On the Cooling and Freshening of Antarctic
Intermediate Water in the South Pacific Ocean
Between 1970 and 2018
By Alexis Racioppi
Senior Honors Thesis
Environment, Ecology, & Energy Program University of North Carolina at Chapel Hill
Global warming is a driver of sometimes subtle shifts in climate patterns. One such area where subtle changes occur is deep ocean water formation, where alterations in processes regulating water density may impact surface conditions that determine water mass density (e.g., precipitation/evaporation, warming/cooling). The Antarctic Intermediate Water mass (AAIW) forms as an amalgamation of multiple water masses at the Antarctic Polar Front and spreads northward into all ocean basins. This feature of the AAIW makes it especially interesting for identifying how climate change impacts deep water formation, as the presence of the AAIW in all of the southern ocean basins allows scientists to use it as an early warning sign of climate change. Modern climatic processes, such as increased precipitation and rising surface ocean temperatures, influence the conservative temperature and salinity characteristics of the waters contributing to the volume of the AAIW, thus influence the characteristics of the AAIW itself. The AAIW should retain the signatures of temperature and salinity change at the time of its formation as it sinks and travels throughout the South Pacific Ocean during thermohaline circulation. In March and April of 2018, the expectation of preserved temperature and salinity change was tested aboard the Sea Education Association vessel SSV Robert C. Seamans. This senior honors thesis research sought to calculate rates of temperature and salinity change in the South Pacific AAIW from 1970 to 2018 on multiple time scales at both the upper boundary (700-800 m) and core (950-1050 m) of the water mass. CTD data collected during the six-week SEA cruise from Lyttelton, NZ, to Pape’ete, French Polynesia, were compared to historical cast data from the 2018 version of NOAA’s World Ocean Database to derive these temperature and salinity trends. For the sampled region of the South Pacific in my study, the AAIW was found to be experiencing freshening at both upper boundary and core depth zones, with increased
The long-term overturning cycle of the oceans means that climate changes since the
industrial revolution and from centuries ago can simultaneously impact the conditions of ocean
waters from surface to depth (Gebbie and Huybers, 2019). Surface condition shifts in the North
Atlantic and Southern Ocean, where much of the volume of the global deep sea is derived, are
particularly relevant to understanding change in subsurface water masses. This includes the
Antarctic Intermediate Water (AAIW), the creation of which is dependent on thermohaline
circulation of waters forming in the poleward oceans.
Thermohaline circulation is a global ocean current pattern that describes the constant
movement of deep seawater between ocean basins and its upwelling to return as surface waters
(Figure 1; Stommel and Arons, 1960; Ramsdorf, 2006). This process depends on the
temperature and salinity of different water masses. Together, temperature (thermo-) and salinity
(-haline) determine the density of seawater. More saline water is denser than water containing
less salt; water that has a lower temperature is denser than warmer water. These two parameters
combine to yield a range of seawater densities, thus dictating how different water masses interact
and travel. During thermohaline circulation, slow-moving currents force the horizontal transport
of water while convective processes drive the vertical transport of unique water parcels, where
Figure 1. Global overturning (thermohaline) circulation schematic shown from a Southern Ocean perspective. Figure from: Talley, 2011.
In today’s oceans, global thermohaline circulation begins in the North Atlantic where
cold, salty (dense) water formed at the surface sinks, becoming a deep ocean water mass that
travels south along the seafloor towards Antarctica (Figure 2). This movement is slow
(centimeters per year), but eventually the North Atlantic Deep Water (NADW) encounters the
Antarctic Divergence Zone (55˚S), where surface waters separate due to Ekman transport
influences on both the Antarctic Coastal Current (pushing water towards Antarctica) and the
strong Antarctic Circumpolar Current (pushing water away from Antarctica)(Gordon, 1971).
This divergence allows the long-traveled NADW to upwell, reaching the surface and moving
either towards Antarctica, joining the Antarctic Circumpolar Water to later be subducted as
Antarctic Bottom Water (AABW), or away from Antarctica, northward via Ekman transport
obtains the name Antarctic Surface Water, and encounters the sub-polar low around 60-50˚S
latitude. This atmospheric low pressure zone is characterized by high levels of precipitation
(Segar, 2018), which drives further freshening while slight warming occurs by heat transfer from
the air above. Eventually, the Antarctic Surface Water mass reaches the Antarctic Polar Front, a
convergence zone marking the northern boundary of the Antarctic Circumpolar Current (Gordon,
1971). Here, it meets and mixes slightly with warmer (less dense) Subantarctic Surface Waters.
