past
and present of sediment and carbon biogeochemical cycling
models. Biogeosciences 1: 11–32, doi:10.5194/bg-1-11-2004
MILLIMAN, J. D., K. L. FARNSWORTH, P. D. JONES, K. H. XU, AND L.
C. SMITH. 2008. Climatic and anthropogenic factors affecting
river discharge to the global ocean, 1951–2000. Glob. Planet.
Change 62: 187–194, doi:10.1016/j.gloplacha.2008.03.001
426
Liu et al.
MUCCI, A. 1983. The solubility of calcite and aragonite in seawater
at various salinities, temperatures, and one atmosphere total
pressure. Am. J. Sci. 283: 780–799, doi:10.2475/ajs.283.7.780
MUKHOPADHYAY, S. K., H. BISWAS, T. K. DE, S. SEN, AND T. K.
JANA. 2002. Seasonal effects on the air–water carbon dioxide
exchange in the Hooghly estuary, NE coast of Bay of Bengal,
India. J. Environ. Monit. 4: 549–552, doi:10.1039/b201614a
NEAL, C. 2002. Calcite saturation in eastern UK rivers. Sci. Tot.
Environ. 282: 311–321, doi:10.1016/S0048-9697(01)00921-4
OFFICER, C. B. 1979. Discussion of the behaviour of the
nonconservative dissolved constituents in estuaries. Estuar.
Coast. Mar. Sci. 9: 91–94, doi:10.1016/0302-3524(79)90009-4
ORTEGA, T., R. PONCE, J. FORJA, AND A. GÓMEZ-PARRA. 2005.
Fluxes of dissolved inorganic carbon in three estuarine
systems of the Cantabrian Sea (north of Spain). J. Mar. Syst.
53: 125–142, doi:10.1016/j.jmarsys.2004.06.006
———, ———, ———, AND ———. 2008. Benthic fluxes of dissolved
inorganic carbon in the Tinto–Odiel system (SW of Spain). Cont.
Shelf Res. 28: 458–469, doi:10.1016/j.csr.2007.10.004
QIAO, S. Q., X. F. SHI, A. M. ZHU, Y. G. LIU, N. S. BI, X. S. FANG,
AND G. YANG. 2010. Distribution and transport of suspended
sediments off the Yellow River (Huanghe) mouth and the
nearby Bohai Sea. Est. Coast. Shelf Sci. 86: 337–344,
doi:10.1016/j.ecss.2009.07.019
RAYMOND, P. A., J. E. BAUER, AND J. J. COLE. 2000. Atmospheric
CO2 evasion, dissolved inorganic carbon production, and net
heterotrophy in the York River estuary. Limnol. Oceanogr.
45: 1707–1717, doi:10.4319/lo.2000.45.8.1707
REDFIELD, A. C., B. H. KETCHUM, AND F. A. RECHARDS. 1963. The
influence of organisms on the composition of seawater, p.
26–77. In M. N. Hill [ed.], The sea (2). Interscience.
REGNIER, P., A. MOUCHET, F. WOLLAST, AND F. RONDAY. 1998. A
discussion of methods for estimating residual fluxes in strong
tidal estuaries. Cont. Shelf Res. 18: 1543–1571, doi:10.1016/
S0278-4343(98)00071-5
TSUNOGAI, S., M. NISHIMURA, AND S. NAKAYA. 1968. Complexometric titration of calcium in the presence of larger amounts
of magnesium. Talanta 15: 385–390, doi:10.1016/00399140(68)80247-4
VAN OOST, K., AND OTHERS. 2007. The impact of agricultural soil
erosion on the global carbon cycle. Science 318: 626–629,
doi:10.1126/science.1145724
VANDERBORGHT, J. P., R. WOLLAST, M. LOIJENS, AND P. REGNIER.
2002. Application of a transport-reactive model to the
estimation of biogas fluxes in the Scheldt estuary. Biogeochem. 59: 207–237, doi:10.1023/A:1015573131561
WANG, H. J., Z. S. YANG, N. S. BI, AND H. D. LI. 2005. Rapid shifts
of the river plume pathway off the Huanghe (Yellow) River
mouth in response to water–sediment regulation scheme in
2005. Chin. Sci. Bull. 50: 2878–2884, doi:10.1360/982005-1196
———, ———, Y. SAITO, J. P. LIU, AND X. X. SUN. 2006.
Interannual and seasonal variation of the Huanghe (Yellow
River) water discharge over the past 50 years: Connections to
impacts from ENSO events and dams. Glob. Planet. Change
50: 212–225, doi:10.1016/j.gloplacha.2006.01.005
———, ———, ———, ———, ———, AND Y. WANG. 2007.
Stepwise decreases of the Huanghe (Yellow River) sediment
load (1950–2005): Impacts of climate change and human
activities. Glob. Planet. Change 57: 331–354, doi:10.1016/
j.gloplacha.2007.01.003
WOLLAST, R. 2003. Biogeochemical processes in estuaries, p.
61–67. In G. Wefer, F. Lamy, and F. Mantoura [eds.], Marine
science frontiers for europe. Springer.
YANG, D., AND OTHERS. 2004. Analysis of water resources
variability in the Yellow River of China during the last half
century using historical data. Water Resour. Res. 40: W06502.
ZHAI, W. D., M. H. DAI, W. J. CAI, Y. C. WANG, AND Z. H.
WANG. 2005. High partial pressure of CO2 and its maintaining
mechanism in a subtropical estuary: The Pearl River estuary,
China. Mar. Chem. 93: 21–32, doi:10.1016/j.marchem.2004.
07.003
———, ———, AND X. H. GUO. 2007. Carbonate system and CO2
degassing fluxes in the inner estuary of Changjiang (Yangtze)
River, China. Mar. Chem. 107: 342–356, doi:10.1016/
j.marchem.2007.02.011
ZHANG, L. J., L. XUE, M. Q. SONG, AND C. B. JIANG. 2010.
Distribution of the surface partial pressure of CO2 in the
southern Yellow Sea and its controls. Cont. Shelf Res. 30:
293–304, doi:10.1016/j.csr.2009.11.009
———, M. XUE, AND Q. Z. LIU. 2012. Distribution and seasonal
variation in the partial pressure of CO2 during autumn and
winter in Jiaozhou Bay, a region of high urbanization. Mar.
Pollut. Bull. 64: 56–65, doi:10.1016/j.marpolbul.2011.10.023
———, L. WANG, W. J. CAI, D. M. LIU, AND Z. G. YU. 2013.
Impact of human activities on organic carbon transport in the
Yellow River. Biogeosciences 10: 2513–2524, doi:10.5194/bg10-2513-2013
Associate editor: Markus H. Huettel
Received: 15 May 2013
Accepted: 29 October 2013
Amended: 18 November 2013