Being the denser water mass, the Antarctic Surface Water subducts (sinks) beneath the
Subantarctic waters at this boundary, settling around 1000m depth as the relatively cold and
fresh, newly named, Antarctic Intermediate Water (Sloyan and Rintoul, 2001; Figure 2). The
majority of AAIW by volume forms near the tip of South America and is then distributed into
every ocean basin by the Antarctic Circumpolar Current (ACC) as it speeds unhindered around
Antarctica (Talley, 2011).
In this story of deep ocean circulation and the creation of AAIW, salinity and temperature
play a large role determining the vertical stratification for deep ocean water masses. Salinity and
temperature are both “conservative properties” of water, meaning that they are not influenced by
biological, chemical, or physical processes once they are set (Segar, 2018). In the ocean, a water
mass (e.g. NADW, AABW, AAIW, etc.) refers to a volume of water with defined lateral and
depth boundaries, having a specific location of formation and nearly uniform salinity and
temperature values throughout. The temperature and salinity of a water mass are established at
the air-sea interface and retained over its lifetime. It is possible to differentiate between water
masses based on these conservative properties, as well as track their historical variations
(England 1991). Temporal changes in key properties are possible following shifts in prevailing
atmospheric and surface ocean conditions and/or decadal to long-term climate at the location of
water mass formation, each of which determine the temperature and salinity signatures of
newly-created waters destined for the deep sea. Thus, these deep waters carry the signature of the global
climate at the time of their formation as well as some influence of local climate change as they
transport heat across the surface of the planet (Rhamstorf, 2006). Though it is a fundamental
process in heat transport, similar to atmospheric circulation, the entire circuit of thermohaline
circulation requires about 1,000 years, enabling a long climatic historical signal to be carried
within deep ocean waters (Gebbie and Huybers, 2019).
Other factors also control global ocean density patterns, including the amounts of
precipitation, evaporation, and ice formation/melt that occur at any location on the ocean’s
surface resultant from local climate. These processes can have dramatic impacts on surface water
salinity, and thus the characteristics of locally-forming water masses (Rhein et al., 2013).
ice formation yield increased sea surface salinity (Segar, 2018). The main process that affects
surface water temperature, thus heat content of deep waters, is the temperature of the atmosphere
immediately above; locally, air is warm or cool based on solar input at its geographic position
and radiative influences from the Earth’s surface. The upper ocean either absorbs atmospheric
heat energy or loses it via conduction and convection, the latter common in the polar regions
with strong, persistent wind regimes (Sekma et al., 2013).
In recent years, changes in ocean salinity and temperaturepatterns have been observed in
many parts of the world, including surface regions where deep water masses are known to form.
On average, global oceans to 700m depth warmed at a rate of 0.015 degrees Celsius from 1970
to 2010 with more prominent increases in the North Atlantic (Rhein et al., 2013). Additionally,
statistically significant salinity changes were detected in 43.8% of the world’s surface ocean
from 1950 to 2008 (Durak and Wijffels, 2010), with patterns of salinity increase and decrease
resembling the mean salinity field. These surface water changes are reflected in deep water
masses like the AABW, which has become 0.06 degrees Celsius warmer and 0.004 0.001 psu
fresher per decade from 1994 to 2016, potentially as a result of glacial calving in Southern Ocean
(Menezes et al., 2017). Interestingly, this freshening rate increases to 0.008 0.001 psu when
only considering data from the study’s most modern decade (2007 to 2016), suggesting a
possible acceleration of the freshening rate in recent years (Menezes et al., 2017). The NADW
likewise showed changes at its formation region, exhibiting an estimated cooling trend from
1955 to 2005 (Mauritzen et al., 2012) and a freshening rate of -0.01 psu per decade from the
1960s to the 2000s (Dickson et al., 2002).
Sediment data have documented the influence of changing surface ocean conditions on
Atlantic after atmospheric warming (Broecker, 2000; Rahmstorf, 2006). Such dynamics could
have sufficiently weakened thermohaline circulation that a lack of heat transport to the North
Atlantic allowed Arctic glaciers to rapidly form (Rahmstorf, 2006; Clark et al., 2002). This
ocean-atmosphere coupled shift has been suggested as a mechanism for the Little Ice Age, which
followed a slight warm period (Broecker, 2000). Occurring less than 1000 years ago, the
remnants of cooling from the Little Ice Age may still be detectable in old Pacific waters (Gebbie
and Huybers, 2019). The potential for such dramatic climate shifts following ocean salinity and
temperature variations highlights the importance of monitoring ongoing changes of these
parameters in today’s water masses.
The AAIW is an especially interesting water mass to study because it migrates into in all
sectors of Southern Hemisphere oceans, from the northern edge of the Antarctic Polar Front to as
far north as 20 degrees South (England and Santoso, 2002; Figure 3). In general, AAIW waters
reside at depths of 700-1300m, with average salinity readings of 34.1-34.6 psu and average
temperatures of 2.4-7 degrees Celsius (Emery, 2001; England and Santoso, 2002; Schmidtko and
Johnson, 2012; Menezes, 2017). However, like other extensive water masses, it too has
experienced temperature and salinity changes. In the Indian Subtropical Gyre near 32˚S,
freshening (-0.13 psu total from 1930-1990) and cooling (-0.33˚C total from 1930-1990) of the
water mass were observed (Bindoff & McDougall, 2000). A cooling trend has likewise been
found for the AAIW in the southwest Pacific (-0.4˚C from 1970 to 1990; Johnson and Orsi,
1997) and salinity trends at 50˚S in the southwest Atlantic (-0.02 psu over 40 years; Curry et al.,
2003) and along 17˚S in the South Pacific (-0.064 psu from 1930-1995; Wong et al., 1999) agree
that the AAIW has experienced freshening. However, opposite trends were observed in the
salinity (0.0041 psu) were reported at the water mass’ central depths (32.5˚S; Schnieder et al.
These conflicting trends in salinity and temperature changes highlight geographic
variations in climate-driven properties for the AAIW and underscore the need to observe and
understand water mass changes on local scales, especially where little inter-basin mixing occurs
once waters are distributed via the ACC. Though inter-basin mixing is minimal for the AAIW, it
is predicted that subtle density differences within the water mass may result in different depth
zones of the AAIW exhibiting their own distinct temperature and salinity characteristics that
experience non-cohesive temporal change.
Though past research on the AAIW documented changing levels of salinity and
temperature, data on these trends must be recorded regularly and locally, as well as analyzed at
multiple depths within in the water mass in order to understand zonal differences. This research
aimed to quantify and contextualize rates of AAIW temperature and salinity change from
1970-2018 in its boundary and core depth zones by addressing three objectives:
1. To determine and compare decadal rates of temperature and salinity change (1970-2018) to
more recent rates on a shorter time scale (2000-2018).
2. To compare rates of temperature and salinity change with and without the addition of
isolated SEA cruise S-278’s 2018 hydrocast data.
3. To compare rates of temperature and salinity change for the AAIW’s upper boundary and
core depth zones.
In 2018, the Sea Education Association’s sailing research vessel, SSV Robert C. Seamans,
followed the S-287 cruise track (Figure 4) beginning in Lyttelton, NZ (March 31st departure) and
ending in Pape’ete, Tahiti, French Polynesia (sampling concluded April 25th in the eastern
Pacific Subtropical Gyre). This month-long research cruise was part of the SEA Semester
Program and consisted of students from around the world working together with research faculty
to learn field techniques and conduct research on the future of the oceans and climate.
Between 0930 and 1200 each day aboard ship, weather permitting, a hydrographic
carousel with CTD (Conductivity, Temperature, Depth) probe (SBE19PlusV2; SeaBird
Electronics, Bellvue, WA) was lowered into the ocean and continuously collected salinity,
cast yielded data in averaged 5-m bins from the sea surface to depths up to 1500m. At the time of
each cast, the latitudinal and longitudinal position was recorded.
Using CTD casts along the S-278 cruise track, the AAIW was identified by locating
where the water mass’ characteristic salinity and temperature minima and maxima (34.1-34.6 psu
and 2.4-7˚C) appeared in a Temperature-Salinity diagram (Emery, 2001; England and Santoso,
2002; Schmidtko and Johnson, 2012; Menezes, 2017; Figure 5). This process also permitted
identification of the depth range of data within each profile belonging to the AAIW, enabling the
identification of the water mass’ core depth zone (950-1050m)
Figure 4. Map of S-278 CTD casts along a 2018 cruise track from Lyttelton, NZ to Pape’ete, French Polynesia. The red box outlines the area from which historical data was collected for time series analysis. *Box spans 32.5 – 33.5 ˚S and 160 – 170 ˚W.
Historical data was compiled from the World Ocean Database (2018 version) using the
Select and Search tool on NOAA’s National Centers for Environmental Information website
repository was queried by geographic location (bounds: West -170, East -160, North -32.5, and
South -33.5) and date (1970-2018). The red box created by these coordinates (Figure 4) was
selected as the acceptable geographic range for historic data comparisons because it was the
narrowest range of latitude and longitude that contained at least one cast per time period, as well
as an S-278 cast. Depth (meters), temperature (˚C), and salinity (psu) data were extracted for all
World Ocean Database water column profiles (sourced from CTDs, expendable
bathythermographs (XBTs), and profiling floats) that met these time and location criteria.
All extracted historical data were compiled in a Microsoft Excel document with S-278
data that met the same geographic requirements (027-HC). In this file, cast data were first
separated into five decadal time period bins (1970-79, 1980-89, 1990-99, 2000-09, 2010-18) and
eight bins of shorter, more recent intervals, hereafter “fine” bins, (2000-03, 2004-05, 2006-07,
2008-09, 2010-11, 2012-13, 2014-15, and 2016-18).
For each time bin, all data were run through two sets of filters, each corresponding to a
different target depth range. Quality control filters for temperature and salinity were applied to
remove any outliers well outside the established AAIW ranges. The first set of filters focused on
the AAIW boundary zone (700-800m) and included temperatures between 5.9-7.2˚C and
salinities of 34.3-34.5 psu. All remaining data were used to calculate average temperature,
salinity, and respective standard deviation values for each time bin; rates of change were also
determined for both parameters within the 700-800m depth zone, calculated separately for the
decadal and fine time bins. Data from the geographically relevant S-278 cast (027-HC) were
filtered using the same procedures. In all cases, S-278 temperature and salinity data were
included in the 2018 year for one set of rate calculations, and then copied to their own time bin
separate and highlight the latest observations from the massive amount of data that existed before
the SEA Semester voyage in 2018.
Second, the data were filtered for the 950m –1050m depth range, selected to represent the
“core” of the AAIW located using the T-S plot peak generated with S-278 data. Here, quality
control filters included temperatures between 4.45-5.71˚C and salinities of 34.3-34.42 psu. Filter
adjustments from the first target depth to the second were made in order to accommodate the
water mass’ characteristics at different zones. Once again, the remaining data were used to
calculate average temperature, salinity, and respective standard deviation values for each set of
time bins. As before, rates of change for both parameters were determined for the core
(950-1050m) depth zone, calculated separately for the decadal and fine time bins. Filtered S-278 data
(027-HC) for this depth zone yielded its own average temperature and salinity values which were
included in a second set of rate calculations for each set of time bins (decadal and fine).
Rates of change for temperature and salinity were compared between the two target depth
zones and within each target depth between the two timescales. In addition, rates of change
excluding S-278 data were compared to rates including S-278 data in all cases.
This ocean hydrography dataset contained 645 historical and modern water column
profiles binned into two series of time intervals. In the series of decadal bins, the number of
profiles in each time bin ranged from one profile in the 1970-79 and 1980-89 bins to 12 during
1990-99, 149 from 2000-09, and a maximum of 482 between 2010-18. When broken into smaller
time bins for the years with greater available data (2000-2018), the number of profiles in each
the 2012-13 bin. The decadal 2010-18 bin and fine 2016-18 bin each contained data for a single
modern profile measured during SEA’s S-278 cruise in 2018.
The AAIW was located using T-S plots with bounds of 2.4-7.2 ˚C and 34.3-34.5 psu. The
core depths of the water mass were identified using the temperature and salinity values at the
peak of the AAIW T-S curve (Figure 5). The values found at this peak consistently occurred
between 950-1050m depth. CTD depth profile measurements from the S-278 cruise and
historical ocean database resulted in the determination of rates of salinity and temperature
change in two depth regions of the AAIW over multiple time periods and scales.
For the core depth zone and the upper boundary depth zone, time binned data is presented
in a series of box plots for salinity (Figure 6) and temperature (Figure 7). Salinity showed a
decreasing trend for the 700-800m boundary depth zone in both time bin configurations (Figure
6a, 6b). The fine time scale rate of salinity change for this depth zone was 1.59x the speed of the
decadal time scale (Table 1). For each plot in Figure 6, rates of change were determined once
excluding the white S-278 box and once including the white S-278 box in order to separate the
most recent data from the earlier part of the 2010-18 or 2016-2018 time periods. With the
inclusion of the S-278 CTD data, the rate of decadal salinity change for the upper boundary zone
increased to 1.18x its previous value while the fine time series rate increased to 1.22x its
previous value (Table 1). Rates of salinity change are consistently of lower magnitude in the
core depth zone than in the upper boundary zone (Table 1). The 950-1050m core depth zone also
shows a trend of decreasing salinity in both time bin configurations (Figure 6c, 6d). The fine
time scale rate of salinity change for this depth zone was 1.2x the speed of the decadal time scale
(Table 1). With the inclusion of the S-278 CTD data, the decadal rate of salinity change for the
core depth zone increased to 1.13x its previous value while the fine time series rate increased to
1.28x its previous value (Table 1). It is important to note that all S-278 data represent only one
CTD cast, while the historical data incorporates multiple casts within a time range and is an
average. Additionally, the middle 50% of data points for the core depths are generally spread
over a narrower range of salinity values than the middle 50% of data points for the boundary
depths of the AAIW. According to these calculations, both depth ranges of the AAIW examined
Sequential changes in mean salinity at upper boundary and core depths are summarized at
the decadal scale (Table 3, 4) and at finer scale over the most recent 20 years (Tables 5,6;mean
values shown as “x” symbols in Figure 6).
Figure 6. AAIW salinity change over time was analyzed at the decadal (a, c) and fine (b, d) scales for two depth ranges. Boxes represent data within the 25-75th percentiles while whisker
extent represents the rest of the data values, excluding outliers represented as dots beyond the whiskers, within each time period. Horizontal lines within boxes indicate median values while “x” symbols indicate mean values. For each panel, the rightmost box (white) represents S-278 data (2018).
00-03 04-05 06-07 08-0 10-11 12-1 14-15 16-18 S278 04-05 06-0 08-09 10-1 12-13 14-1 16-18 S278
00-03 04-05 06-07 08-0 10-11 12-1 14-15 16-18 S278 04-05 06-0 08-09 10-1 12-13 14-1 16-18 S278
70-79 80-89 90-99 00-09 10-18 S278
Table 1. Rates of Salinity Change (psu/yr)
Time scale 700-800m 950-1050m
Decade -0.0017 -0.0015
Fine -0.0027 -0.0018
Decade S-278 -0.002 -0.0017
Fine S-278 -0.0033 -0.0023
Temperature showed a decreasing trend for the 700-800m boundary depth zone in both
time bin configurations (Figure 7a, 7b). The fine time scale rate of temperature change for this
depth zone was 8.44x the speed of the decadal time scale (Table 2). As with salinity, for each
plot in Figure 7, rates of change were determined once excluding the white S-278 box and once
including the white S-278 box. With the inclusion of the S-278 CTD data, the rate of decadal
temperature change for the upper boundary zone increased to 3.33x its previous value while the
fine time series rate increased to 1.70x its previous value (Table 2). For rates of temperature
change that are in the same direction (negative/decreasing), rates are of a lower magnitude in the
core depth zone than in the upper boundary zone (Table 2). The 950-1050m core depth zone also
shows a trend of decreasing temperature on a fine time scale, but increasing temperature on a
decadal scale (Figure 7c, 7d). With the inclusion of the S-278 CTD data, the decadal rate of
temperature change for the core depth zone increased to 1.33x its previous rate of warming while
the fine time series rate decreased to 0.61x its previous rate of cooling (Table 2). Once again, it
is important to note that all S-278 data represent only one CTD cast, while the historical data
incorporates multiple casts within a time range and is an average. Additionally, the 950-1050m
core zone is shown to have consistently lower temperature values than the 700-800m upper
boundary zone (Figure 7). According to these calculations, the upper boundary zone of this
AAIW segment are cooling on both time scales, while the core zone appears to warming on a
Sequential changes in mean temperature at upper boundary and core depths are
summarized at the decadal scale (Tables 3, 4) and at finer scale over the most recent 20 years
(Tables 5,6; mean values shown as “x” symbols in Figure 7) after extracting mean values from
the appropriate plots in Figure 7.
Figure 7. Temperature change over time was analyzed at the decadal (a, c) and fine (b, d) scales for two depth ranges.
00-03 04-05 06-07 08-0 10-11 12-13 14-15 16-18 S278 04-05 06- 08-09 10-1 12-13 14-1 16-18 S278
00-03 04-05 06-07 08-0 10-11 12-13 14-15 16-18 S278 04-05 06- 08-09 10-1 12-13 14-1 16-18 S278
70-79 80-89 90-99 00-09 10-18 S278
Table 2. Rates of Temperature Change (˚C/yr)
Time scale 700-800m 950-1050m
Decade -0.0009 0.0024
Fine -0.0076 -0.0021
Decade S-278 -0.003 0.0032
Fine S-278 -0.0129 -0.0013
Table 3. Decadal temperature and salinity averages of the AAIW from 700-800m.
Time Period Avg. Temperature (˚C) Standard Deviation (Temperature) Avg. Salinity (psu) Standard Deviation (Salinity) Number of data points
1970-79 6.55 0.17 34.39 0.012 41
1980-89 6.59 0.16 34.45 0.013 101
1990-99 6.55 0.22 34.38 0.019 777
2000-09 6.54 0.22 34.37 0.017 1568
2010-18 6.53 0.20 34.35 0.013 5868
S-278 6.42 0.16 34.33 0.008 20
Table 4. Decadal temperature and salinity averages of the AAIW from 950-1050m.
Time Period Avg. Temperature (˚C) Standard Deviation (Temperature) Avg. Salinity (psu) Standard Deviation (Salinity) Number of data points
1970-79 5.00 0.21 34.36 0.003 41
1980-89 5.05 0.15 34.41 0.005 101
1990-99 5.11 0.24 34.34 0.009 777
2000-09 5.06 0.24 34.34 0.011 1568
2010-18 5.11 0.24 34.32 0.008 5868
Table 5. Fine time scale temperature and salinity averages of the AAIW from 700-800m. Time Period Avg. Temperature (˚C) Standard Deviation (Temperature) Avg. Salinity (psu) Standard Deviation (Salinity) Number of data points
2000-03 6.60 0.18 34.38 0.013 985
2004-05 6.60 0.18 34.38 0.013 71
2006-07 6.57 0.13 34.37 0.015 13
2008-09 6.43 0.25 34.36 0.018 499
2010-11 6.40 0.16 34.35 0.011 187
2012-13 6.40 0.16 34.35 0.010 216
2014-15 6.49 0.16 34.35 0.011 895
2016-18 6.55 0.20 34.35 0.014 4550
S-278 6.42 0.16 34.33 0.008 20
Table 6. Fine time scale temperature and salinity averages of the AAIW from 950-1050m.
Standard Deviation (Temperature)
Number of data points
2000-03 5.08 0.22 34.34 0.012 985
2004-05 5.16 0.24 34.35 0.009 71
2006-07 5.14 0.14 34.34 0.010 13
2008-09 4.99 0.26 34.34 0.009 499
2010-11 5.00 0.23 34.32 0.011 187
2012-13 5.00 0.20 34.32 0.006 216
2014-15 5.08 0.21 34.32 0.006 895
2016-18 5.13 0.25 34.32 0.008 4550
The AAIW forms at the Antarctic Polar Front as upwelling NADW meets Antarctic
Surface Water and freshens before subducting underneath warmer, less dense surface waters
settling between those waters and denser, underlying deep water masses. This formation process
determines the ways in which AAIW temperature and salinity may be impacted by both
long-term (hundreds of years) and short-long-term (decades) climatic processes as it is created. In fact,
much of the data from this study supports the assertion that two separate mechanisms may be
acting to induce observed variations in the AAIW from 1970-2018. The first leverages the
centuries-long deep water memory of surface water cooling during the transition from the
Medieval Warm Period to the Little Ice Age (Gebbie and Huybers, 2019). The second is a more
recent, anthropogenically-induced climatic shift to a warmer lower atmosphere, thus warmer
surface oceans, inducing freshwater influxes to the surface ocean via precipitation and possibly
ice melt (Rhein et al., 2013).
Both of these processes appeared to influence the results pertaining to the first objective
of this study: to determine and compare decadal rates of temperature and salinity change
(1970-2018) to more recent rates on a shorter time scale (2000-(1970-2018). Overwhelmingly, the AAIW in
this segment of the South Pacific experienced freshening (decreased salinity) on both time scales
(Figure 6; Table 1). These findings agree with those reported by Bindoff and McDougal (2000),
Curry and colleagues (2003) and Wong and colleagues (1999) that the AAIW is freshening,
though the rates of core freshening determined by those investigations are, respectively, faster in
the Indian Subtropical Gyre, and slower in the southwest Atlantic and South Pacific at 17˚S than
the rates of change calculated here (Table 1). Freshening in multiple basins was expected for the
AAIW (Rhein et al., 2013) and geographical constraints may explain the differences between
(Figure 3) is in a position such that its subsequent transport by the ACC delivers it first to the
Atlantic basin, then the Indian basin, and last the Pacific basin. Thus, it is possible that salinity
differences between ocean basins are influenced by the time the AAIW has spent in circulation,
interacting with the many waters that contribute to the ACC in varying geographic locations
(Talley, 2011; Figure 1), before being delivered to a specific basin. These varying geographic
inputs support the observation of differing freshening rates in different AAIW locations; near
33˚S in the southwest Pacific (this study) rates were larger in magnitude than those found at
more northerly Pacific latitudes (Wong et al., 1999) or in the southwest Atlantic (Curry et al.,
2003), whereas freshening rates at a comparable latitude in the Indian Subtropical Gyre (Bindoff
and McDougal, 2000) are on the same order of magnitude to those reported here.
Moving forward from the general trend of increasing salinity, this study found that recent
freshening rates for the AAIW are consistently greater in magnitude than decadal rates of
freshening (Table 1). This pattern has also been observed in the NADW by Menezes and
colleagues (2017); although the NADW contributes to AAIW formation, intermediate water
freshening in the South Pacific cannot be attributed to changes in NADW. Thermohaline ocean
circulation is too slow for the influence of 1994-2016 North Atlantic freshening to influence the
Southern Ocean. A more plausible explanation for the freshening trends found in this study relies
on the mechanism of anthropogenic climate change since the industrial revolution. Increases in
atmospheric water vapor drive the delivery of more precipitation to the subpolar lows (Seidel,
2002) and rising air temperatures hasten ice along portions of the Antarctic coastline, together
causing water masses that incorporate the heavily influenced Southern Ocean surface waters, like
the AAIW, to freshen. As atmospheric temperatures continue to warm, further freshwater
vapor)-driven feedback loop, could explain the freshening increase observed in this study (Rhein et al.,
When considering temperature trends for this study’s first objective, the AAIW had a
consistent cooling pattern on both decadal and recent time scales in the upper boundary depth
zone, while the core of the water mass exhibited warming on a decadal scale and cooling on a
more recent fine time scale (Table 2). The trend of AAIW cooling is supported by the findings
of Bindoff and McDougal (2000) and Johnson and Orsi (1997); Bindoff and McDougal (1994)
link it to warming and freshening in the surface waters of Antarctic Surface Water causing a
vertical displacement of density surfaces in the water column. This anthropogenic warming of
the surface oceans (Rhein et al., 2013) could explain why cooling trends in the AAIW upper
boundary zone appear to be faster and why the core depths switch from warming to cooling in
modern years (Table 2). These trends are further supported by observations at core AAIW
depths by Schnieder et al. (2005) in the southeast Pacific. A recent study has also suggested that
sub-surface Pacific waters may still be cooling as a result of North Atlantic surface water cooling
during the transition to the Little Ice Age (Gebbie and Huybers, 2019). This mechanism of deep
ocean water masses preserving and transporting centuries-past cooling trends could thus serve as
a possible explanation for the cooling AAIW. It perhaps serves as a baseline upon which the
previously discussed modern cooling trends compound to produce faster modern rates of cooling
in the water mass than the decadal scale trends.
The fingerprints of anthropogenic climate change (i.e. increasing precipitation over the
subpolar lows and global surface warming) are further supported as a mechanism for the above
trends when considering the results of objective two: to compare rates of temperature and salinity
the isolated S-278 data to the mix consistently increased the magnitude of calculated freshening
rates for the AAIW in this study (Table 1). As S-278 cruise represented the most recent data, its
effect as a unique data point on salinity trends supports well-documented hastening
anthropogenic climate change in the modern era, namely increased high latitude precipitation
and polar ice melt over the past decade (Rhein et al., 2013). Temperature trends continued the
pattern of increased modern cooling in the upper boundary depths when the isolated S-278 cast
data were included, again supporting that this trend is influenced by accelerating surface
warming due to anthropogenic climate change (Rhein et al., 2013). In the core zone, however,
the addition of modern S-278 data increased the magnitude of decadal warming and decreased
the magnitude of more recent freshening. Together with the trends of increased cooling in the
upper boundary zone at both time scales and only modern cooling in the core zone, it is
suggested that a warm temperature anomaly likely appears in the S-278 data. These data could
also be the reason for the decreased modern freshening that the core depths experienced in this
study (Table 2).
As only a single S-278 CTD cast at a single location was used for computing these rates, it is
possible that this anomaly could have been reduced by utilizing a larger 2018 dataset. Using
more than a single cast for the time period would have produced more meaningful and
statistically robust insights into the most recent AAIW trends. The 1970-79 and 1980-89 time
bins for decadal analysis also contained only one CTD cast. As with the 2018 time bin, trends
calculated using these two decadal bins would be more statistically robust had there been more
geographically relevant data to include in each time period.
The final objective of this study was to compare rates of temperature and salinity change
upper boundary zone experienced faster freshening in all cases relative to the AAIW core (Table
1). This trend was fitting for a water mass for which salinity plays such an important role in
determining density (Durack and Wjiffles, 2010); freshening decreases seawater density. Thus, it
is expected that the freshest AAIW water should remain at the upper water mass boundary while
the more saline, but still freshening, waters would be at deeper core depth zones. Waters within
the AAIW that experience the most freshening will be less dense and sink less deeply than those
experiencing less freshening. Larger freshening rates would be expected to be observed in the
upper boundary of the water mass. This study was limited by the S-278 modern cast data, which
did not extend to the lower AAIW boundary layer, and it would be interesting to investigate this
salinity trends across all depths of the water mass. According to the observations reported here,
the lower boundary depth zone of the water mass would be expected to also exhibit freshening
but at a slower rate than the waters above it.
Furthermore, the upper boundary depth zone of the AAIW was observed to experience equal
or greater rates of cooling than the core depth zone, especially considering that, on a decadal
scale, the core depth zone had a warming trend (Table 2). The within-water mass temperature
discrepancy renders it hard to tease apart the mechanisms driving change. One possible
explanation could be linked to the already existing temperature differences in the two depth
zones. The AAIW’s core depth zone has consistently lower average temperatures than the upper
boundary depth zone (Figure 7), suggesting that within the AAIW formation region near the
Polar Front, these core waters are subducted earlier (i.e. farther south) than the upper boundary
waters sourced from a warmer, more northerly portion of the formation region. In this case, the
proportions of sinking water contributing to each of the examined depth zones may have shifted
proportionally largest volume origin. The core depths may therefore show temperature changes
associated with more poleward waters (more warming) while the upper boundary depths reflect
changes on the subtropical edge of the Polar Front (more cooling) (Rhein et al., 2013). Once
again, it would be interesting to extend this analysis to the lower boundary depths of the AAIW,
where colder waters indicating an even more southward subduction could potentially be
experiencing even stronger decadal warming trends than the core depth zone.
Global Implications of Cooling and Freshening AAIW
The cooling and freshening of the AAIW have the potential to impact large scale
thermohaline circulation and its ability to transport heat on Earth and maintain the global heat
budget. While water mass cooling implies increasing density and the freshening implies
decreasing density, the density of the AAIW is believed more suspectable to the latter change of
freshening (Durack and Wjiffles, 2010). With the strong and accelerating freshening trends
found in this study and other scientific studies cited earlier, it is possible that the overall density
of the AAIW is decreasing. As the AAIW depends on its density to subduct beneath surface
waters near the Polar Front and settle at intermediate depths, a decreased density could limit its
future ability to do so; limited AAIW formation could in turn influence global thermohaline
circulation velocities and volumes. Sinking of one water mass is what allows upwelling of
another water mass during this circulation. Thus, an interruption in one section of the
thermohaline circulation, however remote it may seem, could have consequences thousands of
kilometers away. An interruption may cause related slow-downs in deep water mass movements
around the planet and possibly result in an intensification of stratification in the global ocean
Any such slowdown of worldwide deep ocean currents would vastly disrupt the global heat
budget, which relies on the deep ocean to remove thermal energy from the atmosphere and store
it for 1,000 years before being ventilated in upwelling zones. With inhibited or halted deep ocean
circulation, the ocean cannot effectively remove thermal energy from the Earth’s surface; this
could result in even further atmospheric warming and hysteresis (Broecker 2000; Rahmstorf,
2006). This grim situation is currently unlikely (Rhein et al., 2013), but gradual trends towards
such a state are not out of range of possibility. These findings demonstrate consideration for such
eventual impacts of anthropogenic climate change are warranted. Clearly the mechanisms
driving climate change have both natural and anthropogenic sources, as shown in the long-term
and short-term trends in rates of AAIW temperature and salinity change. Historical effects of
climate transitions surrounding the Little Ice Age are found in the salinity signals of NADW.
However, climate change in more recent centuries is driven by human activities which intensify
Though these findings support the potential for more than one mechanism forcing temperature
(cooling) and salinity (freshening) change in the South Pacific AAIW, the influence of each
process cannot be distinguished conclusively by the results of this or other studies conducted to
date. However, reported observations do clearly indicate that rates of salinity and temperature
change in the AAIW increased in more recent years. Furthermore, this study concluded that it is
essential to analyze changes in these parameters at multiple depths within a water mass, as not all
depth zones of a water mass may experience change in the same way. In order to better
understand how such changes in the AAIW would influence thermohaline circulation and the
global heat budget, it is imperative that these trends continue to be monitored across the AAIW,
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This project was supported by Sea Education Association faculty, staff and students onboard the SSV Robert C. Seamans during SEA Semester S-278; data collection and analysis would not have been possible without their help. Thanks to the David and Vicki Craver Trust Fund to JEC for financially supporting my presentation of this material at the AGU 2020 Ocean Sciences Meeting. Thank you also to the UNC Marine Sciences Department for bestowing the Hill Fund Award to support the original thesis project that was put on hold due to COVID-19 restrictions. Many thanks to Honors Carolina and the Michael P. and Jean W. Carter Research Fund for their financial contribution to the original thesis project, and again to Honors Carolina for supporting the transition to a new thesis project amid pandemic challenges.